Altered Volcanic Rocks

Erratum There are several instances of a typographic error in Figure 4.2 (page 79). The references to 'Y/Zr! in the caption and 'Y/Zr' in labels on the diagram are incorrect. They should be 'Zr/Y' in all cases.

AlteredVolcanicRocks A guide to description and interpretation

CathrynGifkins WalterHerrmann RossLarge

Published by the Centre for Ore Deposit Research University of Tasmania, Australia

UTAS

Published by CODES Centre for Ore Deposit Research, University of Tasmania, Private Bag 79, Hobart, Tasmania, Australia 7001 An ARC Special Research Centre

© Centre for Ore Deposit Research, 2005

National Library of Australia Cataloguing-in-Publication Data Gifkins, Cathryn. Altered volcanic rocks : a guide to description and interpretation. Bibliography. Includes index. ISBN 1 86295 219 1. 1. Rocks, Igneous. 2. Hydro thermal alteration. I. Herrmann, Walter, 1951- . II. Large, Ross R. III. University of Tasmania. Centre for Ore Deposit Research. IV. Title.

552.2

another Pongratz Production 2005

Copy editing: Im'press: clear communication Index: Word Wise and Im'press: clear communication Printed in Australia by the Printing Authority of Tasmania

Ill

I CONTENTS PREFACE ACKNOWLEDGEMENTS INTRODUCTION 1

|

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS

1.1

Submarine volcanic successions Volcanic facies Volcanic facies associations Evidence for submarine environment of emplacement Alteration in submarine volcanic successions Devitrification Alteration processes Characteristics inherited from volcanic facies Geology of the Mount Read Volcanics Stratigraphy of the Mount Read Volcanics Submarine facies associations and architecture Post-depositional alteration processes Mineral deposits and prospects Geology of the Mount Windsor Subprovince Stratigraphy of the Seventy Mile Range Group Submarine facies associations and architecture Post-depositional alteration processes Mineral deposits and prospects

1 1 2 2 2 4 4 6 7 9 10 11 11 12 12 13 14 14

|

DESCRIBING ALTERED VOLCANIC ROCKS

15

2.1 2.2

2.6

Frequently asked questions Alteration nomenclature Mineral-based alteration nomenclature Compositional alteration nomenclature Generic alteration nomenclature Descriptive nomenclature — alteration facies Alteraction facies — the recommended method Alteration mineral assemblage Tools for mineralogical determination Alteration intensity Qualitative estimates of alteration intensity Quantitative estimates of alteration intensity An integrated approach to alteration intensity Alteration data sheets

15 19 19 20 20 20 22 23 24 25 25 26 33 36

|

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS

37

3.1

Alteration textures Replacement textures

37 37

1.2

1.3

1.4

2

2.3 2.4 2.5

3

v.vii viii ix 1

iv |

CONTENTS

Infill textures Dissolution textures Static recrystallisation textures Dynamic recrystallisation textures Deformation textures Pseudotextures Pseudoclastic textures False polymictic texture False matrix-supported texture False coherent textures Alteration distribution Alteration zonation patterns Regional diagenetic zones Regional metamorphic zones Regional, deep, semi-conformable altered zones Local contact metamorphic or hydrothermally altered halos Local hydrothermally altered halos around ore deposits Vein and fracture altered halos Overprinting relationships and timing of alteration Method Overprinting textures

41 41 52 52 52 54 54 63 63 63 63 64 64 64 66 66 67 67 69 70 70

|

GEOCHEMISTRY OF ALTERED ROCKS

73

4.1

Lithogeochemistry Sampling and analytical methods Closure in composition data Chemostratigraphy Mass transfer techniques Rare-earth-element geochemistry related to alteration Mineral chemistry Principles Applications Stable isotopes Theoretical background Isotopic applications in alteration studies

73 73 78 79 81 87 87 87 88 92 92 92

|

SEAFLOOR-AND BURIAL-RELATED ALTERATION

97

5.1

Alteration related to sea-floor processes and burial Physical conditions Definitions Hydration Palagonite Perlite Diagenesis (glass to zeolite facies) Diagenetic minerals Diagenetic zones Genesis of diagenetic minerals and zones Regional metamorphism (zeolite to amphibolite facies) Transition from diagenesis to regional metamorphism Burial metamorphism Burial metamorphic facies Burial metamorphic zones Zeolite facies Genesis Diagenetic alteration in the Hokuroku Basin Geological setting Alteration facies and zones Genesis of altered zones Data sheets

3.2

3.3 3.4

3.5

4

4.2

4.3

5

5.2

5.3

5.4

5.5

97 98 98 98 99 100 102 102 105 108 115 115 115 115 115 116 116 118 118 119 120 122

CONTENTS | V

6

5.6

Diagenetic alteration in the Mount Read Volcanics Geological setting Alteration fades and zonation Genesis of alteration fades Data sheets

128 128 128 128 133

|

SYNVOLCANIC INTRUSION-RELATED ALTERATION

139

6.1

The role of intrusions in generating hydrothermal systems Subseafloor regional hydrothermal systems Regional altered zones assodated with intrusions Recharge zones Discharge zones Deep, semi-conformable altered zones Altered zones as part of a regional hydrothermal system Altered zones within intrusions Deuteric alteration Hydrothermal alteration Contact altered halos around intrusions Contact altered zones Genesis of contact altered zones Contact altered zones associated with the Darwin Granite Geological setting Alteration fades and zonation Genesis of the alteration system Data sheets

140 140 141 141 141 142 147 148 ' 148 148 149 149 153 154 155 155 156 157

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS

163

6.2

6.3

6.4

6.5

7

|

7.1 7.2

7.3 7.4

7.5

7.6

7.7

Common features of VHMS deposits Hydrothermal alteration halos associated with VHMS deposits Footwall alteration pipes Stratabound altered footwall zones Altered hanging wall zones Chemical reactions and mass changes Alteration box plot trends in altered footwall zones The genesis of footwall alteration pipes Significance of pyrophyllite and kaolinite in VHMS systems Metamorphism of altered zones The spectrum of volcanic-hosted deposits and associated alteration patterns Hydrothermal alteration related to the spectrum of deposits Comparisons between Archaean, Palaeozoic and Cainozoic VHMS alteration systems Australian Palaeozoic VHMS alteration halos Japanese Cainozoic VHMS alteration halos Canadian and Australian Archaean VHMS alteration halos Comparisons : Hellyer: a massive elongate polymetallic lens Geological setting Alteration fades and zonation Ore genesis Data sheets Rosebery: a polymetallic sheet-style deposit Geological setting Alteration facies and zonation Genesis of the ore lenses and alteration system Data sheets Western Tharsis: a hybrid Cu-Au VHMS deposit Geological setting Alteration facies and zonation

163 164 164 166 167 168 169 170 174 174 174 176 178 178 179 179 180 181 182 182 183 184 194 194 195 195 196 202 202 202

Vi | CONTENTS

Ore genesis Data sheets 7.8 Henty: a volcanogenic gold deposit Geological setting Alteration fades and zonation Ore genesis Data sheets 7.9 Thalanga: a polymetallic sheet-style deposit Geological setting Alteration facies and zonation Ore genesis Data sheets 7.10 Highway-Reward: a pipe style Cu-Au VHMS deposit Geological setting Alteration facies and zonation Ore genesis Data sheets

203 204 212 212 212 213 214 221 221 222 222 223 232 232 232 232 233

|

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS

8.1

Principles of discriminating between diagenetic, hydro thermal and metamorphic alteration facies Diagenetic facies Metamorphic facies Hydrothermal alteration facies Exploration vectors and proximity indicators Mineral zonation Major element lithogeochemistry Alteration indices Mass change vectors Mineral chemistry vectors Isotopic vectors

241 241 242 242 243 243 243 244 245 245 246

REFERENCES INDEX

251 271

8.2

| |

:

241

I vii

PREFACE

Altered volcanic rocks is principally for hands-on geologists, our fortunate colleagues who practise in mineral exploration and mining geology, and the students who may in the future play in those professional fields. We began designing and writing this book in mid 2001 after struggling for several decades to come to terms with a variety of alteration styles in ancient submarine volcanic successions. We realised that although a large number of company and research geologists were working on similar rocks there was no existing text to help guide us through the complexity of altered volcanic rocks. The so-called volcanic rocks we deal with in ancient volcanic successions and around ore deposits frequently bear little resemblance to their fresh counterparts, which are studied in undergraduate igneous petrology and volcanology courses. It is typically only with long experience that geologists develop the confidence and skills to be comfortable working with altered volcanic rocks, to interpret the original volcanic facies, unravel complex alteration histories and determine their significance in terms of mineral deposit prospectivity, particularly in ancient and deformed successions. The topic and content of the book were inspired by problems that we have faced, and in many cases overcome, while working on industry-related volcanic facies, alteration geochemistry and economic geology research projects, particularly in the Mount Read Volcanics. Many of the ideas presented in this book come from the results of CODES research projects, which have been run in collaboration with industry partners and the Australian Research Council (ARC) over the last 15 years. In particular, AMIRA-ARC Linkage project P439 (Studies of VHMS-related alteration: geochemical and mineralogical vectors to ore) provided an

enormous amount of data, case studies and expertise. Some of the results of this project have previously been published as a special issue in Economic Geology (Gemmell and Herrmann, eds., A special issue devoted to alteration associated with volcanichosted massive sulfide deposits, and its exploration significance, August 2001, v. 96, no. 5). We were encouraged by the wide acceptance and success of the CODES publication by Jocelyn McPhie, Mark Doyle and Rod Allen (1993) Volcanic textures: a guide to the interpretation of textures in volcanic rocks, which has been a major factor in improving the description and interpretation of volcanic facies over the last decade. The advance we have made in Altered volcanic rocks is to integrate observations and data on volcanic facies and textures with alteration mineralogy and geochemistry at both regional and local scales in order to provide a multidisciplinary method for the study and discrimination of different alteration types: diagenetic, metamorphic and hydrothermal alteration. We hope that this book will help to equip geologists working in altered and deformed successions with the skill and confidence to interpret the original volcanic facies and encourage the use of altered rocks as discriminants and vectors in mineral exploration. This book may not provide all the answers, but if it gives readers the courage to tackle the study of altered rocks, embrace the problems and pursue the answers it will have been worthwhile.

Cathryn C. Gifkins Wally Herrmann Ross R. Large

viii |

| ACKNOWLEDGEMENTS

While preparing this book, we were fortunate to have valuable support, assistance and advice from many people. We extend our sincere thanks to those people whose discussions and/or reviews of various chapters have helped shape this book. Chapters were peer-reviewed by Stuart Bull, David Cooke, Mark Doyle, Kim Denwer, Allan Galley, Bruce Gemmell, Anthony Harris, Jocelyn McPhie, Andrew Rae, Mike Solomon and Fernando Tornos. Valuable discussions were also held with Ron Berry, Stuart Bull, Jocelyn McPhie, Phil Robinson and Mike Solomon. Although samples and photographs used herein are principally from the authors' collections, we also made use of hand specimens and thin sections from the School of Earth Sciences rock catalogue at the University of Tasmania, and samples and photographs from colleagues. Thank you to those people who contributed: Sharon Allen, Stuart Bull, Kate Bull, Tim Callaghan, Cari Deyell, Bruce Gemmell, George Hudak, Karin Orth and Jocelyn McPhie. We also thank Izzy von Lichtan, Curator at the School of Earth Sciences, for her help in finding and returning hundreds of catalogue samples.

Andrew McNeill very kindly provided a long projection of the Rosebery ore lenses. Tim Callaghan assisted with core specimens and whole-rock geochemical data from Mount Julia. Jon Huntington and Melissa Quigley at CSIRO provided HyMap® images of the Mount Lyell field. We are infinitely grateful for the hard work of the production team. Karin Orth and Simon Stephens helped with sample preparation. Mike Blake and Karin Orth assisted with photography. Rose Pongratz and Izzy von Lichtan prepared the bibliography and checked references. June Pongratz provided expert drafting, design and desktop publishing, and was incredibly tolerant of the endless revisions. Final editing was by Impress: clear communication and indexing by Word Wise and Impress: clear communication. We also appreciate our families, friends and colleagues who have been very understanding of our commitment to this project over the last three years. Thank you for your support and patience.

I ix

r

INTRODUCTION

This book is about the processes and products of alteration in submarine volcanic successions, although many of the concepts presented here can be applied to altered volcanic rocks from almost any environment. Its emphasis is on hydrothermal alteration associated with volcanic hosted massive sulfide (VHMS) deposits. Few volcanic rocks in submarine settings are entirely unaltered, and in hydrothermal environments all rocks are altered to some degree. Recognising, describing and interpreting altered volcanic rocks is not always easy, but the results can have important implications for volcanology, petrology and ore deposit studies, and can improve and accelerate success in mineral exploration. Determining prealteration characteristics and discriminating between primary volcanic, magmatic and secondary alteration features requires knowledge of the alteration processes and their products. Valuable base-metal, gold and silver deposits exist in a variety of modern and ancient submarine volcanic successions. Many of these deposits are surrounded by, or spatially related to, extensive altered zones that record the passage of mineralising hydrothermal fluids and fluid-wall rock reactions. Research into the textural, mineralogical and compositional effects of alteration around VHMS deposits has shown that they can be quantified and used as effective exploration tools for discriminating deposit styles and guiding exploration towards mineralised zones.

An introduction to alteration Guilbert and Park (1986) defined alteration as any change in the mineralogical composition of a rock brought about by physical or chemical means, especially by the interaction with hot or cold aqueous solutions or gases. Alteration typically encompasses mineralogical changes and changes in the rock texture and composition. Components of rocks, including ore metals, can be dissolved, replaced or recrystallised. New minerals may precipitate and isotopic ratios may change. Porosity and permeability may be reduced or increased. Primary volcanic textures are overprinted, and may be destroyed and replaced by new 'false' textures, or enhanced. The resulting altered rock is described as the 'alteration fades' (e.g. Riverin and Hodgson, 1980).

Thus, alteration involves complex modification of a rock. Furthermore, a rock may undergo several episodes of syn- to post-depositional alteration, not all of which are related to mineralising hydrothermal systems. Each alteration episode is influenced by the existing texture and composition of the rock, and may also overprint and modify that texture and composition. As a result the characteristics of altered rocks are highly variable. In ancient volcanic rocks it is a challenge to determine host volcanic facies, unravel complex alteration processes and interpret their significance in terms of mineral prospectivity. That challenge is the focus of this book.

How the book is organised Altered volcanic rockshzs two main themes, which are organised into eight chapters: (1) it describes the basic principles behind recognising and describing altered volcanic rocks; and (2) it discusses the different alteration processes that are common in submarine volcanic successions and their products. Chapter 1 introduces the concepts of alteration in submarine volcanic successions and summarises the main alteration processes and volcanic facies. It outlines the regional geology of two of the most productive Australian submarine volcanic successions: the Cambrian Mount Read Volcanics in western Tasmania and the Cambro-Ordovician Mount Windsor Subprovince in Queensland. This book principally employs examples from these two successions, and includes descriptions of other ancient submarine volcanic successions for comparison. Chapter 2 discusses alteration nomenclature, mineralogy, intensity and indices, and the principles of alteration facies. It proposes an integrated multidisciplinary approach to description and classification. The main elements of alteration facies — mineral assemblage, intensity, texture, distribution, zonation and timing — are described in Chapters 2 and 3. Chapter 4 outlines geochemical methods used in alteration studies and their applications. It emphasises whole-rock lithogeochemistry, mineral chemistry and stable isotope analysis. Chapter 5 concentrates on regional alteration styles including hydration, diagenesis and metamorphism associated with burial, Chapter 6 on intrusion-related alteration styles,

X I INTRODUCTION

and Chapter 7 on hydrothermal alteration and mineralisation associated with VHMS deposits. We present short case studies for these different alteration styles, emphasising hydrothermal alteration associated with a variety of VHMS deposits including Rosebery, Hellyer, Henty and Thalanga. These case studies incorporate pictorial data sheets that present the mineralogical, textural and compositional characteristics of each of the main alteration facies. They combine volcanic facies analysis and alteration mineral assemblages, textures, intensity and geochemistry to interpret the features of different alteration styles. Chapter 8 outlines the methods for discriminating between the products of mineral deposit-related hydrothermal alteration and other alteration processes, and identifying favourable altered zones for mineral exploration. It also discusses geochemical vectors that may guide explorers towards mineralised rock within these altered zones.

.

Significance of altered volcanic rocks to mineral exploration Economic geologists are particularly interested in alteration because hydrothermal mineral deposits are commonly hosted by altered rock. Hydrothermally altered zones around mineral deposits provide much larger targets for mineral exploration than the deposits themselves. The mineral assemblages, and in some cases the chemical composition, of the altered rocks may provide indications of the proximity of an ore deposit, and thus vectors towards mineralised rock. In addition, mineralogical, textural and compositional studies of alteration facies can provide important constraints on the timing, physical and chemical conditions, and origins of hydrothermal systems and related mineralisation (Barnes, 1979). The texture and distribution of alteration facies can also be used to infer changes in porosity, permeability and fluid pathways in the host succession. The results of alteration studies are commonly incorporated into ore deposit models used in mineral exploration. Thus, the identification and interpretation of alteration facies is, and should be, a routine part of exploration for hydrothermal mineral deposits.

11

1 | ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS

This chapter describes submarine volcanic successions, the common processes of alteration that occur in these successions, and provides two examples of ancient submarine volcanic successions, that have been variably altered and mineralised: the Mount Read Volcanics and the Mount Windsor Subprovince. In submarine volcanic environments, the coincidence of magmatic fluids, heat and abundant seawater generates hydrothermal convection. As a consequence, submarine volcanic successions may host important hydrothermal mineral deposits, commonly referred to as volcanic-hosted massive sulfide (VHMS) deposits. VHMS deposits are a significant source of zinc, copper, lead, silver and gold, and continue to be a target for base-metal exploration. They range in size from less than one million tonnes to over 200 million tonnes, and commonly contain high metal grades. For example, the Hellyer deposit in western Tasmania produced 16.2 Mt at 13.9% Zn, 7.1% Pb, 0.4% Cu, 168 g/t Ag and 2.5 g/t Au in its nine years of operation. VHMS deposits occur mainly in submarine rift environments particularly back arc and mid ocean rifts; however, they can occur in a variety of other submarine environments including continental rifts, oceanic basins or plateaux, and arc-continent or continent-continent collision zones. They are one of the few classes of ore deposits that exist throughout the geological record from early Archaean to Recent.

by sedimentary processes. In addition, volcanic units may be emplaced into the succession as synvolcanic intrusions. This section summarises the main volcanic facies that occur in submarine volcanic successions. For a more detailed discussion of submarine volcanism, volcanic textures, facies and their interpretation, readers are referred to McPhie et al. (1993) Volcanic textures: a guide to the interpretation of textures in volcanic rocks.

Volcanic facies

For descriptive purposes, volcanic facies are divided into two main textural types: coherent and volcaniclastic. Coherent facies consist of solidified magma and are commonly characterised in volcanic rocks by aphyric (fine grained or glassy) or porphyritic textures, where porphyritic refers to evenly distributed euhedral crystals (phenocrysts) in a finegrained or glassy groundmass (McPhie et al., 1993). Volcaniclastic facies are those composed mainly of volcanic particles (Fisher, 1961). Volcanic particles are crystals, crystal fragments, shards, pumice clasts, scoria clasts and dense volcanic clasts, which may be produced by primary volcanic (pyroclastic and autoclastic) or sedimentary (weathering and erosion) processes. Volcaniclastic facies include a spectrum of facies: primary volcanic facies, syneruptive volcanic facies generated by coeval eruptions and deposited from sedimentary processes, and volcanogenic sedimentary facies that show evidence of temporary storage and reworking prior to deposition (McPhie et al., 1993). 1.1 | SUBMARINE VOLCANIC Primary volcaniclastic facies result from volcanic processes SUCCESSIONS of clast formation, transport and deposition and include Submarine volcanic successions are significantly different from hydroclastic, pyroclastic and autoclastic facies. Hydroclastic subaerial volcanic successions, as the processes of eruption, facies is a general term for facies, typically comprising blocky transport, emplacement, and post-emplacement alteration glassy particles, produced by magma-water interactions, may be strongly affected by the presence of water. Typically, whether by explosive steam generation or by non-explosive submarine volcanic successions comprise a wide variety of quench fragmentation of magma (Fisher and Schmincke, coherent and volcaniclastic facies intercalated with mixed 1984; Hanson, 1991). Pyroclastic facies comprise volcanic provenance and non-volcanic sedimentary facies (Fig. 1.1). particles (pyroclasts) that were generated by explosive The volcanic facies may be derived from intrabasinal, extra- eruptions and deposited by primary volcanic processes, by basinal or basin-margin eruptions in submarine or subaerial fallout, flow or surge. Autoclastic facies comprise volcanic settings. Eruption styles may be effusive or explosive and the particles generated by in situ non-explosive fragmentation of products may remain in situ or be redeposited or reworked lava or magma (autobrecciation and quench fragmentation).

2 | CHAPTER 1

Autobrecciation occurs when the more viscous parts of a moving lava respond in a brittle fashion to locally higher strain rates, and fragment into blocky clasts (Fisher, I960). Quench fragmentation occurs in situ where hot lava or magma comes into contact with water, ice or water-saturated sediment (Rittmann, 1962; Pichler, 1965; Yamagishi, 1987). The resulting autoclastic deposits - autobreccia, hyaloclastite or peperite - typically comprise dense blocky or splintery clasts, but they may be pumiceous and have fluidal shapes (Fisher, 1960; Pichler, 1965; Busby-Spera and White, 1987; Gifkins et al., 2002). Syneruptive volcaniclastic fades are composed dominantly of unmodified volcanic clasts that were fragmented by volcanic process such as explosive eruptions, autobrecciation or hydration, but were transported and deposited by sedimentary processes (McPhie et al., 1993; McPhie and Allen, 2003). They can occur directly from eruption when clasts bypass initial deposition as primary deposits and are delivered directly to sedimentary transport and deposition systems, such as subaqueous eruption-fed water-supported gravity currents or water-settled fall (e.g. White, 2000; McPhie and Allen, 2003). They may also occur indirectly by rapid remobilisation and redeposition during or shortly after the eruption (Fisher and Schmincke, 1984; Cas and Wright, 1987; McPhie et al., 1993). Unconsolidated volcanic debris may be remobilised by: the slumping and sliding of gravitationally unstable rapidly accumulated clastic debris; explosive eruptions; local uplift; syn-depositional faulting; and extrusion and intrusion of magma. Volcanogenic sedimentary fades (epiclastic volcanic, Fisher,

1960) contain volcanic particles derived from the posteruptive erosion and reworking of pre-existing volcanic facies and may include a significant proportion of non-volcanic particles (McPhie et al., 1993). In submarine volcanic successions, volcaniclastic facies are dominated by in situ autoclastic and syneruptive volcaniclastic facies where the particles were derived from either autoclastic fragmentation or explosive eruption. Most volcanic and nonvolcanic clastic deposits were emplaced by water-supported density currents (i.e. high- and low-concentration turbidity currents, debris flows and grain flows) and as fallout from suspension in the water column.

Evidence for submarine environment of emplacement VHMS deposits occur in submarine volcanic successions, thus exploration for new deposits is restricted to submarine successions. However, there are few volcanic or sedimentary facies that unequivocally constrain the host depositional environment. A subaqueous setting (marine or lacustrine) may be interpreted based on the presence of: water-supported massflow deposited facies; hemi-pelagic, biogenic, biochemical and chemical sedimentary facies; pillow lavas; and quench fragmented volcaniclastic facies. Also seawater-related diagenetic alteration facies (e.g. widespread albite alteration facies) can suggest a submarine environment. Without fossil evidence, differentiating between marine and lacustrine settings is difficult as few facies are restricted to either environment. Facies with tidal and wave tractional structures, such as bimodal-bipolar ripples, are submarine, whereas lacustrine settings may be indicated by the presence of evaporites. Hummocky cross stratification is more common in, although not restricted to, marine settings. Carbon-oxygen isotope signatures of carbonates can be used to support marine or lacustrine environments. Although bedforms, sedimentary structures and some sedimentary deposits help us to interpret a marine environment, they are of little help in constraining the water depth. Water depth may be an important consideration for mineral exploration as recent research suggests that Au-rich VHMS deposits are restricted to shallow water environments (e.g. Hannington et al., 1999; Hannington and Herzig, 2000; Herzig et al., 2000). Shallow water environments are typically dominated by the tractional processes of tidal and wave action and result in characteristic sedimentary structures and bedforms. In contrast, deep water environments, below storm wave base, generally lack tractional currents: sediment distribution and deposition mainly occurs through the actions of turbidity currents, debris flows and the process of suspension sedimentation. Water depth and depositional setting may be more accurately constrained by the presence of fossiliferous limestone or sedimentary facies that contain marine fossils intercalated in the volcanic succession.

Volcanic facies associations A facies association is a collection of facies that are spatially, mineralogically, compositionally or texturally related, and that may also be genetically related (Cas and Wright, 1987). There are three common types of volcanic units represented by facies associations in submarine volcanic successions (Fig. 1.1): lavas, synvolcanic intrusions (cryptodomes, sills and dykes) and syneruptive volcaniclastic facies. Lavas and synvolcanic intrusions comprise associations of coherent and autoclastic facies. The syneruptive volcaniclastic facies can be divided into two principal categories, those dominated by non- to poorly-vesicular blocky lava clasts and related to the submarine emplacement of lavas and lava domes, and others that contain abundant pumice or scoria clasts produced by explosive eruptions. In addition, there is a wide variety of volcanogenic sedimentary facies.

1.2 | ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS After emplacement, volcanic facies are commonly subjected to a variety of alteration processes (Fig. 1.2). Alteration occurs when existing components become unstable under changing physical and chemical conditions, and alter to more stable minerals. Volcanic glass, which is the main component of many volcanic facies, is a metastable solid with the structure of a liquid (Carmichael, 1979). It is undercooled to the point where extreme viscosity has prevented crystallisation. As a result, volcanic glass readily devitrifies to minerals that are more stable under surface conditions; generally clay minerals, zeolites, carbonates, feldspar, quartz and oxides (Carmichael, 1979; Henley and Ellis, 1983; Fisher and Schmincke, 1984;

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 3

FIGURE 1.1 | Facies model of a submarine basin in which a variety of coherent and clastic volcanic facies are intercalated with sedimentary facies. The volcanic facies include primary coherent and autoclastic facies, syneruptive and post-eruptive volcaniclastic facies derived from submarine and subaerial eruptions. Many of the volcanic facies associations are laterally discontinuous. Common facies associations represent (A) lavas and lava domes composed of coherent and autoclastic facies; (B) synvolcanic sills and cryptodomes; (C) syneruptive volcaniclastic facies derived from explosive and effusive submarine eruptions; (D) volcanogenic sedimentary or resedimented volcaniclastic facies derived from pre-existing deposits; (E) syneruptive volcaniclastic facies derived from subaerial explosive eruptions; (F) mixed provenance sedimentary facies; and (G) marine sedimentary facies.

FIGURE 1.2 | Facies model showing the distribution of different styles of altered zones in a submarine volcanic succession that hosts VHMS deposits. See Figure 1.1 legend for the patterns denoting volcanic and sedimentary facies.

4 | CHAPTER 1

Friedman and Long, 1984; Cerling et al., 1985). Alteration of volcanic glass involves not only devitrification, but changes in texture, composition, porosity and permeability, and may simultaneously affect both the chemistry and circulation of pore fluid in the volcanic succession (Noble, 1967; Dimroth and Lichtblau, 1979; Fisher and Schmincke, 1984; Noh and Boles, 1989; Torres et al., 1995). Understanding alteration requires a range of skills that include recognising alteration minerals, textures, paragenesis, distribution, zonation, intensity, mineralogical and chemical changes associated with alteration, pathways and mechanisms for fluid migration, and fluid origin. These characteristics are related to the alteration processes and to the characteristics of the host volcanic succession.

Devitrification The cooling history of volcanic facies may involve primary crystallisation and later devitrification. Primary or hightemperature crystallisation refers to crystallisation of magma resulting in phenocrysts, microcrysts and microlites. In contrast, devitrification refers to crystallisation of glass at low temperatures (i.e. below the glass transition temperature: Lofgren, 1971a). High-temperature devitrification accompanying first cooling may produce spherulites, lithophysae and micropoikilitic or snowflake texture (e.g. Lipman, 1965; Anderson, 1969; Lofgren, 1971b; Bigger and Hanson, 1992; McArthur et al., 1998) and is not considered to be alteration. Low-temperature devitrification results in the gradual conversion of glass to fine-grained granular crystalline aggregates, which may happen over time as a result of alteration during changing physical conditions or in response to interaction with fluid. Devitrification may be accompanied by changes in whole-rock composition (Lipman, 1965; Lofgren, 1971a; Friedman and Long, 1984).

the presence of fluid (seawater, magmatic fluid or a mixture of both). There are gradations from isochemical metamorphism to metasomatism with increasing compositional changes. The different alteration processes, hydration, diagenesis, metamorphism and local hydrothermal alteration, are all part of this continuum in submarine volcanic successions (Fig. 1.3). The effects of each alteration process may be difficult to distinguish. Hydrothermal alteration, diagenesis and metamorphism can result in similar mineral assemblages and textures. In addition, in many cases, different alteration processes, such as diagenetic and hydrothermal alteration, are contemporaneous and their products may be inseparable (Iijima, 1974, 1978; Ohmoto, 1978; Reyes, 1990; Utada, 1991;Paradisetal., 1993). In Chapters 5, 6, and 7 of this book, the common alteration processes are grouped into those related to burial, intrusions and VHMS deposits (Fig. 1.4). Thus, burialrelated alteration styles include hydration, diagenesis and burial metamorphism. Alteration styles associated with intrusions are hydrothermal alteration within intrusions, contact metamorphism and hydrothermal alteration, and regional hydrothermal alteration. Included below is a brief introduction to each of the common alteration processes that operate in submarine volcanic settings.

Hydration of volcanic glass Hydration of glass involves the absorption of external water into glass and modification of the glass structure, either during cooling or at ambient temperatures (Ross and Smith, 1955; Friedman and Long, 1984). Hydration does not directly produce new minerals, but can form perlitic fractures or palagonite in basaltic glass and it can facilitate subsequent alteration (see Chapter 5). Compositional changes

Alteration processes Alteration may result from regional or local processes. It can occur as a result of the interaction with hydrothermal fluid, as a result of changing physical (mainly temperature and pressure) conditions during burial, in association with the emplacement of intrusions, or a combination of all these processes. Submarine volcanic facies, especially glassy facies, are readily altered during hydration, diagenesis, hydrothermal alteration, metamorphism and tectonic deformation. Hydrothermal alteration is defined as the alteration of rocks or minerals by the reaction of hydrothermal fluid with pre-existing solid phases (Henley and Ellis, 1983). Hydrothermal fluid is a hot aqueous solution or gas, with or without demonstrable association with igneous processes. Hydrothermal alteration usually results in significant changes in rock texture, mineralogy and composition. Alteration is either metasomatic or isochemical. Metasomatism involves changes in mineralogy, texture and composition, whereas isochemical alteration (or metamorphism) involves mineralogical and textural changes only. In submarine volcanic successions, almost all alteration involves some degree of metasomatism, which is facilitated by

FIGURE 1.3 | This cartoon depicts the continuum between isochemical and hydrothermal alteration and shows the alteration processes common in submarine volcanic successions. They are positioned based on the relative degrees of chemical exchange for each process.

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 5

accompanying hydration include gains in H 2 O, and minor losses in silica and alkalis (Noble, 1967; Friedman and Long, 1984; Mungall and Martin, 1994).

Diagenesis Diagenesis encompasses the changes that occur in response to changing temperature and pressure during burial. During diagenesis of volcanic facies, significant textural and mineralogical changes can be produced by precipitation of cement, dissolution and replacement of original components, especially glass, and compaction (Fisher and Schmincke, 1984; Marsaglia and Tazaki, 1992; Torres et al., 1995). In theory, diagenesis in submarine volcanic successions is a metasomatic process involving minor chemical exchange between the host facies and trapped modified seawater at low temperatures (<250°C). Transitions between diagenesis and metamorphism, and diagenesis and hydrothermal alteration have not been rigorously defined and are discussed in Sections 5.4 and 6.2.

Regional metamorphism Regional metamorphism involves pervasive, mainly isochemical, mineralogical and textural changes in response to increasing pressure and temperature (Yardley, 1989). During metamorphism, H 2 O and CO2-bearing fluids are generated by dehydration and decarbonation reactions (Rose and Burt, 1979).

Contact alteration associated with intrusions Contact alteration refers to the changes caused by the rise in temperature of the host rock immediately surrounding an

FIGURE 1.4 | Common alteration processes and their products.

intrusion, which may be accompanied by the circulation of heated pore fluids around and within the intrusion. Contact alteration can be isochemical (i.e. contact metamorphism) or metasomatic (i.e. hydrothermal alteration). Contact metamorphism typically results in recrystallisation of existing minerals or components and minor remobilisation of elements (Yardley, 1989). The effects of hydrothermal alteration may include major changes in texture, mineral assemblage and whole-rock composition on a scale of centimetres to kilometres.

Hydrothermal alteration related to VHMS deposits Two styles of hydrothermal alteration are commonly related to VHMS mineralisation: (1) local alteration halos around ore deposits; and (2) regional hydrothermally altered zones. Regional hydrothermally altered zones are commonly spatially and genetically associated with large intrusions and hence in this book are discussed in Chapter 6 - intrusion-related alteration styles. Local hydrothermally altered halos around VHMS deposits result from the reaction between the host facies and the mineralising hydrothermal fluid (Sangster, 1972; Franklin et al., 1975; Riverin and Hodgson, 1980; Green et al., 1983; Urabe.et al., 1983). These altered halos are commonly zoned, reflecting changes in the composition, pH and temperature of the hydrothermal fluid with time, or the extent of reaction with the host facies (Rose and Burt, 1979; Lydon and Galley, 1986; Schardt et al., 2001). The nature of altered halos around VHMS deposits depends on the host volcanic facies, host-rock composition, timing of the hydrothermal alteration relative to the emplacement or deposition of facies, structures, fluid pathways, distribution pattern of the ore, and chemical and physical conditions of the hydrothermal fluid (Large, 1992). Thus altered halos around VHMS deposits exhibit a wide variety of geometries, sizes,

6 | CHAPTER 1 I CHAPTER 1

-

"

alteration mineral assemblages, intensities, compositions, and zones (Chapter 7). Thick, extensive (up to 20 km), pervasive sub-horizontal semi-conformable altered zones, referred to as regional hydrothermal alteration zones or deep semi-conformable alteration zones (Section 6.2), have been recognised in some volcanic successions hosting VHMS deposits (Gibson et al., 1983; Galley, 1993). Many of these regional hydrothermally altered zones are spatially associated with, and possibly genetically related to, intrusions (Galley, 1993; Brauhart et al., 1998). They are interpreted to result from large-scale convection of seawater through permeable volcanic successions at elevated geothermal gradients (Spooner and Fyfe, 1973; Munha and Kerrich, 1980; Baker, 1985). Reactions between the volcanic successions and modified seawater have involved Na-, Si-, Ca-Fe-, K-, or Mg-metasomatism, and the leaching of ore-forming metals (Munha and Kerrich, 1980; Gibson et al., 1983; Baker, 1985; Galley, 1993). Syntectonic hydrothermal alteration Hydrothermal alteration can also be synchronous with tectonic deformation: syntectonic hydrothermal alteration. In this case, the hydrothermal fluid may be modified seawater, magmatic water or volatiles released during metamorphism, or a combination of these, and migrates principally along faults and shear zones. Contemporaneous deformation may modify or destroy pre-existing textures and create new textures or foliations. Many VHMS deposits, such as Rosebery, Hercules and Mount Lyell (western Tasmania), have been affected by later tectonic deformation and modified by syntectonic hydrothermal fluids (Walshe and Solomon, 1981; Khin Zaw and Large, 1992). Although syntectonic hydrothermal fluids were not responsible for the formation of these VHMS ores, they can be critical to the subsequent formation of other styles of ore deposits in the submarine volcanic successions, such as mesothermal gold deposits. Detailed discussion of syntectonic hydrothermal alteration is not dealt with in this book.

Characteristics inherited from volcanic facies Volcanic components and facies with different compositions and textures behave differently during the initial stages of low-temperature alteration. Some components react more rapidly than others and their composition may influence the composition of the alteration mineral assemblage. For example, Marsaglia and Tazaki (1992), in their study of diagenetic trends in volcaniclastic sandstones of the IzuBonin Arc, found that black mafic crystalline fragments were unaltered, brown intermediate to mafic glassy fragments showed evidence of dissolution, and colourless rhyolitic fragments were altered to clay minerals. These differences were due to the variable reactivity of the components and the proportions of volcanic glass to crystalline facies. Generally glass is the most reactive component, followed by olivine —> magnetite, titanomagnetite and ilmenite —* pyroxene and amphibole —> biotite —* Ca-plagioclase —* microcline, sanidine and orthoclase —* quartz, apatite, rutile

and zircon (Browne, 1978; Reyes, 1990). Alteration rates for different minerals vary considerably because of mineral structure and composition. Silicate materials with an extensive cross-linked (e.g. tetrosilicate) structure react slowly; whereas those silicates with poorly connected fabric tend to react rapidly and uniformly (Casey and Bunker, 1990). Felsic volcanic facies typically have a higher proportion of glassy to crystalline facies than mafic facies. This is because the viscosity of silicic magmas (71-77% SiO2) inhibits diffusive crystal growth and thus produces thick bodies of glass, whereas the low viscosity of basaltic magmas favours crystallisation (Friedman and Long, 1984). Glassy facies or facies that contain glassy fragments are likely to be more rapidly altered, and to form different mineral assemblages, than those that are crystalline (Lee and Klein, 1986). In addition, mafic glasses are more rapidly altered than felsic glasses (e.g. Whetten and Hawkins, 1970; Fisher and Schmincke, 1984; Friedman and Long, 1984). The rate of alteration is related to the glass's viscosity, which in turn is a function of the composition (particularly SiO2 and H2O) and temperature (Friedman and Long, 1984). Increased SiO2 decreases the rate of alteration, whereas increased MgO, CaO and H 2 O increases the rate. Thus the higher SiO2 content of rhyolitic glasses retards reaction (Hawkins, 1981). The primary composition can influence the alteration mineralogy mainly because the dissolution of glass liberates alkalis and silica, which are consumed by subsequent reactions. Hence, highly silicic volcanic facies result in the crystallisation of Si- and Na-rich minerals, such as opal, quartz, tridymite, cristobalite and Na-zeolites (Sheppard et al., 1988). In contrast, mafic glasses, such as those on pillow rims, alter to Ca-, Fe-, Mg- and Mn-rich minerals such as smectites, phillipsite, oxides and chlorite. Volcaniclastic facies, particularly pumice-rich facies, initially have very high porosities. In volcaniclastic facies, the inter- and intra-particle pore space controls the porosity and permeability and thus grain size, type and sorting influence the distribution of early alteration facies. Early diagenetic alteration in well-sorted pumice breccias, although commonly patchy, is pervasive. In poorly sorted polymictic volcaniclastic facies the porosity and permeability are initially much more variable and diagenetic facies typically have complex distribution patterns. Coherent facies have much lower porosity and permeability, factors that are controlled by fractures produced by quenching, flowage and hydration. Alteration in coherent facies typically progresses as fronts that move outward from fractures into the less altered domains (e.g. the alteration of perlite, Noh and Boles, 1989). In addition, patchy or domainal alteration styles in volcaniclastic facies may also be related to variations in the quenching and hydration of glassy clasts (Surdam, 1973; Boles and Coombs, 1977; Marsaglia and Tazaki, 1992). Hydrothermal experiments on rhyolitic glass indicate that at high temperatures (>200°C), rhyolitic glass does not recrystallise but instead acts as Na-K ion exchanger. Quenched glass fixes K+, whereas slowly cooled glass fixes Na+. Therefore variations in cooling history may explain why some glassy fragments alter more readily to particular minerals than others (Marsaglia and Tazaki, 1992).

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 7

1.3 | GEOLOGY OF THE MOUNT READ VOLCANICS Many of the examples of altered volcanic rocks and alteration systems discussed in the following chapters come from the Middle to Late Cambrian Mount Read Volcanics in western Tasmania. The Mount Read Volcanics are a submarine succession of rhyolitic to basaltic volcanic and intrusive rocks with variable proportions of intercalated sedimentary rocks. They are interpreted as the products of post-collisional volcanism associated with arc-continent collision (Berry and Crawford, 1988; Crawford and Berry, 1992). The succession occurs in a 200 x 20 km area that extends from Elliott Bay in

the south through Queenstown and Rosebery to Deloraine in the north (Fig. 1.5). The volcanic succession was deposited in a series of troughs separated by areas of Proterozoic basement (Corbett and Lees, 1987; Corbett, 1992; Crawford and Berry, 1992). The Mount Read Volcanics host gold, silver and base-metal massive sulfide (VHMS) ore deposits at Hellyer, Que River, Rosebery, Hercules, Henty and Mount Lyell (Fig. 1.5). The mineral district is referred to as the Mount Read province. The Mount Read Volcanics have undergone diagenetic and hydrothermal alteration, metamorphism, at least two phases of deformation, and intrusion by Cambrian and Devonian granites (Corbett and Lees, 1987; Corbett, 1992).

FIGURE 1.5 | Distribution of the principal lithostratigraphic units and major massive sulfide deposits in the central Cambrian Mount Read Volcanics, western Tasmania. Modified after Corbett (1992,2002).

8 I CHAPTER 1

FIGURE 1.6 | Stratigraphic correlation diagram showing the major lithostratigraphic units in the Mount Read Volcanics to the west (A) and east (B) of the Henty fault. The sections are located at (1) Hellyer-Que River, (2) Pinnacles, (3) Hollway, (4) Mount Black, (5) Rosebery-Hercules, (6) White Spur, (7) Hall Rivulet Canal, (8) Murchison Gorge, (9) Henty, (10) South Henty, (11) Anthony Road, (12) Comstock-Lyell, (13) Lynchford, (14) Jukes-Darwin. Sections are modified after Fitzgerald (1974), Corbett (1979,1992, 2001), Cox (1981), Komyshan (1986), Coutts (1990), Allen (1991), Dugdale (1992), Waters and Wallace (1992), Jones (1993,1999), McKibben (1993), Herrmann and MacDonald (1996), McPhie (1996), White and McPhie (1996,1997), Callaghan (2001), Gifkins (2001) and Wyman (2001). (A) To the west of the Henty fault, the Central Volcanic Complex interfingers with, and is conformably and disconformably overlain by, the Dundas and Mount Charter Groups of the western volcano-sedimentary sequences (Corbett and Lees, 1987; Corbett, 1992). Immediately overlying the Central Volcanic Complex is a variety of small-volume sedimentary (Black Harry beds and Animal Creek greywacke) and volcanic units (rhyolite and pumice breccia). These units are overlain by the Que-Hellyer Volcanics, a succession of calc-alkaline to shoshonitic, intermediate to mafic lavas and volcaniclastic units (Corbett and Komyshan, 1989; Waters and Wallace, 1992). The Que-Hellyer Volcanics host the Que River and Hellyer ore deposits, and extend via Sock Creek to Burns Peak and Pinnacles (i.e. the Brown's tunnel sequence). The Que River Shale overlies the Que-Hellyer Volcanics and is similar to mudstone in other lithostratigraphic units of the Mount Read Volcanics. The Southwell Subgroup overlies the Que River Shale and is lithologically similar to the White Spur Formation and the Rosebery hanging-wall volcaniclastic units, comprising quartz-bearing volcaniclastic mass-flow units interbedded with black mudstone and Precambrian basement-derived turbidites (Corbett, 1992; McPhie and Allen, 1992). Overlying the Southwell Subgroup is the Mount Charter Group in which the upper Mount Cripps Subgroup is a correlate of the Tyndall Group (Corbett, 1992).

Although the primary textures, mineralogies and whole-rock compositions have been modified to various degrees, volcanic textures are generally well preserved. Locally, two regional tectonic cleavages have been recognised; however, the axial planar S2 Devonian cleavage is the dominant cleavage. S2 strikes north, dips steeply and varies from a weak, spaced cleavage to an intense, pervasive, anastomosing cleavage in the most strongly deformed rocks adjacent to faults and in phyllosilicate-rich altered zones.

The geology of the Mount Read Volcanics has been described in detail by Campana and King (1963), Corbett (1981, 1986, 1992, 1994), Corbett and Lees (1987), Corbett and Solomon (1989), Pemberton and Corbett (1992), McPhie and Allen (1992) and Crawford et al. (1992).

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 9

B

(B) To the east of the Henty fault, the Eastern quartz-phyric sequence overlies the Sticht Range Beds, interfingers with the Central Volcanic Complex and is conformably overlain by the western volcano-sedimentary sequences (Farrell Slates). The southern Central Volcanic Complex is flanked to the west by the Yolande River Sequence, part of the western volcano-sedimentary sequences. To the east, it interfingers with the Eastern quartz-phyric sequence and is overlain by the Tyndall Group (Corbett, 1992) and locally by andesite and basalt lenses that occur between Henty and Queenstown (Anthony Road andesite, Crown Hill andesite, Howards basalt, Spillway basalt, and Lynchford basalt). The Cambrian Murchison and Darwin granites intruded the succession in the Mount Murchison and Mount Darwin areas (Corbett and Lees, 1987; Corbett, 1992). The Tyndall Group is the youngest lithostratigraphic unit. It extends north-south along the eastern margin of the succession where it unconformably overlies Tyennan basement, Sticht Range Beds, southern Central Volcanic Complex, Eastern quartz-phyric sequence, western volcano-sedimentary sequences, and the Darwin Granite (Corbett and Lees, 1987; White and McPhie, 1997). The Owen Conglomerate overlies the Mount Read Volcanics both conformably and unconformably.

Stratigraphy of the Mount Read Volcanics The stratigraphy of the Mount Read Volcanics can be divided into (Figs 1.5 and 1.6): Sticht Range Beds, Eastern quartzphyric sequence, Central Volcanic Complex, western volcanosedimentary sequences, and the Tyndall Group and correlates (Corbett, 1992). These lithostratigraphic units comprise compositionally and texturally diverse coherent volcanic and volcaniclastic facies intercalated with sedimentary rocks, which are distinguished and mapped on the basis of the dominant facies. The principal volcanic facies associations are lavas, synvolcanic intrusions and syneruptive volcaniclastic units (McPhie and Allen, 1992). Lavas and synvolcanic intrusions

are predominantly calc-alkaline rhyolites and dacites with locally abundant andesites and basalts (Crawford et al., 1992). The volcaniclastic facies associations include a variety of primary and secondary volcaniclastic facies including thick extensive syneruptive pumice- and crystal-rich units and in situ and resedimented hyaloclastite (McPhie and Allen, 1992). Sedimentary facies include black mudstone, and graded, bedded sandstone of mixed volcanic and metasedimentary Precambrian basement provenance (McPhie and Allen, 1992). Regional stratigraphic relationships between the lithostratigraphic units are complex and laterally variable (Fig. 1.6). The Mount Read Volcanics are conformably and

10 | CHAPTER 1

unconformably overlain by the Owen Group, a thick sequence of Late Cambrian-Early Ordovician siliciclastic, shallow marine to fluvial conglomerate and sandstone (Corbett, 1992; White, 1996). VHMS deposits occur in a variety of volcanic facies in three of the main lithostratigraphic subdivisions of the Mount Read Volcanics (McPhie and Allen, 1992; Pemberton and Corbett, 1992; Waters and Wallace, 1992). In particular, VHMS deposits are interpreted to occur: (1) at the top of the Central Volcanic Complex, close to large felsic volcanic centres (Rosebery, Hercules and Mount Lyell); (2) associated with proximal facies of andesite-dacite volcanoes in the western volcano-sedimentary sequences (Hellyer and Que River); and (3) in the base of the Tyndall Group (Henty and Comstock) (Corbett and Solomon, 1989; Halley and Roberts, 1997).

range of clasts including quartz + feldspar porphyry, feldsparphyric rhyolite, pumice, granite and massive sulfide clasts.

Tyndall Group and correlates The Tyndall Group is the youngest lithostratigraphic unit in the Mount Read Volcanics. It extends along the eastern margin and locally along the western side of the succession (Fig. 1.5). The Tyndall Group varies in thickness from 350 to 1300 m and comprises distinctive quartz + feldspar crystalrich sandstone, volcanic breccia and volcanic conglomerate intercalated with minor rhyolitic welded ignimbrite, felsic to intermediate lavas and intrusions, and non-volcanic sedimentary rocks including limestone, mudstone and sandstone (White and McPhie, 1996, 1997).

Sticht Range Beds Cambrian granites The Sticht Range Beds are a thin (<500 m) basal succession of interbedded basement-derived sedimentary rocks and volcaniclastic rocks that occur along the eastern margin of the Mount Read Volcanics (Fig. 1.5: Baillie, 1989).

Eastern quartz-phyric sequence The Eastern quartz-phyric sequence is a 2.5 km-thick succession of quartz + feldspar-phyric lavas, synvolcanic intrusions and volcaniclastic units limited to the eastern margin of the Mount Read Volcanics (Fig. 1.5: Polya, 1981; Polya et al., 1986; Pemberton et al., 1991; Corbett, 1992).

Central Volcanic Complex The 3 km-thick Central Volcanic Complex dominates the central part of the Mount Read Volcanics between Mount Darwin and Mount Block (Fig. 1.5: Corbett, 1979). It consists of feldspar-phyric rhyolitic and dacitic lavas, synvolcanic intrusions and pumiceous volcaniclastic units (Corbett, 1979, 1992; Corbett and Lees, 1987; Corbett and Solomon, 1989; Pemberton and Corbett, 1992; Gifkins, 2001). Andesites and basalts are locally intercalated with the felsic succession (Crawford et al., 1992). Quartz + feldspar-phyric intrusions and tholeiitic basalt and dolerite dykes (Henty dyke swarm) occur throughout the northern Central Volcanic Complex (Corbett and Solomon, 1989; Crawford et al., 1992).

Western volcano-sedimentary sequences The western volcano-sedimentary sequences include the Yolande River Sequence, Dundas Group, Mount Charter Group and Henty fault wedge sequence (Corbett, 1992). These sequences are thick (>3 km) mainly sedimentary successions of quartz + feldspar-phyric volcaniclastic facies, mixed provenance sandstone and mudstone intercalated with rhyolitic, andesitic and basaltic lavas and synvolcanic intrusions, (Corbett and Lees, 1987; Corbett, 1989; McPhie and Allen, 1992). The volcaniclastic facies contain a diverse

Five Cambrian granitoids (commonly referred to as 'granites') have been recognised in western Tasmania: the Murchison, Darwin, Elliott Bay, Dove and Timber Tops granites (Leaman and Richardson, 1989). Cambrian granites may also occur at depth in a belt that extends along the eastern margin of the Mount Read Volcanics between Mount Darwin and Mount Murchison (Large et al., 1996). They are typically medium grained quartz + K-feldspar + plagioclase + biotite + hornblende + apatite + zircon ± rutile granite or granodiorite (McNeill and Corbett, 1992). They intrude the western volcano-sedimentary sequences, Central Volcanics Complex and Eastern quartz-phyric sequence and are unconformably overlain by the Tyndall Group in the Murchison and Darwin areas (Corbett, 1992). They are interpreted to be subvolcanic intrusions genetically related to the host volcanic succession (Solomon, 1981).

Submarine facies associations and architecture The Mount Read Volcanics were deposited in a predominantly below wave-base, moderate to relatively deep submarine setting. This interpretation is supported by the presence of trilobite and other marine fossils, fossiliferous limestone, turbidites and black pyritic mudstone in the sedimentary facies association (Jago et al., 1972; McPhie and Allen, 1992). The presence of massive sulfide ore deposits, very thick volcaniclastic mass-flow units, hyaloclastite, peperite and pillow lava in the volcanic facies association are also consistent with a subaqueous environment (Corbett, 1992; McPhie and Allen, 1992; Waters and Wallace, 1992). The essential elements of the facies architecture in the Mount Read Volcanics are a variety of volcanic and sedimentary facies associations that include lavas, lava domes, synvolcanic intrusions and diverse volcaniclastic units (McPhie and Allen, 1992). Lavas and sills occur separately or in clusters in the succession (McPhie and Allen, 1992). The four common types of volcaniclastic facies associations in the Mount Read Volcanics are: (1) very thick (tens of metres), massive to graded beds of rhyolitic to dacitic pumice breccia; (2) very thick, massive

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 11

to diffusely stratified units of crystal-rich (feldspar, quartz, clinopyroxene) sandstone; (3) thick to very thick, massive to graded beds of polymictic volcanic conglomerate or breccia; and (4) massive or laminated shard-rich siltstone (McPhie and Allen, 1992). Many of these volcaniclastic facies contain a high proportion of crystals, crystal fragments, shards and pumice clasts, which are interpreted to be juvenile pyroclasts transported by water-supported gravity flows (McPhie and Allen, 1992). The sedimentary facies association comprises conglomerate, sandstone, interbedded turbiditic sandstone and mudstone, mudstone, and fossiliferous carbonate (Selley, 1997; Large et al., 2001a; McPhie and Allen, 2003). These facies are of non-volcanic and mixed provenance, and include pelagic marine sediment, meta-sedimentary and ultramafic (bonninite, gabbro, peridote) rock fragments derived from the Precambrian basement, and volcanic clasts and crystals. There are regional variations in the proportion of volcanic versus sedimentary facies, the types of volcanic facies and the dominant magma composition. Volcanic facies associations locally dominate the stratigraphy (e.g. at Rosebery) whereas, elsewhere, volcanic facies are intercalated with, or subordinate to, sedimentary facies (e.g. in the hanging wall at Hellyer). Parts of the volcanic succession are dominated by the products of effusive, intrabasinal eruptions, such as the andesitic lavas and domes in the footwall of the Hellyer deposit (McPhie and Allen, 1992). In contrast, other areas are dominated by volcanic facies generated by explosive eruptions, such as the crystal and pumice-rich volcaniclastic units of the White Spur Formation (McPhie and Allen, 1992; McPhie and Allen, 2003). The volcanic facies associations also display marked regional variations in composition. Rhyolite and dacite dominate much of the succession (^90% of the mapped area of the central Mount Read Volcanics: Gifkins and Kimber, 2004); however, intermediate to mafic volcanic facies are locally important at Hellyer and between Henty and Queenstown (Corbett, 1992; Crawford et al., 1992; Large et al., 2001a).

Post-depositional alteration processes Formerly glassy or partly glassy volcanic rocks dominate the Mount Read Volcanics. These rocks have textures and compositions that reflect subsequent modification by a variety of processes, which include: hydration, diagenesis, hydrothermal alteration, regional and contact metamorphism, and deformation.

Two regional Cambrian diagenetic zones (albite zone and epidote zone, Section 5.6) are preserved in the northern Central Volcanic Complex (Gifkins, 2001). Locally, hydrothermal alteration and mineralisation was synchronous with diagenesis. In addition, altered halos developed around thick synvolcanic intrusions and Cambrian granites (Eastoe et al., 1987; Large et al., 1996; Gifkins, 2001). The Mount Read Volcanics were faulted during the Middle to Late Cambrian and more extensively deformed, folded and faulted during the Early to Middle Devonian (Corbett and Lees, 1987; Crawford and Berry, 1992). Regional metamorphism to lower greenschist facies produced assemblages of quartz, albite, sericite, calcite, chlorite, tremolite-actinolite, K-feldspar, epidote and biotite, and was contemporaneous with the Devonian deformation (Corbett, 1981; Green et al., 1981; Walshe and Solomon, 1981; Corbett and Solomon, 1989). During the Late Devonian to Early Carboniferous, contact metamorphism was associated with the intrusion of I- and S- type granites (Corbett and Lees, 1987; Williams et al., 1989; Corbett, 1992).

Mineral deposits and prospects In 2004, the Mount Read province contained two major operating base-metal-sulfide mines (Rosebery and Mount Lyell), one gold mine (Henty), and a number of exhausted and smaller sub-economic deposits or prospects. Published resource estimates are listed in Table 1.1. The Hellyer polymetallic massive sulfide deposit was a classic mound-shaped ore body discovered by a combination of geophysical, geological and geochemical exploration techniques in 1982 (Sise and Jack, 1984). Production commenced in 1989 and mining was complete by 2000. The high grade and simple geometry of the Hellyer ore body made it a profitable operation although metal recoveries were low due to the fine grainsize of the sulfides. The Que River polymetallic massive sulfide deposit was a small deposit comprising five steeply dipping ore lenses (four Zn + Pb rich and one Cu rich). It was discovered in early 1974 by airborne electromagnetic and soil geochemical exploration (Webster and Skey, 1979). Production from the Que River deposit occurred from 1981 to late 1991. Rosebery is the largest polymetallic massive sulfide deposit in western Tasmania. It comprises at least 16 separate stacked ore lenses over a strike length of 1.5 km. The deposit was initially discovered in 1893 when prospectors traced gold and

Table 1.1 | Tonnages and grades of massive sulfide deposits in the Mount Read province (Data from Mineral Resources Tasmania, Pasminco Mining and Exploration Goldfields P/L, and Gemmell and Fulton, 2001: in situ values based on average metal prices in 2000).

tx10 6

Zn wt%

Pb wt%

Cu wt%

Ag g/t

Au g/t

In situ value USS billion

Hellyer

16.2

13.9

7.1

0.4

168

2.5

3.99

Massive lens

Past producer

Que River

3.1

13.5

7.5

0.6

200

3.4

0.81

Stratabound sheet

Past producer

Rosebery

32.1

14.7

4.5

0.58

146

2.3

7.76

Stratabound sheets Current mine

Hercules

2.7

15.9

5.1

0.4

159

2.54

0.71

Stratabound sheets Past producer

Henty-Mount Julia

2.2

-

-

-

-

12.1

0.23

Sheet-like

Current mine

Mount Lyell field

311

0.04

0.01

0.97

7

0.31

6.83

Disseminated

Current mine

Deposit

Form

Status

in stratigraphic order, the Puddler Creek, Mount Windsor,

ALTERATION IN SUBMARINE VOLCANIC SUCCESSIONS | 13

Rollston Range Formation Trooper Creek and Rollston Range formations (Henderson, 1986; Paulick and McPhie, 1999). These have a total thickness of at least 14 km (Henderson, 1986) and generally young to the south. The four formations respectively represent initial continent-derived sedimentation and minor rift-related mafic volcanism, succeeded by voluminous eruptions of crustally derived rhyolitic magmas, abruptly followed by mixed maficfelsic volcanism derived from subduction-modified mantle, and culminating in deep-water fine-grained sedimentation (Stolz, 1995). The stratigraphic relationships and lithologies are summarised in Figure 1.8.

Puddler Creek Formation The Puddler Creek Formation is the oldest formation in the Seventy Mile Range Group and consists mainly of massive to laminated lithic sandstone, greywacke and siltstone of mixed continental and volcanic derivation. Minor altered trachyandesitic to trachytic coherent volcanic rocks are intercalated with clastic rocks in the upper few hundred metres. The volcanic rocks have geochemical signatures indicating an alkali intraplate association related to lithospheric thinning and incipient back-arc basin development (Stolz, 1995). The formation is up to 9 km thick in the western part of the belt and is partly stoped out by the Ravenswood Batholith in the east.

The uppermost formation of the Seventy Mile Range Group, the Rollston Range Formation, is poorly exposed, except in the southern central part of the belt where it has a minimum thickness of 1 km. It consists of Early Ordovician, fossiliferous, thinly bedded sandstone and siltstone of largely volcanic provenance. Minor intervals of felsic lava and volcaniclastic units exist locally (Berry et al., 1992).

Submarine facies associations and architecture The volcanic facies associations in the vicinity of the HighwayReward, Liontown and Thalanga base-metal sulfide deposits are known from several detailed deposit scale studies (Hill, 1996; Miller, 1996; Doyle, 1997; Paulick, 1999; Paulick and McPhie, 1999; Doyle and McPhie, 2000). Recent regional studies of volcanic facies and lithostratigraphy contribute to an improved understanding of the volcanic facies associations (Simpson and McPhie, 1998; Simpson, 2001); however, much of the regional data is currently confidential or unpublished. In summary, the Mount Windsor and Trooper Creek Formations comprise deep submarine volcanic facies that include pyroclasts, probably from both subaerial and submarine explosive eruptions. Lithofacies such as sparse microbialitic ironstones (Simpson and McPhie, 1998) indicate shallow marine settings for the Mount Windsor Subprovince.

Mount Windsor Formation The Mount Windsor Formation is a 0.4 to 5 km-thick succession of subaqueous rhyolitic volcanic rocks dominated by thick lavas, domes and high-level intrusions with subordinate volcaniclastic breccias. Isotopic evidence (Stolz, 1995) suggested the magmas were derived from the melting of continental crust.

Trooper Creek Formation The Trooper Creek Formation is a 0.5 to 2 km-thick succession of highly variable basaltic-andesitic, dacitic and rhyolitic coherent and brecciated volcanic rocks intercalated with abundant volcanogenic siltstone, and minor calcareous meta-sedimentary rocks. It is internally heterogenous and there are major lateral variations in the proportion of volcanic and sedimentary facies. Base-metal sulfide deposits and exhalative siliceous ironstones occur at various stratigraphic levels (Fig. 1.8: Duhig et al., 1992). Stolz (1995) suggested that the Trooper Creek Formation volcanic rocks were derived from a melted subduction-modified, sub-arc mantle wedge that was erupted during back-arc extension. Decreased volcanic activity, or an increase in clastic sedimentation, appears to have coincided with the change from exclusively rhyolitic volcanism in the preceding Mount Windsor Formation. FIGURE 1.8 | Simplified stratigraphic column for the Seventy Mile Range Group. Modified after Large (1992) and Paulick and McPhie (1999).

14 I CHAPTER 1 Table 1.2 | Tonnages and grades of massive sulfide deposits in the Mount Windsor Subprovince (data from Berry et al., 1992; Large, 1992: in situ values based on average metal prices in 2000). Cu wt%

Ag g/t

Au

2.6

1.8

69

0.4

1.02

Stratabound sheet

Past producer

-

6.2

-

1.5

0.47

Subvertical pipes

Current mine

6.2

2.2

0.5

29

0.9

0.18

Stratabound sheet

Prospect

1

10

0.4

0.6

8

0.2

0.13

Stratabound sheet

Prospect

Waterloo/Agincourt

0.4

19.7

2.8

3.8

94

2

0.13

Stratabound sheet

Prospect

Magpie

0.3

15

2

2

30

1

0.06

Stratabound sheet

Prospect

Deposit

tx10 6

Zn wt%

Pb wt%

Thalanga

6.6

8.4

Highway-Reward

3.7

-

Liontown

1.8

Handcuff

Post-depositional alteration processes Some of the least-altered felsic coherent facies of the Seventy Mile Range Group commonly have relict spherulitic or perlitic textures due to devitrification and hydration (Paulick and McPhie, 1999; Doyle, 2001). Pseudoclastic textures, attributed to domainal devitrification and subsequent diagenesis of coherent rhyolites, are prominent in the Mount Windsor Formation at Thalanga and probably elsewhere. Zones of intense hydrothermal alteration comprising quartz + sericite + chlorite + pyrite + carbonate assemblages partly enclose the major sulfide deposits. They vary in style from the broadly stratabound zone extending laterally beneath the Thalanga deposit, to the discordant concentric zones enveloping the Highway-Reward sulfide pipes. Early diagenetic and hydrothermal alteration facies were overprinted by regional deformation coeval with extensive intrusion of Mid-Late Ordovician gneissic granitoids (Berry et al., 1992). This deformation produced relatively low-pressure regional metamorphic assemblages that range from prehnite grade in the east, to upper greenschist grade in the west, and a near vertical axial planar cleavage. Subsequent intrusion of postkinematic Siluro-Devonian plutons in the central and eastern parts of the subprovince produced contact metamorphic aureoles with assemblages up to amphibole-hornfels grade. Despite the multiple alteration processes, well preserved primary volcanic textures that enable detailed interpretations of facies associations occur away from zones of hydrothermal alteration and mineralisation (e.g. Simpson and McPhie, 1998). Even in intensely hydrothermally altered rocks, the

g/t

In situ value US$ billion

Status

Form

existence of resistant primary components such as quartz phenocrysts allow volcanic facies interpretation (e.g. Paulick and McPhie, 1999).

Mineral deposits and prospects The Mount Windsor Subprovince contains two major basemetal sulfide deposits and several small sub-economic deposits and historical prospects. Published resource estimates are listed in Table 1.2. The known deposits are all in the Trooper Creek Formation. The two largest deposits, Thalanga and Highway-Reward, exist at the base and near the top of the formation respectively (Fig. 1.8). A gossanous outcrop led to the 1975 discovery of the Thalanga deposit and its eventual development for open pit and underground mining. Production between 1990 and 1998 amounted to 4.7 Mt from an estimated total resource of 6.6 Mt. Thalanga mine was not highly profitable, mainly because of ore dilution in underground mining and stability problems related to the thin ore lenses. The Highway-Reward Cu-Au deposit consists of two discordant, vertical pipe-like bodies of massive pyrite about 200 m apart. Originally discovered in a surface road-metal scrape in 1953 (Beams et al., 1998), the Highway-Reward deposit has been the subject of intense but sporadic exploration. Open pit mining of small oxide and supergene high-grade Cu-Au resources occurred during the late 1980s and from 1997 to the present. The deeper hypogene parts of the sulfide pipes remain undeveloped.

I 15

2 I DESCRIBING ALTERED VOLCANIC ROCKS

This chapter addresses some of the common problems that we face when studying altered volcanic rocks. Recognising and describing the characteristics of the altered rocks is an important step towards understanding the processes of alteration. Alteration involves complex modifications of the pre-existing rock and can encompass mineralogical, textural and compositional changes. Resolving these complex relationships is dependent on a systematic multidisciplinary descriptive approach incorporating aspects of volcanology, ore deposit geology, petrology and geochemistry. Unfortunately relatively few studies adequately integrate these datasets. Studies of ore deposits generally describe the characteristics of the host rocks (i.e. lithology, petrology, geochemistry and alteration) in separate sections of manuscripts. In many cases, particularly in unpublished company reports, the geochemical data and petrographic descriptions are in appendices, discouraging integration and interpretation. The integration of physical or textural observations and geochemical data is a powerful tool in the study of altered rocks. The physical characteristics and immobile element concentrations of altered volcanic rocks can help to identify the original rocks, where relict primary minerals and textures are inconclusive (e.g. Paulick and McPhie, 1999; Barrett et al., 2001). Physical and chemical changes that occurred during alteration may help to determine the degree of alteration (i.e. alteration intensity), the style of alteration (i.e. isochemical versus metasomatic), and to discriminate between alteration processes such as diagenesis, metamorphism and hydrothermal alteration (e.g. Offler and Whitford, 1992; Gifkins and Allen, 2001). In addition, this integrated approach can lead to the development of vectors to guide explorers toward ore deposits (e.g. Large et al., 2001c). This chapter compares alternative schemes of alteration nomenclature and presents a multi-variable system for describing and naming alteration facies. The different elements of this descriptive approach to nomenclature are explained in detail in subsequent sections and chapters (i.e. alteration mineral assemblage in Section 2.4, alteration intensity in Section 2.5, alteration textures in Section 3.1 and 3.2, distribution and zonation in Section 3.3, and timing in Section 3.5). It also explains alteration indices, and the physical and geochemical techniques for determining the intensity of alteration. Alteration data sheets, which visually

combine mineralogical, textural and chemical data for altered volcanic rocks, are introduced. These alteration data sheets are used in Chapters 5, 6 and 7 to present examples of alteration facies associated with various alteration processes and different VHMS deposits.

2.1 | FREQUENTLY ASKED QUESTIONS Most geologists are introduced to the basic principles of hydrothermal alteration when they are students. However, they typically have a limited knowledge of how to recognise, characterise and interpret altered rocks. Common questions are: • Was the rock altered? • What was the nature of the alteration? • Was the rock hydrothermally altered? • How do we name the altered rock? • What was the original rock? To address these questions we need to make some simple observations, which include the recognition of primary minerals and textures, alteration colour, mineral assemblage, texture and intensity, overprinting relationships, and alteration distribution patterns. These observations can be made at a variety of scales: map, outcrop, hand-specimen and thinsection scales. Was the rock altered? Few volcanic rocks in submarine settings are entirely unaltered and most altered rocks are easily recognised as such. The most effective method of determining if a rock is altered is by comparing it with other samples from the same unit. Observed differences in mineral assemblage, texture and colour may indicate a spectrum from fresh, or least-altered, to significantly altered samples. Some indicators of alteration in submarine volcanic facies may be: • absence of glass • colour differences • presence of abundant minerals that typically form during alteration, such as clays, zeolites, chlorite, micas, kaolinite,

16 | CHAPTER 2

tourmaline, apatite, alunite, epidote, carbonates and quartz association between a distinctive mineral assemblage and sulfides presence of halos around veins, faults, intrusions and mineralised rock lack of, or only partial preservation of, primary textures hardness of the rock: if not silicified, altered volcanic rocks tend to be softer than unaltered volcanic rocks, which are typically glassy or crystalline, hard and brittle degree of deformation: clay- or phyllosilicate-altered rocks are commonly more deformed than unaltered or leastaltered rocks because deformation-related strain is typically partitioned into softer altered rocks.

What was the nature of the alteration? Characterising the nature of the alteration can be challenging. Nevertheless, systematic descriptions of alteration mineral assemblages, alteration textures, preservation of relict minerals and textures, patterns of distribution and overprinting relationships combined with interpretations of alteration indices and compositional changes provide important information for subsequent classification and genetic interpretation. To ensure that the data are meaningful, systematic schemes for core logging and sample description should be employed. Figures 2.1 and 2.2 are examples where detailed observations in drill core and hand specimen have established the characteristics of the altered rocks. Was the rock hydrothermally altered? Hydrothermal alteration can be discriminated from metamorphism and diagenesis, which are typically regionally extensive processes that result in weakly altered rocks and preserve delicate volcanic textures. In contrast, hydrothermal alteration styles, especially those associated with mineralisation, are local in their distribution, have variable intensity (from weak to intense) and generally destroy primary textures. Discriminating accurately between different alteration styles (Section 8.1) requires knowledge of: the host rock; alteration intensity; distribution; timing; mineralogical, textural and chemical changes; and comparison with changes related to diagenesis, metamorphism and hydrothermal alteration, which have been documented in well-preserved, geologically young, submarine volcanic successions.

How do we name the altered rock? Typically, rocks that are only weakly to moderately altered, in which primary textures and minerals can be easily recognised, are given precursor names (e.g. quartz-phyric pumice breccia). In contrast, rocks that are intensely altered, in which few primary textures or minerals are recognisable, are given alteration names (e.g. quartz-augen schist or massive chlorite rock). This is similar to metamorphic rocks where low-grade rocks are given precursor names — the prefix 'meta-' is assumed — and pervasively deformed and metamorphosed rocks are given metamorphic names.

What was the original rock? Outcrops and hand specimens of ancient volcanic rocks rarely exhibit clear evidence of their modes of eruption and emplacement. In many cases, the best we may hope for is to recognise features that help distinguish coherent volcanic facies from volcaniclastic facies. The simplest approach to recognising the primary rock is to move out of the altered zone and examine unaltered rock. However, ancient volcanic successions rarely contain unaltered rocks. As a result we rely on the preservation of relict textures and minerals in altered rocks to provide a guide to the interpretation of the primary volcanic facies. Relict textures are original pre-alteration features that have not been destroyed by alteration. Relict textures are most likely to be visible in polished hand specimens with the aid of a hand lens and in thin sections cut parallel to the tectonic foliation. There are a small number of volcanic, devitrification and hydration textures, and components or structures that usually survive diagenesis, moderate hydrothermal alteration, low-grade metamorphism and deformation - these are particularly helpful in deciphering the primary volcanic facies. For example, porphyritic texture, spherulites, lithophysae, micropoikilitic texture, perlite, flow banding, columnar joints and pillows are all characteristic of coherent volcanic facies. Volcanic components such as pumice and scoria clasts, glass shards, accretionary lapilli and non-vesicular volcanic bombs or blocks, as well as bedding and cross stratification, are characteristics of volcaniclastic facies. For a more detailed discussion of volcanic, devitrification and hydration textures, components and structures that help to determine the host volcanic facies, readers are referred to McPhie et al. (1993) Volcanic Textures: a guide to the interpretation of textures in volcanic rocks.

Primary crystals and crystal fragments are found in a wide variety of volcanic facies and can also be helpful indicators of the host volcanic facies. Whole crystals and crystal fragments in volcanic facies are mainly derived from porphyritic magmas. Crystals may be liberated from magmas during volcanic processes (explosive eruption or auto fragmentation) or by surface sedimentary processes. The shape and distribution of crystals in an altered volcanic rock can be used as a guide to whether the primary facies was coherent or clastic. In pyroclastic facies, angular and broken crystal fragments are much more common than whole euhedral crystals. In autoclastic facies, whole crystals and clusters of jigsaw-fit crystal fragments are common. In contrast, coherent volcanic facies typically, but not necessarily, contain very few broken crystal fragments. The distribution of crystals and crystal fragments in volcaniclastic facies may be random, related to size or density sorting, or concentrated in particular clasts, clusters or lenses. Crystal-bearing coherent facies are porphyritic; they contain evenly distributed euhedral crystals in a fine-grained or glassy groundmass. Although relict textures and primary crystals can be used as a guide, the discrimination of coherent and volcaniclastic facies in altered volcanic rocks is not trivial. In originally glassy volcanic rocks, alteration may produce convincing pseudotextures such as pseudobreccia, false polymictic texture, false thin-bedded texture and pseudomassive texture (Section 3.2: Allen, 1988).

DESCRIBING ALTERED VOLCANIC ROCKS I 17

FIGURE 2.1 | Part of a drill core graphic log using a modified standard logging sheet, which incorporates volcanic and alteration fades descriptions. This drill core, EHP319, is from western Tasmania and includes a thick interval of Central Volcanic Complex rocks. Abbreviations: So = bedding, S, and S2 = tectonic foliations, LCA = long core axis, cc = calcite, chl = chlorite, fsp = feldspar, qtz = quartz, ser = sericite and gb = graded bedding.

18 | CHAPTER 2

FIGURE 2.2 I This sample description - for a rock sample from 245.2 m depth in drill core EHP319 - shows the main descriptive fields for alteration studies. Abbreviations: S2 = regional cleavage, chl = chlorite, fg = fine grained, fsp = feldspar, hem = hematite, plag = plagioclase, py = pyrite, qtz = quartz and ser = sericite.

DESCRIBING ALTERED VOLCANIC ROCKS

Crystal assemblages may reflect the primary volcanic composition; they are relicts of original magmatic mineral assemblages. For example, a rock containing abundant quartz crystals was probably derived from a quartz-phyric magma, which was likely of rhyolitic composition (Table 2.1). In cases of low temperature (<200°C), weak to moderate intensity alteration, the alteration mineral assemblage may also be a guide to the primary composition of the volcanic fades. Alteration minerals rich in Fe, Mg and Ca are common in mafic volcanic rocks; K- and Na-rich minerals in felsic rocks. Typical alteration minerals in mafic rocks are chlorite, epidote, calcite, palagonite, zeolites, albite, micas, actinolitetremolite and clays (Table 2.2). In contrast, common alteration minerals in felsic rocks are quartz, micas, feldspars, zeolites, cristobalite, opal and clays, especially montmorillonite and kaolinite (Table 2.2). At temperatures above 200°C and at high water-rock ratios, the alteration mineral assemblage formed is less dependent on primary host composition and more on the fluid composition, temperature, permeability and pressure (Browne, 1978; Henley and Ellis, 1983; Reyes, 1990). Generally, consideration of a combination of field relationships, relict textures and mineral assemblages will enable interpretation of coherent versus clastic, and felsic versus mafic volcanic facies, in all but the most intensely altered volcanic rocks. Beyond that we must resort to lithogeochemical techniques (Chapter 4).

2.2 I ALTERATION NOMENCLATURE A variety of approaches have previously been taken to the classification of alteration and altered rocks, particularly hydrothermal alteration associated with different styles of mineral deposits. Common methods of alteration nomenclature are mineral based, compositional, generic, or use terminology that reflects a combination of mineralogical and textural characteristics (e.g. alteration facies). Discussions of alteration nomenclature appear in Meyer and Hemley (1967), Rose and Burt (1979), Beane (1982), Titley (1982), Guilbert and Park (1986) and Thompson and Thompson (1996).

Mineral-based alteration nomenclature Classifying altered rocks in terms of mineral assemblage was discussed in detail by Creasey (1959). Mineral-based classification involves field and petrographic observations, in some cases supported by other analytical techniques (e.g. microprobe, X-ray diffraction and SWIR spectroscopy). It is based on direct observations and provides the simplest nongenetic approach to naming alteration and altered rocks. There are two levels of mineral-based alteration nomenclature: (1) terminology based on the dominant

TABLE 2.1 | Summary of the common volcanic rock compositions, their chemical classification (SiO2 content) and likely phenocryst minerals. SiO2 contents for unaltered modern subduction-related volcanic rocks are from Ewart (1979).

Rhyolite

>69

K-feldspar (orthoclase)

± quartz ± plagioclase ± biotite ± muscovite ± amphibole ± pyroxene ± fayalite

Dacite

63-69

Na-plagioclase

± quartz ± biotite ± amphibole ± pyroxene

Andesite

52-63

Na- or Ca-plagioclase + biotite or amphibole or pyroxene

± quartz ± K-feldspar ± olivine

Ca-plagioclase + pyroxene

± olivine ± hornblende

Basalt

<52

|

TABLE 2.2 | Common alteration minerals that replace glass and magmatic minerals in volcanic rocks. Alteration minerals are from Schwartz (1959), White and Sigvaldason (1962), lijima (1978), Hay (1978), Honnorez (1978), Brey and Schmincke (1980), Tucker (1987) and Utada (1991).

Silicic volcanic glass

Zeolites (mordenite, clinoptilolite, laumonite, analcime, heulandite), cristobalite, opaline silica, quartz, calcite, clays (montmorillonite, smectite, mixed-layer clays)

Mafic volcanic glass

Palagonite, nontronitic clays, smectite, calcite, chlorite, epidote, Ca-rich zeolites, Fe/Ti/Mn-oxides

Magnetite, ilmenite and titano-magnetite

Pyrite, leucoxene, titanite, pyrrhotite, hematite

Pyroxene, amphibole, olivine and biotite

Chlorite, illite, quartz, calcite, pyrite, anhydrite

Plagioclase

Calcite, albite, adularia, wairakite, quartz, anydrite, chlorite, illite, kaolin, montmorillonite, epidote, sericite

Anorthoclase, sanidine and orthoclase

Adularia, albite, sericite

Quartz

Microcrystalline quartz

19

20 | CHAPTER 2

mineral; and (2) the use of the complete or abbreviated alteration mineral assemblage. Some authors also use negative mineral-based names, such as K-feldspar-destructive alteration (e.g. Gustafson and Hunt, 1975). The simplest method of alteration nomenclature uses the dominant or most recognisable mineral phase in the altered rock. Examples of this are albitic, which is dominated by albite; silicic, dominated by quartz; chloritic, dominated by chlorite; and sericitic, dominated by sericite. In addition, the terms chloritisation, sericitsation, silicification and carbonitisation are common in VHMS literature in reference to the processes of alteration (e.g. Sangster, 1972; Paradis et al., 1993). They, like the terms alteration and mineralisation, are widely misused (Solomon, 1999). Deciding which mineral is dominant in an altered zone is not always a straightforward task. Several minerals may be obvious and their proportions may vary. In addition, common alteration minerals, such as sericite, can occur as the dominant mineral in several different mineral assemblages that have different origins, timing and economic significance. In fact, sericitic assemblages are probably the most abundant and widespread of all alteration assemblages. They are present in aluminous rocks in nearly all types of hypogene alteration associated with ore deposits (Meyer and Hemley, 1967). Along with sericite, carbonates, chlorite and quartz are among the most widespread alteration minerals (Meyer and Hemley, 1967). Alternatively, more detailed alteration mineral assemblages can be used; either complete assemblages of all the visible alteration minerals, or abbreviated assemblages of the most abundant and distinctive minerals. Minerals are usually listed in order of decreasing abundance; thus a mixture of 60% sericite, 35% quartz and 5% pyrite becomes the sericite + quartz + pyrite alteration assemblage. This nomenclature has the advantage of clearly defining the alteration assemblage. However, some confusion may exist where mineral assemblages contain identical or similar minerals in different abundances. For example, sericite + quartz + pyrite is easily confused with sericite + chlorite + quartz + pyrite. Dana's Textbook of Mineralogy (Dana, 1957) defined sericite as 'fine scaly muscovite united in fibrous aggregates and characterized by its silky lustre'. The term has since been widely used to refer to all fine-grained pale-coloured micas, and indeed almost any fine-grained aggregates of pale-coloured layer-lattice minerals (Whitten and Brooks, 1972), particularly in hydro thermally altered and low-grade metamorphic rocks. White mica is the preferred term to avoid the ambiguity in sericite where compositional differences such as sodic muscovite, muscovite and phengite, may be important (e.g. Yang, 1998). In this book, we always use sericite in the loose sense, referring to fine-grained pale-coloured micas of undetermined composition. We use the non-specific alternative term: white mica, where compositions are known (e.g. in discussions of mineral chemistry in Sections 4.2 and 8.2).

Compositional alteration nomenclature Chemical methods of assessing hydrothermally altered rocks (i.e. lithogeochemistry) have been applied in mineral

exploration, especially around VHMS and porphyry Cu deposits, leading to the classification of alteration by compositional changes that occurred during alteration (Hemley and Jones, 1964). Examples of this include Nametasomatism or soda-metasomatism, Mg-metasomatism and K-metasomatism (e.g. Hemley and Jones, 1964) or Kenriched alteration, Ca-enriched alteration and Mg-enriched alteration (e.g. Elliott-Meadows and Appleyard, 1991), and Na-depleted alteration (e.g. Date et al., 1979, 1983). There are several problems with compositional alteration nomenclature: (1) it becomes increasingly complicated where more than one element is mobilised during alteration, which is almost always the case in the alteration of volcanic rocks; (2) it cannot be applied in the field, as it requires a detailed knowledge of the addition and removal of elements; and (3) the character of the chemical alteration can only be accurately determined if a least-altered protolith can be unequivocally identified (Section 4.1).

Generic alteration nomenclature A number of generic terms, such as advanced argillic, intermediate argillic, phyllic or sericitic, potassic, propylitic, skarn and greisen, have been applied to common alteration mineral assemblages or groups of assemblages (Table 2.3: Meyer and Hemley, 1967). These terms are widespread in the geological literature, however they are not always clearly defined or uniformly applied by different authors, and are less precise than alteration assemblages. Many workers apply generic terms based on the occurrence of indicator minerals rather than complete alteration assemblages, with the result that the terms are not always distinguishable in their usage (Rose and Burt, 1979). To apply these terms rigorously, alteration mineral assemblages for specific host rocks need to be identified and correlated. Generic alteration nomenclature tends to reflect detailed work on altered rocks associated with particular deposit types or geothermal systems, specifically porphyry, skarn, mesothermal vein and epithermal deposits (Table 2.4). In each case, the generic classification conveys a sense of the mineralogical composition and implies knowledge of alteration processes and environment of formation. Although generic classification of altered rocks surrounding an ore deposit can be useful, the reliance on understanding the environment of formation can cause problems and may incorrectly imply genetic processes. Thus, generic classification is best avoided during the early stages of recognition, description, mapping and interpretation of altered rocks in favour of a more rigorous and descriptive classification.

Descriptive nomenclature — alteration facies The term alteration facieswas first proposed by Creasey (1959) in an attempt to standardise the subdivision of hydrothermally altered rocks in a similar manner to metamorphic facies, which are assemblages of co-existing metamorphic minerals that characterise particular pressure and temperature regimes during metamorphism (Yardley, 1989). Creasey's concept of three chemically and mineralogically distinctive alteration

DESCRIBING ALTERED VOLCANIC ROCKS | 21 TABLE 2.3 | Generic alteration terms based on common alteration mineral assemblages. Modified after Creasey (1959), Meyer and Hemley (1967), Lowell and Guilbert (1970), Rose (1970), Gustafson and Hunt (1975), Rose and Burt (1979), Beane and Titley (1981), Guilbert and Park (1986), Beane (1982), and Thompson and Thompson (1996). Forsimplic,*, skarn implies a limestone or dolomite host rock.

Argillic

Kaolinite (or halloysite, metahalloysite or dickite) + montmorillonite ± sericite (or muscovite) ± chlorite

Porphyry Cu, high-sulfidation epithermal, lowsulfidation epithermal, geothermal

Advanced argillic

Pyrophyllite + kaolinite (or dickite) ± quartz ± sericite ± andalusite ± diaspore ± alunite ± topaz ± zunyite ± enargite ± tourmaline ± pyrite ± chalcopyrite ± hematite

Porphyry Cu, high-sulfidation epithermal, lowsulfidation epithermal, geothermal

Intermediate argillic

Chlorite + sericite ± kaolinite ± montmorillonite ± illitesmectite ± calcite ± epidote ± biotite ± pyrite

Porphyry Cu, high-sulfidation epithermal

Porphyry Cu

Phyllic (or sericitic)

Sericite + quartz + pyrite ± biotite ± chlorite ± rutile ± leucoxene ± chalcopyrite ± illite (Note: K-feldspar absent)

Sericitic (or phyllic)

Sericite + quartz + pyrite ± K-feldspar ± biotite ± calcite ± dolomite ± chlorite ± andalusite ± chloritoid ± albite ± pyrrhotite

Porphyry Cu, low-sulfidation epithermal, geothermal, VHMS , sediment hosted massive sulfide

Propylitic (or saussuritization)

Epidote (or zoisite or clinozoisite) + chlorite + albite ± carbonate ± sericite ± montmorillonite ± septachlorite ± apatite ± anhydrite ± ankerite ± hematite ± pyrite ± chalcopyrite

Porphyry Cu, high-sulfidation epithermal, lowsulfidation epithermal, geothermal

Potassic

K-feldspar (orthoclase) + biotite + quartz ± magnetite ± sericite (or muscovite) ± albite ± chlorite ± anhydrite ± apatite ± rutile ± epidote ± chalcopyrite ± bornite ± pyrite

Porphyry Cu

Greisen

Muscovite (or sericite) + quartz + topaz ± tourmaline ± fluorite ± rutile ± cassiterite ± wolfranite ± magnetite ± zunyite ± K-feldspar

Porphyry Cu, porphyry Sn

Calcic skarn (or tactite)

Pyroxene + garnet + wollastonite ± epidote (or zoisite) ± actinolite-termolite ± vesuvianite ± pyrite ± chalcopyrite ± sphalerite

Porphyry, skarn

Magnesian skarn

Forsterite + diopside + serpentine + talc ± actinolitetremolite ± calcite ± magnetite ± hematite ± chalcopyrite ± pyrite ± sphalerite

Porphyry, skarn

Retrograde skarn

Calcite + chlorite ± hematite ± pyrite

Porphyry, skarn

Quartz + pyrite + hematite

Sedimented-hosted Au, VHMS

Skarn

Jasperiod

fades — propylitic, argillic and potassium silicate facies — was abandoned for a wide variety of generic and non-generic terms. Subsequently, Riverin and Hodgson (1980) proposed that alteration facies be used as a descriptive term to refer simply to altered rocks that could be identified during the course of mapping or in hand specimen. They described a 'spotted facies' characterised by a well-developed spotted texture due to large, strongly altered, cordierite porphyroblasts, and a 'silicified facies' that lacked spots and was typically grey in colour and siliceous in appearance. More recently, the concept of alteration facies has been expanded to incorporate other descriptive elements, particularly alteration mineral assemblages (e.g. Gibson et al., 1983; Elliott-Meadows and Appleyard, 1991; Paradis et al., 1993; Tiwary and Deb, 1997; Brauhart et al., 1998; Gifkins and Allen, 2001). Examples are 'mottled quartz-epidote' and

'silicifi cation alteration facies' described by Gibson et al. (1983) in the Amulet Rhyolite of Noranda, Canada, and 'domainal feldspar-quartz-sericite', 'fracture-controlled chlorite-sericite' and 'stylolitic chlorite-sericite-hematite alteration facies' described by Gifkins and Allen (2001) in a regional study of alteration in the Mount Read Volcanics, western Tasmania. The advantage of characterising alteration in terms of alteration facies is that it is a purely descriptive scheme in which the basic criteria used to classify the alteration can be recognised and established in the field or in hand specimen. More importantly, by using a combination of textural and mineralogical terms, alteration facies convey the general appearance of an altered rock. Also, the descriptive variables in the alteration facies provide information that is critical to subsequent genetic interpretations of the alteration process (e.g. diagenetic, metamorphic, or hydro thermal).

22 | CHAPTER 2 TABLE 2.4 | Examples of different alteration nomenclature (i.e. dominant mineral, abbreviated mineral assemblage, compositional and generic terminology) applied to altered rocks in a variety of ore deposit environments.

VHMS deposits Silicic Chloritic Sericitic Albitic Carbonate

Porphyry deposits Kaolinitic Pyrophyllitic Kaolinitic Sericitic Feldspathic Biotitic Chloritic

Epithermal deposits Silicic Al unite K-mica or kaolinite Chloritic Sericitic Sediment-hosted deposits Silicic Silicic Tourmaline Carbonate Sericitic Albitic

Quartz + sericite + pyrite ± chlorite ± K-feldspar Chlorite + pyrite + sericite ± quartz Sericite ± quartz ± chlorite ± pyrite Albite + sericite ± quartz Dolomite/siderite/ankerite ± quartz ± sericite ± chlorite ± pyrite

Si-metasomatism Mg-metasomatism K-enrichment Na-depletion Ca, Mg, or Mn-metasomatism

Not used in VHMS literature

Kaolinite + montmorillonite ± sericite + chlorite Pyrophyllite + kaolinite ± quartz ± sericite Kaolinite + chlorite + sericite ± montmorillonite ± illitesmectite ± calcite ± epidote ± biotite Sericite + quartz + pyrite ± chlorite ± biotite K-feldspar ± biotite ± quartz ± sericite ± albite ± anhydrite ± epidote Biotite + K-feldspar + magnetite ± quartz ± albite ± anhydrite Chlorite + epidote + albite ± carbonate + sericite ± montmorillonite ± pyrite

K, Ca, Mg-metasomatism K, Ca, Mg-metasomatism K, Ca, Mg-metasomatism

Argillic Advanced argillic Intermediate argillic

Na, Ca, Mg-metasomatism K-metasomatism

Phyllic Potassic

K-metasomatism

Potassic

Ca-Mg-metasomatism

Propylitic

Quartz ± chalcedony ± alunite ± barite ± pyrite Alunite + kaolinite/dickite + quartz/cristobalite ± pyrophyllite ± diaspore ± pyrite ± topaz ± andulusite Kaolinite/dickite + illite-smectite ± quartz ± pyrite Chlorite + calcite + epidote + albite ± pyrite Sericite + illite-smectite ± quartz ± calcite ± dolomite ± pyrite

Si-enrichment Ca, Mg, Na-depletion

Silicic Advanced argillic

K, Ca, Mg, Na-metasomatism Ca, Mg-metasomatism K-metasomatism

Intermediate argillic Propylitic Argillic

Quartz + pyrite + hematite Quartz ± muscovite ± carbonate + pyrite + pyrrhotite Tourmaline ± muscovite ± quartz ± pyrrhotite Ankerite/siderite/calcite + quartz ± muscovite ± pyrrhotite Sericite + chlorite + quartz ± pyrrhotite ± pyrite ± albite Albite + chlorite + muscovite ± biotite

2.3 | ALTERATION FACIES THE RECOMMENDED METHOD We advocate a multi-faceted, descriptive approach to studying altered volcanic rocks. Different alteration facies can be defined not only on the basis of their mineral assemblage and texture, but also on distribution, intensity and composition (or compositional changes). This approach to describing and naming alteration facies is similar to the nomenclature scheme adopted by McPhie et al. (1993) for volcanic facies and the descriptive scheme for diagenetic calcite used by Folk (1965). The four alteration variables are: mineral assemblage, texture, distribution and intensity (Fig. 2.3). However, because it is not always practical to provide information on all four variables, we suggest that the alteration mineral assemblage and at least one other variable be used. Ideally, descriptive names for alteration facies follow the formula: intensity + distribution + texture + mineral assemblage. The intensity variable (Section 2.5) provides information on the degree of mineralogical, compositional and textural

Jasperiod Tourmalinite

modification (i.e. subtle, weak, moderate, strong or intense). It is determined from petrographic descriptions in combination with compositional data (e.g. Na2O) and alteration indices. The distribution variable (Sections 3.3 and 3.4) refers to the mappable extent of the alteration facies and its relationship to host facies or components, structures, mineralised rock, veins and other alteration assemblages (i.e. local or regional; footwall or hanging wall; stratabound, pipe or plume). The texture variable (Sections 3.1 and 3.2) refers to the alteration texture that is superimposed on the rock, and is typically described in hand specimen and/or thin section. It may incorporate the shape, form, grainsize or fabric in the altered rock (e.g. pervasive, selective or vein halo). The alteration mineral assemblage (Section 2.4) is expressed as an abbreviated alteration mineral assemblage in which the minerals are listed in order of decreasing abundance (e.g. the assemblage feldspar > quartz > sericite becomes feldspar + quartz + sericite). This approach produces alteration facies names such as weak, regional, selective, chlorite + sericite alteration facies or strong, massive, footwall, quartz + sericite alteration facies.

DESCRIBING ALTERED VOLCANIC ROCKS | 23

FIGURE 2.3 | Descriptive names for alteration facies.

2.4 | ALTERATION MINERAL ASSEMBLAGE Mineral assemblage refers to specific, and usually characteristic, observed mineral associations that may be in equilibrium or disequilibrium. An equilibrium mineral assemblage is a group of minerals formed at the same time, lacking any indication of disequilibrium, such as replacement or veining textures, and hence interpreted to have formed due to the same process

and under the same fluid-rock conditions (Hemley and Jones, 1964). Disequilibrium or metastable mineral assemblages are common and caution must be exercised in equating coexistence with stable equilibrium (Meyer and Hemley, 1967; Rose and Burt, 1979). In general, alteration is a process of re-equilibration. The pre-existing mineral constituents in a rock become unstable under changed physicochemical conditions (e.g. the addition of hydrothermal fluid) and progressively alter to a new stable mineral assemblage, with or without

24 | CHAPTER 2

metasomatic chemical changes. The alteration process may be only partially completed and may result in a disequilibrium assemblage containing a mixture of the pre-existing and new alteration minerals. Indeed, disequilibrium assemblages are typical of altered volcanic rocks. Common examples, at low metamorphic grade, are domainal devitrification of felsic glass and incipient sericitisation of feldspar crystals. Subsequent overprinting alteration may complicate disequilibrium assemblages. Volcanic rocks commonly retain relicts of primary minerals (especially as phenocrysts) and alteration minerals from several stages of diagenetic, metamorphic and/or hydrothermal alteration (e.g. Fig. 2.4). Equilibrium assemblages may be attained in zones of intense hydrothermal alteration or metamorphism, but primary equilibrium assemblages are rarely preserved in ancient volcanic rocks. This is true even in least-altered rocks. When mapping altered rocks it is important to recognise disequilibrium assemblages and correctly attribute minerals to the various processes of formation. Equally important is an understanding of the effects and constraints that earlier alteration facies, at various scales, may impose on subsequent processes.

Stage 1: Hydration Glassy plagioclase-phyric coherent rhyolite with perlitic fractures. Fracture surfaces are coated with clay minerals.

Stage 2: Diagenetic alteration Partly clay + zeolite-altered plagioclase-phyric coherent rhyolite. The alteration facies distribution is controlled by the perlitic fracture pattern.

Stage 3: Hydrothermal alteration

Tools for mineralogical determination The first steps in the identification of alteration minerals are to apply the three essential field tools: the practised geological eye, the hand lens and the scriber. These are frequently adequate for useful descriptions of alteration mineral assemblages, mapping of altered zones and interpretation of styles or processes of alteration. Large-scale features must not be overlooked; these provide the geological context that is critical for interpretation of processes. Simple chemical field tests, such as the use of dilute hydrochloric acid for discriminating carbonates, and sodium cobalt nitrite for staining K-feldspar, can also be useful. However, when alteration minerals occur as fine-grained masses additional instrumental techniques, such as microscopic petrography, short wavelength infrared spectrometry, X-ray diffraction and micro-analyses, may be necessary to identify them. In many situations, such as mineral exploration, the practising geologist must rely largely on field skills, perhaps augmented with limited laboratory work to substantiate and assist in developing an 'eye' for particular mineral assemblages.

Moderately sericite + quartzaltered plagioclase-phyric coherent rhyolite. Pervasively developed sericite + quartz has replaced all glass and previously altered domains. Some clay-altered relicts have been altered to sericite. Plagioclase phenocrysts are partly altered to sericite.

Stage 4: Hydrothermal alteration Intense chlorite + pyritealtered plagioclase-phyric coherent rhyolite. Vein-halo chlorite + pyrite associated with cross-cutting chlorite + carbonate veins has overprinted and destroyed earlier clay and sericite + quartz alteration assemblages and textures.

FIGURE 2.4 | Cartoons of the microscopic textural and mineralogical evolution of an originally glassy plagioclase-phyric coherent rhyolite. Overprinting hydration, diagenesis and two stages of hydrothermal alteration are visible in the final rock.

Polished slabs The identification of primary and alteration minerals, textures, and overprinting relationships can often be facilitated by the careful examination of polished slabs using a hand lens or simple binocular microscope. Polished slabs can be made from drill core or hand specimens that have been sawn to produce a relatively flat surface, which is subsequently ground smooth using a diamond lap. At this stage many coarser minerals and textures will be visible on the wet surface. The resolution of finer features can be improved by polishing the slab surface on a rotating metal lap with 220 to 400-grit zinc or iron powder and then finer powder (with water) on a glass plate. In the absence of polishing equipment, it is sometimes beneficial to

buff the sawn surface with wet sandpaper. Steel wool can be used to clean tarnished sulfides. Polished slabs are the cheapest and most readily available tools for the field geologist. Petrography Examination of standard 75 x 25 mm thin sections or polished thin sections with a polarising microscope is an excellent and relatively inexpensive method of mineral identification. It is the best way of resolving small-scale spatial relationships between minerals to assist determination of alteration reactions, paragenesis and likely processes.

DESCRIBING ALTERED VOLCANIC ROCKS | 25

Petrography is most effective if carried out by the person who maps and samples the rocks. This requires access to preparation facilities and a polarising microscope. This is not a practical solution for mineral explorers, but is still widely applicable in academia. The alternative is to send selected samples to a consultant petrographer with complete details of the geological context and the underlying objectives. Many professional petrographers are unapologetic petrologists, principally interested in petrogenesis and not enthusiastic about the obscuring effect of alteration. Therefore, it is imperative that the client informs the petrographer of the importance of alteration mineral assemblages.

X-ray diffraction X-ray diffraction (XRD) is the definitive method for the identification of all crystalline minerals, including opaque minerals and structural polymorphs with similar chemical compositions. Modern powder diffractometers can provide semi-automated analysis and computerised semi-quantitative mineral identifications from a small amount of powdered sample (Berry et al., 1983). Like the electron microprobe, these machines are mainly used as research tools. Commercial quantitative XRD is not widely available and is relatively expensive, currently around $75-95 per sample (AMDEL) in Australia.

Short-wavelength infrared spectroscopy The development of portable field instruments like the PIMA (portable infrared mineral analyser), has increased the use of short-wavelength infrared (SWIR) spectroscopy in mineral exploration and related research (Thompson et al., 1999). The technique identifies phyllosilicates, hydroxylated silicates, carbonates and sulfates in most types of dry geological samples and can also provide information on crystallinity and compositional variations in some minerals, such as clays, white mica and chlorite. These minerals, particularly phyllosilicates, are common constituents of alteration mineral assemblages and may be difficult to discriminate by other field or optical methods. Portable SWIR analysis has significant limitations in resolving complex mineral assemblages, analysing dark samples with significant opaque components and in identifying aspectral anhydrous minerals, such as quartz. It is an empirical method and does not supersede precise determinative methods such as X-ray diffraction. Nevertheless, portable SWIR has practical advantages including rapid infield analyses of up to 30 samples per hour and no sample preparation other than drying. It has many applications in the recognition and mapping of altered zones in a variety of mineral deposit styles. Thompson et al. (1999) listed recently published SWIR studies in epithermal, Archaean greenstone, VHMS, uranium, evaporite and regolith environments. Electron microprobe Micro-analysis of mineral grains by electron microprobe has become the standard tool for studies of mineral chemistry over the past few decades. It has fine resolution, down to a few microns diameter, and provides quantitative analyses of elements with atomic numbers greater than four (Be) at concentrations of greater than about 0.01 wt% (Berry et al., 1983). Major element data can be used to estimate the molecular formulae of unidentified minerals and investigate spatial variations in mineral composition. Non-destructive analyses are made on standard polished petrographic thin sections or polished grain mounts. However, the electron microprobe is an expensive laboratory instrument; it requires a skilled operator and the sample throughput is low. Consequently, it is essentially a research tool. It is rarely applied in alteration studies, mineral exploration or mapping, but is potentially useful for the verification of mineral identification and spatial compositional variations interpreted by other means.

2.5 | ALTERATION INTENSITY Alteration intensity is an indication of how completely a rock has reacted to produce new minerals and textures, and is independent of the alteration process. The alteration intensity does not reflect the new mineral species, only their abundance. It is closely linked to textural and compositional changes because it reflects the extent to which pre-existing textures and minerals (relicts of the original volcanic facies) are preserved, and the degree of metasomatism (Rose and Burt, 1979). Alteration intensity can be estimated both qualitatively and quantitatively.

Qualitative estimates of alteration intensity Qualitative estimates of the alteration intensity summarise the textural and mineralogical changes that occurred during alteration. They are based on the abundances of new alteration minerals, the degree of destruction of pre-existing minerals, the pervasiveness of alteration textures, and/or the degree of preservation of pre-existing textures. Although these features can be estimated to some extent in hand specimen, they are commonly estimated petrographically. Many geoscientists apply terms such as least altered, weakly altered, moderately altered, strongly altered and intensely altered to describe alteration intensity; however, these terms are subjective and are rarely well defined. Simmons and Christenson (1994) determined alteration intensity by measuring the percentage conversion of primary to secondary minerals, such that a weakly altered rock contained 0-33% alteration minerals; moderately altered 33-67%; and strongly altered 67-100%. Alteration intensity can also be measured by independently estimating the addition of new minerals in the groundmass and the destruction of primary phenocrysts such as plagioclase. In contrast, Guilbert et al. (in Guilbert and Park, 1986) proposed that alteration intensity be described in terms of both the growth of new alteration minerals and the destruction of pre-existing textures. Their two-part alteration intensity scale incorporates estimates of the susceptibility of minerals to alteration and the pervasiveness of alteration minerals. Mineral susceptibility is the degree to which minerals in the rock are altered, S1-S10 (vol%), whereas pervasiveness is the degree to

26 [ CHAPTER 2

which alteration minerals permeate the entire rock, PI—P10 (vol%). Despite these attempts to quantify alteration intensity, it is still applied subjectively by most geologists. For this reason we prefer to avoid a numerical system and retain the descriptive terms subtle, weak, moderate, strong and intense. The term least altered is reserved for rocks that are less altered than their counterparts in the same environment. Leastaltered rocks may be weakly to moderately altered, especially in hydrothermal environments where all rocks are altered to some degree. Here we define subtle, weak, moderate, strong and intense alteration based on the extent of growth of new alteration minerals, the destruction of primary minerals and textures, and pervasiveness of alteration textures (Table 2.5). Typically with increasing intensity of alteration, primary minerals are progressively replaced, new minerals are more pervasively distributed, primary textures are less consistently preserved, and new textures are developed (Fig. 2.5). For example, Gustafson and Hunt (1975) noted that with increasing intensity of hydrothermal K-silicate alteration at the El Salvador porphyry deposit in Chile, there is an increasing degree of replacement of plagioclase phenocrysts by Kbearing phases until the phenocrysts are obliterated. With progressively more intense alteration, the mafic phenocrysts are replaced, the groundmass becomes coarser grained with K-feldspar overgrowths, magnetite and hematite disappear, and the abundance of veins increases. Estimates of alteration intensity that incorporate textural changes are biased towards texturally destructive alteration styles such as feldspar-destructive hydrothermal alteration. It is important to recognise that under some circumstances alteration, particularly carbonate alteration and some forms of silicification, can enhance some primary or pre-existing textures (e.g. Fig. 2.6: Titley, 1982; Allen, 1988). For example, carbonate nodules preserve delicate shard textures in the Hercules footwall, western Tasmania (Fig. 2.6A: Allen, 1997), and shards are preserved in quartz nodules and quartz + chlorite (± muscovite) altered zones in the Gossan Hill footwall, Western Australia (Fig. 2.6B and C: Sharpe and Gemmell, 2001). Although these alteration styles preserve pre-existing textures, they may still be recognised as intensely altered because of the pervasiveness of the new mineral assemblage. Colour contrasts related to overprinting alteration assemblages or different mineral habits within an assemblage can enhance textures, such as clast margins, whereas another alteration assemblage with lower colour contrast and of equal intensity may preserve textures just as well but textures may be less discernable. The pervasiveness of alteration textures and the degree of preservation of pre-existing textures are dependent on the resilience of the pre-existing textures, the intensity and style of alteration (Doyle, 2001; Gifkins and Allen, 2001).

Quantitative estimates of alteration intensity Alteration indices Alteration indices are simple, multi-component or normalised ratios of lithogeochemical composition data. They are usually calculated from composition data expressed as weight percentages (wt%) or parts per million (ppm), although in some cases molar proportions are used. They are geochemical representations of hydrothermal mineral assemblages designed to facilitate discrimination of alteration styles, quantification of alteration intensity, and exploration vectors. Alteration indices have been widely applied in research and exploration forVHMS deposits (Ishikawaetal., 1976; Large etal., 2001a). They have also been used to a lesser extent in sediment hosted Zn-Pb-Ag deposits (Large et al., 2001a) and Archaean lode Au deposits (Eilu et al., 1997; Bierlein et al., 2000). Simple ratio indices, especially of molar proportions, are generally easily related to mineralogical changes (Eilu et al., 1997). However, that is not the case for some more complex indices where changes in the index value could be due to changes in one or more of three or four components, and thus related to several mineral phases. Stanley and Madeisky (1996) noted that some empirically determined alteration indices are not universally effective outside the district where they were initially developed, tend to generate many false anomalies, or may fail to identify significant altered zones, because losses of one component may cancel out gains in another. Alteration indices are formulated by placing proportions of components that were gained during alteration in the numerator and components that were lost in the denominator, thus producing the highest values in the most intensely altered rocks. In developing new indices, it is therefore useful to first apply mass transfer techniques to determine the components gained and lost. Because alteration indices are ratios, they are less affected by closure than composition data (closure is discussed in Section 4.1). They respond only to changes in the concentrations of those components used in the index, but not to all other components of the rock. Nevertheless, and contrary to the opinion of Eilu et al. (1997), alteration indices are not independent of closure because each component of composition data is affected by closure. Hence, major components that dominate igneous rock compositions, such as SiO2 and A12O3 (which is also relatively immobile), are rarely used in alteration indices for volcanic rocks. Simple indices are ratios formulated from two components of analytical data. For example, the S/Na2O ratio of Large et al. (2001a). S/Na2O shows high contrast in VHMS alteration systems, typically with values less than 0.1 in least-altered rocks and values several orders of magnitude greater in sulfidebearing, intense, proximal altered footwall zones (Fig. 2.7). Multi-component and normalised indices have two or more components added together in either or both the numerator and denominator of the index. The alkali based K2O/ Na 2 O + K2O + CaO index of Date et al. (1983) is a typical example. It has a common structure for alteration indices: the components of the numerator are also in the denominator. This has a normalising effect of limiting the possible range of values from zero to one. The normalisation in some indices involves multiplication by a factor of one hundred to

DESCRIBING ALTERED VOLCANIC ROCKS | 27 TABLE 2.5 | Descriptive alteration intensity terms (subtle, weak, moderate, strong and intense) defined on the extent of growth of new alteration minerals, the destruction of primary minerals and textures, and pervasiveness of alteration textures.

Subtle

Phenocrysts and free crystals of feldspar, quartz, and mafic minerals (amphiboles, pyroxenes etc.) were unaffected by alteration, Plagioclase may have been dusted with sericite, carbonate or hematite.

New minerals have coated the surfaces of existing phenocrysts, fractures and clasts, and infilled open space (fractures, vesicles, pore space, etc.). Glass has been devitrified.

Primary volcanic, devitrification and hydration textures are clearly visible with little or no modification,

Minor replacement/recyrstailisation (micro- or cryptocrystalline, overgrowths, poikilitic, microlitic, spherulitic, variolitic and perlitic) and infill textures,

Weak

Feldspar has been partly replaced by albite, sericite, carbonate, hematite and/or epidote. Mafic minerals have been partly replaced by Mg- and Fe-rich minerals, such as chlorite, epidote and Fe-oxides.

Patchy or domainal and disseminated selective alteration styles. Alteration commonly nucleated on existing minerals, clasts or fractures and interstitial in glomerocrysts.

Good preservation of most textures (original groundmass, matrix textures and phenocrysts). Delicate textures such as shards, pumice clasts, perlite and the fine fibrous textures in spherulites show some modification. J

Replacement, dissolution, recrystallisation, deformation and infill textures. Most common textures include: pseudomorphs, cleavage and rim texture, core and zonal texture, core and rim texture, skeletal texture, overgrowths, micro- or cryptocrystalline, dissolution vugs, stylolites, poikilitic, foliations, fiamme, and infill textures.

Moderate

Feldspar has been partly to completely replaced by feldspar, sericite, carbonate, epidote, quartz and/or magnetite, with outlines still visible. Mafic minerals commonly completely pseudomorphed. Minor recrystallisation or replacement of quartz.

Patchy or domainal and disseminated alteration styles. Individual domains may have been texturally destructive (i.e. chlorite alteration of pumice clasts). Selective alteration of individual clasts, groups of clasts or minerals. Vein-halo alteration.

Most textures modified and/or destroyed by alteration. Delicate textures commonly destroyed or substantially modified. Coarser groundmass textures (perlite, spherulites, amygdales and flow banding) and clasts partially to completely recrystallised but still clearly visible in domains.

Replacement, dissolution, recrystallisation, deformation, infill and pseudotextures. In particular: pseudomorphs, partial pseudomorphs, overgrowths, disseminated nodules, spheriods, micro- or cryptocrystalline, dissolution vugs, stylolites, fiamme, porphyroblasts, poikiloblasts, poikilitic, hornfelsic and augen textures, and foliations and lineations.

Strong

Feldspar has been completely replaced by chlorite, sericite, carbonate and/or opaques (although outlines still partly visible) and quartz partly replaced or recrystallised.

Domainal selective to pervasive. Vein-halo alteration.

Primary volcanic, devitrification and hydration textures almost completely destroyed (regardless of grainsize). Pervasive replacement of groundmass, matrix and phenocrysts. Sparse relict fiamme, amygdales and clast outlines preserved.

Replacement, dissolution, recrystallisation, deformation, infill and pseudo textures. Including: pseudomorphs, nodules, spheroids, micro-or cryptocrystalline, dissolution vugs, stylolites, fiamme, porphyroblasts, poikiloblasts, poikilitic, granoblastic, decussate, hornfelsic and augen textures, and foliations and lineations.

No primary minerals remain. Sparse outlines after primary minerals may still be visible.

Transgresses textural facies and unit contacts and primary textures. Pervasive, typically homogenous, alteration on a local scale. Veinhalo alteration.

All original rock textures including phenocrysts have been destroyed, Weak pseudomorphs or outlines of coarse phenocrysts may be visible. Primary rock type indeterminate.

Replacement, dissolution, recrystallisation, deformation, infill and pseudo textures. Including: nodules, spheriods, micro- or cryptocrystalline, rare stylolites, granoblastic, decussate and hornfelsic textures, and foliations and lineations.

Intense

,

28 | CHAPTER 2

DESCRIBING ALTERED VOLCANIC ROCKS | 29

A. Bubble-wall shards Delicate bubble-wall and platy shards (S) have been preserved within a carbonate nodule in the proximal, carbonate zone beneath the Hercules VHMS deposit. The carbonate nodule comprises quartz + calcite + chlorite-altered pumice breccia. Plane polarised light. Sample MR96-57, Cambrian Hercules Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Hercules footwall, western Tasmania.

B. Pumice shards Delicate tube pumice clasts (P) have beeen preserved in intensely quartz + chlorite (± muscovite)-altered pumice breccia from the footwall to the Gossan Hill VHMS deposit. The tube vesicles have been coated in thin films of chlorite and filled with quartz, and vesicle walls have been altered to quartz. Plane polarised light.

Sample 138752, Archaean Golden Grove Formation, Gossan Hill footwall, Western Australia.

C. Shards This quartz nodule (Q) from the footwall, quartz + chlorite (± muscovite) zone contains delicate shard textures. Plane polarised light. Sample 138795, Archaean Golden Grove Formation, Gossan Hill footwall, Western Australia.

FIGURE 2.6 | Photographs of intensely altered pumice breccias with delicate primary textures.

FIGURE 2.5 | Pairs of hand-specimen and thin-section photographs of increasing intensity of alteration in rhyolitic feldspar-phyric pumice breccia in the Hercules footwall, northern Central Volcanic Complex, western Tasmania. (A) Hand-specimen and (B) thin-section photographs of subtle, domainal, albite + sericite- and sericite + chlorite-altered pumice breccia (sample MR96-63) showing excellent preservation of volcanic textures. Plagioclase crystals are partly replaced by albite. In albite-rich domains, tube vesicles and clast margins are lined with sericite and albite + quartz altered. In contrast, pumice clasts and shards in the chlorite-rich domains are pervasively sericite + chlorite altered. The Al = 40 and CCPI = 26. (C) Hand-specimen and (D) thin-section photographs of weak, domainal, albite + sericite- and sericite + chlorite-altered pumice breccia (sample MR96-54). Volcanic textures are well preserved in the albite-rich domains and poorly preserved in the chlorite-rich domains. Plagioclase crystals (P) are sericite ± albite ± opaques altered and have albite overgrowths or nodules (alb), which locally preserve delicate vesicular textures. Elsewhere vesicles are coated in sericite and filled with albite. Pumice walls are albite + quartz altered and sericite ± chlorite + hematite fiamme and stylolites are abundant. The Al = 58 and CCPI = 37. (E) Hand-specimen and (F) thin-section photographs of moderate, pervasive, albite + sericite-altered pumice breccia (sample MR96-48) with partly preserved pumice textures and plagioclase crystals. Sericite fiamme (F) and sericite + hematite stylolites are abundant. Nodules or overgrowths of albite occur around calcite and albite + hematite-altered plagioclase crystals (P). The Al = 70 and CCPI = 38. (G) Hand-specimen and (H) thin-section photographs of strong, pervasive, quartz + sericite + pyrite-altered pumice breccia (sample MR96-50). Primary volcanic textures are faint, with sparse sericite-altered pumice clasts or fiamme (F). Plagioclase crystals (P) are polycrystalline-quartz ± pyrite altered. The Al = 98 and CCPI = 64. (I) Hand-specimen and (J) thin-section photographs of intense, schistose, quartz + sericite + pyrite-altered pumice breccia (sample MR96-46). No relict plagioclase or volcanic textures are preserved in thin section: in hand specimen irregular lenses of sericite resemble fiamme (F). This alteration facies is pervasive and strongly foliated. The Al = 99 and CCPI = 30.

30 | CHAPTER 2

FIGURE 2.7 | West-east 1700mN section through the K-lens of the Rosebery VHMS deposit, western Tasmania, showing geology and contoured S/Na2O data.

produce a potential range from zero to one hundred, which is convenient for quantification of alteration intensity. The classic example is the Alteration Index (AI) of Ishikawa et al. (1976): AI =

100(MgO + K2O) MgO K2O CaO + Na 2 O

Originally devised as a measure of intensity of sericite and chlorite alteration associated with the Kuroko-VHMS deposits, it is useful in many types of plagioclase-destructive hydrothermal alteration systems. In some cases where there are large differences in magnitudes between components, some components are multiplied by appropriate factors to adjust their effect in the index. An example is the AI mark 4 index,

AI mark 4 =

100(FeO + lOMnO) FeO + lOMnO + MgO + (SiO2/10)

which quantifies alteration in siliciclastic dolomites (Large et al., 2000). Molar proportion alteration indices are said to be more easily related to the stoichiometry of alteration reactions and hence to alteration assemblages (e.g. Eilu et al., 1997). The extra step in converting composition data to molar proportions of oxides or elements is easily achieved in computer spreadsheets but it significantly complicates manual calculations. Some examples of molar indices are the 3K/A1 sericitisation index and the CO 2 /CaO carbonation index used in exploration for lode Au deposits (Davies et al., 1990). The ACNK index of Hodges and Manojlovic (1993) used the molecular proportions of Al 2 O 3 /(CaO + Na 2 O + K2O) to quantify intensity of alteration related to metamorphosed massive sulfide deposits at Snow Lake, Manitoba.

DESCRIBING ALTERED VOLCANIC ROCKS | 31

The AI-CCPI alteration indices and box plot The well-known Alteration Index (AI) was developed in the Kuroko VHMS deposits, Japan, to represent the principal components gained (MgO and K2O) during chlorite and sericite alteration, and those lost (Na2O and CaO) during the breakdown of Na-plagioclase and volcanic glass (Ishikawa et al., 1976). The AI has since been widely used in VHMS mineral exploration to provide quantitative estimates of the intensity of alteration. It typically increases to maximum values in the proximal hydrothermal zones beneath massive sulfide lenses (e.g. Saeki and Date, 1980). The AI ranges from 0 to 100. High (> 60) values reflect high MgO and K2O contents relative to CaO and Na 2 O, and may be related to intense hydrothermal sericite and chlorite alteration. In contrast, low (< 30) AI values reflect high CaO or Na 2 O contents that may be due to intense albite or calcite alteration more typical of regional diagenetic alteration or metamorphism. For example, at Hellyer AI increases from 35 to 95 from the margin to the centre of the alteration pipe directly below the ore deposit (Fig. 2.8A and B: Gemmell and Large, 1992; Large et al., 2001a). There is a strong inverse relationship between AI and Na concentration (e.g. Fig. 2.8A and D) because loss of Na, and sometimes loss of Ca, is the major chemical change involved in the breakdown of plagioclase. In many studies Na depletion is used instead of AI as the principal measure of alteration intensity (Date et al., 1983). The Ishikawa alteration index has two major limitations (Large et al., 2001a). Firstly, it does not take carbonate alteration into account, even though this type of alteration can be significant in some VHMS alteration systems. Where Ca-carbonates are present they cause a decrease in AI, even where plagioclase destruction is extreme, because CaO is in the denominator. Secondly, the AI effectively measures plagioclase destruction but does not differentiate chloritefrom sericite-altered rocks. Variations in relative proportions of chlorite and sericite or spatial relationships between chlorite and sericite zones may be important guides to exploration in some VHMS alteration systems. A geochemical index to quantify the variation would be an improvement on subjective visual estimates. The chlorite-carbonate-pyrite index, CCPI =

100(FeO +MgO) FeO+MgO+Na2O+K2O

was developed to reflect the prominence of chlorite, FeMg carbonates, and pyrite, which are common minerals in the proximal altered zones of many VHMS deposits (Large et al., 2001a). High values of CCPI reflect high FeO and MgO contents, suggesting intense alteration to Fe- or Mgrich minerals such as chlorite, Fe-Mg-bearing carbonates (dolomite, ankerite or siderite), pyrite, magnetite or hematite. However, the CCPI of least-altered rocks is dependent on primary composition and magmatic fractionation. Mafic rocks with high primary FeO and MgO contents typically have CCPI values greater than 50, whereas more evolved felsic rocks have lower CCPI values between 10 and 50. Thus the CCPI is not well suited to the study of altered mafic rocks.

hangingwall Y volcaniclastic unit \

FIGURE 2.8 | Alteration intensity in the altered footwall zones at the Hellyer VHMS deposit, western Tasmania. (A) Schematic cross-section of the altered footwall zones and variations in alteration intensity in these zones as measured by (B) Alteration Index (Al), (C) Chlorite-carbonate-pyrite index (CCPI), and (D) Na2O. Modified after Gemmell and Large (1992) and Large et al. (2001a).

Used in conjunction with the AI, particularly graphically on x-y bivariate plots with AI as the x-axis, the CCPI provides an effective means of discriminating sericite-, chlorite- and carbonate-rich altered zones. Furthermore, the AI-CCPI bivariate plot, termed the Alteration box plot by Large et al. (2001a), discriminates these VHMS-related hydrothermal alteration assemblages from diagenetic albite- or albite + Kfeldspar-bearing assemblages. Feldspar, phyllosilicate, carbonate and several other alteration mineral compositions plot around the margins of the Alteration box plot (Fig. 2.9). Albite plots at the lower left, K-feldspar and pure muscovite at the lower right, chlorite at the top right, and carbonates along the upper margin et cetera. Calcite plots at the top left corner (although CCPI is indeterminate for pure calcite, the merest trace of Fe or Mg will result in CCPI = 100), magnesite at the top right and the Ca-Mg carbonates spread between them according to AI

32 | CHAPTER 2

0

10

20

30

40

50

60

70

80

90

100

Al (Ishikawa Alteration Index)

FIGURE 2.9 | Al - CCPI Alteration box plot for least-altered samples from the Mount Read Volcanics, western Tasmania (modified after Large et al., 2001a). The data are classified according to Ti/Zr ratios, where rhyolites have Ti/Zr <10, dacites 10-20 and andesites and basalts >20, and show the effect of magmatic fractionation on the CCPI.

determined by Mg/Ca ratios. Similarly, Mn-carbonates (except pure rhodochrosite, which is indeterminate in both indices) plot along the top margin of the Alteration box plot; their positions determined by the inevitable minor concentrations of Ca, Mg and Fe. Siderite, pyrite and Fe-oxides have CCPI values of 100, but are indeterminate for Al and thus plot as a line, rather than a point, along the top. Large et al. (2001a) found that least-altered rocks in the Mount Read Volcanics have an Al range of 20 to 65 and a CCPI range of 15 to 85 (Fig. 2.9). A compilation of 1734 geochemical analyses for unaltered volcanic rocks from various modern volcanic arcs shows a slightly smaller Al range of 20 to 60 and a slightly greater CCPI range of 10 to 90 (Fig. 2.10). Hence, least-altered volcanic rocks plot within a rectangle near the middle (somewhat left of centre) of the AI-CCPI bivariate plot. This is the least-altered 'box' that inspired the term box plot. The extent and position of the least-altered box may vary for data from different districts, according to the diversity of primary compositions. Fluid-dominated pervasive hydrothermal alteration tends to produce simple equilibrium assemblages of only a few phases. Therefore, intensely altered samples tend to plot outside the least-altered box and towards the positions of the dominant alteration minerals. For example, unaltered calc-alkaline rhyolites plot in a box towards the centre of the Alteration box plot; with increasing intensity of alteration, altered samples plot progressively further away from the unaltered box (Fig. 2.11). The relative direction of movement away from the unaltered box is controlled by the alteration assemblage and hence by the alteration process (Large et al., 2001a). Large et al. (2001a) defined 10 different mineralogical trends on the Alteration box plot. Six of these trends relate to common VHMS hydrothermal alteration styles and four are associated mainly with diagenetic alteration (Fig. 2.12).

0

10

20

30

40

50

60

70

80

90

100

Al (Ishikawa Alteration Index)

FIGURE 2.10 | AI-CCPI Alteration box plot for 1734 analyses of rocks from modern volcanic arcs; most are assumed to be unaltered. Geochemical data are from Aleutian, Andean, Indonesian and Scotian volcanic arcs, and were compiled by A.J. Stolz (electronic communication, 1998). These data (classified by SiO2 content) show the effects of magmatic differentiation on CCPI, and to a lesser degree on Al. Mafic rocks have high CCPI and low to moderate Al values because of their high Fe, Mg and high Ca contents, respectively. In contrast, felsic rocks have low CCPI because of their low Fe, Mg and high K contents, and high Al due to their low Ca and high K contents.

It is important to note that neither Al nor CCPI includes SiO2; thus the Alteration box plot does not provide a direct measure of the intensity of quartz or silica alteration. As outlined in Section 7.2, silica-altered rocks are important around some VHMS deposits, including the silicic core zone in the Hellyer alteration pipe (Gemmell and Large, 1992), the stockwork zones of some Kuroko deposits (Shirozu, 1974), and in some Cyprus-type alteration pipes (Lydon, 1984). The Alteration box plot is a powerful tool for relating lithogeochemical data to mineral assemblages and alteration intensity in VHMS systems, particularly in felsic volcanic rocks. It has obvious applications in mineral exploration for recognising favourable alteration styles, delineating altered zones and providing vectors to ore within large altered systems. A similar dual index approach may be useful for other deposit types, with different indices designed to highlight specific alteration assemblages. Element concentrations and mineral abundances Element concentrations can also be used as guides to alteration intensity. Many alteration studies have used Na depletion as a measure of hydrothermal alteration intensity (e.g. Franklin et al., 1975; Date et al., 1979; 1983; Hashiguchi et al., 1983; Ashley et al., 1988). Typically unaltered modern arc calcalkaline rhyolites have Na 2 O values between 3 and 5 wt% (Barrett et al., 1993; Stolz et al., 1996). Rhyolites with greater than 5 wt% Na 2 O are normally albitised, whereas rhyolites with less than 3 wt% Na 2 O reflect feldspar-destructive alteration styles (e.g. sericite, chlorite, pyrite, K-feldspar and carbonate). The lower the Na 2 O content the more intensely hydrothermally altered the rock; thus, Na 2 O typically

DESCRIBING ALTERED VOLCANIC ROCKS I 33

0

10

20

30

40

50

60

70

80

90

100

Al (Ishikawa Alteration Index) FIGURE 2.11 | Al - CCPI Alteration box plot for rhyolites in the northern Central Volcanic Complex, western Tasmania. Samples are of rhyolitic pumice breccias from the Rosebery and Hercules footwalls. With increasing intensity of footwall chlorite + sericite ± pyrite alteration, Al and CCPI values increase and samples plot in the upper right of the Alteration box plot.

decreases towards the centre of VHMS alteration systems (e.g. Fig. 2.13). Alteration mineral abundances quantitatively estimated from whole-rock composition data can also provide measures of alteration intensity. For example, Large et al. (2001b) found that calculated mineral percentages closely approximated the petrographic estimates of alteration mineral abundances in samples from Rosebery. By plotting the calculated mineral abundances down hole they showed that diagenetic minerals, such as albite, decrease in abundance and hydrothermal minerals, such as sericite, quartz, chlorite and Mn-carbonate, increase in abundance with proximity to ore (Fig. 2.14). Mineral abundances can be calculated as percentages from the whole-rock analyses by the least-squares method outlined in Herrmann and Berry (2002). A free copy of the MINSQ (least-squares spreadsheet method for calculating mineral proportions from whole-rock major element analyses) is available to download from the University of Tasmania's Centre for Ore Deposit Research website <www.codes.utas. edu.au>.

An integrated approach to alteration intensity A combined compositional and descriptive approach to estimating alteration intensity can also be used. In fact, the Alteration box plot is most powerful when used in combination with petrographic and/or other instrumental mineralogical studies, such as X-ray diffraction (XRD) or short wavelength infra-red spectral analysis (e.g. PIMA). Used in this way the box plot reveals trends in the data, from least altered to intensely altered, which can be related to alteration processes and hence exploration targets.

FIGURE 2.12 | Schematic AI-CCPI Alteration box plots showing the 10 alteration trends recognised by Large et al. (2001a). These provide a tool for graphically discriminating prospective from non-prospective altered zones and/or systems. (A) The six trends marked by arrows on this box plot are typical of hydrothermally altered rocks associated with VHMS deposits. Trend 1: sericite alteration at the margins of the hydrothermal alteration halo in felsic volcanic rocks. Trend 2: footwall sericite + chlorite ± pyrite alteration in felsic and mafic volcanic rocks. Trend 3: chlorite ± sericite ± pyrite alteration, typical of footwall, chlorite-dominated zones in either felsic or mafic volcanic rocks. Trend 4: chlorite + carbonate alteration typically developed proximal to massive sulfide lenses in the footwall of either felsic or mafic host rocks. Trend 5: sericite + carbonate alteration in the proximal hanging wall to ore deposits or along strike in the host rocks. Trend 6: K-feldspar + sericite, an uncommon trend developed locally within footwall felsic volcanic rocks. (B) The four trends marked by arrows on this box plot are mainly attributed to diagenetic processes and are unrelated to mineralisation. Trend 7: albite + chlorite alteration, typical of low temperature seawater-volcanic rock interaction. Trend 8: epidote + calcite ± albite alteration common in intermediate and mafic volcanic rocks. Trend 9: K-feldspar + albite alteration. Trend 10: paragonitic sericite + albite alteration.

34 | CHAPTER 2

FIGURE 2.13 | Contoured Na2O data for the west-east 1700 mN section through the K-lens of the Rosebery VHMS deposit, western Tasmania. Modified after Large etal. (2001b).

In the northern Central Volcanic Complex (Mount Read Volcanics), detailed petrographic descriptions were combined with compositional data to assess the range of AI and CCPI values for subtly to intensely altered rhyolites (Table 2.6 and Fig. 2.15). The least-altered rhyolites were subtly altered and have comparable alteration indices to unaltered modern arc rhyolites (AI = 30-60 and CCPI = 10-40). Intensely altered rhyolites have mid to high alteration indices (AI = 40-100 and CCPI = 28-100). AI ranges for subtly, weakly, moderately, strongly or intensely altered rhyolites, dacites, andesites and basalts are similar. In contrast, the CCPI is influenced by the

primary composition. For this reason the Alteration box plot should never be used independently as a method of classifying the alteration system, but should be integrated with the primary geochemical and/or petrographic data. Using a combination of alteration mineral assemblage and composition data also enables separation of rock samples into least-altered, diagenetically altered and hydrothermally altered samples (Gifkins and Allen, 2001; Large et al., 2001a). The trend from subtly to intensely hydrothermally altered rocks associated with VHMS deposits is characterised by increases in both CCPI and AI, and decreases in Na 2 O (Table 2.7).

TABLE 2.6 | Alteration indices for altered rhyolites in the northern Central Volcanic Complex, western Tasmania. The broad range in Al and CCPI values reflects different alteration styles. For example, strongly altered rhyolites with low Al values probably reflect diagenetic alteration, whereas high Al values reflect hydrothermal nyuruinermai alteration. aiieiauon.

TABLE 2.7 | Alteration indices, Na2O contents and approximate mass changes for hydrothermally altered rhyolites from the footwalls to the Rosebery and Hercules VHMS deposits, western Tasmania. AI and CCPI increase and Na2O decreases with increasing intensity of alteration.

Alteration intensity Subtle

Alteration intensity

AI

CCPI

Na2O

Mass changes

(wt%)

(g/100g)

AI

CCPI

Na2O (wt%)

30-55

10-32

3.5-5

Subtle

30-55

10-32

3.5-5

<1

2-5.5

Weak

40-60

15-45

2-A

<10

Moderate

40-75

10-55

1-2

5-30

0.5-4.5

Strong

70-90

30-90

0.5-1

15-60

0-3

Intense

90-100

28-100

0-0.5

15-100

Weak

25-60

Moderate

10-75

Strong

5-90

Intense

40-100

15^5 10-55 28-90 28-100

1-6

DESCRIBING ALTERED VOLCANIC ROCKS | 35

FIGURE 2.14 | Variations in calculated mineral abundances in samples from DDH 120R through K lens of the Rosebery VHMS deposit, western Tasmania. Increasing intensity of hydrothermai alteration towards the ore lens corresponds with increasing proportions of chlorite, Mn-carbonate and calcite, and decreasing concentrations of quartz and albite.

36 | CHAPTER 2

0

10

20

30

40

50

60

70

80

90

100

Al (Ishikawa Alteration Index) • intense A moderate • subtle • strong o weak FIGURE 2.15 | Bivariant plots of rhyolite samples from the northern Central Volcanic Complex, western Tasmania, which have been classified qualitatively. (A) AI-CCPI Alteration box plot for subtly, weakly, moderately, strongly and intensely altered rhyolites. With increasing intensity of alteration rhyolites plot away from the subtle box in all directions depending on the alteration composition and hence on the processes. (B) Al versus Na2O for subtly, weakly, moderately, strongly and intensely altered rhyolites. With increasing intensity of alteration rhyolites plot away from the subtle field.

2.6 | ALTERATION DATASHEETS A practical way of integrating different alteration data, constructing descriptive names and defining alteration facies is to use alteration data sheets. These visually display a combination of different data types for a specific rock or alteration facies on one page. They present all the relevant data for a particular rock sample or alteration facies together, and graphically illustrate relationships between texture, mineral assemblage, composition and alteration zonation, particularly with respect to an ore body. Alteration data sheets act as 'flash cards', incorporating the distinctive physical and

chemical characteristics of the altered rock or alteration facies, providing quick reference to the data collected and facilitating interpretation. Data sheets are used in Chapters 5, 6 and 7 to illustrate the dominant alteration facies or zones associated with each of the case studies. The information that is included on the data sheets may vary because the relevant or available data varies in different volcanic successions and in different deposits. Where appropriate, data sheets may incorporate: • sample number • location information • geographical or geological feature • formation or group • succession • coordinates • map, cross-section, or alteration zonation model showing the location of the sample or alteration facies • volcanic facies characteristics • descriptive name for the volcanic facies (see McPhie et al. (1993) for guidelines) • relict primary minerals • composition (e.g. rhyolitic, dacitic, andesitic or basaltic) estimated from relict primary minerals and/or geochemical data • lithofacies characteristics • relict textures • interpretation of the volcanic facies and application of genetic nomenclature (e.g. volcanogenic sedimentary deposit, resedimented mass-flow or turbidite deposit, syneruptive mass-flow deposit, autobreccia, hyaloclastite, peperite, pyroclastic-flow deposit, pyroclastic-fall deposit or pyroclastic-surge deposit) • alteration facies characteristics • descriptive name for the alteration facies • alteration mineral assemblage • alteration textures • distribution or zonation of alteration facies • alteration intensity • relative timing • interpretation of the alteration process (i.e. diagenetic, metamorphic, hydration, intrusion-related hydrothermal alteration, proximal or regional hydrothermal alteration and mineralisation, syntectonic hydrothermal alteration) • photographs of distinctive features of the alteration facies in outcrop, drill core, hand specimen or thin section • composition data (whole-rock, mineral-chemistry or isotope analyses) • chemical characteristics such as important mass changes, alteration indices (e.g. Al and CCPI) and immobile element ratios (e.g. Ti/Zr) • Alteration box plot, with the sample highlighted • other significant compositional plots, such as Ti/Zr-SiO 2 bivariant plot, mass change bar graph, SWIR spectra et cetera.

I 37

3 | COMMON ALTERATION TEXTURES AND ZONATION PATTERNS

This chapter describes common alteration textures, pseudotextures, and alteration distribution and zonation patterns in submarine volcanic successions, which can be observed at a variety of scales: map, outcrop, hand specimen and thin section. It also discusses the use of overprinting relationships in determining the paragenetic sequence. Alteration textures, patterns of distribution and zonation, and overprinting relationships are fundamental elements in describing and interpreting alteration facies (e.g. Fig. 2.3). Alteration textures can aid determination of equilibrium mineral assemblages, alteration intensity, and overprinting relationships. Alteration facies distribution and zonation patterns can be used to interpret patterns of fluid flow, changes in physicochemical conditions and development of alteration systems. Superimposed alteration patterns and overprinting textures are important for determining paragenesis involving multiple stages of alteration and hence for understanding evolution of the system over time.

3.1 | ALTERATION TEXTURES Typically alteration encompasses mineralogical and textural changes. Textural changes are changes in the shape, form, grainsize and orientation of grains within the rock and can be texturally destructive, preserve relicts of pre-existing textures or enhance textures (McPhie et al, 1993; Doyle, 2001). Alteration textures are those that are superimposed on the rock by the processes of alteration (i.e. by hydration, dissolution, diagenesis, hydrothermal alteration, metamorphism and deformation). Changes in texture during alteration may involve: the precipitation of minerals along fluid pathways; creation or infilling of pore space; the dissolution and replacement of earlier minerals and glass by subsequent minerals; and recrystallisation. There are five common types of alteration textures that occur in volcanic facies (Tables 3.1 and 3.2): (1) replacement textures, (2) infill textures, (3) dissolution textures, (4) recrystallisation textures, and (5) deformation textures. In addition, the combined effects of a number of different overprinting alteration facies can result in false or pseudotextures (De Rosen-Spence et al., 1980; Allen, 1988).

Furthermore, there are two types of textures that are common in and unique to volcanic rocks, which, although not alteration textures, influence subsequent alteration, especially the development of pseudotextures. These are hightemperature devitrification textures (i.e. spherulites, varioles, lithophysae and micropoikilitic texture) and perlitic fractures (Figs 3.1 and 3.2). The formation and alteration of perlite is described in detail in Section 5.2.

Replacement textures Most alteration forms by replacement, because pre-existing mineral phases and glass become unstable during changed geothermal conditions and are readily substituted by new, more stable minerals. Replacement is the process of practically simultaneous solution and deposition of a new mineral of partly or completely different composition either in a preexisting mineral or an aggregate of minerals (Lindgren, 1933). Although mineral exchange is essentially simultaneous, replacement may occur in stages, where intermediate products form, at least temporarily, before the final alteration TABLE 3.1 | Types of textural changes that occur during alteration. Replacement (metasomatism)

Existing minerals or glass are replaced by one or more new mineral species

Infill

A mineral or minerals are precipitated from solution into open space

Dissolution

Existing minerals or glass are leached and removed by solution with or without replacement

Static recrystallisation

Recrystallisation of existing minerals to new grains, and/or a change in morphology of the same mineral species or composition

Dynamic recrystallisation

Recrystallisation of existing minerals to new grains and/or a change in morphology and/or orientation of the same mineral species or composition

Deformation

Existing component or texture is rotated, milled, broken, compressed, modified, distorted or fractured

38 | CHAPTER 3 TABLE 3.2 | Common macroscopic and microscopic alteration textures in volcanic rocks.

Replacement

Pervasive

Pervasive, selective, massive, disseminated, microcrystalline, cryptocrystalline

Selective

Disseminated

Disseminated, pseudomorph, overgrowth, cleavage and rim texture, core and zonal replacement texture, microcrystalline, cryptocrystalline, spheroid, nodule, concretion

Domainal

Infill

Dissolution

Static

Vein halo

Pervasive, selective, disseminated, pseudomorph, overgrowth, cleavage and rim texture, core and zonal replacement texture, microcrystalline, cryptocrystalline, spheroid, nodule, concretion

Incomplete infill

Crustiform, fibrous, prismatic, spherulitic

Massive infill

Microcrystalline, prismatic

Layered or banded infill

Crustiform, colloform, comb, botryoidal

Stylolitic foliation

Stylolites, solution seams

Corrosion vug

Open pore space ± infill textures (prismatic, fibrous and massive)

Pervasive (hornfels)

recrystallisation

Equigranular, granoblastic, granophyric, decussate

Seledive

Porphyroblastic, idioblastic, xenoblastic, poikiloblastic, intergrowths, overgrowths, reaction rims, polycrystalline grains

Dynamic

Foliation

Slaty cleavage

recrystallisation

Cleavage, mineral alignment, granoblastic, porphyroblastic, poikiloblastic

Schistosity Layering (gneissosity) Lineation

Differential layering, microcrystalline, granoblastic, granophyric Aligned, strained, bent, kinked, flattened, twinned and broken grains (crystals or clasts), cleavage

Deformation

Cataclasite

No foliation, porphyroblastic, microcrystalline

Mylonite

Foliation, granular

Foliation Lineation Augen structure

minerals. For example, relict radiating fibrous textures locally preserved in feldspar-altered pumice and perlite clasts in the Mount Read Volcanics, western Tasmania, suggest that an intermediate phase between felsic glass and feldspar, possibly fibrous zeolites, occurred (Fig. 5.11: Gifkins and Allen, 2001). Replacement can range from the conversion of specific mineral phases or domains to new minerals (selective alteration, Fig. 3.3B, C and D), to complete replacement of a rock to a completely new mineral assemblage (pervasive alteration, Fig. 3.3A). Where alteration occurs dominantly by diffusion, it may affect a large volume of rock. Elsewhere it may occur along well-defined fluid pathways (vein-halo alteration, Fig. 3.3E) with its effects restricted to a scale of millimetres to metres (Titley, 1994). It is worth noting that the terms pervasive, selective or vein-halo depend on the scale of observation. For example, vein-halo alteration can appear pervasive when viewed in thin section.

Cleavage, aligned, strained, bent, kinked, flattened, twinned and broken grains (crystals or clasts), fiamme, eutaxitic Aligned, strained, bent, kinked, flattened, twinned and broken grains (crystals or clasts), cleavage Cleavage

Pervasive Pervasive alteration is extensive alteration that has completely changed the rock composition and texture at scales that range from millimetres to kilometres (Rose and Burt, 1979; Titley, 1982). Pervasive alteration is distributed without regard for pre-existing textures and can result in disseminated, massive microcrystalline or cryptocrystalline microscopic textures (Fig. 3.4A).

Selective Selective alteration converts only specific pre-existing phases to new mineral phases (Titley et al., 1978; Titley, 1982). The original rock texture may be only slightly modified during selective alteration because only certain components in the host (e.g. minerals, volcanic glass or clasts: Fig. 3.3B, C and D) are preferentially replaced, and others are left relatively unaltered (Rose and Burt, 1979).

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 39

A. Spherulites and obsidian in rhyolite Pink, isolated spherulites and densely microspherulitic flow bands are enclosed in black obsidian in this flowbanded rhyolite. Spherulites are radiating aggregates or bundles of acicular and fibrous crystals. They vary in shape from spherical to bow-tie shaped sheafs and axiolitic bundles, and are commonly composed of feldspar or intergrowths of alkali feldspar, plagioclase, cristobalite or tridymite and clinopyroxene (Lofgren, 1971b). Spherulites are typically the product of hightemperature (above the glass-transition temperature) devitrification of silicic glass (Lofgren, 1971a). Sample NG4, recent Ngongotaha lava dome, Hendersons quarry, Rotorua, New Zealand. B. Lithophysae in rhyolite This red albite + quartz + hematite-altered, flow-banded quartz + plagioclase-phyric rhyolite contains abundant spherulites and lithophysae. The lithophysae are filled with layered quartz. Sample from the Lower Devonian Snowy River Volcanics, Flukes Knob area, Victoria.

C. Varioles in basalt Dark spots in this basalt outcrop are varioles: radial or sheaf-like aggregates of plagioclase and pyroxene, olivine or iron oxides, and are similar to spherulites, but only occur in mafic facies (cf. Fowler et al., 1987; Williams etal., 1982). Shirakawa quarry, Miocene Green Tuff Belt, Odate, Japan.

D. Micropoikilitic texture in thin section The groundmass of this rhyolite is densely micropoikilitic; comprising patches of optically continuous quartz, which enclose variably oriented laths of sericitised albite. Poikilitic and micropoikilitic texture (snowflake texture) comprise an optically continuous crystal enclosing numerous randomly oriented inclusions of a different composition (Anderson, 1969). The boundaries between the micropoikilitic quartz domains in this sample are highlighted by concentrations of sericite. Cross polarised light. Sample 133921, Cambrian Mount Black Formation, Central Volcanic Complex, Mount Read Volcanics, Mount Black, western Tasmania. FIGURE 3.1 | Examples of high-temperature devitrification textures.

40 | CHAPTER 3

A. Altered macroperlite Relict macroperlitic factures in this coherent dacite are enhanced by dark grey sericite + chlorite-altered zones along and adjacent to the perlitic fractures. The arcuate shape of the fractures is preserved in some areas. The perlite cores are pink albite + quartz + sericite altered. Cambrian Mount Black Formation, Central Volcanics Complex, Mount Read Volcanics, Pieman Road, western Tasmania.

B. Relict perlite in thin section The formerly glassy groundmass of this rhyolite preserves perlitic fractures. Perlitic fractures are a network of fine typically concentric, arcuate fractures that enclose glassy or originally glassy cores. Here, the perlitic fractures are filled with dark, mixed-layer smectite/chlorite and the groundmass adjacent to the fractures is clinoptilolite altered. The perlitic cores are partly glassy and partly smectite altered. Plane polarised light. Sample J6-737 m, Miocene Nishikurosawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

C. Banded perlite This finely flow-banded, plagioclase-phyric rhyolite contains an intersecting fracture network of sub-parallel long fractures linked by short cross fractures (banded perlite) superimposed on the flow-banded texture. Sample KB257, Siluro-Devonian rhyolite, Ural Volcanics, Ural Ridges area, New South Wales.

D. Banded perlite in thin section In thin section, concentrations of sericite ± hematite mark the relict perlitic fractures. The pale flow bands comprise a fine-grained mosaic of feldspar + quartz, whereas the darker bands consist of sericite + feldspar + quartz + chlorite. Disseminated fine-grained hematite occurs throughout the groundmass. Plane polarised light. Sample KB257, Siluro-Devonian rhyolite, Ural Volcanics, Ural Ridges area, New South Wales.

FIGURE 3.2 | Examples of perlite.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 41

Two textural types of selective alteration occur: disseminated alteration (or selective-pervasive alteration), which refers to the replacement of selective pre-existing phases throughout the host rock; and domainal alteration, which refers to the alteration of patches, pods, or groups of clasts within the host rock (Fig. 3.3F, G and H). In addition, selective alteration may result in concentrically zoned alteration facies within clasts or alteration halos around clasts (Fig. 3.31, J, K and L). Selective alteration can result in a patchy or mottled appearance (e.g. Allen, 1988). Common microscopic selective replacement textures are pseudomorphs, partial pseudomorphs (cleavage and rim texture, core and zonal texture, core and rim texture and skeletal texture), overgrowths on pre-existing components, and spheroids or nodules (Fig. 3.4B to L: Dimroth and Lichtblau, 1979; Craig and Vaughan, 1981; Ineson, 1989). Carbonate and zeolite nodules are common in submarine volcaniclastic facies and can have a wide variety of grainsizes from 0.2 to greater than 20 mm (Fig. 3.41 to L: Franklin et al., 1975; Lees, 1987; Khin Zaw and Large, 1992; Hill and Orth, 1994; Allen, 1997).

Vein halo Vein-halo alteration involves the replacement of either the whole rock (pervasive alteration: Fig. 3.3E) or specific preexisting phases (selective alteration) in restricted areas, such as the halos around veins, intrusion contacts, or at stratigraphic contacts. Alteration progresses in fronts, moving out from fractures or contacts into the adjacent wall rock. Vein-halo alteration has also been termed infiltration metasomatism, vein-veinlet, reaction rim, vein-wall-rock, vein-envelope, veinlet-controlled and fracture-controlled alteration (e.g. Titley et al., 1978; Titley, 1982; Thompson and Thompson, 1996; Doyle, 2001).

Infill textures Infill or open space-filling textures result from the precipitation of new mineral phases from solution into open spaces or cavities such as pore spaces, vesicles, inter-clast space, vugs and fractures (Taylor, 1992). Infill textures are characterised by well-developed crystal faces, zoned crystals and mineral banding (Craig and Vaughan, 1981). Silicate, carbonates, oxides, sulfates and sulfides all occur as void fill in altered volcanic rocks. Infill results from the precipitation of minerals from solution. The first mineral to be deposited forms a crust on the cavity walls and grows inwards, generally with the development of inward facing crystal faces. Common infill textures include incomplete infill, massive infill, and layered or banded infill (Fig. 3.6: Taylor, 1992). These textures can include fibrous, prismatic, spherulitic or equant crystal shapes and exist on a range of scales from micrometres to metres (Dimroth and Lichtblau, 1979; Taylor, 1992).

Incomplete infill Incomplete or partial infilling of veins and cavities or dissolution of void fill can leave an open vug in the centre (e.g. Fig.3.5A). In many cases, the resulting infill texture contains well-formed crystals that project into this vug.

Massive infill Massive infill textures result from the continuous deposition of a mineral or aggregate of minerals until the cavity is filled (e.g. Fig. 3.5B). Massive infill commonly contains well-formed crystals, especially quartz, feldspar, fluorite, cassiterite, galena, sphalerite and chalcopyrite crystals. Massive, microcrystalline forms also exist (Taylor, 1992).

Layered infill Layered or banded infill textures result from the deposition of a succession of minerals inwards from the cavity or fracture wall (Bateman, 1951). Layered infill textures do not generally contain well-formed crystals, such as comb texture (e.g. Fig. 3.5H), but vary from thin layers of individual minerals to crustiform bands or colloform textures.

Dissolution textures Dissolution textures are common in altered volcanic rocks (Allen, 1990; Allen and Cas, 1990; Marsaglia and Tazaki, 1992; Gifkins and Allen, 2001). They form from the corrosion or leaching of pre-existing phases (either glass or mineral phases), with or without minor replacement by new mineral phases (Fisher and Schmincke, 1984). For example, leaching of rhyolitic glass is commonly accompanied by crystallisation of muscovite or clay minerals that absorb leached ions from solution (Karkhanis et al., 1980). Dissolution textures include corrosion vugs or dissolution pits, stylolites and solution seams (e.g. Fig. 3.6: Pettijohn, 1957).

Corrosion vugs Dissolution or corrosion of volcanic glass or pre-existing minerals can create open cavities or oversized pores (Fig. 3.6A to F) in which infill can occur synchronous with dissolution or after dissolution (Hay, 1963; Sheppard et al., 1988). In some cases pseudomorphs of minerals or originally glassy particles, such as glass shards, form by dissolution and precipitation (Riech, 1979; Sheppard etal., 1988). Riech (1979) recognised infill textured zeolites and calcite within clinopyroxenes, and proposed that clinopyroxenes were corroded during diagenesis, creating an open void that was subsequently filled with zeolites and calcite. Similarly, Hay (1963) recognised partial to complete dissolution of glass shards followed by the precipitation of authigenic minerals, especially zeolites, in the new cavities as well as in original pore space.

42 | CHAPTER 3

A. Pervasively altered rhyolite Intense, pervasive, fine-grained K-feldspar + quartz alteration has completely replaced the groundmass and plagioclase phenocrysts in this rhyolite. Sample 143286, Central Volcanic Complex, Mount Read Volcanics, Mount Darwin, western Tasmania.

B. Selectively altered phenocrysts Sericite has selectively altered the coarse prismatic feldspar phenocrysts (F) in this latite. The pale greengrey, fine-grained groundmass is moderately and pervasively phengite + chlorite + ankerite altered, and the amygdales are quartz filled. Sample 144369, Ordovician Lake Cowal Volcanics, JuneeNarromine Volcanic Belt, Endeavour 42 prospect, New South Wales.

C. Selectively altered pumice clasts Large pumice clasts (P) in this sample of crystal- and pumice-rich volcaniclastic breccia have been selectively altered to orange albite + quartz, whereas the finer grained matrix has been altered to green sericite + chlorite + albite. The domainal alteration style enhances its clastic appearance. Sample 131993, Cambrian Mount Julia Member, Tyndall Group, Mount Read Volcanics, Comstock, western Tasmania.

D. Selectively altered matrix In this andesitic volcaniclastic breccia, the matrix is moderately and selectively epidote altered. In contrast, the plagioclase-phyric clasts (C) are weakly chlorite + sericite altered. Sample 144805, Ordovician Mingelo Volcanics, JuneeNarromine Volcanic Belt, Peak Hill, New South Wales.

FIGURE 3.3 | Examples of replacement textures in hand specimen.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS I 43

E. Vein halo Red albite altered zones are restricted to 5 mm halos or selvages adjacent to quartz + actinolite + pyrite veinlets in this feldspar porphyritic dacite. Sample TH386 271.1 m, Cambro-Ordovician Trooper Creek Formation, Seventy Mile Range Group, Mount Windsor Subprovince, Thalanga, Queensland.

F. Banded, domainal alteration facies Diffuse and discontinuous pink and green bands in this massive crystal-rich volcaniclastic sandstone are defined by domains of albite + quartz ± chlorite, and chlorite + sericite + magnetite alteration facies, respectively. The bands are not obviously consistent with grainsize or component variations; they alternate on a 2—10 cm scale, are laterally extensive (10—20 m) and are commonly, but not exclusively, bedding parallel. Sample 131982, Cambrian Mount Julia Member, Tyndall Group, Mount Read Volcanics, Lyell Comstock, western Tasmania.

G. Patchy, domainal alteration facies The domainal, green epidote + quartz and grey albite + quartz + hematite alteration facies are distributed in irregular patches with diffuse margins in this coherent plagioclase-phyric dacite. Sample M142, Cambrian Mount Black Formation, Central Volcanic Complex, Mount Read Volcanics, Tullah, western Tasmania.

H. Domainal alteration facies in pseudobreccia Domainal red albite + quartz and dark green epidote + sericite + albite alteration facies in this sample of macroperlite gives it a pseudo-polymictic and -clastic texture. However, the red apparent clasts have diffuse margins and identical phenocryst populations to the apparent matrix. Sample 147550, Cambrian Mount Black Formation, CentralVolcanics Complex, Mount Read Volcanics, Pieman Road, western Tasmania.

FIGURE 3.3 | Examples of replacement textures in hand specimen, cont.

44 | CHAPTER 3

I. Zonation within clasts The andesite and basalt clasts in this polymictic volcanic breccia are concentrically zoned, with sericite + quartz + calcite-altered rims, and chlorite-altered cores. Some of the larger clasts have an additional quartz + sericite + chlorite-altered core zone. The matrix has been moderately and pervasively quartz + sericite + calcite ± chlorite altered. Cambrian Que-Hellyer Volcanics, Mount Charter Group, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer, western Tasmania.

J. Zonation within clasts Clasts in this basaltic pebble conglomerate display heterogenous alteration facies, and some clasts are internally zoned. The basalt clast (B) has a fine-grained, pale green sericite-rich rim, and darker sericite + chloritealtered core. Sample 134632, Cambrian Red Lead Formation correlate, Dundas, Kapi Creek, western Tasmania.

K. Zonation within pillows This metamorphosed, amphibolite-grade lava-pillow has a typical triangular, draped shape and is concentrically zoned. The central red zone is coarse-grained, scapolitepoor and albite + hematite + sericite ± epidote altered. The average grainsize decreases, and scapolite grainsize and abundance increases, in consecutive zones towards the rim. Biotite + calcite + hornblende + microcline + scapolite + epidote + quartz comprise the inter-pillow matrix. Proterozoic Corella Formation, Mary Kathleen Group, Malbon River, northwest Queensland.

L. Altered halos around clasts Orange albite + quartz alteration facies is distributed in a halo around a massive, albite-altered dacite clast (C) in this crystal- and lithic-rich volcaniclastic sandstone. The more pervasive green-grey domain is sericite + chlorite + quartz + albite altered. Sample 132090, Cambrian Mount Julia Member, Tyndall Group, Mount Read Volcanics, Anthony Road, western Tasmania.

FIGURE 3.3 | Examples of replacement textures in hand specimen, cont.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 45

A. Microcrystalline texture in thin section The groundmass of this rhyolite is a microcrystalline mosaic of quartz + feldspar + sericite. Quartz phenocrysts (Q) have been recrystllised. Microcrystalline texture (aphanitic) is a fine-grained granular texture where the individual crystals can be distinguished in thin section. In contrast, cryptocrystalline texture (phaneritic) is where the crystals are too minute to be distinguished even with the aid of a microscope (Williams et al., 1982). Cross polarised light. Sample 133318, Cambro-Ordovician Mount Windsor Formation, Seventy Mile Range Group, Thalanga, Queensland.

B. Pseudomorphs in thin section The plagioclase phenocrysts in this sericite + quartz + tourmaline-altered andesite were pseudomorphed by tourmaline, and subsequently almost completely replaced by blue-green chlorite. Pseudomorphs are crystals or aggregates of crystals that preserve the shape of a pre-existing mineral or particle (e.g. glass shard or pumice clast) (Spry, 1976). Plane polarised light. Sample 145199, Ordovician Forest Reefs Volcanics, Molong Volcanic Belt, Black Rock, New South Wales.

C. Pseudomorphs in thin section This thin section of plagioclase + quartz + pyroxenephyric rhyolite shows an illite pseudomorph after pyroxene. Plagioclase phenocrysts have been altered to K-feldspar and the groundmass comprises a fine-grained mosaic of K-feldspar + quartz + chlorite + smectite. Cross polarised light. Sample KB495, Siluro-Devonian Coan rhyolite, Mount Hope Volcanics, Coan Gonn Peak, New South Wales.

D. Cleavage and rim texture in thin section The plagioclase crystals in this basalt have been selectively altered by sericite along cleavage planes. Cleavage and rim textures occur by selective alteration of mineral grain boundaries and cleavages. It is common in plagioclase, in which montmorillonite, sericite or calcite form along the cleavage planes (Sales and Meyer, 1948). Plane polarised light. Sample SVD87a-104.9 m from the Cambrian Sterling Valley Volcanics, Mount Read Volcanics, Sterling Valley, western Tasmania.

FIGURE 3.4 | Examples of replacement textures in altered volcanic rocks.

46 I CHAPTER 3

E. Core and zonal texture in thin section Zones within plagioclase phenocrysts in this subtly, smectite + calcite-altered diorite have been selectively altered to sericite. These incomplete pseudomorphs, termed core and zonal texture, are particularly common in zoned feldspar, amphibole and mica crystals where the cores, or one or more zones in zoned minerals, are altered (Barker, 1990). In plagioclase crystals, like those pictured here, the calcic zones are typically altered to calcite or sericite (Sales and Meyer, 1948). Plane polarised light. Sample 152958, Pliocene-Pleistocene Luise Volcano, Lihir Island, New Ireland Province, Ladolam epithermal Au mine, Papua New Guinea. F. Core and zonal texture in thin section In this example of core and zonal texture, the core zones of plagioclase phenocrysts have been altered to sericite. The groundmass of this plagioclase + clinopyroxenephyric basalt was subtly and pervasively smectite + calcite-altered. Plane polarised light. Sample 152830, Pliocene-Pleistocene Luise Volcano, Lihir Island, New Ireland Province, Ladolam epithermal Au mine, Papua New Guinea.

G. Overgrowth texture in thin section A discontinuous K-feldspar overgrowth encloses a hematite-altered plagioclase phenocryst (P) in this strongly and pervasively albite + quartz + sericite-altered pumice breccia. K-feldspar nucleated on the plagioclase phenocryst, spread outwards filling vesicles, and replaced vesicle walls in the pumice clasts. Overgrowth textures are mineral rims that may be composed of one or more crystals of similar or different minerals. Cross polarised light. Sample 133814 from the Cambrian Hercules Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Hercules footwall, western Tasmania.

H. Altered nodules in pumice breccia Blue-green celadonite nodules have overprinted pumice clasts in this polymictic volcanic breccia. These nodules are composed of fine-grained aggregates of celadonite ± opal CT ± quartz, preserve uncompacted tube and round vesicle pumice textures, and are surrounded by pervasive smectite + mordenite + calcite alteration facies. Sample FK2, Miocene Byobu-iwa Member, Tokiwa Formation, South Fossa Magna, Green Tuff Belt, Fujikawa River, Japan.

FIGURE 3.4 | Examples of replacement textures in altered volcanic rocks, cont.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS I 47

I. Carbonate spheroids Large dolomite spheroids are enclosed in the strongly chlorite + quartz + dolomite-altered matrix of this formerly plagioclase-phyric andesite. Nodules and spheroids are spherical domains of alteration, which may comprise radiating aggregates of fibrous crystals, fine internally concentric structures, or mosaics of anhedral grains, with or without cores (Allen, 1997; Hill andOrth, 1995). Sample 135756, Cambrian Que-Hellyer Volcanics, Mount Charter Group, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer footwall, western Tasmania.

J. Carbonate spheroids in thin section In thin section, the dolomite spheroids display concentric zones and a coarse, radiating, fibrous texture. This compositional zonation in the spheroids probably indicates multiple stages of carbonate alteration (cf. Hill and Orth, 1995). Plane polarised light. Sample 135756, Cambrian Que-Hellyer Volcanics, Mount Charter Group, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer footwall, western Tasmania.

K. Carbonate spheroids Carbonate spheroids are concentrated in individual beds in this strongly chlorite + carbonate + pyrite-altered laminated volcaniclastic sandstone. The larger spheroids, which are up to 2 mm in diameter, have coalesced. Sample 138601, Archaean Mb5 Golden Grove Formation, Luke Creek Group, Murchison Volcanics, Golden Grove, Western Australia.

L. Carbonate spheroids in thin section Thin section examination shows these carbonate spheroids are supported in a fine-grained quartz + sericite + carbonate matrix. Plane polarised light. Sample 138601, Archaean Mb5 Golden Grove Formation, Luke Creek Group, Murchison Volcanics, Golden Grove, Western Australia.

FIGURE 3.4 | Examples of replacement textures in altered volcanic rocks, cont.

48 | CHAPTER 3

A. Incomplete infill in fractures Incomplete filling of fractures in this altered diorite left sub-planar vugs. The pyrite fill has botryoidal surfaces, representing rounded shapes of either spherulitic radiating aggregates of fibrous crystals or fine-concentric internal structures (cf. Jensen and Bateman, 1981). This massive plagioclase-phyric diorite has been pervasively K-feldspar + pyrite (>quartz + illite) altered, and dissolution of primary mafic minerals produced a fine, spongy, porous texture. Sample 152959, Pliocene-Pleistocene Luise Volcano, Lihir Lsland, New Ireland Province, Ladolam epithermal Au mine, Papua New Guinea.

B. Massive infill Pale green epidote altered halos surround massive chlorite-filled amygdales (A) in this basalt sample. Sample 144753, Ordovician, Junee-Narromine Volcanic Belt, Boundary Prospect, Lake Cowal, New South Wales.

C. Layered infill stringer vein This banded vein consists of successive layers, from the vein wall to centre, of quartz, quartz and intergrown chalcopyrite and pyrite, and dolomite. A thin dolomite vein has overprinted the stringer vein at an oblique angle. These veins are hosted in strongly and pervasively sericite + chlorite + albite + pyrite-altered andesite. Cambrian Que-Hellyer Volcanics, western volcanosedimentary sequences, Mount Read Volcanics, Hellyer footwall, western Tasmania.

D. Layered infill in amygdales Amygdales (A) in this basalt clast, from a basaltmudstone peperite, contain concentric layers of quartz and calcite, which have grown inwards from the vesicle walls. The basalt groundmass has been pervasively sericite + chlorite + calcite altered. The clast grainsize decreases towards the clast rim, to the left of the field of view in this photograph. Sample 76836, Cambrian Que-Hellyer Volcanics, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer, western Tasmania.

FIGURE 3.5 | Examples of infill textures in altered volcanic rocks.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 49

E. Layered infill texture in vesicles Vesicles (V) in this plagioclase-phyric pumice clast have been filled with roughly concentric layers of tan-coloured mordenite, dark smectite and clear clinoptilolite. The zeolites occur in clusters or aggregates of fine, radiating fibres. The originally glassy vesicle walls (W) have been replaced by mordenite + K-feldspar ± smectite. Plane polarised light. Sample OH8-387 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

F. Layered infill texture in amygdales Amygdales in this subtly altered perlitic rhyolite have been filled with bands of fine-grained montmorillonite and unknown radiating fibrous minerals. Pale polarised light. Sample J6-737 m, Miocene Nishikurosawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

G. Layered infill in amygdales The amygdales in this palagonite-altered trachytic basalt clast from a crystal- and lithic-rich pumice breccia are filled with layers of montmorillonite and fibrous zeolites. Plane polarised light. Sample OH8-387 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

H. Comb texture This example of comb texture shows layers of prominent sparry quartz + amethyst ± carbonate crystals projecting inwards from the vein or cavity wall. Sample T5> Cretaceous, andesite, Fresnillo epithermal district, Mexico.

FIGURE 3.5 | Examples of infill textures in altered volcanic rocks, cont.

50 | CHAPTER 3

A. Dissolution vugs This hand specimen of polymictic breccia has a spongy or vuggy porous texture due to the dissolution of primary mafic igneous minerals and glass. It has been intensely and pervasively adularia + illite + pyrite altered with illite replacing plagioclase crystals, secondary K-feldspar in the altered matrix, and disseminated pyrite. Sample 152726, Pliocene-Pleistocene Luise Volcano, Lihir Island, New Ireland Province, Ladolam epithermal Au mine, Papua New Guinea.

B. Dissolution vugs in thin section In thin section, irregularly shaped, empty, corrosion or dissolution vugs (V) are conspicuous in the matrix and clasts. Some vugs cut across clast margins. Plane polarised light. Sample 152726, Pliocene-Pleistocene Luise Volcano, Lihir Island, New Lreland Province, Ladolam epithermal Au mine, Papua New Guinea.

C. Filled dissolution vug in thin section Corrosion vugs, created by the dissolution of volcanic glass or pre-existing minerals, are commonly filled by subsequent mineral precipitation from solution. Successive layers of montmorillonite and zeolite have filled an irregular vug (V) in this thin section. The vug occurs in the matrix and in a basalt clast, crossing the clast-matrix contact. Plane polarised light. Sample OH8-387 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

D. Vuggy quartz The prominent features in this quartz-rich sample are the corrosion vugs, which were generated by the dissolution of pumice clasts and crystals from this pumice and lithic tuff. Miocene rhyodacitic pumice and lithic tuff, Pierina Au-Ag deposit, Peru.

FIGURE 3.6 | Examples of dissolution textures in altered volcanic rocks.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS I 51

E. Kaolinite + dickite-altered andesite Large (up to 4 mm), blocky feldspar phenocrysts have been kaolinite altered in this sample of massive, coherent andesite. Miocene andesitic lava, Pierina Au-Ag deposit, Peru.

F. Vuggy quartz In this sample of vuggy quartz, which is equivalent to the previous kaolinite + dickite-altered andesite, the feldspar phenocrysts have been dissolved resulting in blocky vugs. The groundmass is composed dominantly of quartz. Miocene andesitic lava, Pierina Au-Ag deposit, Peru.

G. Stylolite in thin section Stylolites (S2) in this rhyolitic pumice breccia have concentrated fine-grained opaques and sericite ± chlorite. The stylolites define the compaction foliation and are crenulated by the dominant regional cleavage {S2) defined by alignment of sericite in the subtly albite + quartz + sericite-altered matrix. Sample 147422, Cambrian Kershaw Pumice Formation, Central Volcanics Complex, Mount Read Volcanics, Rosebery, western Tasmania.

E. Solution seams in thin section These analcime-filled solution seams occur in a smectiterich fiamme, extending from the damme terminations into the shard- and crystal-rich matrix of a crystal-rich pumice breccia. They are interpreted to have formed by dissolution and precipitation under the influence of lithostatic load during diagenesis. Plane polarised light. Sample FK7, Miocene Wadaira Tuff Member, Tokiwa Formation, South Fossa Magna, Green Tuff Belt, Wadaira, Japan.

FIGURE 3.6 | Examples of dissolution textures in altered volcanic rocks, cont.

52 | CHAPTER 3

Vuggy silica (quartz) alteration facies is characterised by fine-grained, microcrystalline quartz and abundant open vugs or pores, which may be partly infilled (e.g. Fig. 3.6D, E and F). It is common in high-sulfidation epithermal systems and results from the extensive leaching of all phases, except SiO2 and TiO 2 , from volcanic rocks by hot acid solutions (White and Hedenquist, 1990).

Stylolites Stylolites are common in altered volcaniclastic rocks (Allen, 1990; Allen and Cas, 1990; Marsaglia and Tazaki, 1992; Gifkins and Allen, 2001). They are surfaces of dissolution associated with strain (pressure solution). They are roughly planar surfaces that exhibit mutual column and socket interdigitation and may branch. Stylolites result from mechanical compaction and removal of elements by diffusion and precipitation (Merino et al., 1983). They indicate volume loss and may form parallel or sub-parallel to bedding during burial, or at high angles to bedding during folding. Stylolites often contain a residue of insoluble material and minerals precipitated from solution. RecrystaUisation, dissolution, grain growth, grain orientation, pressure twinning, fracturing and residual accumulation of minerals along stylolites are common (Amstutz and Park, 1967). Irregular, anastomosing, bedding-parallel stylolites have been recognised in originally glassy volcanic facies, especially pumice breccias, in the Mount Read Volcanics (Allen, 1990; Allen and Cas, 1990; Gifkins and Allen, 2001). These are seams that concentrate fine-grained opaques and sericite at the margins of originally glassy clasts, along tube vesicle walls in pumice clasts and in the matrix (Fig. 3.6D: Gifkins, 2001). These stylolites are interpreted as diagenetic compaction and dissolution fabrics that formed by the dissolution of soluble components, particularly glass, and by the precipitation of clays and Fe-oxides as a result of pressure during burial (Allen, 1990; Allen and Cas, 1990; Gifkins, 2001).

Solution seams Solution seams are non-sutured, discontinuous mineralfilled seams that may form during diagenesis as a result of stress-related dissolution of soluble components and reprecipitation (Merino et al., 1983). Analcime-filled solution seams in pumice-rich rocks from the Green Tuff Belt (Japan) are anastomosing and roughly parallel to bedding. They occur in the fine-grained matrix and within fiamme and partly compacted pumice clasts (Fig. 3.6E: Gifkins et al., in press).

Static recrystaUisation textures RecrystaUisation is the transformation of a mineral or glass to a new grainsize, morphology or orientation of the same mineral species or minerals of the same composition (i.e. neomorphism, Folk, 1965). Pre-existing minerals recrystallise to a new grainsize in an attempt to assume a more stable form by minimising the ratio of the surface area to the volume during changed physical conditions (Yardley, 1989).

RecrystaUisation textures are produced by changes in the size, shape and arrangement of minerals in a rock. With increasing temperature, recrystaUisation generally involves the change from fine to coarse grainsize (aggrading), except for static recrystaUisation to hornfels and dynamic recrystaUisation where large, strained grains are replaced by a mosaic of tiny, unstrained crystals (Folk, 1965). Minerals may be directed or randomly orientated (non-directed: Spry, 1976). Directed textures occur where recrystaUisation is accompanied by stress (dynamic recrystaUisation). Non-directed textures occur where the pressure is equal in all directions (static recrystaUisation). Common macroscopic and microscopic recrystaUisation textures include mineral overgrowths, porphyroblasts, poikoblasts, and hornfelsic, granoblastic, granophyric and decussate textures (Fig. 3.7).

Dynamic recrystaUisation textures Directed fabrics or textures are common to regional metamorphic rocks where recrystaUisation is accompanied by stress (dynamic recrystaUisation). These textures include subgrains (granoblastic, porphyroblastic and poikiloblastic textures), foliations, layering and lineations. In common usage the term foliation is non-genetic and describes any planar, spaced or pervasive fabric in the rock. Foliations may form during diagenesis, metamorphism or tectonic deformation. Planar foliations are due to the preferred orientation of minerals, particularly micas, aligned perpendicular to the maximum compression direction (Yardley, 1989). Planar foliations are subdivided on the basis of grainsize and overall appearance of the altered rock. These include slaty cleavage, schistosity and gnessic layers. In fault zones or zones of intense ductile shear, two characteristic textures occur: cataclastic and mylonitic textures. Cataclase refers to a fine-grained fault gouge breccia with an unfoliated matrix (Sibson, 1977). Ideally, cataclasis is mechanical fragmentation without recrystaUisation, however this rarely occurs in nature (cf. Sibson, 1977). Mylonite is a term used for strongly foliated fine-grained rocks in which the grainsize has been reduced by recrystaUisation (Bell and Etheridge, 1973).

Deformation textures Deformation textures are stress activated and develop in response to overburden pressures or regional tectonic stress. Because volcanic deposits typically have high initial porosities they are easily modified by mechanical compaction during burial, tectonic deformation or, in the case of pyroclastic deposits, welding (Peterson, 1979; Allen, 1988; Branney and Sparks, 1990). Textural modification associated with compaction is essentially a result of increased pressure causing the re-arrangement and deformation of grains and reduction of intergranular pore space (Deelman, 1975). Deformation textures result from the rotation, brittle fracturing, flattening and distortion of existing grains or fabrics, especially clasts that were previously altered to soft minerals (McBride, 1978; Galloway, 1979; Craig and Vaughan, 1981; Branney and Sparks, 1990).

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 53

A. Porphyroblasts The large, strongly altered, cordierite porphyroblasts give this rhyolite a distinctive coarsely spotty texture. This texture inspired the terms dalmatianite, which was applied by early workers in the Noranda Camp, and the spotted fades, which was applied by Riverin and Hodgson (1980), for the cordierite-altered Amulet rhyolite and Millenbach andesite. The porphyroblasts are enclosed in a groundmass of chloritised biotite + sericite + quartz. Porphyroblasts are metamorphic crystals that are surrounded by a much finer grained matrix of other minerals (Spry, 1976). These large minerals have formed at the expense of the matrix and are the metamorphic equivalent of phenocrysts. Archaean Amulet Rhyolite Formation, Noranda, Abitibi greenstone belt, Amulet Upper A deposit, Canada. B. Porphyroblasts in gneiss This biotite + garnet gneiss is characterised by spotty 5 mm diameter garnet porphyroblasts in a medium- to fine-grained quartz + feldspar + biotite groundmass. The garnet porphyroblasts commonly have a biotite rim. A gneiss is a rock with coarsely differentiated layering denned by the segregation of minerals of different composition (typically dark and light minerals) in medium- to coarse-grained, granular rocks. Layering forms parallel to the tectonic foliation and in this case deviates around the garnet porphyroblasts. The precursor is interpreted to have been a felsic volcaniclastic rock. Sample 154061, Proterozoic Potosi gneiss, Harp prospect, Broken Hill Block, New South Wales. C. Porphyroblasts in thin section This sample of garnet hornfels, from the contact zone between rhyolite and a diorite intrusion, contains euhedral garnet porphyroblasts in the biotite + muscovite + quartz groundmass. Porphyroblasts, like these, with well-developed crystal shapes are idioblastic or euhedral, whereas those with poorly developed crystal shape are xenoblastic or anhedral (Yardley, 1989). Plane polarised light. Sample 140868, Cambro-Ordovician Mount Windsor Formation, east Thalanga, Mount Charter, Queensland.

D. Poikiloblast in thin section This thin section of amphibolite displays an amphibole poikiloblast with quartz and biotite inclusion trails. Poikiloblasts are porphyroblasts that contain numerous inclusions that may or may not show a preferred orientation (Barker, 1990). Poikiloblastic texture is analogous to poikilitic or micropoikilitic texture. Usually the inclusions are minerals that also occur in the matrix (Yardley, 1989). In this sample, the biotite inclusion trails display snowball rotation indicating syntectonic growth of the amphibole. Plane polarised light. Sample 3215, Proterozoic Corella Formation, Mary Kathleen Group, Malbon River, northwest Queensland.

FIGURE 3.7 | Examples of static recrystallisation textures in altered volcanic rocks.

54 | CHAPTER 3

Common deformation textures include intergranular textures and fabrics such as foliations, lineations, and augenstructure, and intragranular textures such as strained, bent, kinked, flattened, twinned and broken grains (crystals or clasts), as well as irregular grain contacts (e.g. Fig. 3.8A, B and C: Deelman, 1976; Spry, 1976). Deformation can modify pre-existing textures such as volcanic, hydration and devitrification textures (e.g. Fig. 3.8D to H). Fiamme and eutaxitic deformation textures are unique to volcanic facies (e.g. Fig. 3.9: Ross and Smith, 1960; Allen and Cas, 1990; McPhie et al., 1993; Gifkins et al., in press). Fiamme and eutaxitic textures are characteristic of, but are not restricted to, welded ignimbrites (e.g. Fig. 3.9A and B: Ross and Smith, 1960; Smith, I960), welded pyroclastic fall deposits (e.g. Sparks and Wright, 1979), welded autobreccia (e.g. Sparks et al., 1993) and pyroclastic deposits that have undergone secondary welding as a result of contact with hot lava or intrusions (e.g. Ross and Smith, I960; Christiansen and Lipman, 1966; Schmincke, 1967; McPhie and Hunns, 1995). Similar fiamme and eutaxitic texture also occur in non-welded altered pumice-rich rocks (e.g. Fig. 3.9C and D: Fiske, 1969; Allen, 1988; Branney and Sparks, 1990; Gifkins et al., in press) and felsic lavas (e.g. Pichler, 1981; Allen, 1988). The terms fiamme and eutaxitic texture are used herein to describe the rock texture and not to imply any particular origin. Fiamme are flame-like, glassy or devitrified lenses, which define a pre-tectonic foliation (cf. McPhie et al., 1993). Fiamme may have a wide variety of sizes (0.5 mm to 1 m), length to height ratios (up to 40:1), shapes (e.g. flame-like, bow tie, irregular branching and blocky) and internal textures (aphyric, porphyritic, vesicular or stylolitic) (Gifkins et al., in press). Eutaxitic texture is the pre-tectonic foliation defined by the parallel alignment of fiamme (cf. Fritsch and Reiss, 1868; Ross and Smith, I960; Smith, I960). Eutaxitic texture typically imparts a blotchy or streaky appearance to the rock due to the colour contrast between the darker fiamme and paler matrix (e.g. Fig. 3.9A and C).

3.2 I PSEUDOTEXTURES The incomplete destruction of primary textures and the combined effects of a number of different overprinting alteration styles (polyphase alteration) can result in significant textural modification and the development of false textures or pseudotextures (De Rosen-Spence et al., 1980; Allen, 1988; McPhie et al., 1993). Pseudotextures are alteration textures that modify or obscure primary volcanic textures and often lead to incorrect interpretation of the primary volcanic facies. Allen (1988) described examples of altered silicic lavas and autobreccias from Benambra, New South Wales, that have the remarkably deceptive appearance of welded and non-welded pyroclastic facies and thinly bedded tuffaceous rocks. Pseudotextures can be subdivided into pseudoclastic textures (pseudobreccia, pseudogranular, false thin-bedded volcaniclastic) or false pyroclastic textures (false shards, false pyroclastic or eutaxitic: Fig. 3.10). However, strong pervasive alteration can also produce false massive textures that resemble either massive volcaniclastic or coherent facies (Allen, 1988; McPhie et al., 1993; Doyle and Huston, 1999;

Doyle, 2001). Polyphase and patchy alteration of monomictic volcaniclastic facies can also result in false clast-supported and false polymictic textures.

Pseudoclastic textures The most common pseudoclastic textures are pseudobreccia and false pyroclastic texture (also referred to as false eutaxitic texture). Other pseudoclastic textures include false thinbedded volcaniclastic and pseudogranular textures. Pseudobreccias have the appearance of breccias, but form as a result of alteration of coherent facies (Carozzi, 1960; Allen, 1988). In outcrop they resemble coarse-grained, monomictic or polymictic, clast- to matrix-supported breccias comprising angular to sub-rounded clasts in a fine-grained matrix (Fig. 3.10A, BandC). False pyroclastic textures occur in both coherent facies and in situ hyaloclastite. In outcrop and hand specimen they may have a eutaxitic texture and contain abundant fiamme (e.g. Fig. 3.10D). In thin section they appear to contain splintery and arcuate fragments, which may closely resemble pyroclastic glass shards (false shards: Fig. 3.10E). Both pseudobreccia and false pyroclastic texture develop as a result of two-phase alteration of fractured (perlitic or quench fractured) coherent or autoclastic facies and domaincontrolled alteration of nodular devitrification textures in coherent facies (e.g. Fig. 3.11: Allen, 1988). Networks of intersecting quench and/or perlitic fractures may control polyphase alteration in the fractured glassy parts of coherent lavas and intrusions because they are permeable pathways for fluid flow. Initially glass immediately adjacent to the perlitic fractures is altered, then, as the fractures are filled, replacement fronts migrate away from the perlitic fractures towards the core. This may either obscure the continuity of the perlitic fractures or, if alteration is incomplete, enhance the perlitic fractures. False shard textures develop either due to the preservation of less altered, relatively siliceous slivers between two or more fractures, or as altered segments of the fractures themselves (e.g. Fig. 3.10E: Allen, 1988). The shape of false shards is a function of the shape of the fracture network. For example, cuspate false shards are produced from classical perlite, whereas those resembling flattened or welded shards result from banded perlite (e.g. Fig. 3.2D). False clasts develop where altered perlitic glass is partly overprinted by a subsequent alteration phase, thereby preserving isolated relicts of the earlier phase. Alternatively, the earlier phase may be incomplete, leaving isolated kernels of glass that are subsequently altered to a different mineral assemblage. Pseudobreccia may also result from domainal or selective alteration of nodular devitrification textures: spherulites and lithophysae (e.g. Fig. 3.10F and G). Spherulites and lithophysae are typically recrystallised to quartzo-feldspathic compositions, whereas the interstitial originally glassy domains are altered to phyllosilicate-rich assemblages (Allen, 1988). Consequently, the originally glassy and crystalline domains differ in alteration mineralogy and colour, and the spherulites appear as rounded siliceous clasts in a fine-grained phyllosilicate-rich matrix. Pseudogranular or sandy textures resemble well-sorted sandstones. These result from the recrystallisation of densely

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 55

A. Augen schist This sample of quartz-augen schist comprises large lenticular quartz-rich domains (Q) enclosed in a strongly foliated, sericite + quartz ± chlorite-altered matrix. Augen texture is common in deformed, strongly porphyritic coherent and crystal-bearing volcaniclastic rocks. This augen schist probably resulted from the superposition of a strong regional cleavage on an altered pumice breccia. The cleavage anastomoses around competent silicified pumice clasts. Sample 040617, Cambrian western volcano-sedimentary sequences, Mount Read Volcanics, Rosebery hanging wall, western Tasmania.

B. Broken crystals in andesite This deformed andesite contains broken plagioclase phenocrysts in a strongly foliated, sericite + chlorite + magnetite ± epidote-altered groundmass. Broken or fractured grains may result from mechanical pressure during tectonic deformation (McBride, 1978). Typically, feldspar crystals have been fractured along their cleavage planes, whereas quartz crystals have developed conchoidal fractures (Taylor, 1950; Sippel, 1968). Cross polarised light. Sample 144387, Ordovician Lake Cowal Volcanic Complex, Junee-Narromine Volcanic Belt, Lake Cowal, Gateway Prospect, New South Wales.

C. Deformed grains In this amphibolite grade volcaniclastic siltstone, the quartz grains are deformed polycrystalline grains with undulose extinction, and elongated parallel to the regional cleavage. Cross polarised light. Sample GA9, Early Proterozoic Supra crustal succession, Bergslagen mining district, Garpenberg, Sweden.

D. Deformed clasts and pillows Pillow fragments and clasts in this basaltic hyaloclastite were deformed and stretched parallel to the regional foliation. The clast shapes are irregular and difficult to recognise as pillow or hyaloclastite fragments. Despite this, many clasts preserve an internal zonation. Amphibolite, Proterozoic Corella Formation, Mary Kathleen Group, Malbon River, northwest Queensland.

FIGURE 3.8 | Examples of deformation textures, deformed clasts and pre-existing textures in altered volcanic rocks.

56 | CHAPTER 3

E. Deformed clasts Lens-shaped siliceous clasts (Q in this volcaniclastic breccia have been rotated and stretched into the strong tectonic cleavage. The fine-grained matrix has been foliated and chlorite + sericite + quartz altered. Sample 133520, Cambro-Ordovician Trooper Creek Formation, Seventy Mile Range Group, Mount Windsor Subprovince, central Thalanga, Queensland.

F. Folded pumice clast This sample of rhyolite-, pumice- and crystal-rich breccia contains a folded tube pumice clast with an axial planar cleavage (Sj) defined by aligned sericite. The pumice clast has been albite + quartz + sericite altered. Plane polarised light. Sample KB304B, Siluro-Devonian Ural Volcanics, Ural Ridges area, New South Wales.

G. Deformed relict perlite Relict perlitic fractures in this jigsaw-fit andesitic breccia are elongate and flattened, especially adjacent to competent phenocrysts. The groundmass has been sericite + chlorite + calcite + albite altered and the perlitic fractures chlorite filled. Plane polarised light. Sample 76902, Cambrian Que-Hellyer Volcanics, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer, western Tasmania.

H. Deformed grains In this sample of volcaniclastic sandstone, strongly' deformed feldspar grains and clasts have been rotated parallel to the strong cleavage. Plane polarised light. Sample 133520, Cambro-Ordovician Trooper Creek Formation, Seventy Mile Range Group, Mount Windsor Subprovince, central Thalanga, Queensland.

FIGURE 3.8 | Examples of deformation textures, deformed clasts and pre-existing textures in altered volcanic rocks, cont.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 57

A. Fiamme and eutaxitic texture in welded ignimbrite Dark flame-like obsidian lenses or fiamme (F) are aligned in this sample of subaerial welded rhyolitic ignimbrite. Fiamme are commonly interpreted as flattened pumice clasts. The fiamme in this sample are interpreted to result from the plastic deformation, flattening and sintering together of hot glassy pumice clasts during welding (cf. Smith, I960). The bedding-parallel alignment of flattened, elongate fiamme and glass shards defines the eutaxitic texture. Sample OW7, Pleistocene Owahoroa ignimbrite, Whitianga Group, Coromandel Volcanic Zone, Owharoa Falls, New Zealand. B. Fiamme in thin section In thin section, the former pumice clasts, fiamme (F), lack uncompacted vesicles, have feathery terminations and are enclosed in domains of cuspate and platy shards (5), and quartz, feldspar and biotite crystal fragments. Although some shards have preserved bubble-wall shapes, others were plastically deformed and compacted, especially adjacent to crystals. Plane polarised light. Sample OW11, Pleistocene Owahoroa ignimbrite, Whitianga Group, Coromandel Volcanic Zone, Owharoa Falls, New Zealand.

C. Fiamme and eutaxitic texture in non-welded pumice breccia Dark, plagioclase-phyric, wispy, chlorite-rich fiamme are enclosed in pale domains of quartz + chlorite + pyritealtered pumice clasts, shards and crystal fragments in this non-welded rhyolitic pumice breccia. The beddingparallel alignment of fiamme defines the eutaxitic texture. Alteration and compaction of pumice clasts during diagenesis formed these apparent welding textures. Sample 133809, Cambrian Hercules Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Hercules footwall, western Fasmania.

D. Fiamme in thin section In thin section, the chlorite fiamme (F) have feathery terminations and lack internal textures other than sparse plagioclase phenocrysts and hematite-rich stylolites. The pale quartz + chlorite + pyrite-altered domains contain uncompacted tube pumice clasts (P). Plagioclase crystals are dusted with hematite and sericite. Plane polarised light. Sample 133811, Cambrian Hercules Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Hercules footwall, western Fasmania.

FIGURE 3.9 | Examples of fiamme and eutaxitic texture.

58 I CHAPTER 3

A. Pseudobreccia in perlitic rhyolite Sericite + chlorite-altered perlitic fractures and pink albite-altered perlitic cores result in the pseudoclastic texture in this plagioclase-phyric coherent rhyolite. Sample BBP248-504.7 m from the Cambrian Central Volcanic Complex, Mount Read Voleanics, Boco, western Tasmania.

B. Pseudobreccia in macroperlitic dacite Polyphase alteration of macroperlite in this plagioclasephyric dacite has resulted in dark chlorite + epidote-rich domains along and adjacent to the perlitic fractures. This overprinted and enclosed earlier, pale grey albite + sericite-altered perlitic cores. The chlorite + epidoterich domains resemble the matrix in a matrix-supported breccia. Locally, the arcuate perlitic fractures are well defined. Cambrian Sterling Valley Volcanics, Mount Read Volcanics, Sterling Saddle, western Tasmania. C. False clasts

In this coherent andesite, plagioclase phenocrysts have been extensively replaced by epidote, pyroxenes by chlorite and the groundmass domains by green epidote + chlorite and orange albite. The domainal distribution of the alteration facies gives the andesite a patchy pseudoclastic texture. The false clasts have both sharp and diffuse margins, which are transgressed locally by altered plagioclase phenocrysts. Sample 145147, Ordovician Forest Reefs Volcanics, Molong Volcanic Belt, Cooramilla, New South Wales. D. False pyroclastic texture

The most conspicuous feature of this sample is the wispy dark green chlorite-rich lenses that resemble fiamme. However, these lenses are aligned in the tectonic cleavage and occur in an evenly porphyritic rhyolite. The lenses are interpreted to result from domainal chlorite and sericite + quartz + biotite alteration of the groundmass in a coherent quartz + plagioclase-phyric rhyolite. Sample 140727, Cambro-Ordovician Mount Windsor Formation, MountWindsorSubprovince, centralThalanga, Queensland. E. False shards in thin section Polyphase chlorite + biotite and K-feldspar alteration of perlitic fractures, and their subsequent deformation have resulted in irregular arcuate and platy shapes, which resemble shards (arrows) in this quartz + plagioclase + pyroxene-phyric coherent rhyolite. Plane polarised light. Sample KB499, Siluro-Devonian Coan rhyolite, Mount Hope Volcanics, Mount Hope area, New South Wales.

FfGURE 3.10 | Examples of pseudotextures in altered volcanic rocks.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 59

F. Pseudoclastic texture in devitrified rhyolite This rhyolite contains silicified nodules that are composed of coalesced spherulites in a fine-grained sericite-altered groundmass. These nodular devitrification textures give a clastic appearance to the hand specimen and outcrop. Late Devonian Bunga Beds, Boyd Volcanic Complex, Bengunnu Point, New South Wales.

G. False clasts in thin section The clastic texture in this rhyolite comes from the uneven distribution of strongly chlorite + hematitealtered spherulites in a strongly foliated, chlorite + sericite + feldspar + hematite-altered groundmass. The foliation has wrapped around the altered spherulites, which have preserved fibrous textures and quartz cores. Plane polarised light. Sample KB536D, Siluro-Devonian Mount Hope Volcanics, Boolahbone tank, Mount Hope area, New South Wales.

H. Pseudogranular texture This altered dacite has a fine granular texture in hand specimen and resembles massive sandstone. However, in thin section it has a densely microspherulitic groundmass in which the recrystallised quartz + albite spherulites are separated by cuspate sericite-rich domains. The fine-grained, densely packed spherulites give the hand specimen its sandy texture. Sample MR25, Cambrian Kershaw Pumice Formation, Central Volcanics Complex, Mount Read Volcanics, Mount Read summit, western Tasmania.

I. Pseudogranular texture in thin section Recrystallised micropoikilitic textures in the groundmass of this aphyric rhyolite resemble sand-sized, rounded grains in a well-sorted quartzo-feldspathic sandstone. The margins of the micropoikilitic domains are marked by concentrations of sericite, which enhance the granular texture. Plane polarised light. Sample 147448, Cambrian Kershaw Pumice Formation, Central Volcanics Complex, Mount Read Volcanics, western Tasmania.

FIGURE 3.10 | Examples of pseudotextures in altered volcanic rocks, cont.

60 | CHAPTER 3

J. False thin-bedded volcaniclastic texture Flow banding in this microspherulitic plagioclase-phyric dacite is defined by alternating pale albite + quartz and darker albite + sericite + quartz layers. The planar, repetitive, thin flow banding resembles thin bedding in clastic facies such as tuffaceous siltstones. Sample 76772, Cambrian Que-Hellyer Volcanics, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer, western Tasmania.

K. False thin-bedded volcaniclastic texture in thin section In thin section, this finely flow-banded rhyolite resembles a thin-bedded volcaniclastic facies with fractured plagioclase crystals. However, axiolitic and bow-tie shaped spherulitic textures are locally preserved in the groundmass. Cross polarised light. Sample KB132A, Siluro-Devonian Ural Volcanics, Ural • area, New South Wales.

L. False volcaniclastic texture in thin section Fractured and broken plagioclase crystals, and recrystallised spherulites in the groundmass of this flowbanded rhyolite contribute to its pseudoclastic texture. Cross polarised light. Sample 133837 Cambrian Mount Black Formation, Central Volcanic Complex, Mount Read Volcanics, Mount Read summit, western Tasmania.

M. False polymictic, matrix-supported texture The dark chlorite-rimmed clasts in this plagioclasephyric basaltic breccia appear to be supported in a compositionally different, pale calcite + chlorite-altered matrix. However, the matrix comprises jigsaw-fit, blocky and splintery clasts of perlitic basalt that are identical to the darker clasts. The chlorite-rimmed clasts appear subrounded, because chlorite alteration of the clast margins and adjacent matrix has obscured the blocky and splintery shapes. Sample 76833, Cambrian Que-Hellyer Volcanics, western volcano-sedimentary sequences, Mount Read Volcanics, Hellyer, western Tasmania.

FIGURE 3.10 | Examples of pseudotextures in altered volcanic rocks, cont.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS I 61

N. False polymictic texture Overprinting domainal albite + hematite and epidote alteration facies in this plagiocla.se + hornblende-phyric dacite gives the sample a heterogenous appearance. The abundance of pink, albitised plagioclase phenocrysts in the red and green domains is equivalent, although they appear more abundant in the epidote-altered domains due to the colour contrast between the phenocrysts and epidotealtered groundmass. Colour differences between the two alteration facies and more prominent phenocrysts in the epidote-altered domains obscure the massive, coherent texture and uniform composition of this sample. Sample 147557, Cambrian Mount Black Formation, Central Volcanic Complex, Mount Read Volcanics, Pieman Road, western Tasmania. O. False matrix-supported texture Clasts in this monomictic dacite breccia display jigsaw-fit texture. However, feldspar + quartz + sericite alteration facies has replaced the groundmass adjacent to the quench fractures between clasts, enhancing the matrix, and imparting an apparent matrix-supported texture. The feldspar + quartz + sericite-altered matrix has been more resistant to weathering than the sericite + chlorite-altered clasts and forms ridges on the outcrop. The larger clasts are perlitic, plagioclase + hornblende porphyritic with planar and curviplanar margins typical of clasts produced by quench fragmentation. Cambrian, Mount Black Formation, CentralVolcanic Complex, Mount Read Volcanics, Tullah lakeside, western Tasmania.

P. False matrix-supported texture In this in situ andesitic hyaloclastite, blocky chloritealtered plagioclase + pyroxene-phyric clasts appear to be supported in a pyrite + quartz + sericite-rich matrix. However, thin section inspection reveals relict clasts with jigsaw-fit textures preserved in the false matrix domains. Sample 144710, Ordovician Lake Cowal Volcanics, JuneeNarromine Volcanic Belt, Boundary Prospect, New South Wales.

Q. False coherent texture This albite + quartz + sericite- and chlorite + epidotealtered pumice and rhyolite breccia resembles a coherent, feldspar porphyritic facies. The albite-altered plagioclase crystals, pseudophenocrysts, are evenly distributed in a fine-grained, sericite-rich false groundmass. On the left side of the photograph, pervasive, massive albite + quartz alteration facies obscures the plagioclase crystals in a pseudo-aphyric texture. Sample 147402, Cambrian Kershaw Pumice Formation, CentralVolcanic Complex, Mount Read Volcanics, Rosebery, western Tasmania.

FIGURE 3.10 | Examples of pseudotextures in altered volcanic rocks, cont.

62 | CHAPTER 3

FIGURE 3.11 | Sketches summarising the relationship between primary volcanic, high-temperature devitrification and hydration textures, and pseudotextures. (A) Pseudotextures in classical perlite (after Allen, 1988). False shards may be produced either (i) as siliceous segments between phyllosilicate-altered perlitic fractures or (ii) phyllosilicate-altered sections of the perlitic fractures, (iv) Alternatively, pervasive phyllosilicate alteration can obscure the perlitic fractures resulting in a massive texture. (B) Pseudotextures in flow-banded facies comprising alternating crystalline (spherulitic) and glassy flow bands (modified from Doyle, 2001). (i) Glassy domains are altered to phyllosilicate-rich assemblages and spherulitic bands to quartzo-feidspathic assemblages. Consequently the originally glassy and crystalline domains differ in alteration colour and mineralogy. The spherulites superficially appear as rounded siliceous dasts in a fine-grained phyllosilicate-rich matrix, resulting in a pseudobreccia texture, (v) In parallel flow-banded facies this can result in a false thin-bedded voicaniciastic texture. The crystalline nature of the quartzo-feldspathicaltered bands enhances the granular appearance of the false beds. (C) Pseudotextures from in situ quench fragmented porphyritic glass (modified from Doyle, 2001). Alteration progresses as fronts away from the fractures and into the non-fractured glass, (i) Where phyllosilicate alteration is incomplete the remaining kernels of glass are subsequently altered to quartzo-feidspathic assemblages. The further the alteration progresses away from the fractures, the more matrix-supported the resulting pseudobreccia texture, (ii) Pseudoclastic texture also develops as a result of the complete replacement by the first alteration phase (phyllosilicate alteration) and only partial replacement by the second alteration phase (quartzofeidspathic alteration), thereby preserving isolated relics of the earlier phase, (iii) False polymictic texture develops in quench fragmented porphyritic glass as polyphase alteration results in colour variations between the false dasts and matrix. Varying intensities of alteration enhance the polymictic appearance. Phenocrysts are more prominent in dark phyllosilicate domains than in paler siliceous domains, resulting in an apparent variation in crystal content consistent with a polymictic rock, (iv) Pseudomassive texture develops from pervasive phyllosilicate alteration. (D) Pseudotextures from porphyritic autobreccia or hyaloclastite (modified from Doyle, 2001). (i) and (ii) alteration commences along fractures and in the matrix of autobreccia but gradually progresses into the dasts, greatly enhancing the clastic, matrix-supported appearance. The result is commonly false matrix-supported texture, (iii) Polyphase alteration results in pseudoclasts and pseudomatrix that comprise different alteration mineral assemblages and colours, and therefore appear to have different primary compositions (false polymictic texture). Phenocrysts are more prominent in dark phyllosilicate-altered domains, (iv) Pseudomassive texture occurs where alteration has been extensive and pervasive. (E) Pseudotextures from porphyritic pumice breccia (modified after Gifkins, 2001). False eutaxitic texture results from phyllosilicate alteration of originally glassy pumice dasts or shards within the pumice breccia. The phyllosilicate-altered pumice dasts and shards are flattened by compaction (i) or tectonic deformation (ii) resulting in a texture that resembles eutaxitic texture in welded ignimbrites. (iii) False polymictic texture as a result of colour variations and apparent variations in phenocryst content due to polyphase alteration, (iv) False coherent textures result from the complete and pervasive phyllosilicate alteration and subsequent compaction of originally glassy shards and pumice dasts producing a massive textured rock in which the original dasts are indistinguishable.

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 63

spherulitic or micropoikilitic groundmass of coherent felsic fades (e.g. Fig. 3.10HandI). False thin-bedded volcaniclastic texture resembles thinly bedded and tectonically folded volcaniclastic rocks, and can occur in altered, planar flow-banded and flow-folded lavas and intrusions (e.g. Fig. 3.10J and K). The false thin-bedded volcaniclastic texture is due to the planar and uniform character of flow banding and the apparent textural and compositional differences between flow bands (Allen, 1988). This apparent compositional difference between flow bands results from domainal alteration: originally glassy bands were altered to dark phyllosilicate assemblages and contain well-preserved phenocrysts, whereas the originally cryptocrystalline bands were altered to pale quartzo-feldspathic assemblages and contain altered polycrystalline phenocrysts that are hardly recognisable. Phenocrysts may also be extensively fractured and dismembered during deformation, thereby contributing to the pseudoclastic and finer grained appearance (e.g. Fig. 3.10L). Discrepancies in colour and the apparent relative proportions of phenocrysts enhance the compositional contrast between adjacent bands.

False polymictic texture Domainal phyllosilicate alteration may impart a heterogenous appearance due to colour variation and to variable preservation of phenocrysts. Pseudobreccias and pseudoclastic facies may appear polymictic due to patchy or mottled alteration distribution resulting in colour variations and variable preservation of phenocrysts. This is particularly common in pseudobreccias because the pseudoclasts and pseudomatrix contain different alteration minerals, and hence colours: therefore they appear to have different primary compositions (e.g. Fig. 3.10M and N). In addition, phenocrysts (especially feldspar or quartz) are more prominent in dark phyllosilicate domains than in paler siliceous domains, resulting in an apparent variation in crystal content between different alteration domains consistent with different compositions (e.g. Fig. 3.10N).

alteration of pumice breccia produces a massive textured rock in which the original clasts are indistinguishable (Fig. 3.10Q).

3.3 I ALTERATION DISTRIBUTION Alteration distribution refers to the mappable extent of the alteration facies and its relationship to host rocks, structures, mineralised rock and other alteration mineral assemblages or zones. Although alteration distribution, zonation and textural relationships are easily observed in thin section and hand specimen they are more difficult to assess on a macroscopic scale. This is due to the typically sparse exposure in ancient volcanic successions, structural complications, commonly patchy mode of occurrence of alteration facies in volcanic rocks, and the considerable amount of detailed work required to determine the distribution and zonation. Closely spaced drill holes may adequately delineate the alteration facies distribution and zonation at prospect scales. However, alteration zones may be superimposed and the original zonation patterns modified. It is important to recognise disequilibrium alteration mineral assemblages and overprinting textures when determining zonation patterns, so as to account for any superposition caused by subsequent alteration styles.

Is the alteration facies limited in extent or widespread? Typically the macroscopic alteration distribution is described as either regional or local in extent. Regional alteration styles are widespread, affecting extensive (hundreds of metres to tens of kilometres) expanses of rock. Local alteration styles are limited in extent and can refer to alteration on a scale of centimetres, such as wall-rock alteration associated with a fault or vein, to hundreds of metres, such as the extensive hydrothermal alteration associated with the Rosebery VHMS deposit (western Tasmania) that extends up to 100 m stratigraphically beneath the ore lenses, 10-20 m into the hanging wall and 500 m along strike (Large et al., 2001b).

False matrix-supported texture Alteration along fractures in coherent facies and in the matrix of autoclastic facies (hyaloclastite and autobreccia) results in colour variation between the clasts and fracture selvages or matrix, and greatly enhances the clastic appearance. Where alteration has progressed into the clasts from the clast margins, the clasts appear to be isolated within a matrix (e.g. Fig. 3.10O and P). The result is an apparently matrix-supported breccia that is difficult to recognise as hyaloclastite or autobreccia (Allen, 1988).

False coherent textures False coherent textures, such as pseudomassive texture, are less common than pseudoclastic textures. They result from the complete and pervasive alteration of an originally volcaniclastic facies. For example, pervasive phyllosilicate

Is there a relationship between the distribution of the alteration facies and stratigraphy? Alteration facies can be distributed within individual volcanic units or formations or cut across stratigraphic contacts. Alteration facies that are confined to individual units or formations and are either concordant or discordant within the unit are described as stratabound or semiconformable (Guilbert and Park, 1986; Large, 1992). In contrast, alteration pipes and alteration plumes refer to the distribution pattern or shape of alteration systems or facies that transgress stratigraphic contacts (pipes, e.g. Sangster, 1972; Gemmell and Large, 1992; Doyle and Huston, 1999, and plumes, e.g. Jack, 1989; Gemmell and Fulton, 2001).

64 | CHAPTER 3

Is the alteration fades associated with other alteration fades, structures, mineralised rock or particular units? Locally distributed alteration fades are commonly associated with structures, such as veins or fractures, ore bodies or particular volcanic units: forming halos, envelopes or selvages. They may also occur around isolated clasts within clastic facies. In zoned alteration systems or halos, they are commonly spatially and temporally associated with other alteration facies.

3.4 I ALTERATION ZONATION PATTERNS Alteration zonation is a regular pattern in the spatial distribution of mineral species, mineral assemblages, major or trace elements or textures, and reflects mineralogical and chemical changes that relate to fluid-rock ratios and temperature gradients (e.g. Meyer and Hemley, 1967). In general, the halo of altered rock is divided mesoscopically into altered zones. These are imprecisely defined as the areal extent of different alteration mineral assemblages or facies. Boundaries between hydrothermally altered zones associated with VHMS deposits are commonly sharp, and reflect changes in the abundance of particular alteration mineral species such as chlorite, sericite or quartz. However, in some cases alteration facies overlap, resulting in diffuse boundaries between altered zones (e.g. Titley, 1982; Doyle and Huston, 1999). In reality, altered zones are identified and named using index minerals - the dominant alteration minerals - and zone boundaries - isograds - are drawn at the first appearance of an index mineral characteristic of the next zone. Thus the clinoptilolite + mordenite zone is placed at the first appearance of clinoptilolite or mordenite (e.g. Iijima, 1978). In some cases an altered zone may contain more than one alteration facies. For example, at Thalanga the tremolite + chlorite ± carbonate zone includes intense pervasive stratabound chlorite + tremolite (no carbonate), intense pervasive stratabound chlorite + calcite, and intense stratabound tremolite + dolomite + calcite (no chlorite) facies. The expression of zoning may be limited by exposure and modified by structural and compositional homogeneity, faulting and/or intrusions. There are three scales of zonation that are generally applied to the relative distribution of metals (e.g. Kutina et al., 1965)

TABLE 3.3 I Scales of alteration zonation. Regional zonation

Alteration facies or minerals occur in zones throughout a region

District zonation

Altered zones are associated with a cluster of ore bodies, fractures or intrusions

Local zonation

Altered zones are associated with an individual ore body, fracture or intrusion

Sample-scale zonation

Altered zones are associated with smallscale features such as clasts or minerals in the rock

and that can also be applied to the distribution of alteration facies or minerals (Table 3.3): (1) regional zoning, (2) district zoning, and (3) local or ore body zoning. In addition, altered zones may occur within or surrounding individual clasts in clastic facies. Thus the dimensions of altered zones may vary from a few centimetres to several tens of kilometres (cf. Bohlke et al., 1980; Galley, 1993). These variations in dimension are a function of the size of the alteration system (regional versus local systems), and changes in physical and chemical conditions such as the porosity and composition of the host rock, fluid-rock ratio and composition of the fluid (Rose and Burt, 1979). Different alteration processes result in different zonation patterns. Regional alteration processes such as diagenesis and metamorphism, produce thick, flat-lying, vertically zoned systems. Local hydro thermal systems and contact metamorphism result in footwall and hanging wall altered zones or concentrically zoned altered halos or contact aureoles. Zonation related to faults or fractures is generally parallel to the structure, commonly cross cutting stratigraphic contacts.

Regional diagenetic zones Diagenesis develops in response to increasing temperature with depth during burial; as a result it forms a sequence of flat-lying altered zones (Fig. 3.12: Iijima, 1974, 1978; Fisher and Schmincke, 1984; Utada, 1991). These zones are characterised by mineral assemblages, which reflect the reaction of glass, primary minerals and diagenetic minerals with interstitial pore water and seawater at a particular temperature — anywhere between 0° and 250°C (Hay, 1978; Iijima, 1978; Utada, 1991). Sequences of diagenetic zones are typically between 500 m and 6 km thick (Hay, 1978) and individual altered zones vary from a few metres to several kilometres in thickness (e.g. Hay, 1978; Iijima, 1978; Vierecketal., 1982; Passagliaetal., 1995). The thickness of diagenetically altered zones is dependent on the geothermal gradient, rate of burial, and the porosity and permeability of the volcanic succession. The Miocene Hokuroku Basin, part of the Green Tuff Belt in northern Honshu, Japan, is an excellent example of diagenetic zonation (Fig. 3.12). The Hokuroku Basin contains a 3 to 6 km thick submarine volcanic succession dominated by rhyolitic to dacitic and minor basaltic coherent and clastic volcanic facies (Horikoshi, 1969; Iijima, 1978; Tanimura et al., 1983; Urabe, 1987; Utada, 1991). Diagenetic alteration in the Hokuroku Basin has produced a series of flat-lying diagenetic zones, which grade vertically from fresh glass at the top, to smectites, zeolites and albite at depth (Fig. 3.12: Iijima, 1978).

Regional metamorphic zones Regional metamorphic zones develop due to regional changes in temperature and pressure. Unless metamorphism is related solely to orogenic deformation, progressive burial results in dehydration and increasing metamorphic grade with depth. This produces a vertical sequence of flat-lying, regional metamorphic zones (Fig. 3.13B). It is generally

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 65

assumed that higher-grade metamorphic rocks formerly had mineral assemblages typical of lower-grade zones and were altered progressively as metamorphism proceeded. However, variations in geothermal gradient, variable rates of sedimentation and folding and faulting can modify the progression of metamorphism and the resulting metamorphic zones, in particular, folding and faulting may affect the position of mineralogically determined isograds by promoting reactions or causing local increases in temperature or heat flow (Coombs et al., 1959). Coombs (1954) first recognised vertical metamorphic zoning of low-temperature, high-pressure metamorphic mineral assemblages in a 10 km section of the uplifted Permian to Jurassic Wakatipu Metamorphic Belt, in the New Zealand geosyncline (Fig. 3.13A). The Wakatipu Metamorphic Belt comprises submarine emplaced rhyolitic to basaltic volcaniclastic and volcanogenic sedimentary rocks, which were progressively metamorphosed (Coombs, 1954). There is a continuous gradational sequence of regional metamorphic zones characterised by zeolite, prehnite + pumpellyite, pumpellyite + actinolite and greenschist facies (a biotite zone and chlorite zone) at depth (Fig. 3.15B: Coombs, 1954; Coombs etal., 1959; Landis and Coombs, 1967). These zones have an even thickness over a wide area (>10 km by 300 km) suggesting that the geothermal gradient, estimated at l4-25°C/km, was consistent throughout the Wakatipu Metamorphic Belt (Landis and Coombs, 1967). FIGURE 3.12 | Schematic cross-section of vertically developed diagenetic zones in the central Hokuroku Basin, Japan (after lijima, 1974; Hay, 1978; lijima, 1978).

FIGURE 3.13 | Alteration zonation patterns associated with regional metamorphism. (A) Regional metamorphic zones in the Wakatipu metamorphic belt, South Island, New Zealand, (after Coombs et al.,1959, and Landis and Coombs,1967). (B) Schematic cross-section reconstruction of regional metamorphic zones and isotherms during peak metamorphism of the Wakatipu metamorphic belt, New Zealand (after Landis and Coombs, 1967).

66

|

CHAPTER 3

Regional, deep, semi-conformable altered zones Deep, semi-conformable altered zones are interpreted to be the products of hydrothermai alteration within regional, subseafloor hydrothermai systems (Galley, 1993). These regional hydrothermai systems involve the large-scale convection of modified seawater driven by the emplacement of synvolcanic intrusions into the subsurface (Spooner and Fyfe, 1973; Norton, 1984; Cathles, 1993; Galley, 1993). The upper contacts of sub-surface intrusions are typically sub-parallel to the volcanic strata and hence the overlying isotherms are also semi-conformable (Galley, 1993). The result is a series of vertically stacked, sub-horizontal altered zones, which are characterised by mineral assemblages that reflect reactions between host volcanic facies and modified seawater at temperatures transitional with diagenesis and greenschist facies metamorphism (Galley, 1993; Skirrow and Franklin, 1994). Sequences of deep, semi-conformable altered zones may be up to 20 km wide and between 1 and 2 km thick (Gibson et al., 1983; Galley, 1993; Skirrow and Franklin, 1994). One of the first comprehensive descriptions of deep, semiconformable altered zones is by Gibson et al. (1983). These authors documented a series of vertically stacked altered zones in the Noranda district of the Archaean Abitibi belt, Canada

(Fig. 3.14). The Amulet Rhyolite Formation comprises coherent and clastic basalts and andesites with minor rhyolitic lava domes. The volcanic rocks are variably altered and the distribution of the alteration facies can be related to stratigraphic depth and primary host rock composition. Within the andesites there are four regionally extensive alteration facies, from top to bottom: chlorite + actinolite + albite + epidote + quartz; silicified; mottled epidote + quartz; and chlorite alteration facies (Fig. 3.14B). The upper chlorite + actinolite + albite + epidote + quartz alteration facies is the result of low-grade regional metamorphism, whereas the other three alteration facies are interpreted to be the products of regional deep, semi-conformable alteration (Gibson et al., 1983).

Local contact metamorphic or hydrothermally altered halos Halos or contact aureoles associated with granitoids, thick synvolcanic sills, cryptodomes and dykes reflect changes in the composition and temperature of the magmatic or hydrothermai fluid away from the intrusion, and its interaction with local pore water (Rose and Burt, 1979; Einsele et al., 1980; Yardley, 1989). The result is a progressive sequence of roughly concentric altered zones or metamorphic zones

FIGURE 3.14 | Diagram showing alteration zonation patterns associated with regional hydrothermai alteration. (A) A schematic cross-section through a vertically stacked sequence of deep, semi-conformable altered zones in the Amulet Rhyolite Formation, Noranda district, Abitibi belt, Canada (after Gibson, 1989, in Galley, 1993). (B) A reconstructed section of deep, semi-conformable altered zones in the Amulet Rhyolite Formation at Turcotte Lake (after Gibson et al., 1983).

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS I 67

that correspond with decreasing temperature away from the intrusion contact (e.g. Fig. 3.15). For example, submarine basaltic lavas and breccias of the Upper Triassic Karmutsen Subgroup, Vancouver Island, Canada, have undergone lowpressure, high-temperature contact metamorphism related to the shallow emplacement of the Jurassic Coast Range Batholith (Carson, 1973; Kuniyoshi and Liou, 1976). Intrusion of the Coast Range Batholith resulted in two locally developed contact metamorphic zones: hornblende + plagioclase zone and epidote + actinolite zone, which are superimposed on a prehnite + pumpellyite facies regional metamorphic zone (Fig. 3.15A: Kuniyoshi and Liou, 1976). The hornblende + plagioclase zone is approximately 2600 m wide and the epidote + actinolite zone less than 900 m wide (Kuniyoshi and Liou, 1976).

Local hydrothermally altered halos around ore deposits Local halos of hydrothermally altered rock around mineral deposits result from the interaction between the hydrothermal fluid/s responsible for mineralisation and the surrounding country rock. In the case of VHMS deposits, hydrothermal fluid temperatures range up to 35O°C (Large, 1977, 1992; Urabe et al., 1983; Gemmell and Large, 1992). The migration of hydrothermal fluids and mixing with cold seawater at the margins of the hydrothermal system produces three-dimensional altered zones composed of successively lower temperature mineral assemblages away from the site of mineralisation (Date et al., 1983; Urabe et al., 1983; Large, 1992; Schardt et al., 2001). The shape, dimensions and distribution of fluid circulation (and thus the altered zones) are usually closely related to initial patterns of permeability, porosity and compositional contrasts in the volcanic succession (e.g. Yamagishi and Dimroth, 1985; Large, 1992; Doyle, 2001). Accordingly, the altered halos exhibit a wide variety of shapes and zonation around VHMS deposits, typically with intense proximal footwall and weak hanging wall zones (Large, 1992). In contrast, altered halos around porphyry deposits may have a more concentric zonation, comprising a potassic core zone, an outer propylitic zone and, in some examples, minor phyllic zones (Gustafson and Hunt, 1975). Later hydrothermal alteration and mineralisation zones may overprint these initial zones, such as the widespread phyllic and argillic zones at the El Salvador deposit, Chile and Batu Hijau in Indonesia (Gustafson and Hunt, 1975; McMillan and Panteleyev, 1998; Garwin, 2002). Hydrothermally altered zones associated with VHMS deposits can extend laterally up to 500 m and persist at least 500 m stratigraphically (Iijima, 1974; Utada et al., 1974; Eastoe et al., 1987; Gemmell and Large, 1992; Gemmell and Fulton, 2001). Footwall zones occur either as diffuse stratiform zones (e.g. Rosebery, Fig. 3.16C: Green et al., 1981; Large et al., 2001b) or well-defined and zoned alteration pipes (e.g. Hellyer, Fig. 3.16A: Millenbach, Fig. 3.16D: Franklin et al., 1981; Morton and Franklin, 1987; McArthur, 1989; Gemmell and Large, 1992). Hanging wall alteration is normally of lower intensity, and developed above the thickest part of the ore body either as a diffuse stratabound zone (e.g. Woodlawn and Scuddles) or an alteration plume (e.g. Hellyer,

FIGURE 3.15 | Diagram showing alteration zonation patterns associated with contact metamorphism. (A) Contact metamorphic zones in the northeastern Vancouver Island, Canada (from Kuniyoshi and Liou, 1976). (B) Schematic cross-section of contact-metamorphic zonation.

Fig. 3.16A: Large and Both, 1980; Jack, 1989; Large, 1992; Gemmell and Fulton, 2001). In the Hokuroku district, local hydrothermally altered halos associated with the Kuroko deposits comprise up to four altered zones in a roughly concentric distribution around the ore deposits (Fig. 3.16B: Urabe et al., 1983; Utada et al., 1988; Utada, 1991). From core to margin the altered zones are: sericite + quartz ± pyrite zone; chlorite + sericite + quartz zone; mixed-layer clay (illite-montmorillonite) zone; kaolinite or smectite-montmorillonite zone (Date et al., 1983; Urabe et al., 1983). These altered zones vary in thickness from approximately 16 m to 100 m, transgress stratigraphic boundaries, and grade into and interfinger with regional diagenetic zones such as the clinoptilolite + mordenite zone in Figure 3.16B (Utada, 1970, 1991; Iijima, 1974). In contrast, the mineral assemblages in altered zones associated with Australian VHMS deposits reflect highergrade regional metamorphic overprints. Some have proximal chlorite zones and halos of quartz + sericite, sericite and carbonate zones (Fig. 3.16A and C: Allen, 1988; Large, 1992; Doyle, 2001).

Vein and fracture altered halos Altered halos around fractures or veins are typically limited in width (from millimetres to tens of metres) occurring

68 | CHAPTER 3

FIGURE 3.16 | Schematic cross-sections showing hydrothermal alteration zonation patterns associated with VHMS deposits. (A) Hydrothermally altered zones developed in the footwall alteration pipe and hanging wall alteration plume to the Hellyer VHMS deposit (after Gemmell and Fulton, 2001). (B) Hydrothermally altered halo developed around the Uwamuki group deposits, Kosaka VHMS mine, Japan (after Urabe et al., 1983). (C) Hydrothermally altered halo developed around the K-lens VHMS ore lens at Rosebery (after Large et al., 2001b). (D) Hydrothermally altered zones developed in the footwall alteration pipe and hanging wall alteration plume to an ore lens in the Millenbach VHMS deposit, Canada (after Riverin and Hodgson, 1980).

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 69

either as narrow selvages around individual veins or as wide pervasive fracture-controlled altered zones. Decreasing fluid temperatures and changes in fluid chemistry (especially pH) away from fractures or faults results in alteration zonation parallel to the fracture or vein surface.

3.5 | OVERPRINTING RELATIONSHIPS AND TIMING OF ALTERATION The aim in determining overprinting relationships is to establish a time sequence for mineral growth, which is referred to as a paragenetic sequence (Kutina et al., 1965). Normally paragenesis refers to the growth of minerals from oldest to

youngest. The paragenetic sequence is typically depicted using a horizontal bar chart (e.g. Fig. 3.17) or a space-time diagram, such the schematic diagram by Wilson et al. (2003) that shows alteration and vein stages relative to the intrusive history of the Ridgeway Complex at the Ridgeway Au-Cu porphyry deposit in New South Wales. Unfortunately, interpreting timing relationships between different alteration facies can be difficult and confusing as there are few unambiguous overprinting textures. The superposition of many different alteration phases on the same rock can obscure or complicate alteration textures. In addition, while one mineral is being deposited under certain conditions at one place, other minerals are forming elsewhere under different conditions. Thus the deposition of one mineral may overlap with another in both time and space.

FIGURE 3.17 | Interpretation of the relative timing of alteration mineral assemblages in the northern Central Volcanic Complex, western Tasmania (modified after Gifkins and Allen, 2001). S1 is the bedding-parallel, stylolitic foliation interpreted as a diagenetic compaction and dissolution fabric (Allen, 1990). S2 is the regional tectonic cleavage related to Devonian folding. Assemblages in brackets refer to pre-metamorphic equivalents. For example, early regional sericite is probably a metamorphosed equivalent of early clays. The feldspar + quartz + sericite alteration mineral assemblage may include the growth of early zeolites and the replacement of zeolites and glass by K-feldspar and albite during diagenesis.

70 | CHAPTER 3

Method The relative timing of alteration mineral assemblages is determined by documenting overprinting relationships among the alteration mineral assemblages, mineral deposits, successive volcanic units or intrusions, the compaction foliation and regional tectonic cleavages. This requires detailed examination of textural and mineralogical features that are visible in the field, hand specimen (including drill core) and thin section. Although thin sections can be useful in displaying the inter-relationships between different minerals, they may be too small to show larger cross-cutting features (e.g. veins) and textural features in coarse-grained rocks. Paragenetic determination relies on a representative suite of samples that includes all rock types and alteration facies in the alteration system. Because overprinting relationships are not always straightforward the inspection of a large number of samples is required for a systematic understanding of the overprinting textures.

Overprinting textures Overprinting texture refers to any texture or geometry that can be used to infer that one mineral or group of minerals has formed later than another mineral. Table 3.4 describes the three main types of overprinting textures. There is a wide range of overprinting textures, some of which are described in Sections 3.1 and 3.2; however, the reader is referred to Ramdohr (1980), Craig and Vaughan (1981), Ineson (1989) and Taylor (1996) for more detailed discussions of overprinting textures and their significance. The possible relationships among alteration mineral assemblages, mineralised rocks, successive volcanic units or intrusions, the compaction foliation and regional tectonic cleavages and their implication for the timing of alteration are described in Table 3.5. Recent work (Allen, 1997; Gifkins, 2001; Gifkins and Allen, 2001) in pumice breccias around the Rosebery VHMS deposit has unravelled a complex sequence of hydrothermal and non-hydrothermal alteration mineral assemblages. This work was based on overprinting relationships between alteration mineral assemblages, compaction foliation (S:) and tectonic cleavage (S2: Fig. 3.17). Alteration mineral assemblages that infill primary porosity, preserve delicate uncompacted

TABLE 3.4 | Types of overprinting textures.

Mineral superposition

Where one mineral or a group of minerals can be seen to have nucleated upon a pre-existing mineral or grain; this applies to replacement, infill and recrystallisation textures

Fracture superposition

Where fractures fragment pre-existing minerals or grains, provide space for subsequent infill and focus wall-rock alteration

Foliation superposition

Where foliations or lineations rotate, modify, distort or fracture pre-existing minerals or components

pumice textures and/or are overprinted by Sj and S2 are interpreted to be early, prior to compaction and complete lithification. These include thin films of sericite, chlorite + sericite + hematite and calcite (or their pre-metamorphic equivalents, smectites, calcite and palagonite) that coated all original surfaces — including vesicle walls, plagioclase crystals, shards and fractures — and pre-date the infilling of vesicles by subsequent minerals. Feldspar + quartz + sericite, chlorite + sericite, chlorite + sericite + hematite, chlorite, sericite and quartz + sericite alteration facies have filled and preserved vesicles, and replaced glass, indicating that they (or their premetamorphic equivalents such as zeolites and clays) formed prior to compaction and deformation. Alteration features that define the bedding-parallel compaction foliation, such as chlorite + sericite + hematite and chlorite + sericite fiamme, are crenulated by the regional cleavage (S2) and are interpreted to be pre- to syn-S,. Alteration features that overprint the early alteration mineral assemblages and Sp but which are strongly foliated by S2, are broadly syntectonic or post-Sj pre-S2. A chlorite + epidote alteration assemblage has overprinted the early sericite films and chlorite + sericite fiamme but is strongly foliated by S2 and interpreted to be post-Sj and pre- to syn-S2. Chlorite + calcite + magnetite has replaced and post-dated the chlorite + sericite fiamme and chlorite + sericite + hematite stylolites that define Sv and which are commonly transposed into the S2 cleavage and associated with syn-S2 chlorite veins. Chlorite ± pyrite is associated with shear zones parallel to S2, and undeformed post-S2 brittle fractures and faults. Sericite + calcite + actinolite ± epidote recrystallised earlier alteration mineral assemblages and are aligned parallel to S2. Alteration mineral assemblages that are unaffected or only very weakly affected by the S2 foliation are interpreted to be post-deformation. These alteration mineral assemblages are related to chlorite + quartz + calcite vein infill and associated vein wall-rock replacement and altered halos associated with Devonian granites, such as quartz ± calcite ± tourmaline + magnetite veins. Early Mn-carbonate, chlorite, sericite and quartz + sericite alteration facies are spatially associated with ore at Rosebery (Fig. 3.16C), whereas the other alteration facies are regionally extensive, spatially associated with synvolcanic or Devonian intrusions, veins, faults or shear zones. The overprinting relationships, combined with spatial associations and distributions of the alteration facies, suggest that diagenetic alteration began shortly after eruption in the Cambrian and continued until the transition to regional metamorphism. Hydrothermal alteration associated with the VHMS deposits at Rosebery and Hercules commenced prior to burial compaction, but was synchronous with diagenesis and the intrusion of synvolcanic sills. Peak regional metamorphism was synchronous with Devonian deformation and is overprinted by contact metamorphic assemblages associated with the emplacement of Devonian granites. The interpretation of textures to determine a paragenetic sequence can result in misleading results unless the subsequent effects of metamorphism and deformation upon the textures and minerals are appreciated. For example, in the Que River VHMS deposit galena commonly occurs in the S2 cleavage. As a result, conventional interpretation may conclude that it formed syn-S2 (Devonian). This interpretation would

COMMON ALTERATION TEXTURES AND ZONATION PATTERNS | 71

suggest that galena post-dates the other sulfides (sphalerite and pyrite), which are overprinted by S2. However, geological evidence (Large et al., 1989) combined with lead isotope data (Gulson and Porritt, 1987) indicates that galena deposition occurred on the seafloor at the same time as the other sulfides in the Cambrian. In this case, the textural evidence constrains

the timing of recrystallisation of the galena to syn-S2, whereas the geological and isotopic evidence suggests syn-depositional formation of galena. Misinterpretations of this type are common for soft and easily recrystallised minerals such as sericite, chlorite, clay minerals, galena and chalcopyrite.

TABLE 3.51 Overprinting relationships in altered volcanic rocks and their implications for the relative timing of alteration (modified after Allen and Large, 1996)

Alteration facies to primary volcanic texture Alteration facies truncated by clast margins Clasts with different alteration facies in the same rock Alteration facies infills primary porosity Alteration facies cross cuts primary porosity or clast margins Rock contains relict high-temperature devitrification textures

Pre-fragmentation Pre-frag mentation Pre-lithification Post-lithification Post-devitrification

Alteration facies to successive volcanic units or intrusions Alteration facies truncated by younger less-altered rocks Alteration facies overprints intercalated intrusions Alteration facies overprints younger rocks

Older syn-volcanic/intrusion Syn- to post-intrusion and overlying units Syn- to post-younger rocks

Between alteration facies and diagenetic compaction textures Alteration facies protects primary texture from compaction Alteration facies defines the compaction foliation (e.g. fiamme) Alteration facies overprints the compaction foliation

Pre-compaction Pre- to syn-compaction Syn- to post-compaction

Hydrothermal assemblages and diagenetic facies Hydrothermal facies overprinted by early diagenetic facies Hydrothermal facies overprinted by late diagenetic facies Hydrothermal facies overprints late diagenetic facies

Pre- to early-diagenetic Syn-diagenetic Post-diagenetic

Alteration facies to early bedding-parallel stylolitic dissolution foliation (S1) Alteration facies overprinted by stylolitic foliation Alteration facies overprints stylolitic foliation

Pre- S1 Post-S1

Alteration facies to tectonic foliations and lineations (>S2) Alteration textures deformed and cut by tectonic foliation Alteration textures less deformed than primary volcanic textures Undeformed alteration textures

Pre-cleavage Syn-cleavage Post-cleavage

Alteration facies to metamorphic assemblages and textures Alteration facies overprinted by metamorphic assemblages Alteration facies overprints metamorphic assemblages

Pre-metamorphic Post-metamorphic

Megascopic alteration facies distribution Regional distribution Stratabound in formerly permeable facies Localised in fractured domains of coherent volcanic rocks Restricted to faults and shear zones

Diagenetic or metamorphic Pre- to syn-lithification Post-lithification and post-fracturing Syn- to post-faulting

Overprinting relationships between all alteration assemblages

Provides a paragenetic sequence for different alteration assemblages

72

I 73

4 | GEOCHEMISTRY OF ALTERED ROCKS

This chapter reviews three principal types of geochemical data that are used to characterise, quantify and interpret the processes of alteration: (1) whole-rock lithogeochemistry including major, trace and immobile trace elements; (2) mineral chemistry or composition of individual mineral phases in alteration mineral assemblages; and (3) stable isotope ratios of whole rocks and specific alteration minerals. Geochemistry is the study of the abundance and distribution of chemical elements in the earth (Mason, 1966). It has particular applications for the interpretation of altered rocks. Geochemical data can aid recognition of the type and composition of the precursor rocks before alteration. The chemical compositions of metasomatised or altered rocks and mineral phases reflect compositional changes during alteration and provide clues about the alteration processes. On a utilitarian level, geochemical data are routinely used in mineral exploration to identify broad chemical halos and gradients or vectors toward mineralised zones, and to discriminate between prospective and non-prospective targets.

4.1 | LITHOGEOCHEMISTRY Literally, lithogeochemistry is the determination and study of the abundances and distribution of elements in stones. In this chapter, we understand it to mean the whole-rock chemistry of the 10 or 12 major elements in rocks (usually expressed as oxides) and various groups of trace elements, such as rare earth elements (REE) and immobile elements. Galley's (1995) review of lithogeochemical applications in massive-sulfide exploration also included mineral and isotope chemistry but those methods are considered here in separate Sections (4.2 and 4.3). There are three main applications of lithogeochemistry in mineral exploration (Eilu et al., 1997): (1) identification or discrimination of prospective and nonprospective areas and lithological units (2) recognition of large alteration and geochemical halos to increase the size of exploration targets (3) definition of exploration vectors based on compositional gradients around ore deposits.

Several aspects of lithogeochemistry that attempt to 'see through' the effects of alteration to reveal the compositions of the protoliths, identify the process or type of alteration and quantify the chemical changes associated with alteration have particular relevance to studies of altered rocks. Chemostratigraphy uses immobile elements to identify lithologic correlations, magmatic affinities and geotectonic settings in otherwise unrecognisable altered rocks. Mass-transfer techniques, also based on immobile elements, are used to infer the compositions of unaltered precursors and quantitatively estimate the major element changes that occurred. Rareearth-element geochemistry can facilitate recognition of rock alteration processes (e.g. Eu anomalies in exhalites). Alteration indices may assist in the discrimination of alteration styles or facies, and the quantification of alteration intensity (Section 2.4). If composition data are available, lithogeochemical methods certainly contribute to the interpretation and evaluation of altered rocks. However, they are not quick solutions to all problems and may, if used in isolation, lead to false conclusions. In every case it is important to consider other geological information: field relationships, distribution or zonation of alteration mineral assemblages, macro- and micro-scale textures, mineralogy et cetera. Before describing the lithogeochemical methods, we provide explanations of some common lithogeochemical terms (Table 4.1), brief descriptions of geochemical sampling and analytical techniques, and draw particular attention to the phenomenon of closure in composition data.

Sampling and analytical methods How do we acquire lithogeochemical data? As in many other fields, the quality of interpretation rests on the quality of the data. Rocks are particularly difficult to analyse because of their wide compositional range and the chemical diversity of the elements of interest. An inappropriate analytical technique can lead to an expensive list of useless numbers and to the discredit of the unwary geologist who attempts to use them. This section offers advice on effective methods (they are never inexpensive) and some pitfalls to be avoided.

74 | CHAPTERTABLE 4.1 | Explanation of lithogeochemical terms.

Composition data

Data expressed as part of a whole (e.g. in weight percent, wt%, or parts per million, ppm).

Mass

A quantity of material, whole rock or its components, measured in weight units (g).

Units of mass change

Usually expressed in grams per hundred grams (g/100 g) to avoid confusion with composition data in weight percent, although they are effectively the same.

Absolute mass change (Aa)

Mass changes that are proportional to the whole initial mass of the rock. Usually expressed in grams per hundred grams (g/100 g). It may refer to individual components or the sum of all components (net mass change).

Relative mass change (Ar)

The absolute mass change in a component expressed as a percent proportion of the initial composition of that component in a rock. Relative mass changes distort the perception of chemical processes. For example, the addition of a small quantity (say 0.5 g/100 g) of MnO initially present at 0.5% would produce a 100% relative gain in MnO. In comparison, addition of the same quantity of SiO2 (0.5 g/100 g) to the same rock initially containing 50% SiO2 would represent only 1% relative gain. This gives the false impression that MnO was a vastly more important element of mass change than SiO2. Absolute mass changes are preferable because they are quantitative rather than proportional, and accurately reflect the quantities of materials added to or subtracted from the system.

Net mass change

The sum of positive and negative rock component mass changes relative to the initial mass of the rock. Usually expressed in grams per hundred grams (g/100 g).

Proportion

The amount of a component expressed as a proportion of the whole rock. Usually given in units of weight percent (wt%) for major elements and oxides, and in parts per million (ppm) for trace elements. It is analogous to "concentration" in chemical solutions.

Sampling strategy The number and size of samples to be used in a lithogeochemical investigation depends on many factors. These include the available budget, degree of exposure, geological complexity, compositional homogeneity of rock units or altered zones, the elements of interest and the volume of rock that can be physically removed from the outcrop or drill-core shed. There can be no universally applicable strategy but generally, more is best. Obviously, a good understanding of the geological context is fundamental to any subsequent lithogeochemical interpretation. Smash and grab sampling may lead to inadequate and often meaningless interpretations. Field methods should avoid sampling rocks potentially modified by weathering or other superficial secondary processes, unless these processes are under investigation. The sampler must also recognise potential overprinted alteration facies. For example, a sample containing a dense network of metamorphic quartz veins is of little value if the purpose is to study synvolcanic pervasive alteration phases; it would be better to select a sample without veins. Handling, transport and subsequent storage of lithogeochemical samples should not allow contamination or cross contamination. When sampling drill cores or cuttings stored at a mine site, be aware that mines and concentrators are dusty places, and expect to see the signature of the ore in your trace-element data if the samples are not clean. In all but very coarse-grained rocks, samples of around 1 kg should be adequate (Potts, 1987). In fine-grained or glassy volcanic rocks, samples in the range 200 to 500 g are generally sufficient for major element analyses. However, for low-level trace elements (<1 ppm) the 'nugget effect' may be significant if elements of interest are concentrated in sparsely disseminated grains (e.g. zircon). Potts (1987) provided illustrative tables of the sample weights required for analytical precision, at various concentrations and grainsizes. These

represent the quantities of sample that actually undergo analysis and should not be confused with the amount of pulverised sample from which smaller portions are taken for analysis.

Sample preparation All whole-rock analytical techniques (except neutron activation analysis) require samples to be crushed and ground to fine powder for direct analysis, acid digestion or fusion. There are opportunities here for sample contaminations from wear on the grinding machinery, cross contamination between samples in a batch, and contamination from previous batches of samples put through the mill. Reputable laboratories will use routine procedures to minimise those problems. However, if low-level trace elements are important, it may help to inform the laboratory personnel and request extra care. Jenner (1998) cautioned against the use of new diamond saw blades and recycled coolant for trimming slabs. He suggested that if sawn rock slabs must be analysed, they should be washed in acetone and distilled water. If the rock samples include overprinting phases, such as veins and weathered rinds, these are best cut out before crushing. The practice of crushing to -10 mm and attempting to pick out the rock chips without extraneous materials does not eliminate contamination. Jaw crushers are infamous contaminators. Crevices between their moving parts harbour dust that causes cross contamination, and the wear of plates directly contributes steel particles. Magnetic separation of steel particles is inadvisable because it may lead to selective removal of magnetic minerals. A better alternative is to crush rock samples down to ^5 mm fragments in a hydraulic press fitted with tungsten carbide plates. The crushed rock, including any fine dust, can then be coned and quartered down to about 80 or 100 g for pulverising to a final grainsize of less than 75 (Am in a swing mill.

GEOCHEMISTRY OF ALTERERD ROCKS | 75

Materials used in swing mills include tungsten carbide, chrome steel, alumina ceramic, and agate. All of these are potential contaminators of some elements. Chrome steel is the least expensive, but seriously compromises Cr and lowlevel Fe analyses. Tungsten carbide is popular and durable but will contribute Co and W to the sample, and possibly also Ti, Ta and Nb (Potts, 1987). Agate mills are fragile, expensive and slow but are the least likely to contaminate samples destined for low-level trace element determination.

Precision, accuracy and reference materials Precision is a measure of the repeatability of data within an analytical session, and reproducibility in different analytical sessions over longer intervals. Analysing several duplicate samples within batches and subsequent batches provides an estimate of the precision of a technique. Precision is reported as the co-efficient of variation (CV), which is equivalent to relative standard deviation (RSD) expressed as a percentage of the mean.

where s> is the standard deviation of element i in the samples analysed, and %J is the sample mean. Accuracy is a measure of how close analytical data lie to the 'true' values. It may be evaluated by including in each analytical batch some international geochemical reference materials, or in-house standards for which the true values are relatively well determined. It is important that such reference materials have compositions near the compositional range of the rocks being analysed (i.e. a standard rhyolite is inappropriate when analysing a batch of basalts). For each element accuracy is expressed as a positive or negative percentage difference, relative to the true or accepted values of the standard.

where Cr' is the true or accepted proportion of element i in the reference material, and Ca] is the apparent or analysed proportion of element i in the reference material. Reputable analytical laboratories will routinely run reference materials to calibrate their instruments and maintain an acceptable level of accuracy. It is helpful if the client, upon submittal, gives the analyst an indication of the rock types, mineral assemblages and compositional range of the samples, to enable an appropriate selection of reference materials. However, laboratories do not routinely report their analyses of reference materials for comparison with accepted values. Even if they do, the potential for fudging the results provides the user with a less than satisfactory assurance of accuracy. The solution is for the user to acquire some appropriate geochemical reference materials and include a sample or two, preferably disguised, in each analytical batch. This can involve considerable expense. Potts (1987) suggested that a minimum of 10 reference materials should be submitted with each batch for a full assessment of accuracy. In many cases, the user may not be greatly concerned with the 'correct' values provided that the data are relatively consistent or precise within and between batches. Submitting

representative duplicate samples allows the user to monitor long-term precision and calibration drift. It permits greater confidence in the analytical data, which justifies the additional expense. Keeping the reference materials and duplicates anonymous can be problematic, even if the sample numbers are disguised. A few small packets of powdered or crushed rock stand out prominently in a batch of drill-core samples and analysts have no difficulty in recognising them as standards. They may analyse them with special care, develop their own comparative data over successive batches, and be tempted to adjust any outliers. However, the recognition that the client is prepared to carry out independent quality control generally has a positive effect on analytical practice and discourages complacency at the laboratory.

Limit of detection Limit of detection of an element is commonly understood to be 'the lowest concentration that can be confidently measured by a particular method on an average sample' (e.g. Anonymous, 1997). However, as pointed out in detail by Potts (1987), the levels of confidence are frequently not stated. That obscures the reality that the quoted detection limits are often below the levels at which reliably quantitative measurements are possible. Potts (1987) proposed three new terms for better definition of the much abused 'detection limit': • Lower limit of detection (LLD) for a signal level of three standard deviations higher than the mean background (mean + 3s). This is the lowest level at which an element can be recognised but not quantitatively estimated. • Limit of determination (LOD) representing a level six standard deviations above the background signal (mean + 6s). It is the lowest level at which the signal can be quantitatively measured for a confident analysis. • Limit of quantitation (LOQ) is set at 10 standard deviations higher than background (mean + 10s) to provide extra confidence in the analysis in cases where there are legal, commercial or statutory implications placed on the interpretation of detection limits. According to Jenner (1998) the limit of detection commonly quoted by analysts is equivalent to Pott's (1987) lower limit of detection (LLD; mean + 3s). Jenner (1998) gave the example that if LLD for an element is 0.3 ppm, then data reported between 0.3 and 1 ppm may or may not be significantly different, but the element is recognisably present. Data above 1 ppm may be considered quantitative, or in other words, data at 1.1 and 1.5 ppm can be confidently regarded as different. The consequence is that one cannot place much reliance on data reported at close to the lower limit of detection. A safe rule of thumb is to treat with circumspection any data of less than one order of magnitude above the quoted detection limit. Therefore, select an analytical method that will provide quantitative data at an order of magnitude lower than the threshold of interest.

76 | CHAPTER 4

Analytical techniques The most popular methods for analysis of whole-rock samples are X-ray fluorescence spectrometry (XRF), inductively coupled plasma atomic emission spectrometry (ICP-AES) and inductively coupled plasma mass spectrometry (ICP-MS). These methods offer good precision for a large number of major and trace elements over wide concentration ranges. XRF remains the preferred method for major elements (Robinson, 2001). However, some commercial laboratories have recendy converted to ICP-AES, presumably for its rapid throughput and ability to measure most major and trace elements. ICP-MS is used only for trace element determinations. Neutron activation analysis (NAA) provides low limits of detection for REE, some platinum-group elements and some high-field-strength elements but it requires access to a nuclear reactor, produces radioactive waste and has slow turnaround of weeks or months. The geochemical laboratory at the University of Tasmania's School of Earth Sciences uses a combination of XRF on flux-fused and powdered samples (for major elements and moderate-abundance trace elements, respectively) and solution or laser-ablation ICP-MS for low-abundance trace elements (Robinson, 2001). To obtain the appropriate quality of data, it is important to involve the analysts in the selection of analytical methods. Inform them of the approximate compositional range in the samples and the geochemical objectives of the analyses. X-ray fluorescence (XRF) can be used to analyse up to about 60 elements with atomic numbers greater than 10 (Na upwards) at concentrations from 100% down to a few parts per million (Rollinson, 1993). Detection limits for trace elements in the range 0.5 to 2 ppm are achievable, and precision for major elements approaches less than 1% RSD (Robinson, 2001). However, routine procedures used in commercial laboratories generally result in higher detection limits of 2-10 ppm. Optimum detection limits and precision for trace elements, at concentrations below about 0.2%, are obtained by analysis of 6—10 g undiluted rock powder pressed into a pill. The rock powder must be ground to a grainsize of less than 75 [im to ensure homogeneity in the sample. Major element concentrations are determined on glass discs made by fusing a small amount of powdered rock diluted with lithium metaborate and tetraborate fluxes. The fusion produces a homogenous glass, which enables analysis of the light major elements and minimises X-ray absorption and enhancement matrix effects. The composition of the rock influences the type and dilution factor of the flux to be used, especially in sulfide-, base-metal- and carbonatebearing samples. Accordingly, at the geochemical laboratory in the University of Tasmania's School of Earth Sciences, approximate proportions of S, Fe, Ca, Ba, Cu, Pb and Zn are first determined on the powder pills, along with the trace elements, to enable appropriate selection of fusion fluxes. In sulfide-bearing samples, fusion should be a two-stage process with LiNO 3 in the flux to oxidise sulfur. Initially, the carefully weighed sample-flux mixture is heated and held for 10 minutes at 700°C to ensure that the sulfur is retained as sulfate, and not evolved. It is then heated to 1000°C for a further 10 minutes to complete the fusion and the melt is cast into a 32-mm diameter glass disc.

In the past decade, inductively coupled plasma (ICP) analysis has revolutionised geochemical analysis, particularly for trace elements. There are two separate methods of analysis known as ICP-AES and ICP-MS. Both use inductively coupled high-temperature argon plasma to generate atomic and ionic emissions in the sample. In ICP-AES (atomic emission spectrometry) the spectrum of atomic emissions is measured by an array of photomultipliers. This method is sometimes called ICP-OES (optical emission spectrometry). ICP-MS uses a mass spectrometer to measure ionic particles in plasma-sample gases. ICP-AES provides low limits of detection typically 2—10 ppm for trace elements and 10— 100 ppm for major elements. ICP-MS enables determinations of heavier trace elements at extreme detection limits, up to four orders of magnitude lower than ICP-AES (e.g. REE and HFSE detection limits in the range of 0.1 to 2 ppb). ICP methods are able to measure most elements at low detection limits with high precision using linear calibration over eight orders of magnitude (Robinson, 2001). Up to 50 elements can be analysed simultaneously in a few minutes on samples of less than 100 mg. There are a number of disadvantages to ICP related to the requirement that the rock sample must be dissolved in dilute solution before introduction to the plasma. Ensuring complete dissolution of rocks, including refractory phases such as zircon and REE minerals, and maintaining them in solution without contamination, is a difficult task (Yu et al., 2001). Samples are typically digested by strong acid cocktails in sealed teflon vials for one or two days, dried by evaporation and then the residue is re-dissolved in dilute nitric acid ready for analysis. In some cases, samples are fused with lithiumborate fluxes or sodium peroxide before acid dissolution. It is advisable to test for complete solution by comparing the ICP-MS data with XRF determinations of some relatively immobile elements, such as Zr, Nb, Y and REE (Robinson, 2001). The likely loss of some volatile elements during fusion and evaporation may render the method unsuitable for Hg, Tl, Sb, Se and As.

H 2 O and CO 2 A combination of XRF and ICP-AES or ICP-MS can provide accurate analyses of most major elements and a surfeit of trace elements. However, they cannot determine H 2 O and CO 2 as XRF is limited to elements of atomic numbers greater than 10, and ICP, which measures the samples in solution, obviously cannot determine water. Loss on ignition (LOI) is a poor substitute for H 2 O and CO 2 analyses. For geochemical analyses of altered rocks that contain significant hydrous minerals or carbonates, it is preferable to separately determine H 2 O and CO 2 . Hydrogen and total carbon in rock samples can be analysed by the 'C-H-N elemental analyser', an instrument designed for routine determinations of C, H and N in organic compounds (Potts, 1987). It is a relatively straightforward and inexpensive technique requiring only about 25 mg of finely ground rock powder. On the, generally safe, assumption that altered volcanic rocks do not contain significant organic substances, the total H and C determinations can be recalculated to H 2 O and CO 2 for inclusion in major element composition data.

GEOCHEMISTRY OF ALTERERD ROCKS | 77

Interference Some instrumental analytical methods are susceptible to inaccuracies caused by peak overlap, particularly in mineralised rocks with exotic-metal contents. For instance, in XRF analyses on powdered samples, high concentrations of Ba and Pb interfere with determination of Ti and Zr, respectively. These problems may be minimised by instrument configuration or, in extreme cases, another method of analysis. Therefore, it is important that the analyst is advised which elements are critically important for the geochemical interpretation and the approximate compositional ranges of the rocks submitted, so that the appropriate analytical methods and procedures are applied. For example, if the purpose of analyses is to infer the volcanic precursors of Pb + Zn + Ba-bearing mineralised rocks by immobile element chemostratigraphy, it would help the analyst to know that accurate Ti and Zr analyses are priorities, and that the mineral assemblage includes Pb- and Zn-sulfides, and barite.

Reporting data Major elements in rocks are conventionally reported in weight percent (wt%), mostly as oxides and in order of decreasing cation valency: SiO2, TiO 2 , A12O3, FeO (or Fe2O3 total), MnO, MgO, CaO, Na 2 O, K 2 O, P 2 O 5 , H 2 O + , CO 2 , S and Total. The percentages are relative to the dry sample; analysed after driving off moisture (H2O~) by heating for a few hours at 105°C. H 2 O + in the analysis represents structural water in crystals or glass. All other elements are typically regarded as trace elements although some may be present at greater than the conventional 0.1% cut off. Trace elements are usually reported in alphabetic order in parts per million (ppm), which is equivalent to the SI unit |ig/g preferred by analysts. It is often useful to re-order the trace elements into groups according to their geochemical characteristics (e.g. immobile trace elements, rare earths, etc.) to facilitate plotting. It is important to note that geochemical analyses are usually reported to one decimal place more than can be quoted with confidence (Potts, 1987).

Loss on ignition Hydrogen and carbon, mainly as H 2 O and carbonate, are significant components of some silicate rocks, particularly altered rocks. However, since XRF or ICP methods cannot determine these light elements, it has become common to report loss on ignition (LOI) as a proxy. LOI is determined by igniting a weighed sample at 1000-1050°C for at least 12 hours and then weighing the residue. Expressed in weight percent (as for the major elements), it is often assumed that LOI represents the combined proportions of H 2 O + and CO 2 in the rock. However, for several reasons pointed out by Potts (1987), the determination of LOI has little geochemical value: • The maximum temperature may not be high enough to dehydrate some minerals, such as talc, topaz, cordierite and epidote. • Oxidation of ferrous iron (e.g. in silicates, magnetite and

sulfides) can produce a weight gain, partly offsetting the LOI. Even if FeO and Fe are separately determined, it is not reasonable to assume that they will be completely oxidised in the LOI process. • LOI may include a contribution from volatilisation of alkali metals, sulfur oxides and fluorine; the loss may be only partial and not predictable. For example, in samples containing both sulfide and carbonate, some SO 2 may react with CaO and remain in the ignited sample. The variables affecting LOI and the wide range of compositions in altered and mineralised rocks may lead to unpredictable results that are not readily interpretable. Consequently, LOI is not a reliable substitute for H2O+ and CO 2 analyses, and it may significantly underestimate the evolved volatiles.

Totals The sum of the elements in a major-element analysis is frequently taken as an indication of analytical reliability. Considering the shortcomings of LOI determinations, it is unreasonable to expect that their inclusion in a major element analyses will provide totals close to 100%. Nor should one expect that the error values on determinations of individual elements should cancel each other to produce totals of 100%. Further ambiguity is due to the usual practice of analysing and reporting total Fe as Fe2O3, irrespective of its actual oxidation state. Sums of XRF major element oxides, sulfur and LOI data will exceed 100%. This is because the sulfur is measured twice: in the fused disc XRF and in the LOI determinations. However, simple subtraction of XRF determined sulfur does not solve the problem because the sulfur may not be entirely evolved in the LOI process. Reasonable totals are obtained if H 2 O + and CO 2 are separately determined by another method and summed with XRF major element oxides and sulfur. Low totals do not necessarily indicate erroneous analyses; it may be that some significant elements were not determined. For example, boron constitutes about 10% of tourmalines and some micas contain significant proportions of Li, Fl, Rb and Ce (Deer et al., 1966), and mineralised rocks may contain significant proportions of base metals. Potts (1987) cautioned that, although it is nice to find totals near 100%, it is not a satisfactory test of the quality of analytical data. He considered that the only acceptable way of checking accuracy of modern instrumental analytical techniques, such as XRF and ICP, was to analyse appropriate geochemical reference materials. By applying this quality control practice to all of the methods used to produce a set of major-element data, the user can have reasonable faith in the relative accuracy of the individual elements analysed. Likewise, confidence in the accuracy of individual determinations permits acceptance of odd totals as accumulated errors, or indications of incomplete analyses.

Recalculating to volatile free It is common practice, particularly in petrological literature, to recalculate major-element analyses to 100% 'anhydrous'

78 | CHAPTER 4

or 'volatile free'. The recalculation involves multiplying the proportion of each major-element oxide in an analysis, except LOI, by a factor derived from the formula 100/2 Q, where 2 C| is the weight percent sum of all major oxides (not including LOI) in the analysis. Thus the 10 major oxides are commonly adjusted to sum to exactly 100%. LOI is reported, but not included in the total (e.g. Crawford et al., 1992; Stolz, 1995). The object of recalculation is to remove apparent variations in the proportions of major oxides that may be due to differences in LOI values. This may be a valid approach for studies of petrogenesis and fractionation in unaltered or weakly altered rocks, where alteration was limited to hydration or minor calcite vesicle fillings (Crawford et al., 1992). In these cases, additions of H 2 O + and CO 2 may be the only significant metasomatic changes, although CaO may be added if it is not otherwise derived from the decalcification of plagioclase. It is not unreasonable to 'subtract' the estimated additions due to hydration and carbonation to determine equivalent anhydrous magmatic compositions. In more altered rocks, however, many of the other major elements, particularly Si, Fe, Mg, Ca, Na and S, are also likely to be involved in significant mass changes. Thus recalculating to volatile free would distort the pattern of mass changes, artificially increasing the changes of other elements relative to the constituents of LOI. In cases where the 10 major oxides (SiO2 to P2O5) sum to low totals, volatile free recalculation followed by immobile-element-based mass change estimates could result in actual small net mass losses of some elements appearing as mass gains. Similarly, it may upwardly distort the estimates of net mass change, with implications for interpretation of volume changes. Barrett and MacLean (1994a) recommended using volatile free recalculated major element data for mass change estimates. They neglected to mention whether the same recalculation factor should also, for consistency, be applied to all the trace elements. Failure to similarly adjust the trace elements could lead to small inconsistencies in immobile element ratios (e.g. Ti/Zr, which is usually calculated from major element TiO 2 data). It could also positively skew subsequent mass change estimates based on immobile elements. The volatile free recalculation has a proportionally greater effect on data for altered rocks because they generally have higher LOI than their unaltered precursors. Upward recalculation of all trace elements, as well as majors, can avoid those problems. However, comprehensive recalculation is unnecessary as both chemostratigraphic methods and mass change estimates are based on ratios of elements and hence are unaffected by recalculation. The method used in some studies (e.g. Gemmell and Large, 1992), of recalculating the 10 major oxides but not other major elements, such as sulfur, is also inconsistent. It has the effect of under-estimating mass changes in sulfur relative to the other oxides. In dealing with altered rocks, it is preferable to obtain as near as practical 'complete' major element analyses. The major element suite should include sulfur, H 2 O + and CO 2 and any other elements likely to exceed 0.1% (e.g. base metals). Ensure adequate quality controls and then trust in the accuracy of the

data. They do not require fudging or normalisation at the risk of introducing new errors and misinterpretations. Otherwise, where analyses of altered rocks are limited to the usual 10 major oxides and LOI, it is best to treat LOI as a (somewhat fuzzy) component in its own right, and not to recalculate to artificial 100% totals. This particularly applies to data that are to be used in mass transfer calculations.

Closure in composition data Closure, also known as the constant sum effect, affects all analytical data expressed as proportions of a whole; that is, as composition data (or as 'concentrations' in Stanley and Madeisky, 1996). The total of all elements in a rock analysis must sum to 100%, or one million ppm, and so forth, according to the units of measurement. Chemical mass transfers that change the total mass of a system by adding or removing some elements (e.g. hydrothermal alteration) will affect the proportions of all elements, even those not involved in the mass transfers (Fig. 4.1). Consequently, the apparent differences in composition data between unaltered and altered rocks do not accurately reflect the real material changes, except in systems where there has been no net mass changes Closure particularly obscures mass changes in the major elements such as Si, Al and Fe, which exist in high proportions in primary unaltered rocks. For example, intense silicification of felsic volcanic rocks may not be apparent in composition data (e.g. footwall alteration zone of the Thalanga deposit, Herrmann and Hill, 2001). Closure is less of a problem for elements of low initial proportions. For example, in felsic volcanic rocks, Ca and Na concentrations average only a few weight percent and their depletion associated with plagioclasedestructive alteration is usually evident in the composition data. Trace elements, by definition, are present at low proportions in background rocks. They tend to provide high-contrast anomalies in mineralised and altered rocks, commonly orders of magnitude greater than background levels, and are therefore practically unaffected by closure.

FIGURE 4.1 | Schematic illustration of the effect of closure in composition data (after Eilu et al., 1997). Alteration leading to a net mass gain in the system results in a lower proportion (or concentration) of element i, although the actual mass of element i is not changed (AM, = 0). Mp = initial mass and MA = altered mass of total system.

GEOCHEMISTRY OF ALTERERD ROCKS I 79

Rollinson (1993) presented a detailed discussion of the problem of closure in geochemistry and suggested some solutions. It is not a trivial matter in the study of altered rocks. The techniques for estimating alteration-related mass transfers overcome the problem of closure by quantifying the amounts of elements or oxides gained by or lost from altered rocks.

Chemostratigraphy Major element compositions are routinely used to classify volcanic rocks in terms of petrogenesis and tectonic setting (e.g. Pearce and Cann, 1973). However, the same method is not applicable to altered rocks because many of the major elements, especially Si, Fe, Mg, Ca, Na and K, are relatively mobile during alteration. Consequently, compositional changes related to alteration may considerably outweigh their primary variations. Fortunately, several elements are chemically immobile during most types of alteration and these can be reliably used to classify and correlate altered volcanic rocks. In this context, immobile means elements that are neither added to, nor taken from, the rock during alteration. Immobile elements may be involved in phase changes and perhaps be mobile at millimetre scale, but their mass in the altered rock remains unchanged. Although the proportions (concentrations) of immobile elements may change, due to net mass changes in the size of the system, their inter-element ratios remain the same.

Incompatible elements Incompatible elements are those that tend to be excluded from the lattices of minerals crystallising from magmas and are instead partitioned into the melt phase. Hence, incompatible elements exist at highest proportions in the most evolved felsic rocks. The high-field-strength elements (HFSE) Zr, Y and Nb are generally incompatible, except in some calc-alkaline suites. They have similar low magmatic liquid-solid distribution coefficients and so tend to retain similar inter-element ratios throughout a single magmatic fractionation series, and on x-y bivariate plots form linear trends that project from the zero origin. Subsequent alteration involving net mass gain or loss can change the proportion of incompatible elements in the whole rock but their primary ratios are preserved. As a result alteration trends coincide with the primary fractionation trends. The gradients of these trends vary according to magmatic affinity (MacLean and Barrett, 1993; Barrett and MacLean, 1994a). Samples from different magmatic suites thus produce separate linear trends of magmatic enrichment, which regress to the origins on incompatible-incompatible immobile element plots (e.g. Fig. 4.2). Incompatible-incompatible element ratios and bivariate plots have chemostratigraphic applications even in hydro thermally altered samples in which the major elements may not be reliable discriminants. Incompatible-incompatible element ratios are used to identify magmatic affinities, favourable volcanic suites and terranes.

Immobile elements The high-field-strength elements Ti, Zr, Nb and Y are relatively immobile during hydrothermal, diagenetic and weathering alteration, and during regional metamorphism up to mid-amphibolite facies. Ratios of these immobile elements are the basis of tectono-magmatic discrimination diagrams developed in the 1970s (e.g. Pearce and Cann, 1973; Floyd and Winchester, 1978). Many studies of VHMS deposits have shown that Al, Ti, Zr, Nb, Y, heavy REE (Lu, Yb), Hf, Ta and Th, and in some cases P, Sc, V and Cr, remain essentially immobile during alteration. Their immobility has been documented even in the most intense hydrothermally altered zones directly beneath deposits (e.g. MacLean and Kranidiotis, 1987; Skirrow and Franklin, 1994; Barrett and MacLean, 1994a). Barrett and MacLean (1994a) recognised some mobility of the light REE in proximal, intense chlorite altered zones beneath some deposits and suggested that they may be useful as exploration vectors. Y and Nb show considerable scatter in some datasets and this may be partly attributable to primary variations (Ewart, 1979) or slight chemical mobility in some systems. Analytical precision may also be a factor, particularly for Nb, which typically occurs at low concentrations not much above XRF detection limits. In practice, Ti and Zr are the most reliably immobile elements. They can be inexpensively and accurately analysed by XRF on pressed powder pellets and they exist at easily detectable levels in most volcanic rocks, unlike the heavy REE, Sc, Nb, Ta, Hf and Th, which generally exist at less than 20 ppm.

Compatible elements Compatible elements have high magmatic distribution coefficients (>1) and are preferentially taken up by mineral

FIGURE 4.2 | Schematic Y-Zr (incompatible-incompatible) plot used for determination of magmatic affinities in altered volcanic rocks (modified after MacLean and Barrett, 1993).

80 | CHAPTER 4

phases crystallising from magma. Consequently, compatible elements are depleted from the melt phase (this is opposite to the enrichment in incompatible elements). Thus, the relative proportion of compatible and incompatible elements in residual melts changes as fractionation proceeds. Batches of magma that are successively tapped off and emplaced as eruptive or intrusive units will have successively smaller compatible-incompatible element ratios. Aluminium, Ti, Cr, Sc and V are generally compatible during crystallisation and immobile during alteration (Barrett and MacLean, 1994a). Bivariate plots of immobile compatible-incompatible element data for least-altered samples from a particular magmatic affinity should show smooth fractionation trends, generally with negative slopes (Fig. 4.3). Subsequent net mass gains and losses imposed by alteration will produce, on compatible-incompatible immobile element plots, separate alteration lines for each chemically distinct rock unit (Fig. 4.3). This property is particularly useful for the discrimination and correlation of initially homogenous but subsequently altered volcanic units that may be otherwise unidentifiable. Chemostratigraphic fingerprinting is commonly done with Ti/Zr ratios and TiO2—Zr bivariate plots. However, in tholeiitic suites Ti enrichment parallels the typical Feenrichment trend at the mafic end of the fractionation series (MacLean and Barrett, 1993). In other words, Ti is initially incompatible in tholeiitic fractionation series up to about the composition of basaltic-andesite. On a TiO2— Zr plot, the mafic end of the tholeiitic series has a positive fractionation trend, essentially similar to the alteration trends superimposed by subsequent metasomatic net mass changes (Fig. 4.3). Therefore, TiO 2 -Zr is not a reliable discriminant of altered mafic tholeiites. A reasonable substitute is A12O3— Zr, which has a near linear, slightly negative trend in tholeiites (MacLean and Barrett, 1993). Titanium becomes compatible in tholeiitic magmas more evolved than basaltic-andesite and the TiO 2 -Zr fractionation curve then has a negative slope. Two examples of chemostratigraphic discrimination and correlation in the Mount Read Volcanics are presented in Figures 4.4 and 4.5.

FIGURE 4.3 | Schematic TiO2-Zr (compatible-incompatible) plot showing the negative curvilinear fractionation trend typical of co-genetic calc-alkaline volcanic suites (after Barrett and MacLean, 1994a). Mafic tholeiites may show a positive trend up to the composition of basaltic-andesite, because of TiO2 incompatibility in the early stage of magmatic differentiation.

It is worth emphasising that immobile element ratios must be used with discretion in chemostratigraphic discrimination and correlation. This method is most effective in rocks with primary compositional homogeneity, such as coherent lavas and sills, and possibly in some massive syneruptive volcaniclastic units such as pumice breccias. Processes of magma generation, crystallisation and fractionation determine the immobile element ratios of magmas. Therefore, immobile element ratios are likely to be uniform within single coherent eruptive or intrusive emplacement units. However, volcaniclastic debris may be subject to unhomogenising processes. Both lateral and vertical compositional variations may occur in volcaniclastic units as a result of: • mechanical sorting of components of different densities, such as clasts, pumice, scoria, crystals (e.g. Ti-oxides and zircon) and glass shards, during eruption and transport • winnowing of glass shards from turbidity currents • mixing of debris from a variety of volcanic sources in variable proportions or • incorporation of extraneous clasts into the base of volcaniclastic mass flows. Gifkins (2001) showed that thick, graded, rhyolitic pumice breccia units in the Central Volcanic Complex, western Tasmania, have Ti/Zr ratios that vary from -5 near the crystal- and lithic-rich bases, to -9 in the fine-grained, shard-rich tops of the units. This is consistent with an increased concentration of zircon crystals towards the base of the units but may also reflect the abundance of felsic clasts in the basal portions. In contrast, graded rhyolitic volcanic breccia units in the Rosebery hanging wall, western Tasmania, display the opposite trend. Large et al. (2001b) suggested that decreasing Ti/Zr ratios towards the top of these units were due to physical fractionation of Zr-poor crystal and lithic components from Zr-bearing originally glassy pumice and shards during emplacement.

Testing immobility It is preferable that element immobility in any system is established, rather than assumed, before proceeding with chemostratigraphic interpretations of altered rocks. This is also an important preparatory step for methods of estimating mass transfers of mobile elements. The simplest test is to plot potentially immobile elements on x-y bivariate diagrams with the origins at zero. If possible, the tests should use data from unaltered and variably altered samples of a single, originally compositionally uniform volcanic unit, such as a coherent lava or sill. If the selected elements are immobile, the data points for a single-precursor system should align in a highly correlated linear trend, which projects to the origin of the plot and through the data points of the least-altered samples. These linear trends, or alteration lines, are due to net mass gains and losses of the mobile elements in the altered rock samples (Fig. 4.6). Typically there is some data scatter due to analytical errors and slight inhomogeneities in the primary rock. However, if both elements are immobile, calculated linear correlation co-efficients (r) for alteration lines should exceed -0.85 (Barrett and MacLean, 1994a). In contrast, elements that were mobile during alteration are

GEOCHEMISTRY OF ALTERERD ROCKS | 81

0

10 20 30 40 50 0 Ti/Zr

25

50

75

100

Alteration Index

100(K2O+MgO) (K 2 0+Mg0+Ca0+Na 2 0) FIGURE 4.4 | Graphic log and down-hole Ti/Zr and Al plots of drill hole NC4, near Lake Newton, western Tasmania. The Ti/Zr data clearly delineates units of different primary volcanic compositions despite effects of strong hydrothermal alteration in rocks intersected in the middle and lower part of the hole. The high Ti/Zr ratios at -200-230 m led to recognition of an altered mafic volcanic breccia unit, which had previously been interpreted as a zone of chlorite-altered felsic volcaniclastic rocks. Another altered mafic unit occurs below 530 m. Quartz and feldspar crystal-rich sandstones in the upper 100 m have a range of Ti/Z ratios, which suggests that they do not have a unique provenance.

readily identifiable by their erratic distribution or near total removal (Finlow-Bates and Stumpfl, 1981).

Mass transfer techniques Mass transfer techniques aim to quantify the amounts of individual elements added to and subtracted from the rock during alteration in order to overcome the distortions of closure that are inherent in composition data. As noted by Barrett and MacLean (1994b), significant mass change anomalies may not be apparent in untreated

compositional data, due to closure. The results of mass transfer calculations are usually easy to relate to mineral assemblages and may reveal clues about the composition, source and temperature of hydrothermal fluids (e.g. Barrett and MacLean, 1994b). Mass change data have been used to infer hydrothermal water-rock ratios (e.g. MacLean and Hoy, 1991). They are also used in the interpretation of whole-rock 618O and REE data (e.g. MacLean and Barrett, 1993). Thus, they may enable discrimination of favourable alteration systems and altered zones within systems. When plotted spatially mass transfer data can be used as quantitative exploration vectors (Section 8.2).

82 I CHAPTER 4

FIGURE 4.5 | Chemostratigraphic correlation diagram of the volcano-sedimentary succession that hosts the Rosebery massive sulfide deposit, western Tasmania. A thin unit of dacitic pumice breccia, between the massive sulfide lens and the feldspar + quartz + biotite porphyry sill intersected in hole 120R, is texturally indistinguishable from the footwall rhyolitic pumice breccias (Ti/Zr = 7-9) but has a distinctive Ti/Zr ratio of between 12 and 14.

FIGURE 4.6 | TiO2-Zr plot of data from a tholeiitic volcanic suite (after MacLean and Barrett, 1993). The data points for least-altered basalt, andesite and rhyolite samples define the magmatic 'fractionation curve'. Two linear arrays of data represent variably altered samples of originally homogenous units of andesite and rhyolite. These highly correlated alteration lines intersect the least-altered data points on the fractionation curve and project towards the origin. Data points on the alteration lines below the fractionation curve represent net mass gains, those above the curve represent net mass losses. The positive slope at the basaltic end of the fractionation curve is due to incompatibility of TiO2 in the early stages of differentiation of tholeiitic magmas. This may lead to confusion of fractionation and alteration trends in mafic tholeiites.

GEOCHEMISTRY OF ALTERERD ROCKS | 83

There are several approaches to estimating mass transfers. These include the mathematically complex graphical method of Gresens (1967); subsequently simplified by Grant (1986) and Huston (1993) to produce the isocon method; the immobile element techniques of MacLean and Barrett (1993), and the Pearce element ratio analysis of Stanley and Madeisky (1996). These techniques all depend on recognition of immobile elements. They also, with the exception of the Pearce element ratio analysis, depend on the identification of precursor-rock compositions. The determination of unaltered precursor compositions is often problematic in the altered, lithologically and structurally complex rocks that host ore deposits. The safest approach is to consider the geological context, established from field relationships and rock textures, in combination with various immobile element tests (e.g. MacLean and Barrett, 1993; Stanley and Madeisky, 1996). This ensures that precursors are appropriately matched to the altered rocks under investigation. The isocon method does not include a procedure for selecting precursors and it commonly produces erroneous results because incorrect geological assumptions are applied. For this reason, we prefer the more rigorous MacLean and Barrett (1993) method, which is also simplest to calculate. Another potential difficulty arises from primary compositional variations in volcanic rocks related to magmatic fractionation or volcaniclastic mixing processes. Huston (1993) recognised this as a deficiency of the isocon method. He suggested examining standard deviations of data from least-altered samples to screen their suitability, estimating errors for the consequent mass changes and recognising that small apparent mass changes may be artefacts of primary inhomogeneities, not due to alteration. The Pearce element ratio analysis method (Stanley and Madeisky, 1996) and the multiple-precursor variant of the MacLean and Barrett (1993) method attempt to overcome the limitations of primary variability. Of these two, we prefer the MacLean and Barrett approach, again for its relative ease of calculation.

The MacLean and Barrett multiple-precursor method involves sorting samples into affinity groups, calculating primary variability trends (fractionation curves) and using immobile elements to synthesise primary compositions for altered samples that may fall between compositions of the available least-altered samples. This method relies on the assumptions that co-genetic groupings can be recognised and separated, and compositional variations are smooth linear or curvilinear trends for every major element. This is generally true for magmatic fractionation, and therefore applicable to volcanic suites, but the variations may be erratic if volcaniclastic, sedimentary or multiple alteration processes were involved (Eilu et al., 1997). Thus, due caution must be exercised when using this method in mixed provenance volcano-sedimentary successions. Filtering sample sets through geological field evidence and petrographic textural observations are obvious fundamental precautions. Sample sets from restricted prospect-scale areas are less likely to have complex primary compositional variations than district-scale sample sets. Peculiar aberrations or significant mass changes in elements otherwise expected to be immobile, such as Al, should arouse suspicion that primary fractionation trends may have been incorrectly modelled.

Single precursor mass transfer technique The MacLean and Barrett method of estimating mass transfers in single-precursor systems incorporates testing of immobility as a fundamental step. The analytical data for potentially immobile elements, from an initially homogenous but variably altered lithological unit, are plotted on x-y bivariate diagrams and linear regressions are calculated. The existence of highly correlated (r >0.85) trends that pass through the origin enables selection of the optimal element to be used as the immobile monitor in the mass change calculations. This graphical process also highlights any outliers or samples

FIGURE 4.7 | Diagram representing the calculation of mass transfers for an altered rock with an initial mass of 100 units (after MacLean and Barrett, 1993). The proportion of immobile element in the altered rock decreases (is diluted) because of net mass gains in mobile element(s). The ratio of immobile elements Z°/Za represents the factor, which multiplied by the proportion of immobile element in the altered rock, would restore it to a mass unit, rather than a proportion of the whole. The same factor applied to all other elements produces the 'reconstructed composition' in mass units. The mass transfers in each element are then calculated by subtracting its primary composition from the reconstructed composition. In this example, a large mass gain in X combined with a small mass loss in Y produces a large net mass gain.

84 | CHAPTER 4

from different precursors, which should be eliminated or treated separately. It identifies any primary compositional inhomogeneities in the rocks, which may require treatment by the multiple-precursor method (MacLean and Barrett, 1993) or Pearce element ratio analysis (Stanley and Madeisky, 1996). These more complex methods are not explained here; the reader is referred to the relevant references for details. The single-precursor mass transfer method proceeds by calculating the ratios of the proportions of an immobile element in the altered and unaltered samples. Each of the mobile element proportions is then multiplied by that ratio to obtain a reconstructed composition (Fig. 4.7). The mass change of each element is found by subtracting its percent

proportion in the precursor from that in the reconstructed composition. The steps below and the flow chart in Figure 4.8 outline the procedure for estimating mass changes in single-precursor systems. Figures 4.9, 4.10 and 4.11 present a worked example based on compositions of Thalanga footwall rhyolites, Queensland. (1) Acquire and tabulate the lithogeochemical analyses (e.g. Figure 4.9). The first part of this section (4.1) presents some guidance on sampling and analytical techniques. Barrett and MacLean (1994a) recommended using major element composition data that are recalculated to 'volatilefree' totals of 100%. However, it is unreasonable to expect Comments related to Thalanga data in Figures 4.9, 4.10 and 4.11. Values below detection replaced by arbitrary small values; e.g. 0.5 x detection limit or zero (Figures 4.9 and 4.10).

Data should preferably be from mappable single-precursor units. This example includes 34 rhyolites from diverse volcanic units. However, the remarkable uniformity of im m obile- el em ent ratios suggests they were essentially co-magmatic and can be treated as a singleprecursor system.

No obvious outliers. Yttrium (Fig. 4.10) typically shows considerable scatter, which may be a result of primary variation.

AI?O3 andZr both have three of four correlation co-efficients >0.85 (Fig. 4.10). Either could be used as the immobile monitor component. Zirconium was selected because of marginally higher average correlation (0.86 vs 0.83).

Appropriate $ symbols in the formula enable it to be filled across and down in the spreadsheet and maintain cell references to Z°, Za and C°.

Large positive and negative net mass changes are mainly due to SiO2 gains and losses, or addition of Fe and S in pyritic samples (Fig. 4.11), Note that because of closure, the SiO2 changes are not obvious in the analytical data.

FIGURE 4.8 | Flow chart for mass-change calculations by the single-precursor method (after MacLean and Barrett, 1993).

GEOCHEMISTRY OF ALTERERD ROCKS | 85

elements of an analysis to total exactly 100%, and the use of recalculated data may produce positive distortions in subsequent mass transfer estimates. In dealing with altered rocks, it is preferable to obtain accurate analyses of all the major elements (including S, CO, and H 2 O + ) and include them in the mass transfer calculations. If S, CO 2 and H 2 O + data are not available, and losses on ignition (LOI) are significant (>2%), then LOI could be included as a separate, albeit loosely denned, component of mass change. (2) Test for immobility. Plot analytical data for potentially immobile elements (e.g. Al, Ti, Zr, Nb, Y, etc.) on x-y bivariate diagrams with the origins at zero (e.g. Fig. 4.10). If possible, use data from variably altered and unaltered samples of a single uniform rock unit. (3) Inspect for outliers and, with geological considerations, decide whether to cull them, or treat the data as a multipleprecursor system. Readers are referred to MacLean and Barrett (1993) for details of the multiple-precursor method. (4) The immobile element data for single-precursor systems should plot on highly correlated linear trends (r >0.85) that pass through the origin and through the data points for least-altered samples. Evaluate the bivariate diagrams A

B

and correlation factors to determine which element consistently occurs in highly correlated trends and is most suitable as an immobile monitor (i.e. was uniform in the primary rock and least mobile during alteration). (5) Select a composition for the precursor. This could be from a single unaltered sample or an average of several unaltered samples. (6) Calculate the absolute mass change for each component using the formula: Aa = [Z7 Z a . Ca ] - C° where Aa is absolute mass change expressed in g/100 g Ca = wt% proportion of component in altered rock C° = wt% proportion of component in precursor Za = proportion of immobile element in altered rock Z° = proportion of immobile element in precursor. The mass changes may be calculated from compositions of individual altered samples or from average compositions of sample groups representing certain mineral assemblages or altered zones. (7) For visual comparison, plot the absolute mass changes for individual elements on a bar graph (e.g. Fig. 4.11).

C

D

E

F

G

H

I

J

K

L

M

N

O

SiO,

TiO2

AIA

FeA

MnO

MgO

CaO

Na2O

K2O

PA

S

COj

Total

1

Whole-rock major and trace element composition data

2 3 4 5

Sample no. * 140802 * 140727

east-altered footwall -noderate, foliated sericite + chlorite

76.40 70.90

0.11 0.10

11.90 14.50

1.64 1.69

0.04 0.04

0.67 3.94

1.42 0.12

2.27 1.05

4.04 3.92

0.02 0.01

0.01 0.00

6 7 8 9 10

* 140808 * 140724

strong, pervasive quartz + sericite + pyrite ± chlorite ntense, pervasive quartz + pyrite

75.70 67.00

0.07 0.05

11.40 6.60

5.14 14.34

0.08 0.02

2.38 1.67

0.00 0.05

0.21 0.06

2.39 1.99

0.01 0.01

0.65 10.61

* 140902 TH394-142 TH144B-34 TH41A-575 TH005-256 TH238-236 TH038-191 TH085-125 TH148-159 TH085A-422 TH270-278 TH270-145 TH085-312 TH018-266 TH038-054 TH085-159 TH085A-348 TH238-194 TH270-220 TH085A-335 TH085-204 TH085-215 TH085A-241 TH085-188 TH270-313 TH061-086 TH41A-713 TH061-157 TH270-381 TH085A-384 Precursor composition

ntense, microcrystalline quartz + K-feldspar east-altered footwall east-altered footwall

84.20 77.39 77.25 76.71 77.80 74.62 77.81 75.25 82.02 69.97 64.37 72.70 72.73 66.82 75.37 75.29 66.93 74.41 63.38 78.04 71.22 67.08 71.59 74.54 72.56 74.39 75.87 78.48 55.36 49.30

0.06 0.09 0.10 0.07 0.07 0.09 0.07 0.07 0.08 0.11 0.12 0.09 0.08 0.08 0.08 0.09 0.10 0.09 0.05 0.06 0.05 0.07 0.07 0.06 0.10 0.11 0.10 0.08 0.10 0.10

7.40 11.85 11.90 10.18 8.83 13.40 10.54 11.13 8.84 13.06 17.56 12.22 12.33 11.26 10.70 12.79 14.57 12.18 7.99 9.48 8.28 10.44 11.03 8.40 12.55 13.75 12.17 11.60 13.90 13.47

0.60 1.13 2.09 0.99 1.90 1.40 1.51 2.32 0.68 2.80 4.73 2.39 5.48 2.46 4.47 1.96 5.70 1.42 13.46 4.45 9.18 8.62 6.91 6.95 4.93 1.56 1.70 0.88 14.36 19.14

0.00 0.02 0.03 0.08 0.06 0.02 0.09 0.06 0.02 0.03 0.07 0.09 0.01 0.13 0.02 0.09 0.08 0.05 0.13 0.04 0.09 0.13 0.06 0.05 0.07 0.03 0.05 0.01 0.14 0.30

0.08 0.76 0.68 4.33 2.14 3.54 3.13 1.90 1.05 1.98 4.35 4.24 1.10 9.88 2.91 2.51 4.70 4.16 5.88 1.76 3.97 6.42 2.53 2.77 3.88 1.19 1.12 0.61 4.90 7.79

0.11 0.46 0.90 1.75 3.08 0.12 0.25 3.03 0.24 0.28 0.01 0.27 0.00 0.01 0.00 0.08 0.00 0.07 0.00 0.02 0.00 0.01 0.08 0.00 0.00 0.34 0.08 0.12 0.11 0.19

0.34 4.79 2.77 1.57 0.84 1.82 0.90 0.59 0.53 0.38 0.24 0.22 0.17 0.12 0.12 0.10 0.09 0.08 0.00 0.16 0.15 0.12 0.11 0.09 0.03 0.95 0.42 0.37 0.18 0.08

5.87 1.82 3.60 2.80 3.06 2.23 3.40 3.68 5.33 3.69 4.15 4.73 3.75 5.08 2.43 4.41 2.91 3.57 0.91 2.45 1.51 1.59 2.82 2.14 2.50 6.33 6.36 7.60 2.40 0.93

0.01 0.01 0.02 0.02 0.05 0.01 0.01 0.02 0.02 0.05 0.02 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.02 0.02 0.02 0.02

0.48 0.00 0.24 0.01 1.33 0.00 0.06 1.39 0.30 2.03 1.08 0.84 3.47 0.53 1.61 0.81 0.86 0.44 7.64 1.57 4.97 2.89 3.59 4.36 0.44 0.00 0.66 0.13 5.92 5.12

77.01

0.10

11.88

1.62

0.03

0.70

0.93

3.28

3.15

0.02

11

12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38

average of 140802, TH394-142 and TH144B-34

{Formula in cell C451 ={$Q$38/$O5"C5)'• . 1 C$38] is ?lled across and down. ' units h \ Placement of $ symbols is critical to \ DL 0.05 i maintain cell references for Zo, Za and Co. \

39 40 41 42 43 44 45 46 47 48 49

Alteration facies

0/

0.08

V_

siOj

140727 140808 140724

moderate, foliated sericite + chlorite ' >• strong, pervasive quartz + sericite + pyrite ± chlorite ntense, pervasive quartz + pyrite

-15 7 43

0.00 98.03 2.75 128 0.00 102.40 8.24 79

14 7

36 44 20 34 33 30 57 30 41 42 28 49 46 36 35 45 27 48 46 39 19 30 17 29 37 29 39 52 47 37 63 44

0.15 0.15 0.99 0.00

0.26

0.00 0.18

0.10

0/

0/

0.01

0.01

0.01

0.01

0.05

0.01

0.01

0.01

0.1

TiO2

AIA

FeA

MnO

MgO

CaO

Na.O

K2O

PA

s

co2

140902

ntense, microcrystalline quartz + K-fe!dspar

72 units

g/100q

9/100g

0 4 24 -1 g/100g

0 0 0 0

3 2 2 -1

-1 -1 -1 -1

-2 -3 •3

-3

0 -1 0 7

0 0 0 0

0 1 19 1

99.15 98.47 99.74 98.51 100.15 97.25 97.77 99.44 99.11 94.38 96.70 97.80 99.13 96.39 97.72 98.14 95.95 96.74 99.45 98.04 99.43 97.38 98.81 99.37 97.07 98.66 98.55 100.08 97.39 96.44

0.49 0.49 0.77 1.38 2.12 2.49 2.08 1.75 1.03 3.39 4.23 2.56 4.07 2.65 3.54 2.50 4.11 2.79 7.60 2.95 5.33 5.02 4.73 4.50 3.17 1.24 1.71 0.76 7.22 7.83

80 9 137 14 141 12 113 12 114 13 140 13 119 13 116 12 114 11 162 17 195 20 130 14 129 15 132 15 119 14 137 16 176 16 138 16 83 9 105 11 84 9 116 12 120 14 92 11 140 16 158 18 157 17 122 13 154 19 145 16

98.91 0.62 141

13

T

Y

40 54

0.05

1 1 0 1

LOI Zr Nb 13 19

0.01

0 0 0 0

S

Q|R

98.52 0.60 146 96.37 2.67 162

0.00 0.10

35

ppm ppm ppm 1 1 1 ^Formula in cell 045 is Tiled • 1 down [ =sum(C45:N45)i

\ Mass Changes

P

Net

0 0 0 0

-15 9 84 76

q/100q q/100q q/100q q/100g q/100q g/100g q/100q q/100q q/100q

FIGURE 4.9 | Example of layout and calculation of mass changes by the single-precursor method for 34 rhyolites from the Thalanga footwall, Queensland. Data from Paulick (1999) and Paulick et al. (2001). Values below limit of detection replaced by zero. Sample numbers with * are featured in the Thalanga data sheets (Section 7.7).

86 | CHAPTER 4

FIGURE 4.10 | Bivariate plots of potentially immobile elements: AI2O3, TiO2, Zr, Nb and Y (Thalanga data from Figure 4.9). (A) TiO2 versus Zr, (B) AI2O3 versus Zr, (C) Nb versus Zr, (D) Y versus Zr, (E) TiO2 versus Nb, (F) AI2O3 versus Nb, (G) Y versus Nb, (H) TiO2 versus Y, (I) AI2O3 versus Y, (J) TiO2 versus AI2O3, (K) Na2O versus Zr, and (L) sulfur versus Zr. These plots facilitate recognition of compositional outliers, which should be excluded from single-precursor calculations, and selection of the least-mobile element to serve as the immobile monitor in mass-change calculations. In this example, Zr and AI2O3 both show consistent highly correlated trends that project to the zero origin. Note the considerable scatters, and hence poorer linear correlations, in plots involving Y and the major elements sulfur and Na2O. These are consistent with primary compositional variations or significant chemical mobility, particularly for sulfur and Na2O.

FIGURE 4.11 | Bar graph showing estimated absolute mass changes of major elements in four samples representing major alteration fades in the Thalanga footwall, Queensland (data from Figure 4.9).

GEOCHEMISTRY OF ALTERERD ROCKS I 87

Rare-earth-element geochemistry related to alteration In the past few decades, trace elements have become basic tools or pathfinders in ore deposit exploration and in petrogenetic interpretation. Immobile trace elements have valuable applications in studies of altered rocks. The rare earth elements (REE) have some special properties during alteration, which may be useful in interpretation and should be understood in order to avoid false conclusions.

Mobility of light REE and effects of net mass change Rare earth elements (with the exception of Eu) are generally incompatible during igneous fractionation. Heavy REE (Lu and Yb) are essentially immobile, whereas the light REE may be variably mobile during alteration (MacLean and Barrett, 1993). Lanthanum is the most likely to be affected and the mobility of the other REE decreases towards the heavy REE (Barrett and MacLean, 1994a). These incremental changes in the lighter REE modify the slopes of chondrite-normalised REE patterns, which may confuse petrogenetic interpretation. Therefore, immobility of REE needs to be established before they are used to infer magmatic affinity. Plotting geochemical data for each rare earth element against a reliably immobile element, such as Zr, is a means of testing for scatter and mobility. If the elements were immobile, REE and Zr data

from altered single-precursor systems or co-genetic volcanic suites should produce highly correlated linear trends on bivariate plots. In systems where REE were chemically mobile, both positive and negative shifts in light REE concentrations have been recorded. Significant mobility of REE appear to occur in proximal altered zones associated with VHMS deposits (MacLean and Barrett, 1993). The greater mobility of the light REE may produce a fan shaped array in chondritenormalised plots with the REE profiles converging towards the heavy, least-mobile REE (e.g. Fig. 4.12A) Even in cases of REE immobility, net mass changes associated with alteration may produce significant vertical shifts in chondrite-normalised REE patterns. The slopes of the REE patterns are retained, but the y-axis magnitudes are modified (downward by net mass gain and upward by net mass loss; Fig. 4.12B).

Europium anomalies in seafloor sediments Recent studies of sediments and hydrothermal precipitates in modern and ancient massive sulfide environments have found them relatively enriched in light REE, particularly in Eu (Barrett et al, 1990; Peter and Goodfellow, 1996; Shikazono, 1999). The explanation is that Eu exists in a divalent state in felsic magmas and hence is compatible in feldspars, unlike the other trivalent REE, which remain incompatible (Rollinson, 1993). The divalent Eu+2 in feldspars may be liberated by subseafloor hydrothermal alteration to sericite or chlorite, whereas the incompatible REE that are concentrated in alteration resistant phases, remain relatively immobile. The liberated Eu+2 is transported by reduced acidic hydrothermal fluids and may ultimately be precipitated by oxidation at the seafloor. Therefore, in felsic volcanic successions, altered VHMS-footwall zones tend to be depleted in Eu. In contrast, positive Eu anomalies exist in seafloor sediments and jaspers, and probably indicate proximity to hydrothermal vents (Barrett et al., 1990). Both phenomena have significance for massive sulfide exploration in modern and ancient submarine volcanic environments. The recognition of positive Eu anomalies in stratiform jasper lenses recently contributed to the discovery of a small satellite massive sulfide deposit at Thalanga in the Mount Windsor Volcanics (Miller et al., 2001).

4.2 I MINERAL CHEMISTRY Principles FIGURE 4.12 | Modifications in REE patterns due to hydrothermal alteration illustrated by REE profiles of variably altered rhyolites from the Ansil and Delbridge deposits, Canada (after Barrett and MacLean, 1994b). (A) Partial leaching of the mobile light REE in the Ansil footwall is reflected in profiles with different slopes converging towards the least-mobile heavy REE. (B) Immobility of all REE is evident in the sub-parallel profiles for rhyolites from the Delbridge footwall. However, net mass changes in mobile major elements have caused changes in the proportions of the REE, producing vertical shifts in the profiles.

Minerals, by definition, are natural homogenous solids of definite chemical composition and definite atomic structure (Dana, 1957). However, many minerals do not have simply defined chemical formulas. Their compositions may lie between limits defined by two or more end-member formulas, effectively forming solid solutions. Mineral crystal structures can accommodate various impurities where atoms and ions

88 | CHAPTER 4

of suitable size and charge can substitute for others in the lattice, occupy interstices in the lattice or be omitted from a proportion of lattice sites. The considerable variety of linked tetrahedral crystal structures in silicate minerals permits a wide range of chemical substitutions and interstitial solid solutions. A frequently cited example is the olivine series in which Mg2+ and Fe2+ ions, having similar charge and size, substitute for each other between the end member compositions of forsterite (Mg2SiO4) and fayalite (Fe2SiO4). The sheet-like structures of phyllosilicates allow a great range of ionic substitutions and interstitial contaminants. For example, muscovite, with the ideal formula of K2Al4[Si6Al2O20] (OH) 4 , commonly contains the isomorphous substitutions of Na, Rb, Cs, Ca and Ba for K; Mg, Fe2+, Fe3+, Mn, Li, Cr, Ti, and V for octahedral Al, and F for OH and tetrahedral cation proportions from Si6Al2 to Si7Al] (Deer et al., 1966). Layered clays (particularly smectites) also accommodate many cationic substitutions and exchanges as well as physical mixtures and inter-layered structures of more than one clay mineral (e.g. smectite-illite). The causes of such variations are both physical and chemical. Temperature, pressure, fluid pH, JO2, and cation solubility may all affect the stability and composition of the minerals formed. Chemical and temperature gradients in hydrothermal alteration systems commonly produce spatial zonation of alteration mineral assemblages. Similarly, they may produce gradational and zonal variations in some alteration mineral compositions. For example, the Mn content of metasomatic pyroxenes associated with Zn skarns generally increases systematically along the fluid pathway, and can be used to identify proximal and distal skarns, and altered zones (Meinert, 1993). The standard methods for determining mineral compositions and crystal structure are electron microprobe and X-ray diffraction analyses. Although Galley (1995) noted that these methods are more readily available than previously, the impracticality of mineral chemistry as an exploration tool, and the complexity and cost of these laboratory techniques means that they are rarely used other than in academic research. The development of field portable short wavelength infra red (SWIR) spectrometers and spectral interpretation software during the last decade, has allowed mineral chemistry to be practically integrated into exploration programs for a variety of deposit types (Thompson et al., 1999). SWIR spectral analysis can determine compositional and crystal-structural variations in white micas, smectites, clays, chlorites, biotites and carbonates (Pontual et al., 1997). These minerals are prominent in altered volcanic rocks and are also prone to significant compositional variations. The technique reliably estimates variations in white mica composition but appears to be less effective at analysing chlorites in typical mixed phyllosilicate assemblages (Herrmann et al., 2001). Carbonates have relatively weak SWIR absorptions, which tend to be obscured in mixed assemblages containing phyllosilicates and are thus less amenable to spectral analysis. Although there are many possible applications in mineral deposit exploration, we do not know of any cases where systematic investigations of mineral chemistry have led to an ore discovery. As pointed out by Simmons and Browne (2000), the extent to which patterns of mineral distribution

and chemical variations can be used in exploration depends largely on whether they are related to a single phase of hydrothermal activity that produced equilibrium mineral assemblages.

Applications The three main applications of mineral chemistry in alteration studies are: (1) interpretation of the processes and physicochemical conditions of alteration (2) discrimination or identification of metasomatic alteration and mineralisation styles from mineral compositions (3) determination of spatial variations and exploration vectors to ore.

Interpretation of processes The compositional and crystal-structure variations in some minerals are diagnostic of particular alteration processes because of physicochemical influences on mineral stability and composition. Thus, mineral chemistry may be used to infer the geological environment in which an alteration mineral assemblage formed. However, these kinds of compositional variations may not always be universally applicable; they may require orientation testing to determine their usefulness in different districts or sites. For example, Dill et al. (1997) found that the kaolinalunite deposits in felsic volcanic rocks of western Peru could be classified into hypogene (hydrothermal) and supergene (weathering) types on the basis of the chemical variations in kaolinite. Hydrothermal kaolinite tended to be rich in Ba, Sr and sulfur, whereas weathering-related kaolin clays concentrated Cr, Ti, Nb and REE. This approach has direct applications to mineral exploration because hypogene kaolinite alteration in Peru is associated with high-sulfidation epithermal Au-Ag deposits. Similarly, Yang et al. (1999) alluded to lowand high-crystallinity forms of kaolinite in the Comstock district of Nevada, which they respectively attributed to lowtemperature weathering and higher temperature hydrothermal alteration processes. They suggested that spectral recognition of distinctive kaolinite could be used in satellite-borne remote sensing to detect prospective hydrothermal altered zones.

Discrimination of hydrothermal alteration styles Mineral compositions can be used to identify or fingerprint hydrothermal-alteration mineral assemblages, and possible associations with mineralised rock. This application is useful at early stages of mineral exploration to discriminate between economically favourable and less favourable alteration and mineralisation styles. For example, Zn-skarn assemblages are commonly dominated by pyroxene with varying amounts of garnet, amphibole, bustamite, chlorite and carbonate, which may all be Mn enriched. Manganese-rich pyroxene (johannsenite) is virtually diagnostic of distal Zn skarns (Meinert, 1983) and could be used as an index mineral in exploration for this type of deposit.

GEOCHEMISTRY OF ALTERERD ROCKS | 89

Another example that is more relevant to submarine volcanic successions, is the recognition of a class of pyriticalteration systems in the Mount Read Volcanics that have some characteristics of high-sulfidation epithermal deposits. These include Western Tharsis and Lyell-Comstock, which contain sub-economic Cu + Au resources (Huston and Kamprad, 2000; Corbett, 2001), and Basin Lake and Chester, which appear to be barren (Boda, 1991; Green and Taheri, 1992; Williams, 2000; Williams and Davidson, 2004). White micas in the central zones of these systems have distinctive, non-phengitic sodic compositions (Herrmann et al., 2001). The recognition of sodic white mica, along with low 634S values in pyrite and the presence of pyrophyllite, enables the discrimination of this type of alteration system from those associated with economic Zn-rich polymetallic VHMS deposits (i.e. Rosebery and Hellyer). The altered footwall zones of Zn-rich polymetallic VHMS deposits contain normal potassic to slightly phengitic white micas. The fact that the compositional variations in white mica can be simply determined by SWIR spectral analysis makes this a practical method for selecting and ranking exploration targets.

Mineral chemistry exploration vectors Many documented studies have shown spatial variations in mineral chemistry in altered zones around VHMS deposits. Chlorite has received the most attention, particularly in the last two decades, and there are few deposits for which no data are available. There has also been significant interest in carbonate and, to a lesser extent, white mica compositions. Chlorite compositions are typically Mg-rich in the proximal altered zones of VHMS deposits. They commonly show systematic distal trends to more Fe-rich compositions. These trends are typically recognisable over several hundred metres, both laterally and stratigraphically into the footwall, away from the ore. Some examples include the Seneca and Corbet deposits in Canada (Urabe et al., 1983), the Arctic deposit in Alaska (Schmidt, 1988), and the Thalanga deposit in north Queensland (Paulick et al., 2001). However, there are many cases where the opposite trend exists and Fe-rich chlorites occur in proximal altered zones. The Aznacollar and Masa Valverde deposits are two examples in the Iberian pyrite belt (Sanchez-Espana et al., 2000). At the Home deposit, Canada (MacLean and Hoy, 1991), chlorites in proximal chlorite-rich zones are more Fe-rich than in the enclosing sericite + chlorite zone (Fig. 4.13). Similarly, at Mattagami Lake (Abitibi belt, Canada) there is a general trend of Fe enrichment in chlorites upwards towards the ore position and outwards from the core of the altered footwall zone (Costa et al., 1983). In northern Turkey, the dacite-hosted deposits of the eastern Black Sea province have altered footwall zones of Mg chlorite and sericite (Cagatay, 1993). In contrast, the western Black Sea ophiolite-hosted pyritic Cu deposits of the Kure district are associated with Fe-rich chlorites and trends of Fe enrichment toward ore. Some deposits exhibit inconsistent patterns of variations in chlorite composition. McLeod (1987) found that Mg chlorites around the Mount Chalmers deposit (Queensland) have a stratigraphic upwards trend of Fe enrichment in the footwall and a sharp reversal to Mg enrichment in the mineralised zone. Two recent regional-

scale studies by Hannington et al. (2003a, 2003b) showed contrasting compositional trends in chlorites associated with VHMS deposits in the Noranda district, Canada, and Kristineberg deposits of the Skellefte district, Sweden. In the Noranda district, the moderately Fe-rich compositions of chlorites (Fe/Fe+Mg 0.4-0.9) associated with sulfide deposits and surrounding district-scale hydrothermally altered zones, contribute to discrimination of prospective and non-prospective volcanic centres. However, chlorites in the Kristineberg district are distinctly Mg rich (Fe/Fe+Mg 0.050.5) and show little variation between proximal and distal alteration facies. These studies also found the chlorites associated with sulfide deposits contained significant Mn (up to 1% MnO) and Zn (up to 0.5% ZnO) suggesting that these could be used as proximity indicators in exploration. McLeod and Stan ton (1984) investigated several eastern Australian VHMS deposits and showed that chlorites in sphalerite-rich ores are relatively Mg rich compared to those in chalcopyrite-rich ores. Furthermore, the Mg/Fe ratios of chlorites are related to Mg/Fe ratios of co-existing phyllosilicates and the Fe content of co-existing sphalerite. Therefore, zonal compositional variations in chlorite may be reflected in other alteration mineral species, such as white mica, which may be more easily measured by SWIR. Importantly, McLeod and Stan ton (1984) concluded that the compositions of chlorites and other phyllosilicates had not been significantly modified by subsequent greenschist facies metamorphism. Variations in chlorite composition have also been used, with some success, as empirical and thermodynamically calculated geothermometers to estimate temperature gradients in hydrothermal systems above 200°C (e.g. Cathelineau and Nieva, 1985; Walshe, 1986). They are sensitive to re-equilibration and therefore not reliable indicators of hydrothermal temperatures in subsequently metamorphosed terrains (Green and Taheri, 1992).

FIGURE 4.13 | Aliv-Mg-Fe cation plot showing trend to Fe-rich chlorite with proximity to the Cu-Au VHMS deposit at the Home Mine, Quebec, Canada (after MacLean and Hoy, 1991). Where Ab = albite, Ep = epidote, Mt = magnetite and Ser = sericite.

90

|

CHAPTER 4

White micas, commonly referred to as sericite, are nearly ubiquitous in massive sulfide-related hydrothermal alteration systems and they can vary considerably from the ideal muscovite formula of K2Al4[Si6Al2O20](OH)4 (Deer et al., 1966). The term phengite refers to white micas in which Fe, Mg and some other cations substitute for Al in octahedral sites and the charge balances are maintained by increased Si/Al ratios in tetrahedral sites. Phengitic micas form solid solutions between the end members of muscovite and celadonite: K2(Mg,Fe2+)2(Al,Fe3+)2Si8O20](OH)4. Barium-rich phengitic micas, in which Ba substitutes for K in inter-layer sites, also exist in some sediment-hosted sulfide and VHMS deposits (e.g. Schmidt, 1988; Jiang et al., 1996; Leistel et al., 1998). At low to moderate temperatures, there may be limited Na substitution for K, with Na/Na+K ratios up to about 0.2. The sodic muscovites generally have low phengite contents. White micas formed at temperature below 300°C may have a significant proportion of vacancies in inter-layer sites normally occupied by K, as well as phengite-like Fe-Mg substitution in octahedral sites. These are commonly termed illites; they form complex solid solutions between three end-members: muscovite, celadonite and pyrophyllite. Yang's (1998) review provides a more detailed description of variations in white mica compositions. Although white mica composition has been examined in a number of massive sulfide-related hydrothermal alteration systems, few studies were systematic enough to evaluate its usefulness as an exploration tool. White micas in the proximal altered zones of the weakly metamorphosed Hellyer deposit, western Tasmania, are more phengitic than the normal muscovites in distal altered zones (Yang, 1998). At the nearby but slightly more metamorphosed Que River deposit, Offler and Whitford (1992) found considerable small-scale compositional variations in mica, even within single samples, due to a complex alteration history. Although the metamorphic phases preserve hydrothermal alteration compositional trends, no convincing vectors were recognised, possibly because of structural complications. Around the Arctic deposit, Alaska, white micas span almost the entire compositional range between muscovite and celadonite (Schmidt, 1988). Metamorphic micas outside the hydrothermal altered zones are highly phengitic. Micas in a variety of proximal alteration mineral assemblages are variably phengitic; some contain up to 0.4 cations of Ba per formula unit and the least-phengitic types are significantly sodic. It is not clear whether these variations are systematic enough to be used as broad exploration vectors. Altered zones in the Iberian pyrite belt also have a confusing variety of mica composition patterns. Plimer and de Carvalho (1982) found that white micas in altered footwall zones around the Salgadinho Cu deposit are phengitic, and appear to show increase in Fe/Fe+Mg ratios towards the mineralised zone. In contrast, in the Rio Tinto deposit the proximal altered zones contain muscovite and the distal altered zones (up to 2500 m from the deposit) contain micas of more phengitic composition (Leistel et al., 1998). The Masa Valverde deposit is associated with Ba-rich muscovites and some ore bodies in the Aljustrel district have extensive halos of sodic white mica (Leistel et al., 1998; Carvalho and Barriga, 2000). With the possible exception of Ba substitution, most of the compositional variations in white micas are semi-quantifiable

by SWIR spectral analysis. This has been demonstrated by several recent studies in the Mount Read Volcanics (Herrmann et al., 2001). Figure 4.14 provides an example of variations in wavelengths of Al-OH bond-related absorption features in SWIR spectra of white mica in samples taken at intervals from a single drill hole. These wavelength variations are directly related to white mica compositional variations. Some alteration systems, particularly those associated with disseminated Cu-Au deposits and/or kaolinite ± pyrophyllite assemblages, exhibit compositional gradients in white mica compositions that are measurable over a few hundred metres. The background white mica compositions are commonly variably phengitic and tend to non-phengitic muscovite or sodic-muscovite in proximal altered zones (e.g. Huston and Kamprad, 2000; Herrmann et al., 2001). Figure 4.15 illustrates variations in wavelengths of Al-OH absorption features, related to white mica composition, spatially around the Western Tharsis deposit. Carbonates are a third group of minerals that can accommodate compositional variations and are common in some VHMS altered zones. Documentation of carbonate compositional trends and zonal distributions is fairly sparse. However, it seems that massive-sulfide-related carbonates are typically Fe-, Mg- or Mn-bearing phases, and background diagenetic or metamorphic carbonates are commonly calcic. Documented examples include the Hokuroku district in Japan (Shikazono et al., 1998), the Rosebery deposit in western Tasmania (Large et al., 2001b), and the South Bay deposit in northwest Ontario (Urabe et al., 1983).

FIGURE 4.14 | Stack of selected SWIR hull quotient spectra of core samples from a diamond-drill hole through the altered zone at the Chester deposit, western Tasmania. Annotations on the left side are depths in metres down the hole. The spectral features are almost entirely attributable to white mica in the alteration mineral assemblages. Note the distinct variation in wavelengths of the Al-OH absorption features at around 2200 nm. These indicate that the hole intersected mineral assemblages containing normal potassic muscovite in the upper part, sodic white mica from about 100 to 250 m and muscovite to slightly phengitic white mica in the lower part.

GEOCHEMISTRY OF ALTERERD ROCKS | 91

FIGURE 4.15 | Cross-section of the Western Tharsis deposit (western Tasmania) showing zonation of wavelengths of AI-OH absorption features in SWIR spectra. The background of 2200-2210 nm, corresponding to slightly phengitic white mica, decreases over a few hundred metres to 2194-2198 nm, attributable to non-phengitic, slightly sodic white mica in the proximal altered zone associated with disseminated pyrite and chalcopyrite.

Inevitably, there are exceptions, such as the deposits of the northern Iberian pyrite belt, which have calcite, ankerite and dolomite in proximal alteration mineral assemblages (Sanchez-Espana et al., 2000). The Mattabi deposit in the Sturgeon Lake area, Canada, is underlain by a funnel shaped siderite-rich altered zone grading outwards to dolomite, which is widespread on a district scale in the footwall and hanging-wall volcanic rocks. Hydrothermal carbonates in the Rosebery-Hercules area, western Tasmania, are conspicuously Mn rich, (Khin Zaw and Large, 1992; Large et al., 2001b). Large et al. (2001b) showed that Mn-siderite and ankerite carbonates in the footwall of the Rosebery deposit increase in Mn content towards ore (Fig. 4.16). Magnesium contents of carbonates in altered footwall zones of the South Bay deposit, Canada, increase steadily towards ore over distances of tens to hundreds of metres (Urabeetal., 1983). Chlorite, white mica and carbonate all have potential as mineral exploration vectors, at least on a prospect or deposit scale. However, the considerable diversity of compositional trends in the published data indicate that exploration vectors need to be empirically established on a district or deposit specific basis, and are not universally applicable.

FIGURE 4,16 | Downhole plot of drill hole 120R illustrating the distribution of Mn-rich carbonates (kutnahorite and manganosiderite-rhodochrosite) in proximity to K-lens of the Rosebery Pb-Zn VHMS deposit, western Tasmania. Magnesiumcarbonates occur in the footwall and in a thin unit of altered pumice breccia immediately above the ore lens; carbonates more than 50 m above ore in the hanging wall sequence are Ca rich.

92 I CHAPTER 4

4.3 I STABLE ISOTOPES

TABLE 4.2 | Natural abundances of H, C, 0 and S isotopes, and standards in common use (data from Rollinson, 1993).

Theoretical background Isotope geochemistry is a diverse and rather specialised science. This section aims to provide a bare outline of aspects that have particular relevance to interpretation of altered volcanic rocks. It includes only a brief introduction to theoretical principles, necessary to grasp the applications. We recommend that interested readers supplement this by referring to other textbooks — Rollinson (1993) provides an excellent working basis. Isotopes are distinct atomic forms of elements that have the same number of protons but different numbers of neutrons in their nuclei. Of the 92 naturally occurring elements, 60 consist of more than one isotope and many of them have two or more stable isotopes. That means that they are non-radioactive, and do not change naturally, or decay, into other radiogenic elements by emissions of subatomic particles from their nuclei. Some natural radiogenic isotopes have important geological uses in geochronology, petrogenesis and metallogenesis, because their rates of decay are constant and measurable. Stable isotopes also have many geological applications, mainly based on their properties of isotopic fractionation. The stable isotopes of the light elements H, C, O and sulfur have received the most attention from geochemists because they are naturally abundant in the hydrosphere and in crustal rocks, not least in altered rocks. Informal isotopic notation uses the chemical symbol of the element preceded by the mass number of the isotope written as a superscript. Thus 17O denotes the oxygen isotope with 17 nucleons, comprising eight protons and nine neutrons. A single isotope, which usually has equal numbers of protons and neutrons, typically dominates the isotopic composition of each element (Table 4.2). Therefore isotopic ratios are very small numbers (e.g. for the average abundances of oxygen isotopes, 18O / 16 O = 0.002). To avoid direct comparison of these unconvincingly small ratios, stable isotopic proportions are expressed in parts per thousand (i.e. per mil, %o) relative to a standard material (i.e. delta form). For example:

Stable isotopes undergo fractionation (or selective partitioning into different phases) according to thermodynamic properties that are related to their differing atomic weights and consequent ionic bond strengths (Faure, 1986; Rollinson, 1993). Fractionation may occur by several physicochemical processes of which the most geologically important are isotopic exchange reactions between phases. The degree of fractionation is controlled by physical and chemical factors, which vary according to the elements and fractionation processes involved. Thus, O-isotopic fractionation is largely dependent on temperature, whereas S-isotopic fractionation is influenced by temperature, pH, / O 2 , and the activities of sulfur and other cations involved with sulfate.

1

H

99.9844

2

D

0.0156

12

C

13

98.89

C

1.11

16

O

99.7630

17

O

0.0375

18

O

0.1995

32

S

95.02

33

S

0.75

34

S

4.21

36

S

0.02

Std mean ocean water (SMOW), ViennaSMOW (V-SMOW) or PDB belemnite.

PDB belemnite

Std mean ocean water (SMOW), ViennaSMOW (V-SMOW) or PDB belemnite.

Troilite in Canon Diablo meteorite (CDT)

Isotopic applications in alteration studies Isotopic studies of alteration mineral assemblages associated with mineralised zones may help to estimate alteration temperatures and water-rock ratios, interpret fluid origins, discriminate between alteration styles and identify altered halos around ore deposits. However, it is worth repeating Ohmoto's (1986) cautionary advice to integrate isotopic studies with geologic, mineralogic and geochemical data. He stated: 'there is more than one process, which may produce the same isotopic characteristics (in an ore deposit) and the same geological process may produce entirely different isotopic characteristics in different conditions. Therefore, isotopic data alone cannot provide a unique answer to any geological problem, especially when the data are limited to isotopes of one element.' Diagenetic and hydro thermal alteration of volcanic rocks invariably involves hydration reactions, between minerals and water, and so the amount and isotopic composition of water are important variables. Apart from geologic variability, sample preparation, isotopic analytical methods and calibration of fractionation factors also introduce significant uncertainties. Experimentally, empirically and thermodynamically determined isotopic fractionation factors provide a confusing diversity of choice for use in isotopic calculations. The Laboratoire de geochimie isotopique at Universite Laval, Quebec, has a comprehensive compilation of fractionation factors from many published sources and is accessible at <www.ggl.ulaval.ca/personnel/beaudoin/labo> (Beaudoin and Therrien, 1999).

Geothermometers The temperature dependency of isotope fractionations between mineral pairs forms the basis of isotope geo-

GEOCHEMISTRY OF ALTERERD ROCKS | 93

thermometry. Provided that the paired minerals formed in equilibrium, that their fractionation factors are known and are significantly different, their original isotopic compositions have been retained and can be separately determined, then a combination of equations can be solved for temperature of formation. This approach is useful with O isotopes, because O is common to, and abundant in, silicates and other alteration phases, such as carbonates and sulfates. It is also applicable to S-isotopic compositions of minerals in complex sulfide and ore assemblages. Hydrogen isotopes are not generally reliable as geothermometers because they are readily modified by subsequent fluid interactions. Furthermore, the mineral fractionation factors are relatively insensitive to temperature and are not well calibrated (Ohmoto, 1986). It is usually difficult to physically separate fine-grained minerals for isotopic analysis and to petrographically demonstrate equilibrium between the analysed phases. However, close agreement between several temperature estimates of two or more pairs of minerals in a single assemblage (e.g. quartz + magnetite, muscovite + chlorite and calcite + chlorite) would inspire reasonable confidence in their isotopic equilibrium and the calculated temperature. In fluid-dominated hydrothermal systems, mineral-water O-isotope fractionation factors can be used to estimate relative temperatures. Although it is difficult to reliably measure the isotopic composition of the water from fluid inclusions (Nesbitt, 1996), an assumed value can provide approximate or relative temperature estimates. This approach is used in the determination of ocean palaeo-temperatures from 618O values of the carbonate shells of marine organisms (Rollinson, 1993) and also has applications in mineral exploration (e.g. Miller et al., 2001).

Fluid origins Natural waters have a broad range of H- and Oisotopic compositions because of fractionation effects in the hydrosphere, lithosphere and mantle (Fig. 4.17). Consequently, it may be possible to infer the source or sources of alteration fluids, and something about their evolution, from their isotopic signature. Fluid inclusions in hydrothermal minerals may permit direct measurement, but commonly the fluid compositions are calculated from isotopic compositions of alteration minerals with known fractionation characteristics, that are assumed to have been in equilibrium with the hydrothermal fluid. Isotopic composition of a single hydrothermal mineral may constrain the fluid composition if independent temperature estimates, such as fluid inclusion data, are available. Otherwise, isotopic compositions of mineral pairs in equilibrium can be used (as outlined above) to deduce temperature, which can then be applied in the mineral-water fractionation relationship to estimate fluid-isotopic composition.

Water-rock ratios Knowledge of water-rock ratios may help to determine the processes of alteration and interpret the hydrology of

FIGURE 4.17 | 6D- and S18O-isotopic compositions of natural waters (from Taylor, 1979; Ohmoto, 1986). SMOW is standard mean ocean water with 6Dand618O values of 0%o.

hydro thermal alteration systems. Water-rock ratios can be estimated from whole-rock O-isotope data or from inferences of mass transfers and solubilities (e.g. Ohmoto et al., 1983). Unaltered mafic volcanic rocks have initial 618O values in the range 6 to 7.5%o, slightly higher than the mantle value of 5.7%o, and unaltered felsic volcanic rocks have values up to about 10%o, (Hoefs, 1973). The 618O values of hydrothermally altered volcanic rocks will differ from initial values, depending on the temperature and mineral assemblage (which affect fractionation), the initial isotopic composition of the water and the quantity of water that reacted with a given amount of rock. The water-rock ratio is usually expressed in atomic proportions of oxygen. Taylor (1979) presented the following equations expressing these relationships in closed and open hydrothermal systems: closed systems

where the superscripts and subscripts i, f, w and r, respectively refer to initial, final, water and rock. These equations can be plotted as curves of the type illustrated in Figure 4.18, which relate 818O / to water-rock ratios. Thus, a measured final 5 18 O r can be used to estimate the amount of water involved in hydrothermal or diagenetic alteration, under assumed (or otherwise determined) values for temperature, whole-rock fractionation factors and the initial isotopic compositions of fluid and rock. However, as discussed in some detail by Ohmoto (1986) and noted by Green and Taheri (1992), natural geologic systems are not likely to be simple isothermic, closed or open systems. Rates of isotopic re-equilibration vary according to temperature, and the isotopic compositions of both rock and

94 | CHAPTER 4 30

fluid change incrementally along the flow path. The final rock 5 I8 O reflects an integrated history of fluid-rock reaction and offers only broad constraints on temperature and water-rock ratio. Green and Taheri (1992) suggested that conditions of diagenesis might approximate a closed system, whereas submarine hydrothermal convection is more analogous to an open system. Water-rock ratios calculated under assumptions of either closed or open systems are likely to represent the minimum values because of kinetic and incremental factors affecting rates of re-equilibration. Natural open systems may require water-rock ratios one or two orders of magnitude greater to achieve equivalent shifts in the isotopic composition of the rock (Ohmoto, 1986). Nevertheless, consideration of water-rock ratios is important in evaluation of whole-rock 618O data. This will be further explained in the following section on isotopic exploration vectors.

Oxygen-isotope exploration vectors FIGURE 4.18 | Curves illustrating relationships between water-rock ratio and final-rock 618O at various equilibration temperatures under parameters of: 618Owatef = 0 %o, S18 Cy = 7 %0, fractionation factor Aw' = (2.68 x K W ) - 3.53 (plagioclase). Solid and dashed lines represent open and closed systems, respectively. Note that re-equilibration with a small amount of water at low temperature can produce a large increase in rock 618O.

Early isotopic studies (e.g. O'Neil and Silberman, 1974; Taylor, 1974) discovered the link between terrestrial epithermal Au-Ag deposits, meteoric-hydrothermal convection and very broad halos of low 618O in volcanic host rocks. These extensive isotopic halos had obvious potential as semi-regional exploration vectors and stimulated further investigations into volcanic successions hosting other deposit types. Among them was the landmark study by Green et al. (1983) on whole-rock O-isotope geochemistry in the host rocks to VHMS deposits in the Hokuroku district, Japan. They found concentric zonation of whole-rock 618O values around the cluster of Fukuzawa ore bodies ranging from 6.7 + 1.3%o

FIGURE 4.19 | Cross-section illustrating the distribution of whole-rock 518O values (black contours), and altered footwall zones around the Fukuzawa deposits, Hokuroku district, Japan (modified after Green et al., 1983).

GEOCHEMISTRY OF ALTERERD ROCKS | 95

in the proximal sericite + chlorite zone and 11.1 + 2.5%o in the surrounding 1—3 km-wide montmorillonite zone, to 16.9 ± 2.7%o in the outer zeolite zone (Fig. 4.19). The 618O anomaly is significantly broader and less variable than elemental geochemical halos; it extends up to 1 km laterally beyond the Na 2 O depletion anomaly and at least 400 m into the hanging wall above the mineralised zone. The wide extent of the whole-rock 518O anomaly is advantageous for regional exploration. It has particular application in deformed terrains where the original mineral assemblages of hydrothermally altered zones has been obscured by subsequent metamorphism, because the hydrothermal whole-rock 618O patterns may still be preserved. This is because regional metamorphism typically involves low water-rock ratios. The observed whole-rock 618O values are consistent with isotopic exchange between the host rocks and large amounts of seawater (0%o, w/r >1) at different temperatures. Isotopic modelling, using a plagioclase fractionation factor as an average felsic volcanic rock value, showed that high 618O values in the diagenetic zeolite zone could be produced by interaction with fluid of virtually any source (magmatic, sea or meteoric) at low temperatures and relatively small waterrock ratios. This may be due to the high fractionation factors between silicates and water at temperatures below 100°C. However, low 818O values of the proximal sericite + chlorite zone are more consistent with conditions of equilibrium with

either meteoric waters at unrealistically low water-rock ratios (-8%o, <0.2) or seawater at 200-300°C and moderate to large water-rock ratios. The submarine volcanic environment and implications of hydrothermal mass transfers favour the latter interpretation. Cathles (1983) carried out detailed thermal, geochemical and isotopic analysis of a hypothetical, but geologically realistic, submarine intrusion-heated convective hydrothermal system. His model produced O-isotopic results that were consistent with 518O data observed by Green et al. (1983) in the Hokuroku district. It predicts that rocks in the shallow substrate become isotopically heavier by reaction with down-welling seawater at low temperatures. Lower isotopic fractionation, due to increased temperatures at depths greater than about 2 km below the seafloor, produces a zone of low rock 5 18 O. As the convective system evolves, and depending on permeability and rate of isotopic exchange, the deeplow 818O zone migrates up through the shallow-high 818O anomaly, to produce a low 518O isotopic anomaly around the vent site (Fig. 4.20). Chapter 8 summarises some other deposit- and districtscale isotopic studies, which illustrate the possible complexities in submarine volcanic successions, but indicate significant potential for whole-rock O-isotope geochemistry in targeting mineral exploration: potential that has not been widely applied outside academic studies.

FIGURE 4.20 | Modelled distribution of changes in whole-rock 618O values due to hydrothermal alteration generated by the convection of fluid around a subseafloor intrusion (after Cathles, 1983). The intrusion is 1 km wide and 3.25 km deep, and emplaced with its top 1.75 km below the seafloor. Re-equilibration with down-welling, low-temperature seawater produces a shallow zone of higher 818O. Increasing temperatures at depth create a sub-horizontal zone of low 818O, which propagates up to the seafloor resulting in the characteristic low 818O surrounded by a halo of positive anomalies. Note that the contours represent shifts from the initial rock 818O values, not the actual rock 818O values.

96

I 97

5 | SEAFLOOR- AND BURIAL-RELATED ALTERATION

This chapter discusses the alteration processes and their products (textures, minerals and zones) that occur immediately after deposition and during burial of volcanic facies in submarine environments. It encompasses the relatively lowtemperature processes of hydration, diagenesis and earJy burial metamorphism. Provided sufficient time, burial-related alteration ultimately results in the lithification of clastic facies in the succession. Oxidation, hydration, dissolution, dehydration, ion exchange, and hydrolysis reactions result in the breakdown of volcanic glass, precipitation of authigenic minerals in pore space, and replacement of glass and magmatic minerals by new minerals. Alteration mineral assemblages may change over time due to changing physical and chemical conditions during burial, and may progress to low-pressure, high-temperature regional metamorphic assemblages at depth (Coombs et al., 1959). The recognition and description of burial-related alteration styles in submarine volcanic successions has implications for exploration and ore genesis studies, because of dramatic changes in porosity and permeability, which result from cementation, compaction and dissolution during diagenesis. These changes influence subsequent fluid pathways and the sites of hydrothermal venting and mineralisation. Although it has been frequently assumed that compositional changes associated with diagenetic alteration are limited, they may involve mass changes of up to 9% (Gifkins and Allen, 2001). Diagenetic and burial metamorphic mineral assemblages and the thickness of altered zones can also be used to determine a basin's thermal history (e.g. Utada, 1991).

5.1 | ALTERATION RELATED TO SEAFLOOR PROCESSES AND BURIAL Distinctive weathering and burial-related alteration processes occur in submarine volcanic successions because of rapid accumulation rates, and the presence of abundant glass and seawater. Silicate glasses are more susceptible than minerals to alteration, because they lack well-developed crystal structures and thus will readily devitrify, dissolve or alter to minerals. Glass fragments are especially prone to alteration because of their reactivity and large surface area to volume

ratio. At elevated temperatures, volcanic glass readily alters in the presence of alkaline fluids, but the rate of alteration is reduced under dry conditions or in the presence of pure water (Lofgren, 1970). For example, hydration and devitrification rates of feJsic voicaniclastic Facies increase one to Eve orders

of magnitude in the presence of seawater (Lofgren, 1970, 1971b). Another important aspect of burial-related alteration in volcanic and igneous rocks is that anhydrous primary igneous minerals that have crystallised at high temperatures (e.g. olivine and pyroxene) become unstable and alter to hydrous minerals at lower temperatures. The extent of these retrograde reactions depends on the availability of water and the rock permeability. The effects of diagenesis and burial metamorphism on thick, proximal volcanic successions are relatively poorly understood and documented, and detailed studies are almost exclusively limited to well-sorted, fine-grained felsic voicaniclastic facies. The Ocean Drilling Program (ODP) in fore-arc and back-arc basins in the western Pacific region has provided important information on the behaviour of volcanic components during early low-temperature alteration and lithification, and the factors controlling the intensity and depth of diagenetic alteration in Miocene to Recent felsic to intermediate sandstones (e.g. Hein and Scholl, 1978; Taylor and Surdam, 1981; Klein and Lee, 1984; Hay and Guldman, 1987; Marsaglia and Tazaki, 1992; Tazaki and Fyfe, 1992; Torres et al., 1995). Studies in mafic volcanic successions have generally been limited to seafloor alteration (e.g. Bonatti, 1965; Hay and Iijima, 1968a; Honnorez, 1978; Zhou and Fyfe, 1989). Limited work in uplifted and eroded ancient submarine successions provides data on diagenetic and burial metamorphic minerals, textures and zones that formed at depths greater than 1 km (e.g. in New Zealand, Coombs, 1954; Coombs et al., 1959; in Canada, Kuniyoshi and Liou, 1976; Starkey and Frost, 1990; in Australia, Smith, 1969; Smith et al., 1982; Gifkins and Allen, 2001; Gifkins et al., in press; and in Japan, Hay and Iijima, 1968a; Seki et al., 1969; Iijima and Utada, 1972; Utada, 1991). Active geothermal regions provide direct measurements of temperatures, alteration mineral assemblages and pore water

98 | CHAPTER 5

chemistry at relatively shallow depths, less than 2 km (e.g. Coombs et al., 1959; White and Sigvaldason, 1962; Viereck et al., 1982). In addition, experimental work on the alteration of natural and synthetic glasses by modified seawater provides estimates of alteration mineralogy, temperature ranges for mineral species, fluid-rock ratios, elemental variations in glass, and variations in fluid chemistry over time. Basalt-seawater experiments were performed by: Hajash (1975, 1977), Keene et al. (1976), Seyfried and Bischoff (1977), Mottl and Seyfried (1977), Seyfried et al. (1978), Hajash and Archer (1980), Seyfried and Mottl (1982), and Ghiara et al. (1993). Rhyolite-seawater experiments were conducted by: Ellis and Mahon (1964), Sakai et al. (1978), Hajash and Chandler (1981), Shiraki et al. (1987) and Shiraki and Iiyama (1990).

Physical conditions Early studies assumed that burial-related alteration mineral assemblages and zonation patterns in submarine volcanic successions were controlled by pressure and temperature conditions. However, it is now believed that the composition and pressure of intergranular fluids and the composition of the primary facies are more important (e.g. Miyashiro and Shido, 1970; Surdam, 1973). Differences in the mineral assemblage, intensity, stratigraphic position and sequence of burial-related altered zones may be explained by variations in: primary rock composition, pore-fluid composition, pore-fluid pressure, geothermal gradient and hence temperature, burial history and sediment accumulation rate, interaction time or age, fluid-rock ratio, porosity and permeability, and tectonic setting (Hay, 1966; Surdam, 1973; Furnes, 1975; Boles and Coombs, 1977; Ratterman and Surdam, 1981; Lee and Klein, 1986; Marsaglia and Tazaki, 1992; Ghiara et al., 1993). Temperatures reached during diagenesis and burial metamorphism are directly related to the geothermal gradient and in submarine settings these range from 0°C at the seafloor to 250°C at a depth of 2-10 km (Alt and Honnorez, 1984; Morrow and Mcllreath, 1990; Alt, 1995b; Torres et al., 1995). In modern volcanic successions, measured geothermal gradients average 40°C/km, although some are as high as 200°C/km (Palmasson et al., 1979; Viereck et al., 1982). High geothermal gradients, associated with magmatism and regions of lithospheric extension such as back-arc basins and rifts, can enhance diagenetic reactions by increasing reaction rates (Boles, 1977; Surdam and Boles, 1979; Torres et al., 1995). The geothermal gradient may have varied in different parts of a geosyncline or basin; it was likely to have been lowest where the sediment was thickest and where sedimentation occurred most rapidly (Coombs et al., 1959). Taylor et al. (1990) proposed that examples of minimal diagenesis in some basins may be explained by rapid sediment accumulation rates that did not allow sufficient time for diagenetic reactions to occur at depth or for the development of porefluid gradients. In addition, magmatism provides heat to the geothermal system, locally increasing the geothermal gradient and compressing isograds near volcanic centres or large intrusions (e.g. Schiffman et al., 1984; Neuhoff et al., 1997). Coeval volcanism, plutonism and rapid burial may establish short-lived elevated geothermal gradients in many

submarine volcanic successions. The result is low-pressure, high-temperature diagenesis and metamorphism, and the suppression of some facies or zones (e.g. pumpellyiteactinolite facies, Patuki ophiolite sequence, New Zealand, Sivell, 1984).

Definitions The term spilite refers to an altered basalt or dolerite, commonly porphyritic and vesicular, in which Ca-plagioclase has been albitised and is accompanied by chlorite, calcite, epidote, prehnite or other low-temperature hydrous minerals typical of greenschist facies (e.g. Cann, 1969; Jolly and Smith, 1972; Grapes, 1976). Spilites are interpreted to result from seawater-basalt interaction during diagenesis on or near the seafloor (Coombs, 1974; Turner, 1980). Similarly, the term keratophyre, although originally restricted to lavas, has been applied to all felsic rocks that contain albite or albiteoligoclase, chlorite, epidote and calcite.

5.2 | HYDRATION Hydration of glass is typically the first stage of alteration of volcanic facies in submarine settings and occurs during lowtemperature (<50°C) seafloor weathering and the early stages of diagenesis. Hydrated glasses (e.g. perlite or palagonite) are very susceptible to alteration (Lipman, 1965). Hydration facilitates subsequent reactions as it increases the alkalinity of the pore fluid, which assists glass dissolution, promotes crystallisation, and may produce perlitic fractures, which further increase porosity and permeability (Lofgren, 1970; Friedman and Long, 1984; Noh and Boles, 1989; Casey and Bunker, 1990). Hydration involves the diffusion of water into solid glass; typically accompanied by a volume change (e.g. reaction R5.1 from Noh and Boles, 1989). As water is rapidly absorbed on to glass surfaces, hydration initially affects the outer surfaces of glassy clasts, lavas or shallow intrusions, margins along fractures in glassy facies, pillow margins, and densely welded pyroclastic deposits. This is followed by the slow diffusion of water into the glass as hydration proceeds inwards along hydration fronts defined by strain birefringence, and changes in glass colour and refractive indices (Ross and Smith, 1955; Friedman et al., 1966; Lofgren, 1971a). The rate of diffusion is dependent on composition and temperature and, hence, the extent of alteration is dependent on the time that glass has been in contact with water (O'Keefe, 1984). Most glasses will not undergo hydration to great thicknesses unless parallel reactions relax the glass structure allowing water penetration (Casey and Bunker, 1990). dacitic glass + nH 2 O -^ perlitic glass + Na+ + (OH)- (R5.1) Hydration increases the H 2 O content of glass, reorganises the glass structure and may form palagonite or silica gels. Changes in the glass structure may include volume changes, and the formation of evenly spaced tiny bubbles and perlitic fractures. Boundaries between glass and hydrated glass are

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 99

typically sharp (Peacock, 1926; Lofgren, 1971a; Fisher and Schmincke, 1984).

Palagonite Palagonite is a dull, resinous, yellow-orange to brown waxlike substance formed from hydrous altered sideromelane (basaltic) glass (Fig. 5.1). It is a mineraloid mixture of relict hydrated glass, nontronite, montmorillonite and other sheet silicates (Hay and Iijima, 1968b; Honnorez, 1969; Jakobsson and Moore, 1986). Eggleton and Keller (1982) described palagonite as a transitional alteration phase between volcanic

A. Gel-palagonite in pillow basalt The sideromelane groundmass of this plagioclase + augite-phyric basalt is altered to yellow-brown palagonite adjacent to the vesicle (V). The gel-palagonite exhibits banding parallel to the vesicle wall and perpendicular contraction cracks. Plane polarised light. Sample 153254, Miocene Waitakere Group, Muriwai, Northland region, New Zealand.

B. Palagonite-altered basalt clast rind The basalt clast in this polymictic conglomerate has a thin palagonitised rind. The plagioclase-phyric clast is concentrically zoned with an unaltered sideromelane core (C), yellow-brown gel-palagonite altered zone (P) and a brown fibro-palagonite rim (R). The conglomerate matrix includes palagonitised basaltic shards and crystal fragments. Plane polarised light. Sample 131562, Tertiary Macquarie Plains volcanics, Bushy Park, Tasmania.

C. Banded palagonite The palagonite-altered rind on this basalt clast displays fine concentric banding. Plane polarised light. Sample 131562, Tertiary Macquarie Plains volcanics, Bushy Park, Tasmania.

FIGURE 5.1 | Photomicrographs of palagonite.

glass and smectite; however, the end product may not always be smectite. There are two main varieties: gel-palagonite and fibropalagonite (Peacock, 1926). Gel-palagonite is isotropic, dark brown and commonly banded, forming directly adjacent to unaltered glass (Peacock, 1926; Zhou and Fyfe, 1989). Fibropalagonite is orange-yellow, transparent and birefringent (Zhou and Fyfe, 1989). Palagonite is widespread in submarine basaltic facies and common around the edges of glassy grains in basaltic tuffs, in pillow rinds, along fractures in glass, and in originally glassy vesicle walls (Moore, 1966; Baragar et al., 1977; Friedman and Long, 1984). Partly altered basaltic pillows typically

1 0 0 | CHAPTER 5

have glassy cores successively surrounded by concentric zones of gel-palagonite and fibro-palagonite (+ smectite), which are enhanced by bands of fine Fe- and Ti-oxides (Fig. 5.2: Dimroth and Lichtblau, 1979; Zhou and Fyfe, 1989). Palagonites have variable compositions with 10-20 wt% H 2 O (Brey and Schmincke, 1980; Eggleton and Keller, 1982; Pichler et al., 1999). Compared with sideromelane, Fe2+ is oxidised, K 2 O, FeO, TiO 2 and Cl may be locally gained, and Na 2 O, Al2O3, SiO2 and CaO lost (Baragar et al., 1977, 1979; Jakobsson and Moore, 1986; Zhou and Fyfe, 1989). However, whole-rock compositions are not significantly changed, except for H 2 O. Palagonitisation is typically accompanied by the growth of authigenic minerals in open pore spaces (Fig. 5.2) and these commonly account for the elements lost from the glass (e.g. Baragar et al., 1979; Jakobsson and Moore, 1986).

Genesis of palagonite Zhou and Fyfe (1989) and others have proposed a twostage solution-precipitation mechanism for palagonitisation of sideromelane based on physical characteristics, chemical changes and the presence of etch or dissolution pits at alteration fronts. The first stage is Ti constant: glass is dissolved and gel-palagonite formed. There is a dramatic reduction in the glass volume due to the loss of greater than 60% of the SiO2, A12O3, MgO, CaO and Na 2 O. The second stage is volume constant: gel-palagonite is replaced by fibro-palagonite, and zeolites begin to fill adjacent fractures and vesicles. CaO and Na 2 O are lost, and K2O and SiO2, Al 2 O 3 and MgO are gained from solution. Titanium and Fe3+ are localised into nearby fracture-filling clay and oxide minerals. The rate of palagonitisation is temperature dependent and doubles with every 12°C increase in temperature (Jakobsson and Moore, 1986). Palagonitisation proceeds rapidly at temperatures above 50°C and up to 150°C (Jakobsson, 1972, 1978). Jakobsson and Moore (1986) noted that palagonitisation of glass varied from less than 40% at 60°C, through 90% at 100°C and was complete at temperatures above 120°C. They also found that both gel- and fibropalagonite occurred below 87°C, but only fibro-palagonite occurred above this temperature. The thickness of palagonite rinds is time and temperature dependent. Palagonite rinds in pillow basalts systematically

increase in thickness with time and doubles for every 8°C temperature increase (Moore, 1966; Jakobsson and Moore, 1986).

Perlite Perlite is a textural term referring to networks of fine fractures or cracks that range from concentric arcuate fractures enclosing cores of glass (classical perlite; e.g. Fig. 5.3A and B) to long sub-parallel fractures linked by short cross fractures (banded or ladder perlite) (Fig. 3.2C and D: Ross and Smith, 1955; Friedman et al., 1966; Allen, 1988). Perlitic fractures are a common feature of glassy rock fragments, felsic lavas and synvolcanic sills, and also occur in the glassy rinds of mafic to intermediate lavas. Felsic perlites typically contain 2-6.5 wt% H 2 O compared with non-hydrated obsidian, which contains a few tenths of one percent (Ross and Smith, 1955; Noh and Boles, 1989). In addition to gains in H 2 O, perlites typically gain K 2 O, and lose Na 2 O and to a lesser degree CaO and SiO2 (Lipman et al., 1969; Fisher and Schmincke, 1984; Noh and Boles, 1989). Iron is oxidised, volatile components Cl2 and F2 may be lost, and 8O 18 isotope values modified by interaction with external fluids (Lipman, 1965; Jezek and Noble, 1978; Cerling et al., 1985). These compositional changes are most intense along the perlitic fractures (Jezek and Noble, 1978; Fisher and Schmincke, 1984).

Genesis of perlite A debate continues over the origin of perlite and the importance of hydration (Ross and Smith, 1955; Friedman and Smith, 1958; Friedman et al., 1966) versus cooling contraction (Marshall, 1961; Yamagishi and Goto, 1992). The formation of perlite is favoured by hydration of rapidly cooled glass (i.e. glass with a high degree of under cooling) either during cooling or later at low temperatures (Friedman etal, 1966; Noh and Boles, 1989; Drysdale, 1991). However, it is also possible that perlitic fractures form in response to strain inherited from rapid cooling contraction, during the conversion of melts to glass, and associated volume changes (Ross and Smith, 1955; Friedman et al., 1966; Davis and McPhie, 1996).

FIGURE. 5.2 | Sequence of palagonite alteration and zeolite cementation stages in phonolitic glass fragments (after Brey and Schmincke, 1980, in Fisher and Schmincke, 1984). (A) Glassy shards, perhaps with montmorillonite rim cements. (B) Hydration and development of perlitic fractures accompanied by partial dissolution and alteration of glass shards to gel-palagonite. (C) Complete dissolution and alteration of hydrated glass shards to gel-palagonite, accompanied by the precipitation of zeolites on to glass surfaces. (D) Alteration of gel-palagonite to fibro-palagonite and precipitation of zeolites into open spaces.

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 0 1

A. Perlite in thin section The glassy groundmass of this quartz latite exhibits classical perlitic fractures comprising intersecting and overlapping arcuate cracks. The perlitic fractures enclose cores of unaltered and locally oxidised glass. Arcuate glassy false shard textures occur where perlitic fractures intersect (arrow). Amygdales have been filled with zeolites. Plane polarised light. Sample ET7-4, Wereldsend Formation, Pilchard Gorge, Etendeka, Namibia.

B. Perlite in partly altered rhyolite Well-developed perlitic fractures are abundant in this partly glassy rhyolite. The perlitic fractures have been lined with fine-grained, dark green to brown smectites, enhancing the fracture pattern. Perlite cores have been partly altered to smectites and zeolites. Amygdales have been filled with cristobalite. Plane polarised light. Sample 147582, Miocene Nishikurosawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

C. Relict perlite and amygdales in altered rhyolite In this diagenetically altered rhyolite, relict perlitic fractures are conspicuous where glass adjacent to the fractures has been altered to dark green mixed layer smectite-chlorite. Elsewhere in the pervasively zeolite altered domains the perlitic fractures have been obscured. The amygdales have been filled with layers of cristobalite and fibrous chlorite. Plane polarised light. Sample J6-735 m, Miocene Nishikurosawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

D. Relict perlite in altered basalt In this hydro thermally altered jigsaw-fit basaltic breccia, perlitic fractures are only weakly discernable due to multiple overprinting alteration facies. The pervasive sericite + quartz + pyrite and nodular carbonate alteration facies obscure the perlitic fracture pattern. Plane polarised light. Sample 76833, Cambrian Que-Hellyer Volcanics, western volcanosedimentary sequences, Mount Read Volcanics, western Tasmania.

FIGURE 5.3 | Photomicrographs of fresh and altered perlite.

1 0 2 | CHAPTER 5

Alteration of perlite Perlite commonly undergoes subsequent alteration to diagenetic mineral assemblages that include smectite, Feoxides, zeolites, K-rich gel-like glass, low-cristobalite, Kfeldspar, chlorite, sericite and carbonate (Noh and Boles, 1989). Alteration begins by dissolution of hydrated glass and crystallisation of smectite, carbonate or Fe-oxides along perlitic fractures (e.g. Noh and Boles, 1989). This commonly accentuates the fracture pattern (e.g. Fig. 5.3B). As alteration progresses, glass dissolution with continued precipitation advances inwards and the perlitic fractures become diffuse and indistinct (e.g. Fig. 5.3C, D and Allen, 1988). Dissolution of remaining glassy cores is succeeded by formation of zeolites, such as clinoptilolite or mordenite, or gel-like glass, which are ultimately replaced by K-feldspar (e.g. Noh and Boles, 1989).

5.3 | DIAGENESIS (GLASS TO ZEOLITE FACIES) Diagenesis encompasses the low-temperature and lowpressure alteration processes that occur during progressive burial of sediments and rocks. It can be defined as the processes (excluding weathering) that change their character and composition, between the moment of deposition, and the onset of metamorphism (Larsen and Chilingar, 1979). Submarine diagenesis involves low-temperature processes, ranging from bottom water temperatures up to crystallisation of unequivocally metamorphic minerals such as laumontite, wairakite, chlorite and pumpellyite (Winkler, 1979; Bohlke et al., 1980). It is impossible to define a unique pressure and temperature range that would characterise the transition between diagenesis and metamorphism, because of the greatly contrasting degrees of mineral stability that characterise different rock types and the wide range of conditions under which the common diagenetic minerals crystallise. Generally, diagenesis in submarine settings occurs at pressures of 0.1 to 10 MPa (1 bar to 1 kbar) and temperatures ranging from 0 to 250°C (Alt and Honnorez, 1984; Morrow and Mcllreath, 1990; Alt, 1995b). Temperatures and pore water salinities increase, and seawater-rock ratios decrease with burial depth (Hanor, 1979; Alt-Epping and Smith, 1997). Submarine diagenesis encompasses compaction, dissolution and leaching of components, precipitation of new minerals, and recrystallisation in response to changes in pressure, temperature and chemical conditions in the subseafloor. New minerals directly replace glass, form mineral overgrowths, fill primary and secondary pore spaces, and form cements, all of which dramatically reduce the porosity and permeability and promote lithification. With increasing diagenesis, porosity and permeability typically decrease. However, reversals in this trend can occur during fracturing or if a major component of the rock becomes under saturated and secondary porosity is formed by dissolution. This can occur where deeply buried sediments are infiltrated by fresh or brackish ground water, or can be due to the release of water of crystallisation from clay minerals (Morrow and Mcllreath, 1990).

The process of dissolution involves corrosion or leaching of pre-existing phases (either glass or mineral phases), with or without minor replacement by new minerals (Morrow and Mcllreath, 1990). It is a complex process involving many distinct reaction steps and pathways. It can modify glass and most primary igneous minerals. Dissolution may ultimately lead to the formation of secondary porosity (e.g. dissolution vugs), replacement of glass and minerals, and development of solution seams or stylolites (Amstutz and Park, 1967; Marsaglia and Tazaki, 1992). Despite changes in mineral assemblage, many pre-existing textures (primary volcanic, high-temperature devitrification and hydration textures) are preserved and sometimes enhanced during diagenesis. Figure 5.4 shows some examples of textures in unaltered volcanic rocks, and their diagenetically altered and in some cases metamorphosed equivalents. Submarine diagenesis may involve multiple stages or episodes of diagenesis (Bohlke et al., 1980; Morrow and Mcllreath, 1990). Diagenesis of most ancient sedimentary successions involved repeated exposure to diagenetic realms as they underwent cycles of subsidence and uplift. Generally, however, the imprint of the first stages of diagenesis is preserved because of the large initial porosity reduction and lithification (Morrow and Mcllreath, 1990).

Diagenetic minerals There are three main types of minerals typical of seafloor weathering and diagenesis in volcanic successions: layered silicates, zeolites and carbonates. Figure 5.5 provides estimates of their formation temperatures.

Layered silicates The layered silicates include clay minerals, mixed-layered minerals, micas, chlorite and prehnite. The common clay minerals in volcanic facies can be divided in to two groups: (1) smectites (e.g. montmorillonite, nontronite andsaponite), and (2) illite group clay minerals (e.g. celadonite, glauconite and illite). Smectites are swelling clay minerals that readily exchange Ca and Na cations. They typically result from the alteration of volcanic grains under alkaline conditions where Mg and Ca ions are available (Deer et al., 1966). Smectites form rims on glass surfaces, replace both felsic and mafic glass, and pseudomorph glass shards and olivine crystals (Sheppard and Gude, 1968; Schmincke and von Rad, 1976; Viereck et al., 1982). Smectites initially forms blebs and web-like arrays on glass surfaces, and become better crystallised as diagenesis proceeds (Hein and Scholl, 1978). The term bentonite refers to felsic tuff that is composed of almost pure smectite (Gary etal., 1974). In contrast, the illite group are K- and Al-rich minerals that typically form in neutral to alkaline conditions from the breakdown of feldspars and micas (Deer et al., 1966). They typically occur as vesicle fill and pseudomorphs of felsic glass shards and pumice (Schmincke and von Rad, 1976; Iijima, 1978). Celadonite and glauconite are less common than illite.

SEAFLOOR-AND BURIAL-RELATED ALTERATION I 1 0 3

A. Flow banding This devitrified flow-banded plagioclase-phyric rhyolite contains alternating dark and light flow bands. The dark bands are dominantly obsidian, whereas the pale bands contain fine spherulites and lithophysae. Sample NG1, < 140 ka Ngongotaha lava dome, Hendersons quarry, Rotorua, New Zealand.

B. This diagenetically altered and metamorphosed flowbanded plagioclase-phyric rhyolite contains alternating orange albite + quartz and grey sericite-rich bands. In thin section, the orange bands contain relict spherulites, whereas the grey bands are microcrystalline. Sample 147481, Cambrian Central Volcanic Complex, Mount Read Volcanics, Mount Block, western Tasmania.

C. Spherulites In thin section, fresh spherulites consist of radial crystal fibres; typically feldspar intergrown with cristobalite, tridymite or clinopyroxene. Many of these spherulites enclose plagioclase phenocrysts and are separated by small cuspate lenses of dark brown obsidian. Plane polarised light. Sample NG4, <14O ka Ngongotaha lava dome, Hendersons quarry, Rotorua, New Zealand.

D. Recrystallised spherulites in this greenschist facies rhyolite are composed of albite, quartz and sericite. Fine sericite trails preserve a radial pattern within the spherulites. The boundaries between the spherulites are marked by concentrations of sericite. Plane polarised light. Sample 147528, Cambrian Central Volcanic Complex, Mount Read Volcanics, Mount Black, western Tasmania.

E. Tube pumice clasts This unaltered, semi-consolidated, dacitic pumice breccia contains glassy tube pumice clasts and plagioclase crystals in a matrix of fine glass shards. The pumice \ clast pictured here displays a fine fibrous texture, which may be preserved during subsequent alteration. Plane polarised light. Sample from the -1 Ma trachydacitic pumice breccias, Efate Pumice Tormation, Vanuatu.

FIGURE 5.4 | Photographs of unaltered and diagenetically altered volcanic textures.

1 0 4 | CHAPTER 5

F. Pumice clasts in this diagenetically-altered and metamorphosed rhyolitic pumice breccia preserve the fine tube vesicle structure. The originally glassy vesicle walls have been altered to albite + quartz + hematite, the vesicles have been lined with sericite and filled with albite. Plagioclase crystals in this sample have been completely replaced by albite and hematite. Plane polarised light. Sample 133815, Cambrian Hercules Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Hercules footwall, western Tasmania.

G. Many tube pumice clasts locally preserve round vesicles adjacent to phenocrysts. In this diagenetically altered pumice breccia, round and tube vesicles adjacent to a cluster of plagioclase phenocrysts have been filled with mordenite. As a result, the vesicles have retained their shapes during burial compaction. Plane polarised light. Sample OH8-369 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

H. Similarly, this pumice breccia, which has been diagenetically altered and metamorphosed to greenschist facies, contains round and tube vesicles adjacent to hematite-altered plagioclase phenocrysts. The vesicles (V) have beeen filled with sericite and albite. Plane polarised light. Sample 147499, Cambrian Kershaw Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, east Hercules, western Tasmania.

I. Palagonitised rinds on clasts The rim of this basalt clast has been altered to orangebrown palagonite. Palagonite has also formed rims around the vesicles in the clast. Plane polarised light. Sample 131562, Tertiary Macquarie Plains volcanics, Bushy Park, Tasmania.

J. The sericite + albite + hematite-altered rim (R) on this basalt clast may be the metamorphosed equivalent of a palagonite-altered rind. Plane polarised light. Sample 147572, Cambrian Sterling Valley Volcanics, Central Volcanic Complex, Mount Read Volcanics, Sterling Valley, western Tasmania.

FIGURE 5.4 | Photographs of unaltered and diagenetically altered volcanic textures, cont.

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 0 5

Carbonates Diagenetic carbonates are dominantly calcite and dolomite. They typically fill originally open spaces such as vesicles, occur as cements in volcaniclastic facies (e.g. Hay, 1977), as spheroids or nodules, and as euhedral crystals replacing palagonite (e.g. Dimroth and Lichtblau, 1979), rock fragments, olivine and plagioclase crystals.

Other diagenetic minerals Other diagenetic minerals include silica phases (e.g. lowcristobalite, opal CT, chert and quartz), Fe-oxides (e.g. hematite), Ti-rich minerals (e.g. leucoxene), anhydrite, pyrite, epidote and feldspars (albite and K-feldspar). These mainly replace glass, primary crystal phases and earlier alteration minerals. Silica phases and feldspars also occur as overgrowths on primary plagioclase and quartz crystals (e.g. Noh and Boles, 1989; Tsolis-Katagas and Katagas, 1989).

Diagenetic zones Diagenetic mineral assemblages commonly show a thick vertical zonation (e.g. Fig. 5.6 and Section 5.5). Diagenetic zones have been described by a number of authors in modern and ancient submarine felsic to intermediate volcanic successions (e.g. Iijima, 1974; Walton, 1975; Iijima, 1978; Ratterman and Surdam, 1981; Sheppard et al., 1988; Williams et al., 1989; Utada, 1991; Passaglia et al., 1995; Ogihara, 1996). Sequences of diagenetic zones are between 500 m and 6 km thick, with individual altered zones varying from a few metres to several kilometres in thickness. This vertical zonation corresponds to progressive mineral reactions that occur in response to changes in pore water chemistry and temperature with depth of burial, and is very similar to burial metamorphism (Coombs, 1954). Some altered zones may be absent or combined. FIGURE 5.5 | Temperature estimates for the growth of common diagenetic and burial metamorphic minerals, and palagonite (data from Thompson, 1971; Seki, 1972; Merino, 1975; Grapes, 1976; Kastnerand Gieskes, 1976; Seyfried and Bischoff, 1979; Bohlke et al., 1980; Munha et al., 1980; Boles, 1982; Viereck et al., 1982; Jakobsson and Moore, 1986; Bish and Aronson, 1993; Ogihara, 1996; Ylagan et al., 1996; Bodon and Cooke, 1998).

Zeolites

Zeolites are hydrous Al-silicates containing Na and Ca (Table 5.1). The most common zeolites in marine settings are clinoptilolite, mordenite, phillipsite and analcime (Marsaglia and Tazaki, 1992). A variety of fibro-radiated and bladed zeolites fill pore spaces, cement volcaniclastic particles and replace glass in altered volcanic facies (Miyashiro and Shido, 1970; Schmincke and von Rad, 1976). Most zeolites precipitate in open space on to smectite or chlorite films or occur as overgrowths on detrital grains such as plagioclase crystal fragments (e.g. Schmincke and von Rad, 1976). Others crystallise directly from glass via dissolution reactions with smectite (e.g. Noh and Boles, 1989) and may pseudomorph glass shards (e.g. Walton, 1975).

Diagenetic zones in felsic volcanic successions Diagenetic zones in felsic volcanic successions can be grouped into four main zones (Table 5.2): (I) partially altered zones, (II) alkali-rich zeolite zones, (III) late-stage zeolite + calcite zones, and (IV) albite zones. At depth Zone IV may pass in to a prehnite + pumpellyite zone, which represents the transition to greenschist facies metamorphic zones (Iijima, 1974, 1978; Utada, 1991). Partially altered zones are characterised by silica and clay minerals, they lack zeolites, contain unaltered and partly altered glass, and unaltered primary minerals such as plagioclase (Iijima, 1974, 1978). Alteration mineral assemblages are dominated by smectites (commonly montmorillonite) + lowcristobalite or opal-CT (Iijima, 1974, 1978; Walton, 1975; Sheppard et al., 1988; Passaglia et al., 1995). Primary pore spaces, such as vesicles, have typically been partially filled with low-cristobalite, glassy clasts have been coated in thin films of smectite, and some originally glassy shards and pumice clasts altered to smectite. Coherent facies were relatively unaltered.

1 0 6 | CHAPTER 5 TABLE 5.1 | Common zeolites and their occurrences in submarine volcanic facies. Zeolite formulas are from Deer et al. (1966).

Analcime

Na[AISi2O6].H2O

A Na-rich late stage zeolite, which replaces earlier alkali zeolites in both coherent and clastic volcanic facies of rhyolitic to basaltic composition (e.g. lijima, 1974; Ratterman and Surdam, 1981; Torres et al., 1995)

Chabazite

Ca[AI2Si4012].6H20 -

Restricted to mafic facies, typically replacing palagonite (e.g. Brey and Schmincke, 1980; Dimroth

thomsonite

NaCa2[(AI,Si)5O10]2.6H2O

and Lichtblau, 1979)

Clinoptilolite

(Na,K)4CaAI6Si30O72.H2O

Occurs as a cement and replaces glass in felsic volcanic facies (e.g. Noh and Boles, 1989; Ratterman and Surdam, 1981; Torres et al., 1995)

Heulandite

(Ca,Na2)[AI2Si7O18].6H2O

Occurs as cements in felsic volcaniclastic facies (e.g. Ratterman and Surdam, 1981)

Laumontite

Ca[AI2Si4012].4H20

A calcic zeolite, which occurs at depth in originally glassy felsic volcanic facies

Mordenite

(Ca, Na2,K2)[AI2Si10O24].7H2O

Only derived from felsic volcanic facies and commonly coexists with smectite and silica phases (i.e. opal, quartz, tridymite and cristobalite) (Noh and Boles, 1989; Ratterman and Surdam, 1981; Sheppard et al., 1988; Sheppard and Gude, 1968; Torres et al, 1995; Tsolis-Katagas and Katagas, 1989; Utada, 1970)

Phillipsite

(1/2Ca,Na,K)3[AI3Si5O16].6H2O

Occurs mainly in basaltic lavas and less commonly in volcaniclastic facies where it replaces basaltic glass and palagonite (Taylor and Surdam, 1981), it commonly contains inclusions of Feoxyhydroxides and smectites (Brey and Schmincke, 1980)

Wairakite

CaAI2Si4012.H20

A common alteration product in basaltic facies at depth in modern geothermal systems (e.g. Boles, 1977; Hay, 1977)

Table 5.2 | Common diagenetic zones and their alteration mineral assemblages for thick submarine volcanic successions.

Zone I: partially altered zone

Zone I: partially altered zone

unaltered glass + smectite (montmorillonite) + low-cristobalite/opal-CT

unaltered glass + palagonite + smectite + illite + low-cristobalite/ adularia + Fe/Mn/Ti oxides + unaltered glass

Zone II: alkali-rich zeolite zone

Zone II: calcic-zeolite zone

(a) clinoptilolite + smectite (montmorillonite) + low-cristobalite/opal-CT

phillipsite/chabzite + phyllosilicate minerals (chlorite, smectite, sericite) + Fe/Mn/Ti-oxides ± K-feidspar

(b) Ca-clinoptilolite + mordenite + smectite + K-feldspar ± quartz Zone III: late stage zeolite + calcite zone

Zone III: late-stage zeolite zone

(a) analcime + heulandite + clacite + phyllosilicate minerals (smectite, chlorite, mixed layer minerals) + K-feldspar ± quartz ± pyrite

analcime± natrolite (± heulandite ± laumontite) + chlorite + Kfeldspar + Fe/Mn/Ti-oxides ± calcite

(b) analcime + laumontite + clacite + phyllosilicate minerals (illite, chlorite, smectite) + K-feldspar + quartz Zone IV: albite zone

Zone IV: epidote zone

albite + phyllosilicate minerals (prehenite, pumpellyite, chlorite, sericite) + quartz ± K-feldspar ± laumontite ± calcite

epidote + chlorite + albite + calcite + sphene ± prehnite

SEAFLOOR-AND BURIAL-RELATED ALTERATION I 1 0 7

FIGURE. 5.6 | East-west schematic cross-sections showing the depth distribution of regional diagenetic zones and local hydrothermal zones associated with the Kuroko deposits in the Green Tuff Belt, Japan (after lijima, 1974,1978). (A) Odate Basin, Hokoroku district. (B) Odate to Hanawa Basin, Hokoroku district. (C) Diagenetic zones in the Neogene and Palaeogene formations of Hokkaido. Traces of cross-sections A and B are shown on the regional map of the Hokuroku Basin (Fig. 5.15).

1 0 8 | CHAPTER 5

These early clay-rich zones are associated with minor initial gains in K2O and Al2O3, and losses in Na 2 O and, to lesser degrees, CaO and SiO2 (Noh and Boles, 1989) Alkali-rich zeolite zones are commonly characterised by assemblages of clinoptilolite + mordenite + smectite (typically montmorillonite, saponite or mixed-layer illite/smectite) + low-cristobalite ± quartz ± opal ± K-feldspar. Opal-CT and low-cristobalite occur in the upper parts of these zones, whereas quartz and K-feldspar occur in the lower parts (Walton, 1975; Sheppard et al., 1988). Plagioclase is rarely altered. In general, clinoptilolite and mordenite has filled pore spaces such as primary vesicles and dissolution voids. Glassy clasts have been coated in montmorillonite or silica rims and completely altered to zeolites (e.g. mordenite), low-cristobalite, quartz, clay minerals and K-feldspar (Iijima and Utada, 1971; Iijima, 1974, 1978). The glassy groundmass of lavas and sills, and the cores of blocky clasts may have been partly altered. Pumice-rich facies in these zones contain dark green, variably flattened phyllosilicate-rich fiamme (e.g. saponite fiamme in pumice breccia in the Hokuroku Basin, Iijima, 1974). Tuffaceous mudstones may be rich in montmorillonite, low cristobalite and quartz. Alteration to alkali-zeolites and phyllosilicate minerals resulted in whole-rock gains of MgO and Fe2O3, and losses of SiO2, Na 2 O, K2O, and variable changes in CaO (Noh and Boles, 1989; Tsolis-Katagas and Katagas, 1989; Passaglia et al., 1995). Late-stage zeolite + calcite zones are characterised by mineral assemblages containing analcime and calcite (Iijima, 1974). In some successions, two late-stage zeolite + calcite zones have been defined (e.g. Iijima, 1978): (a) analcime + heulandite + calcite ± phyllosilicate minerals ± K-feldspar ± quartz ± pyrite, and (b) analcime + laumontite + calcite ± chlorite ± illite ± sericite ± K-feldspar. The phyllosilicate minerals are typically smectites, chlorite and mixed-layer minerals such as illite/smectite, saponite/chlorite or swelling chlorite. These mineral assemblages may also contain relict clinoptilolite and/or mordenite (Iijima, 1974; Walton, 1975; Sheppard et al., 1988). Plagioclase phenocrysts have remained unaltered or have been analcime ± calcite altered. Analcime has replaced mordenite- or clinoptilolite-altered felsic glass fragments and pumice clasts. Saponite, smectite, chlorite and mixedlayer mineral fiamme are typically common in pumice-rich rocks. Calcite may occur as euhedral crystals, concretions or veinlets. Albite zones are commonly characterised by albite + laumontite ± calcite + prehnite ± chlorite ± sericite ± pumpellyite ± quartz ± K-feldspar (Iijima and Utada, 1971; Iijima, 1974). Plagioclase phenocrysts have been extensively albitised, and albite + laumontite have replaced plagioclase crystals, originally glassy shards and pumice clasts, and filled pore spaces. Mass gains in CaO, SiO2, Na 2 O, Sr and Ba in these zones are consistent with seafloor albitisation (Boles and Coombs, 1977; Boles, 1982).

Diagenetic zones in mafic volcanic successions Diagenetic mineral assemblages in mafic volcanic successions contain palagonite, several species of calcic zeolites, Fe/Ti/Mnoxides and abundant clay minerals of the smectite-chlorite series typically distributed in four zones (Table 5.2: Baragar et al., 1979; Zhou and Fyfe, 1989; Utada, 1991). Partially altered zones contain some fresh basaltic glass and have alteration mineral assemblages of palagonite ± Fe/ Mn/Ti-oxides (e.g. maghemite or magnetite) + clay minerals (smectites, illites and mixed-layer minerals) ± low-cristobalite. Compositional changes include major gains of H 2 O, very minor gains of K 2 O, FeO, TiO 2 and Cl and losses of Na 2 O, A12O3, SiO2 and CaO (Baragar et al., 1977, 1979; Jakobsson and Moore, 1986; Zhou and Fyfe, 1989). Calcic-zeolite zones are characterised by phillipsite or chabazite ± chlorite + smectite + Fe/Mn/Ti-oxides + Kfeldspar. Whole-rock gains in MgO in these zones are consistent with the formation of smectite, chlorite and other Mgsilicates during diagenesis (cf. Hajash and Chandler, 1981; Shiraki and Iiyama, 1990). Late-stage zeolite zones are characterised by analcime ± laumontite ± natrolite ± chabazite + heulandite ± mesolite + chlorite + Fe/Mn/Ti-oxides + K-feldspar. Epidote zones are characterised by epidote + chlorite + albite + sphene ± calcite ± prehnite.

Genesis of diagenetic minerals and zones In submarine volcanic facies, dissolution, cementation and lithification begin shortly after deposition (<1 Ma); however, major diagenetic changes develop through a series of recognisable stages over tens of millions of years (Marsaglia and Tazaki, 1992). The paragenesis from glass to smectites to alkali zeolites may be explained by a sequence of hydration and dissolution reactions in most glass bearing rocks (Fig. 5.7). Later reactions involve transitions from less stable to more stable mineral assemblages, such as clinoptilolite to Na-clinoptilolite + mordenite or K-rich gel-like glass to Kfeldspar (Noh and Boles, 1989). There are four stages of clastic diagenesis after initial hydration and oxidation (e.g. Fig. 5.8): (1) formation of clay mineral rims on glassy surfaces, (2) partial to complete dissolution of glass and compaction, (3) precipitation of authigenic minerals, especially zeolites and calcite, in open pore spaces, and (4) alteration and replacement of mineral phases (Hay, 1963; Fisher and Schmincke, 1984; Pichler et al., 1999). Stages two and three may overlap.

Stage 1: coating surfaces The initial stage of diagenesis in volcaniclastic facies is characterised by the precipitation of thin rim cements, which coat all originally glassy surfaces and some crystal surfaces (Fig. 5.9A and B). Rim cements help to preserve shard and clast outlines during subsequent replacement (e.g. Walton, 1975). Rim cements may be accompanied by dissolution of intermediate to mafic glass, alteration of rhyolitic glass to clay

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 0 9

minerals and minor precipitation of calcite or clinoptilolite cements (Fig. 5.9C and D: Marsaglia and Tazaki, 1992; Torres et al., 1995). In felsic volcaniclastic facies, the outer walls of most glass shards and vesicles in pumice clasts are lined with thin films of smectite, calcite, opal or rarely chlorite (e.g. Henneberger and Browne, 1988; Sheppard et al., 1988; Noh and Boles, 1989; Tsolis-Katagas and Katagas, 1989; Marsaglia and Tazaki, 1992; Torres et al., 1995). Only rarely are fine-grained, glassy fragments such as shards and pumice completely replaced by smectite. Smectite rims probably precipitate from alkaline fluids during the dissolution of hydrated glass surfaces. This may follow the reactions:

and high activities of Mg (Hay, 1978; Hajash and Chandler, 1981). The initial stage of diagenesis in basalts involves the palagonitisation of basaltic glass. Palagonitisation initiates on glass surfaces, along fractures and around vesicles in a similar way to the thin smectite coating on rhyolitic glass shards (e.g. Zhou and Fyfe, 1989). Associated with palagonitisation is the precipitation of Fe- and Ti-oxides, which coat all surfaces, thermal contraction cracks, vesicles, and the palagonitisation front (e.g. Dimroth and Lichtblau, 1979).

perlitic glass + 3.88K+ + 0.65H+ + 15.4H2O -» smectite + 9.5gel-like glass + 4.03Na+ + 0.25Ca2+ + 10.55H4SiO4 (Noh and Boles, 1989) (R5.2)

Large-scale dissolution of glass and crystals is typically accompanied and followed by the precipitation of authigenic mineral cements and lithification after a few million years (Marsaglia and Tazaki, 1992; Torres et al., 1995). The dissolution of glass fragments, olivine and amphiboles occurs rapidly at shallow burial depths prior to extensive cementation and lithification (Smith, 1991; Marsaglia and Tazaki, 1992). With increasing depths of burial, dissolution of feldspar microlites and the glassy groundmasses of coherent facies occurs (Marsaglia and Tazaki, 1992). In contrast, plagioclase crystal fragments and phenocrysts undergo only minor dissolution early in the diagenetic history. Elements leached from the glass during dissolution reactions are consumed by the formation of new minerals.

or rhyolitic glass + Mg2+ + H 2 O Na-Ca montmorillonite + SiO2 + Na+ + K+ + Fe2+ (R5.3) The formation of smectite and other Mg-silicates during rhyolite-seawater interaction does not require significant gains in alkalis, but is favoured by high ratios of H/Na and K/Ca,

Stage 2: dissolution of glass and compaction

FIGURE. 5.7 | Flow diagrams showing the successive development of alteration mineral assemblages in volcanic glass during diagenesis. (A) Alteration of silicic glass to day minerals, zeolites and silicates (after Hay, 1978; lijima, 1978; Utada, 1991). (B) Alteration of basaltic glass to palagonite, clay minerals, zeolites and oxides (Honnorez, 1978; lijima, 1978; after Brey and Schmincke, 1980; Viereck et al., 1982; Fisher and Schmincke, 1984).

FIGURE 5.8 | Schematic model for the microscopic textural evolution and reduction in porosity in non-welded pumice breccias during diagenesis (after Gifkins, 2001). (A) Stage 1: thin films of smectite (green lines) coat original surfaces, such as vesicle walls, crystals, shards and lithic clasts. (B) Stage 2: primary porosity is filled, and originally glassy shards and vesicle walls replaced or partly replaced by zeolites (bronze), clay minerals (green) or carbonates. Zeolite or K-feldspar overgrowths (orange) may develop on plagioclase crystals. (C) Stage 3: glass is dissolved, altered to clay minerals and compacted, producing phyllosilicate-rich fiamme. Clays, zeolites and Fe-oxides precipitate synchronously with the dissolution of glass, forming stylolites. After compaction, more stable diagenetic or metamorphic minerals replace any remaining glass and less stable minerals (i.e. Stage 4).

For example, Na released by the hydration and dissolution of rhyolitic glass (reactions 5.2 and 5.3) may be consumed by the precipitation of mordenite in vesicles (e.g. Fig.5.10E) or dissolution vugs (Sheppard et al., 1988). Dissolution may be accompanied by compaction that reduces the pore space geometry by rotating grains, deforming soft grains and crushing grains. This promotes lithification in volcaniclastic facies by pressure welding clasts so that their margins interpenetrate (Taylor et al., 1990; Marsaglia and Tazaki, 1992). Generally, lithification of volcaniclastic facies is associated with the formation of diagenetic clay mineral or carbonate rims, which act as cohesive binders: cements. However, Marsaglia and Tazaki (1992) in their study of modern partially altered sandstones at ODP Site 788 (Japan) suggested that lithification could be related to a combination of compaction and brown glass dissolution. Sandstones in a transitional zone, between the unlithified and cemented zones, appeared to be lithified as a result of compaction and/ or pressure welding with minor cementation by phillipsite and smectite/chlorite rim cements. The rim cements formed where sufficient glass had dissolved to produce favourable conditions for smectite or zeolite precipitation. Compaction may also bend and flatten clasts, particularly pumice clasts. Gifkins etal. (in press) suggested that compaction during burial of clay-altered pumice clasts flattened the clasts and modified tube vesicle structures resulting in beddingparallel phyllosilicate lenses (i.e. fiamme, Fig. 5.9G). They also proposed that stylolites in pumice breccias in the Mount Read Volcanics (western Tasmania) and the Green Tuff Belt (Japan) resulted from the dissolution of soluble components, particularly glass, and the precipitation of clays and Fe-oxides during compaction (Fig. 5.9H).

Stage 3: filling pore space and cementation Precipitation of low-cristobalite and zeolites as pore-fill cements follows the early rim cements in both felsic and mafic volcanic facies (e.g. Klein and Lee, 1984; Zhou and Fyfe, 1989). Zeolites also fill vesicles and dissolution voids in glass, and directly replace glass, forming shard pseudomorphs or altering the glassy cores of perlite (Dimroth and Lichtblau, 1979; Noh and Boles, 1989; Tsolis-Katagas and Katagas, 1989; Passaglia et al., 1995). Many of the zeolites filling vesicles have fibrous radiating textures (e.g. mordenite, Fig. 5.10A and B), which may be partly preserved during subsequent metamorphic recrystallisation (e.g. Fig. 5.IOC and D). Dimroth and Lichtblau (1979) described fibroradial textures defined by a dusting of fine oxides in Archaean basaltic hyaloclastite of the Noranda District (Canada), which suggest the former presence of fibro-palagonite or zeolites. The formation of alkali-rich zeolites as pore-fill cements involves hydration and dissolution of glass by saline, alkaline solutions (Ratterman and Surdam, 1981; Noh and Boles, 1989). For example, the formation of clinoptilolite from hydrated felsic glass in reaction R5-4, which consumes Ca and Si released during reaction R5.1: perlitic glass + 0.1Ca2+ + 0.1H4SiO4 + 0.1 H+ + H 2 O -» clinoptilolite + 0.1K+ + 0.2Na+ (R5.4) The formation of rim and pore-fill cements results in a lithified rock. Cements reduce the porosity, strengthen the grain framework and reduce the amount of compaction. The initial 35-40% porosity of a well-sorted sandstone can be reduced to 15-20% by early clay mineral, carbonate or zeolite cements (Helmold and van de Kamp, 1984). In the clinoptilolite + mordenite zone, Henneberger and Browne (1988) found the porosity of pumice breccias was reduced by half, from 34-47% to 20-50%. Alteration in the quartz + adularia zone further reduced porosity to 4—23%.

SEAFLOOR- AND BURIAL-RELATED ALTERATION I 1 1 1

A. Clay rim cement in pumice breccia Smectite films on all glass and crystal surfaces record the initial stage of diagenesis in this partially altered rhyolitic pumice breccia. Green-brown smectite has coated bubble-wall shards, plagioclase and quartz crystal fragments, and lined vesicles. Some originally glassy shards have been completely replaced by smectite; however, larger clasts are still glassy (G). Sample J6-295 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

B. Clay-lined vesicles Round and elongate vesicles (V) in this pumice clast are coated in irregular, fine-grained, pale brown smectite films. Small vesicles have been completely filled wih smectite, larger vesicles are unfilled, and the originally glassy vesicle walls have been altered to mordenite. Sample FK5B, Miocene Tokiwa Formation, South Fossa Magna, Green Tuff Belt, Odawara, Japan.

C. Pore-filling cements In this pumice breccia sample, calcite cement binds the unaltered glassy and partly calcite-altered tube pumice clasts. Plane polarised light. Sample Y2A, Quaternary Yali pumice breccia, Yali Island, eastern Aegean, Greece.

D. In crossed nicols, the glassy pumice clasts are isotropic and the calcite cement, calcite-filled tube vesicles, and altered shards and pumice clasts are evident.

FIGURE 5.9 | Examples of textures that record the different steps in the evolution of pumice clasts during diagenesis.

1 1 2 I CHAPTER 5

E. Zeolite-filled vesicles in pumice The smectite-lined vesicles (V) in this pumice clast have been infilled with layered fibrous zeolites: mordenite and clinoptilolite. Originally glassy shards have been altered to smectite and vesicle walls to mordenite + smectite. Fine-grained nodules of analcime overprinted the mordenite and smectite altered tube pumice clast (P). Plagioclase and quartz crystals are unaltered. Sample OH8-537 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

F. Clay-altered pumice clast Other pumice ciasts may be completely altered to clay minerals, like this dark green uncompacted smectitealtered pumice. Shards and fine-grained ciasts in the matrix have been altered to smectites (montmorillonite and saponite) + mordenite. Sample OH8-387 m, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

G. Clay-altered and compacted pumice During burial, lithostatic pressure may lead to the flattening of soft clay-altered pumice ciasts. The mixed layer smectite-chlorite fiamme (F) in this pumice and lithic breccia roughly define a bedding-parallel compaction fabric. The fiamme have a fibrous internal texture, wispy terminations and flame-like shapes. Some fiamme are also plagioclase porphyritic. They are interpreted to be diagenetically altered and compacted pumice ciasts. Sample 147583, Miocene Nishikurosawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

H. Dissolution fabrics in pumice breccia The dissolution of glass commonly accompanies compaction during diagenesis. Solution seams and stylolites, like the one pictured here, are interpreted to record the dissolution of soluble components. This stylolite is an anastomosing sutured structure that concentrates clay minerals and oxides. Sample FK7, Miocene Wadaira Tuff Member, Tokiwa Formation, South Fossa Magna, Green Tuff Belt, Wadaira, Japan.

FIGURE 5.9 | Examples of textures that record the different steps in the evolution of pumice ciasts during diagenesis, cont.

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 1 3

A. Fibrous zeolites in vesicles B. In crossed nicols, radial extinction patterns accentuate the Vesicles adjacent to this plagioclase crystal in a pumice clast fibrous nature of the vesicle-filling zeolites. have been lined with smectite and filled with fibrous radiating mordenite. Plane polarised light. Sample 147580, Miocene Onnagawa Formation, Hokuroku Basin, Green Tuff Belt, Odate, Japan.

The vesicles (V) in this altered pumice clast are faintly visible in plane polarised light because they are lined with sericite. Sample 133815, Cambrian Hercules Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Hercules footwall, western Tasmania.

albite-filled vesicles mimic pre-existing fibrous textures.

E. Fibrous feldspar in perlite cores In plane polarised light, perlitic fractures are conspicuous in the groundmass of this altered plagioclase-phyric rhyolite. Sample 147541, Cambrian Kershaw Pumice Formation, Central Volcanic Complex, Mount Read Volcanics, Murchison Highway, western Tasmania.

F. In crossed nicols, overlapping arcuate perlitic fractures are defined by concentrations of sericite and radial fibrous textures are preserved in the extinction pattern of the albite + quartz + sericite-altered perlite cores (C).

FIGURE 5.10 | Photomicrographs of relict fibrous textures in vesicles and originally glassy domains in diagenetically altered fades.

1 1 4 I CHAPTER 5

Stage 4: alteration and replacement of mineral phases

dacitic glass + nH 2 O —> perlitic glass + Na+ + OH~

Later reactions involve the dissolution and replacement of earlier diagenetic phases, remaining gel-like glass, and magmatic minerals such as plagioclase (Torres et al., 1995). With time and/or increasing temperature, zeolite assemblages are especially susceptible to replacement because zeolite crystallisation is controlled by temperature, fluid pressure, and rock and fluid composition. Transitions from unstable zeolites (e.g. phillipsite, clinoptilolite and heulandite) to more stable phases (e.g. mordenite, analcime, laumontite and Kfeldspar) are common. These changes reflect dehydration reactions (Ratterman and Surdam, 1981; Noh and Boles, 1989; Smith, 1991):

This is probably followed by either reaction R5.9 or R5.10, which fixes Mg from seawater.

2.75clinoptilolite + 0.75Na+ + 3.0H4SiO4 -» Na-clinoptilolite + 2.25mordenite + 0.13K+ + 0.27Mg2+ + 0.08H+ + 4.46H 2 O

(R5.5)

Clinoptilolite and mordenite may also react with solution to form analcime, adularia, quartz and calcite (Iijima, 1974). mordenite + 2Na+ + CO 3 2 " —» analcime + 6SiO2 + CaCO 3 + 5H 2 O

(R5.6)

clinoptilolite + Ca2+ + 2HCO 3 " —» analcime + adularia + 5SiO2 + CaCO 3 + CO 2 8H 2 O

(R5.7)

Increasing diagenesis may favour the formation of calcic zeolites, particularly in mafic facies, because of increased Ca/Na activity ratios due to albitisation of calcic plagioclase (Utada, 1991). In addition, chlorite and epidote may crystallise. Chlorite and epidote have been observed as direct alteration products of dacitic to basaltic glasses, and clays at relatively shallow depths (420 m) in modern geothermal regions (White and Sigvaldason, 1962). The development of epidote depends on the availability of Fe3+ and is probably controlled by the earlier formation of palagonite. Where palagonite is absent the reaction of basaltic glass to form chlorite releases Ca, which may be consumed by the precipitation of epidote (Baragar et al., 1979). Chlorite and epidote are also typical of low-temperature metamorphism and the growth of these minerals may bridge the boundary between diagenesis and metamorphism where glass and diagenetic clays are replaced by phyllosilicates. K-feldspar is common in diagenetically altered originally glassy volcanic facies. Munhaetal. (1980) suggested that below 150°C, Na in glass might be exchanged for K in seawater, resulting in precipitation of K-feldspar and K-smectite. However, K-feldspar has not been reported as a direct product of glass alteration, thus some intermediate phases appear to be required. Iijima and Hay (1968), Surdam and Sheppard (1978) and Hay and Guldman (1987) recognised that Kfeldspar replaced earlier mordenite, analcime, clinoptilolite and phillipsite. In contrast, Noh and Boles (1989) proposed that K-feldspar crystallised from silicic glass via a series of hydration and dissolution reactions, which included an intermediate phase of K-rich gel-like glass.

(R5.8)

12.5perlitic glass + 3.88K+ + 0.65H+ +15.4H 2 O -» smectite + 9.5gel-like glass + 4.03Na+ + 0.25Ca2+ + 10.55H4SiO4 (R5.9) 2.79perlitic glass + 0.2Mg2+ + 0.2Fe2+ + 0.32H+ + 5.27H2O —» 1.27smectite + gel-like glass + 1.0Na+ + 0.12Ca2+ + 3.61H4SiO4

(R5.10)

The K liberated during zeolite forming reactions (e.g. reaction R5.4) is then fixed in K-feldspar formation. gel-like glass + 0.5K+ + 0.2H+ +0.2H 2 O 1.3K-feldspar + 0.1 Na+ + 0.1 Ca2 + 0.8H4SiO4 + 0.1 Mg2+ + 0.1Fe2+ (R5.ll) Albite is common as a replacement product of plagioclase and K-feldspar in rhyolitic to basaltic rocks (Munha et al., 1980; Boles, 1982; Torres et al., 1995). Albitisation of plagioclase occurs by dissolution and replacement (Boles, 1982; Shiraki and Iiyama, 1990; Torres et al., 1995). In many cases, albitisation appears to have progressed preferentially along plagioclase grain fractures and cleavage traces, suggesting that fluid films can infiltrate the crystal along lattice defects and planes of weakness, promoting albitisation (Boles, 1982). Microlites in the glassy groundmass of slowly cooled lavas and sills may serve as nuclei for the crystallisation of albite (Dimroth and Lichtblau, 1979). The replacement of plagioclase and K-feldspar by albite could reflect the exchange of K in the rock with Na in seawater at greater depths and temperatures (105-120°C) (Iijima and Utada, 1972; Merino, 1975; Munha et al., 1980; Boles, 1982). The Na and Si required for albite crystallisation are supplied by diffusion from seawater and earlier diagenetic reactions (Boles and Coombs, 1977; Boles, 1982). 2SiO2 + 0.5H 2 O + H+ + Na+ + Ca-plagioclase albite + 0.5Al,Si,O s (OH) 4 + Ca2+

(R5.12)

Albite is typically riddled with minute inclusions of clays, sericite and calcite, which form as by-products and consume Ca and Al released during this reaction.

SEAFLOOR-AND BURIAL-RELATED ALTERATION I 1 1 5

5.4 | REGIONAL METAMORPHISM (ZEOLITE TO AMPHIBOLITE FACIES)

pumpellyite, epidote, albite, K-feldspar, phyllosilicate minerals (smectites, chlorite, celadonite, sericite), calcite, siderite, quartz, apatite, sphene, pyrite and Fe-oxides (Surdam, 1973; Boles and Coombs, 1977).

Transition from diagenesis to regional metamorphism

Burial metamorphic facies

Diagenesis progresses gradually to regional metamorphism with increasing temperature, pressure and, commonly, depth of burial. Many sedimentologists consider that the boundary between diagenesis and metamorphism occurs when a rock has less than 5% interconnected pore space, whereas metamorphic petrologists tend to define it by mineral assemblages that are not stable in sedimentary environments (e.g. Coombs, 1954; Blatt et al., 1972; Winkler, 1979; Turner, 1980; Morrow and Mcllreath, 1990; Bevins and Robinson, 1992). These definitions do not necessarily coincide. Diagenesis and low-grade metamorphism can both have temperatures and pressures in the 200-300°C and less than 1 kbar range. In fact, no clear distinction exists between the processes, pressure and temperature conditions, fabrics (textures) and mineral assemblages of diagenesis and low-grade metamorphism. Diagenesis involves hydration of glass, compaction, dissolution, cementation and minor recrystallisation: metasomatic processes involving minor chemical exchange between the host facies and trapped fluid at low temperatures (up to 250°C). In contrast, metamorphism involves mainly isochemical processes with substantial recrystallisation, but chemical changes that are limited to dehydration and decarbonation (Fyfe et al., 1958). During progressive burial there is a stage when metasomatic reactions occur but the temperatures are generally considered too high for diagenesis; this transitional stage is referred to as burial metamorphism (Coombs, 1954). The transition from diagenesis to burial metamorphism has been studied in andesitic to rhyolitic rocks in New Zealand (e.g. Coombs, 1954; Coombs et al., 1959; Boles, 1974; Boles and Coombs, 1975, 1977), Chile (e.g. Levi, 1970), the USA (e.g. Dickinson, 1962; Sheppard and Gude, 1973) and Japan (e.g. Utada, 1970; Iijima and Utada, 1972; Iijima, 1978), and in basaltic rocks in Canada (e.g. Surdam, 1973; Kuniyoshi and Liou, 1976), Australia (e.g. Smith, 1969; Hellman et al., 1977; Smith et al., 1982) and Iceland (e.g. Viereck et al., 1982).

A metamorphic facies is defined by a group of metamorphic mineral assemblages occurring in spatially associated rock types of diverse chemical composition, which are interpreted to have formed during restricted temperature and pressure conditions (Fig. 5.11). Low-grade facies typical of burial metamorphism in volcanic successions are characterised by hydrous minerals and carbonates, whereas high-grades facies are typically coarser grained and contain anhydrous and CO2-poor minerals. The low-grade facies are (Table 5.3): zeolite, prehnite + pumpellyite, pumpellyite + actinolite, lawsonite + albite + chlorite, blueschist, and greenschist facies (Turner, 1980). Although burial metamorphic facies are widespread, they are commonly heterogenous with patchy and domainal textures (e.g. epidote metadomains of Smith, 1968, 1974, 1977; Smith and Smith, 1976). This domainal style of alteration may reflect mobilisation and local redistribution of elements on a scale of centimetres to metres (e.g. the loss of Fetotal, Mg, Na and K and gain of Ca from epidote metadomains balance losses and gains from the enclosing albite domains, Smith, 1977), rather than the significant addition of elements.

Burial metamorphic zones Burial metamorphism is a progressive process that produces a sequence of regionally extensive metamorphic zones (e.g. Figs 3.13 and 5.12). Metamorphic zones are mappable groups of rocks that have similar metamorphic grade. Adjacent metamorphic zones, like altered zones, are separated by lines of equal grade (isograds), which are delineated by the first appearance of an index mineral or minerals within the same rock type or composition.

Burial metamorphism Burial metamorphism was defined by Coombs (1954) to cover progressive mineral changes that can be directly correlated with increases in temperature and burial depth in thick sedimentary or volcanic successions. It is a form of regional metamorphism that affects thick sedimentary or volcanic successions in subsiding basins, where the basal parts attain low-grade metamorphic conditions without the deformation or folding typical of regional metamorphism. Burial metamorphism, like diagenesis, rarely attains equilibrium mineral assemblages, and penetrative deformation fabrics are absent. Alteration minerals common to burial metamorphism in submarine volcanic successions are: zeolites (heulandite, stilbite, laumontite, analcime), prehnite,

FIGURE 5.11 | Pressure and temperature diagram showing the fields for regional metamorphic facies (after Turner, 1980; Bevins and Robinson, 1992). Boundaries between the fields are gradational. Abbreviations used are PA = pumpellyite + actinolite and PP = prehnite + pumpellyite.

1 1 6 I CHAPTER 5 Table 5.3 | Common burial and regional metamorphic facies and mineral assemblages for submarine volcanic successions (from Coombs, 1954; Humphris and Thompson, 1978; Turner, 1980; Sivell, 1984; Yardley, 1989).

Zeolite

Zeolites (laumonite, analcime, heulandite, stilbite, natrolite, mesolite, wairakite) + mixed-layer clays + quartz + calcite ± muscovite

Prehnite + pumpellyite

Prehnite + pumpellyite ± chlorite + albite + quartz ± epidote ± calcite ± sphene ± rare garnet

Pumpellyite + actinolite

Pumpellyite + actinolite + albite + chlorite + sphene + quartz + muscovite + calcite ± lawsonite

Lawsonite + albite + chlorite Blueschist

Lawsonite + albite + chlorite + quartz ± pumpellyite ± epidote ± actinolite ± sphene + muscovite ± calcite Glaucophanic amphibole + albite + actinolite/phengite + quartz ± epidote ± chlorite ± sphene ± pumpellyite ± stilpnomelane ± calcite

Greenschist

Actinolite + epidote + albite ± chlorite ± calcite ± tremolite ± talc + quartz ± sphene ± magnetite ± biotite

Amphibolite

Hornblende + plagioclase ± epidote ± garnet ± biotite ± quartz + muscovite + calcite ± sphene ± magnetite

Drilling in modern geothermal regions has revealed patterns of low-grade metamorphic zones that can be related directly to temperature and fluid composition at shallow depths (e.g. Fig. 5.13: Coombs et al., 1959; White and Sigvaldason, 1962; Vierecketal., 1982). It is generally assumed that higher-grade rocks formerly had mineral assemblages typical of lower grade zones, which were progressively altered as metamorphism proceeded. Near isograds, index minerals from the lower zones locally overgrow lower grade minerals. For example, in basaltic lavas and tuffs in east Greenland near the boundary between two zeolite facies burial metamorphic zones, the analcime and mesolite + scolecite zones, vesicles are lined with analcime and filled with mesolite (Neuhoff et al., 1997). The generalised sequence of burial metamorphic facies with increasing depth is (e.g. Fig. 3.13): zeolite, prehnite + pumpellyite, pumpellyite + actinolite, lawsonite + albite + chlorite, blueschist, and greenschist facies (Coombs, 1954; Coombs et al., 1959; Turner, 1980). Departures from this pattern are common.

is diagnostic of the zeolite facies as mineral assemblages are sensitive to primary rock composition, fluid composition, burial history and geothermal gradient. Typically the zeolite facies contains heulandite, laumontite, analcime, quartz, albite and smectites + prehnite and pumpellyite (Turner, 1980; Vierecketal., 1982). Fyfe et al. (1958) and Coombs (1954), in studies of Triassic submarine volcano-sedimentary rocks of Tarinagatura Hills (New Zealand), divided the zeolite facies into three mineral assemblages that correlate with increasing depth: (1) heulandite + analcime + quartz ± (montmorillonite + celadonite + sphene), (2) laumontite + albite + quartz ± chlorite, and (3) quartz + albite + adularia. However, at Hokonui Hills, 50 km east of Tarinagatura Hills, Boles and Coombs (1975) found no correlation between mineral assemblage and depth of the zeolite facies rocks. In contrast, Neuhoff et al. (1997) divided the zeolite facies in the Tertiary flood basalts of east Greenland into five diagenetic and burial metamorphic zones based on the index minerals: (1) chabazite + thomsonite, (2) analcime, (3) mesolite + scolecite, (4) heulandite + stilbite and (5) laumonite (Fig. 5.12).

Zeolite facies Genesis Coombs (1954) and Coombs et al. (1959) proposed that regionally extensive zeolite facies in Triassic volcaniclastic rocks of New Zealand bridged the transition between diagenesis and conventional metamorphism. The zeolite facies embraces coexisting assemblages of Ca + Al- and Na + Al-rich zeolites and quartz (Coombs et al., 1959). No single mineral assemblage

The development of the burial metamorphic zones is complex. In many cases, it is progressive through a number of stages that may overlap with diagenesis. However, the transition to burial metamorphism generally involves dehydration and decarbonation accompanied by the release of silica, such as in

FIGURE 5.12 | Cross-section in the Borggraven region, east Greenland, showing the regional extent and vertical distribution of burialmetamorphic zeolite zones (after Neuhoff et al., 1997). The thin discontinuous lines represent the dips of selected lavas in this section.

SEAFLOOR- AND BURIAL-RELATED ALTERATION

FIGURE 5.13 |

Schematic cross-section showing the low-temperature altered

zones and average temperature measurements for the boundaries in the Wairakei geothermal field, Taupo Volcanic Zone, New Zealand (modified from Coombs etal., 1959, after Steiner, 1953; Banwelletal., 1957).

reactions R5.13 and R5.15, below (Coombs, 1954; Boles and Coombs, 1977). Rocks in the Tarinagatura Hills record two stages of zeolite facies alteration: an upper heulandite + analcime diagenetic stage and a lower laumonite + albite + quartz burial metamorphic stage (Coombs, 1954). In the upper part of the succession, volcanic glass has been replaced by heulandite and, less commonly, analcime (Fig. 5.14). These zeolites coexist with newly crystallised quartz and fine-grained smectite. In the lower part, analcime has been replaced by albite and heulandite by laumonite + quartz. Smectite, chlorite, sericite and mixed-layer minerals occur and prehnite and pumpellyite appear as accessory minerals. These stages can be summarised by the reactions: heulandite —> laumontite + quartz + 2H 2 O

(R5.13)

analcime + quartz —> albite + H 2 O

(R5.14)

laumontite + calcite —* prehnite + quartz + H 2 O + CO 2

(R5.15)

Some successions lack textural evidence for shallow diagenetic and zeolite zones as precursors to higher-grade mineral assemblages, suggesting that metamorphism to higher grades is not always progressive (e.g. Neuhoff et al., 1997) or that early textures are destroyed.

FIGURE 5.14 | Down-hole mineral distributions at Tarinagatura Hills, New Zealand (after Coombs, 1954).

|

117

1 1 8 I CHAPTER 5

5.5 | DIAGENETIC ALTERATION IN THE HOKUROKU BASIN The Middle Miocene Hokuroku Basin, part of the Green Tuff Belt in northern Honshu, Japan, is the most frequently cited example of diagenetic zones in a submarine volcanic succession that hosts VHMS deposits. Iijima (1974, 1978), Iijima and Utada (1971) and Utada (1991) described four flat-lying, vertically stacked, zeolite-dominated altered zones that grade into clay-rich hydrothermal zones proximal to the Kuroko VHMS deposits (Fig. 5.6A and B).

Geological setting The Hokuroku Basin is a 30 x 30 km submarine basin containing a 3 to 6 km thick bimodal volcanic succession of calc-alkaline rhyolites and tholeiitic basalts with some locally abundant andesites (Figs 5.6 and 5.15: Dudas et al., 1983; Urabe, 1987). Here, the Nishikurosawa and Onnagawa Formations dominate the stratigraphy.

The lower Nishikurosawa or Hotakizawa Formation is up to 650 m thick and includes intercalated basaltic lavas and breccias, rhyolitic lavas, and laminated mudstone (Tanimura et al., 1983). The upper Nishikurosawa Formation is a thick (<400 m) succession of rhyolitic lavas, domes and interbedded pumice-rich facies and mudstone (Fig. 5.16: Ishikawa, 1983; Urabe, 1987; Yamagishi, 1987). The upper Nishikurosawa Formation hosts the Kuroko VHMS deposits and is conformably overlain by the Onnagawa Formation (Nakajima, 1988). The Onnagawa Formation comprises a sequence of pumice-rich breccia, sandstone, siltstone and black mudstone with abundant felsic synvolcanic intrusions and local basaltic lavas (Fig. 5.16: Ohmoto and Takahashi, 1983; Tanimura et al., 1983; Nakajima, 1988). The lower pumice breccia is extensive and has been correlated across the Hokuroku Basin (Urabe, 1987). The volcanic succession is relatively undeformed but has undergone regional diagenesis and local hydrothermal alteration and mineralisation (Utada, 1970; Tanimura et al., 1983). Generally the stratigraphy has a gentle dip with open, N-S-trending folds (Tanimura et al., 1983).

Noquchi depression

LEGEND Funakawa Formation Onnagawa Formation Nishikurosawa Formation Daijima Formation Rhyolite Andesite Dolerite Quartz diorite Other rocks Drill holes in this study Major mine Township Major fault Railway

FIGURE 5.15 | Geology of the Hokuroku Basin showing the major lithostratigraphic units and inset the distribution of the Green Tuff Belt in Japan (after Sato, 1974; Tanimura et al., 1983). The trace of Ijima's (1974) cross-section A (in Fig. 5.6) is represented by solid line AA', B by the dashed line BB' and our cross-section (Fig. 5.18) by the dotted line BC.

SEAFLOOR-AND BURIAL-RELATED ALTERATION

I

FIGURE 5.16 | Graphic log of drill core OH-8 in the Odate Basin, western Hokuroku Basin (Japan) showing lithology, bedforms, and diagenetic and hydrothermal zones in pumice-rich facies of the Upper Nishikurosawa and Onnagawa Formations.

Alteration facies and zones The regional diagenetic zones are, from top to base (Figs 5.6 and 5.17): (I) partially altered zone, (II) clinoptilolite + mordenite zone (e.g. data sheets HK1 and 2), (III) analcime zone (e.g. data sheets HK3 and 4), and (IV) laumontite + albitezone (Iijima, 1974, 1978; Hay, 1978; Iijima, 1978). The partially altered zone is distributed beneath the Quaternary gravels and overlies the clinoptilolite + mordenite zone in the eastern Odate Basin, western Hokuroku Basin (Fig. 5. 6A). This zone has a maximum thickness of 60 m and is characterised by well preserved volcanic textures, the absence of zeolites and the presence of unaltered glassy pumice clasts and plagioclase crystals (Iijima, 1974). Rocks in this zone are pale grey. Glass shards and parts of pumice clasts

have been altered to montmorillonite and vesicles are empty or have been partly filled with low-cristobalite. The clinoptilolite + mordenite zone is 160—250 m thick and widely distributed in the shallower part of the Odate Basin in the upper Onnagawa Formation. Regionally it overlies the analcime zone, but in the east of the Odate Basin it is repeated below the analcime + heulandite zone (Figs 5.6B and 5.17). In the upper part of this zone, some glass shards and pumice clasts are unaltered (e.g. data sheet HK1). Typically the surfaces of shards, pumice clasts and vesicles have been coated in a thin film of smectite or low-cristobalite. Vesicles were filled sequentially with mordenite and clinoptilolite. Originally glassy shards and clasts have been altered to smectite (montmorillonite, saponite or mixed-layer illite/smectite) ± mordenite. In the lower part of this zone, dark green saponite

119

1 2 0 I CHAPTER 5

FIGURE 5.17 | Schematic cross-section of the western Hokuroku Basin (Japan) showing lithology and altered zones. The locations of the six data sheets are marked on the section.

or mixed-layer smectite/chlorite and pale green mordenite domains are common in the pumice-rich rocks (e.g. data sheet HK2). Plagioclase crystals are typically unaltered. The analcime zone is at least 7 km wide and 150-200 m thick, and occurs in the Upper Nishikurosawa and lower Onnagawa Formations (below the Ml mudstone). It is approximately equivalent to the ore position, grading into the hydrothermal montmorillonite zone and overlying the sericite + chlorite zone (e.g. data sheets HK5 and 6) associated with the Kuroko VHMS deposits (Figs 5.6 and 5.17). In the east of the Odate Basin, an analcime + calcite zone occurs within the clinoptilolite + mordenite zone (Fig. 5.6A: Iijima, 1974). However, it is different from the regional analcime zone, containing disseminated pyrite and calcite concretions and veinlets, and is considered to result from hydrothermal alteration related to Kuroko mineralisation (Yoshida and Utada, 1968; Iijima, 1974). The regional analcime zone is characterised by analcime, which has completely replaced some shards, and rounded crystals of analcime that have replaced clinoptilolite- and mordenite-altered shards and pumice clasts (e.g. data sheets HK 3 and 4). The internal structure of analcime-altered uncompacted tube pumice clasts is not as well preserved as those in the clinoptilolite + mordenite zone. Plagioclase

crystals are unaltered or have been partly analcime and calcite altered. Dark green saponite, saponite + chlorite and chlorite fiamme are common in this zone. They are interpreted as altered and compacted pumice clasts (Gifkins et al., in press). Analcime-filled solution seams are common in the lower part of this zone. The albite + laumontite zone occurs beneath the sericite + chlorite zone at depth in the Odate Basin (Fig. 5.6). It is characterised by albite + laumontite ± calcite ± chlorite + sericite (Iijima and Utada, 1971). Laumontite has replaced plagioclase crystals, originally glassy shards and pumice clasts and filled pore space (Iijima, 1974). Albite has replaced plagioclase crystals.

Genesis of altered zones The four zeolite zones in the Hokuroku Basin are interpreted to have formed during submarine diagenesis of the mainly glassy felsic volcanic succession (Iijima and Utada, 1971; Iijima, 1974, 1978; Utada, 1991). The alteration pattern is interpreted to have formed beneath the seafloor, while sedimentation continued and the rocks were progressively buried. Ptygmatic folds in near vertical calcite veinlets in

SEAFLOOR-AND BURIAL-RELATED ALTERATION

|

Diagenetic zones ZONE!

I | Partially altered zone lacking zeolites ZONE II I I Clinoptilolite + mordenite zone ZONE III — Analcime + calcite zone Hydrothermal zones ^B Montmorillonite and transitional zones Bi Sericite + chlorite zone with plagioclase I | Sericite + chlorite zone lacking plagioclase

f.-' .'•] Gravel I ?• I Dacitic lava I T l Basaltic lava/sill — Massive sulfide ore i i Felsic volcanic facies

FIGURE 5.18 | Model for the development and deformation of the altered zones in the Hokuroku Basin from the Middle Miocene to Recent (after lijima, 1974). (A) Late Nishikurosawa stage (Middle Miocene). (B) Onnagawa stage with deposition of the M1 mudstone. (C) Late Funakawa stage (Late Miocene). (D) Recent.

Zone III are considered to result from compaction that took place when the volcanic facies were most deeply buried. Thus the diagenetic zones probably formed before the end of deposition of the upper Miocene Funakawa Formation, which overlies the Onnagawa Formation (lijima, 1974). The development of the diagenetic zones can be described in four stages (Fig. 5.18): late Nishikurosawa stage, Onnagawa stage, late Funakawa stage and recent (lijima, 1974). During these stages, mineralogical changes progressed with depth as successive reactions between volcanic glass and interstitial modified seawater occurred, originally forming zeolites and then albite as the temperature and pressure increased (Utada, 1991). These mineralogical changes were accompanied by textural and compositional changes. During the late Nishikurosawa stage (Fig. 5.18A) in the middle Miocene, felsic glass was hydrated and began to alter to montmorillonite and low-cristobalite (lijima, 1974). This resulted in a shallow, regionally extensive partly altered zone. Submarine hydrothermal activity associated with VHMS mineralisation also commenced (lijima, 1974). The clinoptilolite + mordenite zone formed during the late Nishikurosawa and early Onnagawa stage (Fig. 5.18A and B), when partly altered facies were buried to a depth

of approximately 100 m (lijima, 1974). This increased the alkalinity of the fluid resulting in the dissolution of felsic glass and the precipitation of alkali clinoptilolite and mordenite in vesicles, interstitial voids and dissolution cavities. Simultaneous reactions altered glassy clasts and tube pumice walls to saponite. Reaction rates were possibly accelerated as a result of hydrothermal fluids circulating within the succession increasing the geothermal gradient and concentrations of Na and K in the fluid. By the end of the Onnagawa stage the hydrothermal sericite + chlorite and montmorillonite zones associated with the ore deposits had formed (lijima, 1974). During the late Funakawa stage (Fig. 5.18C), clinoptilolite and mordenite reacted with solution to form analcime, adularia, quartz and calcite (Ogihara, 1996). As a result, the analcime + heulandite zone was superimposed on the deeper part of the clinoptilolite + mordenite zone (lijima, 1974). The Funakawa stage corresponds to the time of deepest burial and compaction of altered pumice clasts to form fiamme (lijima, 1974). The diagenetic analcime + heulandite zone formed contemporaneously with hydrothermal zones surrounding the Kuroko VHMS deposits (lijima, 1978; Ohmoto, 1978). Since the Funakawa stage, the altered zones have been mildly deformed and eroded (Fig. 5.18D: lijima, 1974).

121

1 2 2 | CHAPTER 5

Subtle, patchy mordenite + smectite-chlorite alteration facies Sample No.

J6-294

Alteration facies

subtle, patchy mordenite + smectitechlorite

Alteration zone

clinoptilolite + mordenite zone

Location

Yoneshiro River

Formation

Onnagawa Formation

Succession

Green Tuff Belt

Volcanic facies

pumice breccia

Relict minerals

plagioclase + quartz

Relict textures

tube pumice clasts, bubble-wall shards, crystal fragments, non-vesicular volcanic clasts

Primary composition

rhyolite

Lithofacies

graded bed

Interpretation

syneruptive, mass-flow emplaced pumice breccia

Alteration minerals

partly glassy, mordenite + saponite + montmorillonite + smectite-chlorite + K-feldspar + pyrite

Alteration textures

saponite films in vesicles, mordenite ± saponite filled vesicles and pore space, smectite-chlorite fiamme, disseminated pyrite

Distribution

patchy

Preservation

excellent

Alteration intensity Timing

subtle early

HK1

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 2 3

Weak, pervasive mordenite + smectite alteration facies Sample No.

OH8-369

Alteration facies

weak, pervasive mordenite + smectite

Alteration zone

clinoptilolite + mordenite zone

Location

Odate city

Formation

Onnagawa Formation

Succession

Green Tuff Belt

Volcanic facies

pumice + lithic breccia

Relict minerals

plagioclase + quartz

Relict textures

tube pumice clasts, bubble-wall shards, crystal fragments, non-vesicular volcanic clasts

Primary composition

rhyolite

Lithofacies

graded bed

Interpretation

syneruptive, mass-flow emplaced pumice breccia

Alteration minerals

mordenite + smectite-chlorite + K-feldspar + calcite + pyrite + glauconite

Alteration textures

smectite films in vesicles, mordenite ± smectite filled vesicles and pore space, mordenite-altered glass shards and vesicle walls, smectite-chlorite fiamme, disseminated pyrite, microcrystalline lithic clasts

Distribution

pervasive

Preservation

excellent

Alteration intensity

weak

Timing

early

Alteration style

diagenetic

HK2

1 2 4 | CHAPTER 5

Subtle, pervasi¥e smectite-chlorite + mordenite + analcime alteration facies Sample No.

OH8-511

Alteration facies

subtle, pervasive smectite-chlorite + mordenite + analcime

Alteration zone

analcime zone

Location

Odate city

Formation

Onnagawa Formation

Succession

Green Tuff Belt

Volcanic facies

pumice breccia

Relict minerals

plagioclase + quartz

Relict textures

tube pumice clasts, crystal fragments

Primary composition rhyolite Lithofacies

massive

Interpretation

syneruptive, mass-flow emplaced pumice breccia

Alteration minerals

smectite-chlorite + mordenite + analcime + sericite + pyrite

Alteration textures

analcime solution seams, smectite-chlorite fiamme, mordenite filled vesicles, analcime replacing mordenite + smectite-altered pumice clasts

Distribution

pervasive

Preservation

good

Alteration intensity

subtle

Timing

early

Alteration style

diagenetic

HK3

SEAFLOOR- AND BURIAL-RELATED ALTERATION | 1 2 5

Weak, pervasive analcime + mordenite alteration facies Sample No.

OH8-537

Alteration facies

weak, pervasive analcime + mordenite

Alteration zone

analcime zone

Location

Odate city

Formation

Onnagawa Formation

Succession

Green Tuff Belt

Volcanic facies

pumice breccia

Relict minerals

plagioclase

Relict textures

tube pumice clasts, bubble-wall shards, crystal fragments

Primary composition

rhyolite

Lithofacies

graded bed

Interpretation

syneruptive, mass-flow emplaced pumice breccia

Alteration minerals

analcime + mordenite + clinoptilolite + smectite-chlorite + pyrite + sericite

Alteration textures

mordenite and clinoptilolite filled vesicles, mordenite and analcime altered vesicle walls, analcime overgrowths on plagioclase crystal fragments

Distribution

pervasive

Preservation

moderate

Alteration intensity

weak

Timing

early

Alteration style

diagenetic

HK4

1 2 6 | CHAPTER 5

Strong, pervasive quartz + sericite alteration fades Sample No.

OH8-794

Alteration facies

strong, pervasive quartz + sericite

Alteration zone

sericite + chlorite zone

Location

Odate city

Formation

Nishikurosawa Formation

Succession

Green Tuff Belt

Volcanic facies

pumice + lithic breccia

Relict minerals

plagioclase

Relict textures

volcanic clasts

Primary composition rhyolite Lithofacies

graded bed

Interpretation

syneruptive, mass-flow emplaced pumice breccia

Alteration minerals

quartz + K-feldspar + sericite + chlorite+ pyrite

Alteration textures

pseudomorphs after plagioclase crystals and clasts in pervasive crystalline matrix, quartz-filled dissolution vugs

Distribution

pervasive

Preservation

poor

Alteration intensity

strong

Timing

early

Alteration style

hydrothermal

HK5

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 2 7

moderate, pervasive sericite + chlorite alteration facies Sample No.

HO20-485

Alteration facies

moderate, pervasive sericite + chlorite

Alteration zone

sericite + chlorite zone

Location

near Fukazawa deposit

Formation

Nishikurosawa Formation

Succession

Green Tuff Belt

Volcanic facies

pumice + lithic breccia

Relict minerals

nil

Relict textures

clasts?

Primary composition rhyolite Lithofacies

graded bed

Interpretation

synemptive, mass-flow emplaced pumice breccia

Alteration minerals

chlorite + sericite + pyrite + montmorillonite

Alteration textures

dissolution vugs after crystals, sericite + chlorite fiamme, disseminated pyrite

Distribution

pervasive

Preservation

poor

Alteration intensity

moderate

Timing

early

Alteration style

hydrothermal

HK6

1 2 8 | CHAPTER 5

5.6 | DIAGENETIC ALTERATION IN THE MOUNT READ VOLCANICS Gifkins (2001) recognised two regionally developed Cambrian diagenetic zones (albite zone and epidote zone) within the northern Central Volcanic Complex in the Mount Read Volcanics. Although the rocks currently have mineral assemblages consistent with greenschist facies metamorphism, diagenetic alteration facies were identified based on combinations of mineral assemblages, overprinting relationships, textures, distribution, alteration intensity and whole-rock geochemistry. The original diagenetic mineral assemblages were inferred from local relict textures and by comparison with younger diagenetically altered volcanic successions (Gifkins and Allen, 2001). Locally, the diagenetic alteration facies merge with hydrothermal alteration facies in the margins of the Rosebery and Hercules VHMS systems (Allen, 1997). Diagenetic and hydrothermal alteration facies are interpreted to have had similar timing (e.g. Fig. 3.20), and the hydrothermal system may have contributed heat and fluid to intensify the diagenetic system (Gifkins, 2001).

Geological setting For a detailed description of the geology of the Mount Read Volcanics refer to Section 1.5. The northern Central Volcanic Complex is exposed in an approximately 30 by 6 km area located north and west of the Henty fault, and east of the Rosebery fault (Fig. 1.5). It includes three compositionally and texturally different formations (Fig. 5.19): the Sterling Valley Volcanics, the Mount Black Formation, and the Kershaw (or Hercules) Pumice Formation (Gifkins, 2001). The Sterling Valley Volcanics (>1.5 km thick) are composed of dacitic to basaltic lavas and sills, and polymictic mafic volcaniclastic facies interpreted as resedimented syneruptive hyaloclastite, autobreccia, pillow lava and scoria. The Mount Black Formation is a laterally extensive (>20 km), thick succession (>1.6 km) of mainly feldspar-phyric massive, flowbanded and autobrecciated lavas, domes, cryptodomes and synvolcanic sills (e.g. data sheets CVC1 CVC5 and CVC6). The Kershaw Pumice Formation, which conformably overlies the Mount Black Formation, is a laterally extensive (>16 km), relatively thick (>800 m) succession dominated by non-welded pumice breccia (e.g. data sheets CVC3 and CVC4), pumice-rich sandstone and shard-rich siltstone, with lesser proportions of pumice-lithic clast-rich breccia and sandstone, and massive, flow-banded and brecciated rhyolitic and dacitic lavas and intrusions (e.g. data sheet CVC2). The upper part of the Kershaw Pumice Formation and the base of the overlying White Spur Formation host the Rosebery and Hercules VHMS deposits. Abundant spherulites, lithophysae, micropoikilitic texture and relict perlite indicate that volcanic rocks in the northern Central Volcanic Complex were initially partly crystalline and partly glassy (Gifkins and Allen, 2001).

Alteration facies and zonation Regionally distributed diagenetic albite and epidote zones formed before, or were synchronous with, stylolitic S, compaction foliation. The alteration intensity is generally weak with volcanic textures and albite-altered plagioclase crystals preserved. Locally the distribution and intensity of the diagenetic alteration facies is patchy, reflecting the complexity of the original volcanic facies (Fig. 5.20). The albite zone is characterised by pervasive albite + quartz + sericite (e.g. data sheets CVC2 and CVC5), domainal albite + quartz + sericite with sericite + hematite ± chlorite (e.g. data sheet CVC3) and pervasive sericite (e.g. data sheet CVC4) alteration facies. It is thick (>2 km) and encompasses the Kershaw Pumice Formation and most of the Mount Black Formation (Fig. 5.21). The albite + quartz + sericite-rich facies are associated with minor increases in SiO2, CaO, Na 2 O and total mass, and decreases in K2O and A12O3 consistent with seafloor albitisation (cf. Boles and Coombs, 1977; Boles, 1982). The sericite + hematite + chlorite alteration facies is associated with minor increases in K2O and A12O3 consistent with the conversion of silicic glass to clay minerals (cf. Noh and Boles, 1989; Passaglia et al., 1995). The abundance of hematite may reflect the oxidation of Fe3+ during alteration of glass to clays (e.g. Klein and Lee, 1984). The epidote zone is characterised by pervasive albite + quartz + sericite, pervasive chlorite + sericite, pervasive chlorite + epidote and domainal chlorite + epidote with albite + quartz + sericite (e.g. data sheet CVC6) alteration facies. The epidote zone is less extensive than the albite zone and is restricted to the Mount Black Formation and Sterling Valley Volcanics at the stratigraphic base of the northern Central Volcanic Complex, adjacent to the Henty fault (Fig. 5.19). In the epidote zone, chlorite + sericite and chlorite + epidote altered felsic rocks have gained MgO, consistent with the formation of smectite, chlorite and other Mg-silicates during diagenesis (cf. Hajash and Chandler, 1981; Shiraki and Iiyama, 1990).

Genesis of alteration facies The epidote zone occurs in the core of the regional anticline in the Sterling Valley, suggesting that it is associated with the deepest stratigraphic level in the northern Central Volcanic Complex: the lower Mount Black Formation and Sterling Valley Volcanics (Gifkins, 2001). The change from the albite zone to the epidote zone with stratigraphic depth is consistent with diagenetic alteration zonation (cf. Iijima, 1974, 1978). Thick (>1 km) diagenetic zones with high-temperature mineral assemblages (albite + quartz + sericite and chlorite + epidote) suggest that they developed in response to a highgrade diagenetic alteration system that involved an elevated geothermal gradient (cf. Utada, 1991). Albite + quartz + sericite, sericite + hematite ± chlorite, and sericite alteration facies are the metamorphosed equivalents of diagenetic alteration facies that coated original surfaces, filled primary porosity and replaced glass in the northern Central Volcanic Complex prior to or synchronous with diagenetic compaction. Thin films of sericite, carbonate and hematite replaced clays that had coated original glassy surfaces at the onset of diagenesis. Albite + quartz + sericite,

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 1 2 9

FIGURE 5.19 | Geology of the northern Mount Read Volcanics in western Tasmania, showing the major lithostratigraphic units and altered zones in the northern Central Volcanic Complex (after Gifkins, 2001). Locations of the six data sheets are marked on the map.

FIGURE 5.20 | Detailed cross-section in the Rosebery hanging wall (western Tasmania) showing the complex distribution of volcanic and alteration fades (after Gifkins and Allen, 2001).

FIGURE 5.21 | Schematic cross-section of the northern Central Volcanic Complex stratigraphy and altered zones, western Tasmania (after Gifkins, 2001).

SEAFLOOR- AND BURIAL-RELATED ALTERATION | 1 3 1 Stage 1: Onset of regional synvolcanic diagenesis Thin films of sericite, hematite and calcite coat original surfaces, including vesicle walls, plagioclase crystals, shards and fractures. These films are the metamorphic equivalents of low-temperature smectite, palagonite and calcite rim cements. This early stage probably involved interaction with seawater trapped in the volcanic succession. Modified seawater may have been expelled from the succession in response to overburden pressure, and migrated towards the seafloor as diffuse unfocused flow.

Stage 2: Diagenesis and synchronous hydrothermal alteration and mineralisation Zeolite or clay mineral cements began to fill primary pore spaces, vesicles and perlitic fractures. Subsequently, these were extensively replaced by K-feldspar or albite and chlorite. Zeolitisation probably occurred at temperatures between 40 and 100°C. Locally, hydrothermal fluids altered the succession. Hydrothermal fluid flow was unfocussed and in places ponded beneath the coherent facies of sills and lavas. The Rosebery and Hercules VHMS deposits and their altered halos are interpreted to have formed during this stage at temperatures greater than 300°C.

Stage 3: Continuing diagenetic alteration and compaction synchronous with deposition of the White Spur Formation Dissolution and alteration of glass to clays, sericite and chlorite occurred synchronous with compaction. Replacement of earlier zeolites by Kfeldspar occurred below 150°C, albitisation of plagioclase phenocrysts and albite replacement of K-feldspar occurred at temperatures between 100 and 190°C. Large volumes of fluid were probably displaced as a result of compaction under the weight of the accumulating White Spur Formation. Rapid and variable sedimentation rates may have over-pressured the pore fluid, promoting lateral fluid flow along permeable layers. Weak hanging wall alteration developed during continued hydrothermal alteration associated with the formation of the Rosebery deposit.

Stage 4: Transition between diagenesis and regional metamorphism More stable, higher-temperature mineral assemblages replaced remaining glass, phenocrysts and early alteration minerals. Chlorite + epidote alteration facies developed at depth in both mafic and felsic volcanic facies: probably at high (>200°C) temperatures.

Stage 5: Devonian metamorphism and deformation Greenschist facies mineral assemblages and tectonic fabrics overprinted diagenetic and hydrothermal alteration facies. Deformation modified preexisting volcanic and alteration textures and produced folds, faults and shear zones. The distribution of syn-S2 alteration facies suggests that metamorphic fluid migration was restricted to regional structures such as faults and shear zones. Mineral assemblages in intermediate and mafic rocks in the Mount Read Volcanics indicate that the peak regional metamorphic temperature was between 370 and 450°C.

FIGURE 5.22 | Model for the post-depositional evolution of the northern Central Volcanic Complex, western Tasmania (after Gifkins, 2001). Schematic crosssections are not to scale.

1 3 2 | CHAPTER 5

chlorite + sericite and sericite + hematite + chlorite replaced zeolites and clays that filled pore space and altered glass, prior to and synchronous with diagenetic compaction. In pumicerich facies, a bedding-parallel stylolitic foliation reflects the dissolution of glass during compaction and fiamme are interpreted as diagenetically altered and flattened pumice clasts (Gifkins et al., in press). Diagenetic alteration involved significant mineralogical and textural changes but only minor changes in composition consistent with the interaction of rhyolitic and basaltic glass with seawater during burial. The chlorite + epidote alteration mineral assemblage may be transitional between diagenesis and burial metamorphism. It developed after lithification and compaction but pre-dated regional deformation associated with peak metamorphism. The chlorite + epidote facies replaced earlier clay or chlorite + sericite-rich facies and filled any remaining pore space. Negligible absolute and total mass changes associated with chlorite + epidote alteration suggest that it grew in response to increasing temperature with increasing depth of burial late in the diagenetic history (Gifkins, 2001).

Mineral assemblages in these diagenetic zones reflect the reaction of glass with interstitial fluid at elevated temperatures. The albite zone probably formed at temperatures between 100 and 190°C (cf. Iijima and Utada, 1971; Thompson, 1971; Merino, 1975; Munha et al., 1980; Boles, 1982). The epidote zone is characterised by chlorite + epidote, chlorite + sericite and albite + quartz + sericite indicating formation at temperatures of at least 200°C (cf. Seki, 1972; Kristmannsdottir, 1976). The regional diagenetic alteration and metamorphism of the northern Central Volcanic Complex can be described in five successive stages (Fig. 5.22): (1) the onset of diagenesis; (2) formation of diagenetic cements, and synchronous hydrothermal alteration and mineralisation; (3) diagenetic alteration and compaction synchronous with emplacement of the White Spur Formation; (4) replacement of early diagenetic minerals and remaining glass by more stable mineral assemblages; and (5) regional Devonian metamorphism and deformation.

SEAFLOOR-AND BURIAL-RELATED ALTERATION | 133

CVC1

Subtle, pervasive albite + quartz + chlorite alteration facies Least-altered rhyoiite Sample no.

133921

Alteration facies

subtle, pervasive albite + quartz + chlorite

Alteration zone

albite zone

Location

Mount Black

Formation

Mount Black Formation

Succession

Central Volcanic Complex

Volcanic facies

massive, plagioclase-phyric rhyolite

Relict minerals

plagioclase

Relict textures

porphyritic, micropoikilitic

Primary composition rhyolite Lithofacies massive Interpretation

coherent facies

Alteration minerals

albite + quartz + sericite + chlorite + hematite albite + quartz ± sericite ± chlorite pseudomorphs of plagioclase, micropoikilitic albite + quartz, interstitial chlorite, disseminated hematite

Alteration textures

Distribution

pervasive

Preservation

excellent

Alteration intensity

subtle

Timing

pre-S2

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

74.58 K2O 0.27 P2O5 13.85 S 2.08 Total 0.01 0.38 Rb 0.13 Sr 3.54 Ba

Photomicrograph (ppl)

4.34 0.03 <0.01 100.32

Cu Pb Zn Th Zr 136 Nb 96 Y 988

2 3 17 22 270 17 41

Al CCPI Ti/Zr

56 22 5.98

1 3 4 | CHAPTER 5

CVC2

Weak, pervasive albite + quartz * sericite alteration facies Sample no.

147407

Alteration facies

weak, pervasive albite + quartz + sericite

Alteration zone

albite zone

Location

120R-438.5 m

Formation

Kershaw Pumice Formation

Succession

Central Volcanic Complex

Volcanic facies

jigsaw fit, monomictic, plagioclase-phyric rhyolite breccia

Relict minerals

plagiociase

Relict textures

porphyritic, perlitic fractures, jigsaw fit clasts

Primary composition rhyolite Lithofacies

massive

Interpretation

in situ hyaloclastite

Alteration minerals

albite + quartz + sericite > chlorite + pyrite > calcite

Alteration textures

albite ± calcite pseudomorphs of plagiociase, microcrystaiiine groundmass, calcite veins, chlorite filled perlitic fractures

Distribution

pervasive

Preservation

moderate

Alteration intensity

weak

Timing

pre-S2

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

74.01 K2O 0.23 P2O5 12.21 S 2.17 Total 0.09 0.49 Rb 2.41 Sr 4.07 Ba

Photomicrograph (xn)

1.82 0.03 0.01 100.61

Cu Pb Zn Th Zr 76 Nb 113 Y 513

2 4 19 12 258 16 36

Al CCPI Ti/Zr

26 29 5.31

SEAFLOOR- AND BURIAL-RELATED ALTERATION

Moderate, domainal albite + quartz + sericite with sencite + hematite ± chlorite alteration facies Sample no.

147410

Alteration facies Alteration zone

moderate, domainal albite + quartz + sencite albite zone

Location

120R-524.5m

Formation

Kershaw Pumice Formation

Succession

Central Volcanic Complex

Volcanic facies

graded, plagioclase-phyric pumice breccia

Relict minerals

plagioclase

CVC3

Relict textures

tube pumice clasts, fiamme, plagioclase crystal fragments, blocky rhyolite clasts Primary composition rhyolite Lithofacies

normally graded

Interpretation

Distribution

syn-eruptive, mass-flow-emplaced pumice breccia albite + quartz + sericite + chlorite + hematite + calcite sericite fiamme, hematite styioiites, albite veins, recrystallised albite + quartz + sericite pumice clasts and matrix, albite + sericite + calcite altered plagioclase domainal

Preservation

poor

Alteration intensity

moderate

Timing Alteration style

Alteration minerals Alteration textures

Geochemistry SiO2 TiO2 AI2O3 Fe2O3

76.08 0.19 10.66 1.67

S Total

pre-S2

MnO MgO CaO

0.09 0.47 2.39

diagenetic

Na2O

4.71

Hand specimen photograph

K2O

0.88 0.03 0.01 99.93

Cu Pb Zn Th

1 4 27 10

Rb Sr

34 144

Zr Nb Y

210 13 37

Ba

280

P2O5

Photomicrograph (xn)

Al CCPi Ti/Zr

16 26 5.43

|

135

1 3 6 | CHAPTER 5

Weak, pervasive sericite alteration facies Sample no.

147552

Alteration facies

weak, pervasive sericite

Alteration zone

albite zone

Location

Pieman Road

Formation

Kershaw Pumice Formation

Succession

Central Volcanic Complex

Volcanic facies

massive, plagioclase-phyric pumice breccia plagioclase

Relict minerals

CVC4

Relict textures

tube pumice ciasts, bubble wall shards, plagioclase crystal fragments, fiamme Primary composition rhyolite Lithofacies normally graded Interpretation Alteration minerals Alteration textures

syn-eruptive, mass-flow-emplaced pumice breccia sericite + albite + calcite + chlorite + hematite sericite fiamme, hematite stylolites, disseminated calcite rhombs, albite + sericite altered pumice ciasts and shards

Distribution

pervasive

Preservation

good

Alteration intensity

weak

Timing

pre-S2

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

70.91 K2O 0.31 P2O5 14.08 S 2.78 Total 0.07 0.77 Rb 1.66 Sr 2.68 Ba

Photomicrograph (ppl)

3.16 0.07 0.01 99.75

Cu Pb Zn Th Zr 124 Nb 87 Y 786

4 2 48 251 13 28

Al CCPI Ti/Zr

48 36 7.41

SEAFLOOR- AND BURIAL-RELATED ALTERATION | 1 3 7

CVC5

Subtle, pervasive aibite + quartz + chlorite alteration facies Least-altered dacite Sample no.

147435

Alteration facies

subtle, pervasive aibite + quartz + chlorite

Alteration zone

aibite zone

Location

MBD4-18.4m

Formation

Mount Black Formation

Succession

Central Volcanic Complex

Volcanic facies

massive, plagioclase + hornblende-phyric dacite

Relict minerals

plagioclase, hornblende

Relict textures

porphyritic, glomeroporphyritic clusters, micropoikilitic

Primary composition

dacite

Lithofacies

massive

Interpretation

coherent facies

Alteration minerals

aibite + quartz + chlorite + epidote

Alteration textures

aibite + quartz micropoikilitic groundmass with interstitial chlorite + epidote, aibite pseudomorphs of plagioclase, epidote + chlorite altered hornblende

Distribution

pervasive

Preservation

excellent

Alteration intensity

subtle

Timing

pre-S2

Alteration style

diagenetic

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na,0

67.53 K2O 0.52 P2O5 14.51 S 4.37 Total 0.06 1.3 Rb 2.38 Sr 3.56 Ba

I Hand specimen photograph

Photomicrograph (ppl)

3.95 0.13 0.01 99.51

Cu Pb Zn Th Zr 102 Nb 242 Y 958

4 4 51 15 216 12 34

Al CCPI Ti/Zr

47 41 14.48

1 3 8 | CHAPTER 5

CVC6

Moderate, domainal chlorite + epidote alteration facies Sample no.

147557

Alteration facies

moderate, domainal chlorite + epidote

Alteration zone

epidote zone

Location

Pieman Road

Formation

Mount Black Formation

Succession

Central Volcanic Complex

Volcanic facies

jigsaw fit, monomictic plagioclase + homblende-phyric dacite breccia plagioclase, hornblende

Relict minerals Relict textures

glomeroporphyritic, perlitic fractures, jigsaw-fit clasts Primary composition dacite Lithofacies

massive

Interpretation

in situ hyaloclastite

Alteration minerals

albite + quartz + chlorite + epidote

Alteration textures

Distribution

microcrystalline groundmass with domainal albite + quartz and chlorite + epidote facies, plagioclase phenocrysts albite or chlorite ± epidote altered, hornblende altered to chlorite + epidote domainal

Preservation

good

Alteration intensity

weak

Timing

pre-S2

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

67.93 K2O 0.59 P2O5 14.32 S 4.57 Total 0.06 1.33 Rb 2.67 Sr 4.97 Ba

Photomicrograph (ppl)

2.05 0.13 0.01 99.65

Cu Pb Zn Th Zr 41 Nb 151 Y 826

10 3 28 201 12 31

Al CCPI Ti/Zr

31 44 17.68

139

6 | SYNVOLCANIC INTRUSION-RELATED ALTERATION

The spatial and genetic associations between intrusions and altered zones are widely appreciated in porphyry and epithermal districts (Lowell and Guilbert, 1970; Titley, 1982; Henley and Brown, 1985). Similar relationships also exist in VHMS districts, where synsedimentary or synvolcanic intrusions are commonly altered and surrounded by halos of altered rocks. In some VHMS districts (e.g. Snow Lake and Sturgeon Lake, Canada), there are spatial associations between synvolcanic intrusions and broad-scale, semi-conformable altered zones and clusters of VHMS deposits in the overlying successions (Spooner and Fyfe, 1973; Campbell et al., 1981; Gibson and Watkinson, 1990; Galley, 1993; Hannington et al., 2003a). It has been suggested that synvolcanic intrusions were heat sources (Spooner and Fyfe, 1973; Ohmoto and Rye, 1974; Solomon, 1976; Cathles, 1977; Franklin et al., 1981; Polya et al., 1986; Galley, 1993; Large et al., 1996), and perhaps also volatile and metal sources (Urabe and Sato, 1978; Stanton, 1990; Yang and Scott, 1996; Hannington et al., 1999) for subseafloor hydrothermal systems that formed altered zones and VHMS deposits. Synvolcanic intrusive sills, cryptodomes, dykes and subvolcanic plutons are volumetrically important in submarine volcanic successions (Polya et al., 1986; McPhie and Allen,

1992; Doyle and Huston, 1999; Galley, 2003). They may be composite intrusions of variable volumes up to 1000 km3, typically emplaced at depths up to 4 km below the seafloor (Nielsen et al., 1981; Galley, 2003; Whalen et al., 2004). Intrusions and intrusion-related altered zones that significantly post-date volcanism are also common in ancient submarine volcanic successions; however, they are not the focus of this chapter. Alteration can occur within intrusions (deuteric and local hydrothermal alteration), locally in the immediate host rocks (contact alteration) or regionally in the host succession (regional hydrothermal alteration) (Fig. 6.1). This chapter describes the role of intrusions in generating regional hydrothermal systems, regional hydrothermally altered zones, altered zones within intrusions and contact altered zones around both small-volume, near-seafloor and larger, deeper intrusions in submarine volcanic successions. The final section presents a case study of contact altered zones associated with the Darwin Granite in the southern Mount Read Volcanics, western Tasmania. The recognition of altered zones related to synvolcanic intrusions can provide insights into fluid-flow and thermal histories of VHMS districts, and thereby assist mineral exploration.

FIGURE 6.1 | A cartoon of the variety of altered zones associated with synvolcanic intrusions. (A) A deuteric altered zone within the top of a large volume intrusion. (B) A fracture-controlled hydrothermally altered zone at the margins of an intrusion and in the surrounding host rocks. (C) Contact-altered zones around synvolcanic sills emplaced into unconsolidated sediment immediately below the seafloor. (D) Concentric contact-altered zones around a large volume intrusion emplaced at depth. (E) Regional hydrothermally altered zones related to emplacement of a subvolcanic pluton. (F) Afootwall alteration pipe beneath a VHMS deposit.

1 4 0 I CHAPTER 6

6.1 | THE ROLE OF INTRUSIONS IN GENERATING HYDROTHERMAL SYSTEMS The most active hydrothermal systems are those related to magma-induced thermal anomalies (Alt, 1999; Butterfleld, 2000). The magma chamber provides heat to overlying strata and active volcanism contributes heat from its eruptive products, intrusions and feeder dykes. The transfer of heat and mass away from the intrusion may occur by either conduction only, or conduction and infiltration. Conduction generally involves only minor diffusion of elements, although Weaver et al. (1990) suggested that at near solidus temperatures vapourphase expulsion may produce local mineral and chemical variations (loss of Na, halogens and REE) in volcanic glass. In contrast, conduction accompanied by infiltration and circulation of hot fluid can remove heat from the magmatic system much faster than conduction alone, and effectively transport elements considerable distances, up to hundreds of kilometres, through the succession. Thermal metamorphism related to the shallowemplacement of synvolcanic intrusions in dry successions typically results in limited alteration with little or no mass transfer. In rare cases, magmatic fluids exsolved from the crystallising magma hydrothermally alter dry host facies. Vapour-phase expulsion of some elemental species as complexes (e.g. fluoride, chloride, hydroxide, sulfide and carbon dioxide) may result in minor losses as glassy rocks devitrify, and glassy clasts may be welded by elevated temperatures in the contact zones (e.g. Christiansen and Lipman, 1966). The effects of intrusions emplaced into water-saturated successions are very different because water mobilises heat and soluble elements. Trapped seawater in submarine volcanic successions is heated by intrusions, initiating convection and metasomatic alteration in the overlying succession. Thus, almost all intrusion-related alteration in submarine volcanic successions involves some degree of metasomatism by magmatic fluid, modified seawater, or both.

Subseafloor regional hydrothermal systems Studies of the petrology, geochemistry and oxygen isotopes of hydrothermally altered volcanic and plutonic rocks from ophiolite complexes provide insight into subseafloor hydrothermal systems, fluid generation and circulation, and

the role of intrusions (e.g. Lydon and Jamieson, 1984; Alt et al., 1986; Gillis and Robinson, 1990; Bettison-Varga et al., 1992; Kelley et al., 1992). The convection cell model for hydrothermal systems and the formation of VHMS deposits is based on observations from VHMS deposits and the upper part of the Cretaceous Troodos Massif in Cyprus, where hydrothermal convection was driven by emplacement of late, high-level gabbro stocks into the fractured and permeable crust (e.g. Spooner et al., 1974; Lydon and Jamieson, 1984; Bettison-Varga et al., 1992). This model involves the circulation of seawater in approximately 10 km diameter cells to depths of 3-5 km within the crust (Fig. 6.2). Initially, increased temperatures in the host succession drive dehydration and decarbonation reactions, and fluids migrate away from the intrusion. Buoyant heated connate seawater rises through the permeable volcanic succession, drawing down cold seawater, which is heated as it descends. In this way, magma drives convective circulation of seawater between the seafloor and the intrusion (Norton, 1984; de-Ronde et al., 1994; Galley, 2003). Fluid flow is focused along joints, fractures and faults formed during extension or in response to intrusive pressures (Bettison-Varga et al., 1992). Alternatively, the multi-tiered convection model involves a high-temperature (450-700°C) cell, which circulates recycled modified seawater in plutonic rocks at depth, overlain by a low-temperature (350-400°C) cell (Gregory and Taylor, 1981; Norton et al., 1984; Alabaster and Pearce, 1985; Kelley et al., 1992). Submarine hydrothermal systems comprise three parts: a down-flow or recharge zone; a high-temperature reaction zone; and an up-flow or discharge zone (Fig. 6.3: Spooner and Fyfe, 1973; Alt, 1999). The locations of the recharge and discharge zones are commonly controlled by faults (Schardt et al., in press). Seawater percolates down through the recharge zone, and is slowly heated and chemically modified by lowtemperature reactions (White, 1970; Gibson et al., 1983; Galley, 1993; Alt, 1999). The reaction zone is a porous reservoir near the heat source where heated seawater reacts with the host rocks, exchanging some elements (Norton, 1984; de-Ronde et al., 1994; von Damm, 1995; Butterfield, 2000; Schardt etal., in press). Hot buoyant hydrothermal fluid (modified seawater) ascends rapidly to the seafloor through the discharge zone, which is characterised by cooling of the fluid, alteration of the host rock, and mineral precipitation (Skirrow and Franklin, 1994; Schardt et al., in press). The rising hydrothermal fluid cools by adiabatic decompression, conductive heat loss, and mixing with cold seawater in the shallow subsurface (Mottl, 1983; Butterfield, 2000). In well-

FIGURE 6.2 | Simple convection cell model for the genesis of the Cyprus VHMS deposits (modified after Heaton and Sheppard, 1977, and Spooner, 1977, in Lydon and Jamieson, 1984).

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 4 1

established hydrothermal systems, the discharge zone may be focused, intensely altered and veined. Surface discharge onto the seafloor may produce high-temperature (150—350°C) features such as black smokers (Goodfellow and Franklin, 1993; Rona et al., 1993). The temperature of the discharging fluids on the seafloor initially increases, and then gradually decreases to ambient temperatures, in a time scale of 100 to 10,000 years (Ohmoto, 1996). Regional hydrothermal systems are interpreted to be related to large volume intrusions, as the volume of circulating fluid in a hydrothermal system theoretically cannot be greater than the volume of the intrusion (Cathles, 1981). However, small-volume near-seafloor intrusions, which are unlikely to generate significant hydrothermal systems, may be related to larger plutons or stocks at depth that were capable of generating hydrothermal convection (e.g. Bettison-Varga et al., 1992).

6.2 | REGIONAL ALTERED ZONES ASSOCIATED WITH INTRUSIONS The products of regional-scale hydrothermal alteration systems in ancient submarine volcanic successions are recorded by cross-cutting recharge and discharge zones, and broad, regional-scale, semi-conformable altered zones or reaction zones (Galley, 1993).

Recharge zones Very little is known about altered zones associated with recharge. They are rarely recognised except in studies of modern crustal alteration beneath mid-ocean ridges (e.g. Mottl, 1983; Saccocia et al., 1994; Alt, 1999), hydrothermal

alteration studies in ophiolites (e.g. Schiffman et al., 1987; Schiffman and Smith, 1988), and studies of O- and S-isotope compositions in ancient hydrothermally altered systems (e.g. Cathles, 1993; Davidson and Kitto, 1997). Rocks in modern recharge zones are pervasively altered at low to moderate temperatures. At less than 150°C, oxidation, the fixation of alkalis (mainly Ca and Na), and Mg-metasomatism produces sericite, hematite and clays (Alt, 1999). At higher temperatures (150-350°C) anhydrite precipitates, alkalis are leached and Mg is consumed by chlorite in the rock (Alt, 1999). Schiffman and Smith (1988) proposed that the distribution of epidosites in the Troodos ophiolite represent areas of high-temperature alteration involving high fluid-rock ratios. Epidosites are granoblastic, fine- to medium-grained rocks, with little or no relict igneous textures, composed of epidote, quart and chlorite. They are inferred to record reaction zones in which circulating modified seawater reacted with host rocks to form metal-rich hydrothermal fluids, and appear diagnostic of the up-welling and deep recharge parts of the hydrothermal system beneath VHMS deposits. Co-incident whole-rock O-isotope patterns support their formation in proximal recharge zones and up-flow conduits beneath VHMS deposits. Regionally extensive, depth-dependent 618O profiles in the sheeted dyke complex reflect oxygen exchange during prograde regional hydrothermal alteration involving diffuse down-welling of cold seawater (Fig. 6.4). However, surfaces of equal whole-rock 618O are not horizontal but nearly vertical in the central epidosite zone. This suggests up-flow of hot modified seawater within the epidosite zone. The spatial association between gabbro intrusions and the epidosite zones in the sheeted dyke complex indicates a genetic link between the emplacement of these intrusions and focused high-temperature hydrothermal up-flow (Richardson et al., 1987; Bettison-Varga et al., 1992).

Discharge ZOneS Discordant footwall alteration pipes and feldspar-destructive zones that directly underlie VHMS deposits are widely interpreted as discharge zones through which metal-bearing hydrothermal fluid ascended to the seafloor (Sangster, 1972; Large, 1977; Lydon, 1984; Galley, 1993; Skirrow and Franklin, 1994; Brauhart et al., 1998). They are characterised

FIGURE 6.3 | Model of an active geothermal system illustrating the recharge, reaction or reservoir and discharge zones. Seawater is drawn down in broad recharge zones or along faults and reacts at increasing temperatures. Hightemperature reactions (>350°C) occur in the reaction zone above a subvolcanic intrusion and hot (>300°C) buoyant fluids rise towards the surface in focused or diffuse discharge zones (modified after Alt, 1995a). Not to scale.

FIGURE 6.4 | Cross-section of the Solea graben, Troodos ophiolite, Cyprus, showing surfaces of equal whole-rock d 18 0. Regionally these surfaces are subhorizontal, but in the central epidosite zone they are nearly vertical indicating up-flow of hot, modified seawater during convection. Modified after Schiffman and Smith (1988).

1 4 2 | CHAPTER 6

by Mg-Fe enrichment and Na-Ca depletion and assemblages that include chlorite, sericite, quartz or rare talc (Lydon, 1984; Eastoe et al., 1987; Skirrow and Franklin, 1994; Brauhart et al., 1998). The characteristics and compositional changes associated with discordant footwall alteration pipes are discussed in Section 7.3. Although discordant altered zones typically cut across the regional, deep, semi-conformable altered zones (Galley, 1993; Brauhart et al., 1998), in some successions, they grade laterally into deep, semi-conformable altered zones (Skirrow and Franklin, 1994; Hudak et al., 2000). Gibson et al. (2000) suggested that whether or not deep, semi-conformable altered zones are cut by or transitional with pipe-like altered zones, depends on whether the host succession (footwall) is dominated by coherent volcanic or volcaniclastic facies respectively, or timing of alteration.

Deep, semi-conformable altered zones Since they were first discussed by Franklin et al. (1981), deep, semi-conformable altered zones have been documented in the footwall beneath VHMS deposits in a variety of districts including: Matagami, Snow Lake, Noranda and Sturgeon Lake districts in Canada; Bersglagen and Skellefte districts in Sweden; Iberian pyrite belt in Spain and Portugal; Troodos Ophiolite Complex in Cyprus; Panorama district in Australia; and the Sirohi district in India (MacGeehan, 1978; Gibson et al., 1983, 2000; Lagerbald and Gorbatschev, 1985; Galley, 1993; Skirrow and Franklin, 1994; Tiwary and Deb, 1997; Brauhart et al., 1998; Bailes and Galley, 1999; Hannington et al., 2003a, 2003b). They have not been documented in eastern Australia, possibly because of structural complexities. Thus the following discussions on deep, semi-conformable altered zones are largely based on Canadian examples. Figure 6.5 depicts the characteristics and typical zonation of deep, semiconformable altered zones in the documented examples.

Deep, semi-conformable altered zones typically extend for up to 20 km laterally and 1-4 km depth beneath paleoseafloors and VHMS deposits (Gibson et al., 1983, 2000; Cathles, 1993; Galley, 1993; Skirrow and Franklin, 1994). They comprise vertically stacked, sub-horizontal altered zones (Galley, 1993; Skirrow and Franklin, 1994). Generally these are (Fig. 6.5): an upper background K-Mg metasomatic zone; a transitional Na-Mg metasomatic zone; a central silicified zone; and a basal Ca-Fe metasomatic and base metal-leaching zone (Galley, 1993). In many systems only one or two of these altered zones are recognised. The alteration minerals in the semi-conformable altered zones reflect the primary host rock composition, bulk-rock composition established during synvolcanic hydrothermal alteration, and the subsequent metamorphic grade (Paradis et al., in press). In greenschist facies metamorphosed felsic rocks, these zones are typically, from base to top: albite or carbonate zone, silica or sericite zone and sericite or chlorite zone (Gibson et al., 1983; 2000). In mafic rocks, the zones are: albite zone, silica zone and clinozoisite/epidote + quartz zone (Galley, 1993; Skirrow and Franklin, 1994; Gibson et al., 2000; Hannington et al., 2003a). At higher metamorphic grades, such as in the Snow Lake District, mineral assemblages can include kyanite, staurolite, sillimanite, chlorite, biotite, quartz, plagioclase cordierite, amphibole, epidote and garnet (Paradis et al., 1993, in press; Bailes and Galley, 1999). The semi-conformable altered zones are interpreted to be synvolcanic because they have undergone the same degree of tectonic deformation as the surrounding rocks, have prograde mineral assemblages, are spatially associated with VHMS deposits, and are commonly truncated by unaltered synvolcanic intrusions (Gibson et al., 1983; Paradis et al., 1993; Skirrow and Franklin, 1994). At Snow Lake, Paradis et al. (1993) recognised that deep, semi-conformable altered zones were superimposed on low-temperature (possibly diagenetic) altered zones and also cross cut by discordant feldspar-destructive zones associated with VHMS deposits.

PROCESS AND COMPOSITIONAL CHANGES

ASSEMBLAGE IN MAFIC ROCKS

ASSEMBLAGE IN FELSIC ROCKS

K-Mg metasomatism + Mg, K, Fe -Na, Ca, Cu, Pb, Zn, Si

Mg clays + chlorite + zeolites + Fe-oxides ± K-feldspar

Mg-clays + zeolites ± cristobalite ± adularia ± analcime ± K-feldspar

Na-Mg metasomatism + Na, Mg -Ca, Fe, Zn.Cu ±Si

Albite + quartz + sericite + Mg-chlorite ± calcite

Albite + quartz + sericite + Mg-chlorite

Silicification or sericitisation + Si, Na -Fe, Mg, Mn,Zn ±Ca

Quartz + albite

Quartz ± albite ± sericite

Ca-Fe metasomatism + Ca - M g , Mn, Na, K ± Fe, Si

Clinozoisite/epidote + quartz ± actinolite ± carbonate

Sericite + quartz ± Mg-chlorite or chloritoid + Fe-chlorite

FIGURE 6.5 | A schematic compilation of regional-scale, deep, semi-conformable altered zones and their characteristics. There is a progression, with increasing depths in submarine volcanic successions, from the background Mg-K metasomatic zone to a transitional Na-Mg metasomatic zone characterised by feldspar alteration, a central silicified zone and a basal Ca-Fe metasomatic and base-metal leaching zone that typically includes epidote or chlorite, After Gibson et al. (1983), Galley (1993), Skirrow and Franklin (1994), and Brauhart etal. (1998).

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 4 3

They suggested that regional hydrothermal alteration postdated the onset of diagenetic alteration, and pre-dated footwall alteration associated with hydrothermal alteration and mineralisation. In some Canadian examples and at Panorama in western Australia, deep, semi-conformable altered zones are gradational with discordant footwall alteration pipes suggesting that regional hydrothermal alteration was synchronous with the VHMS-related alteration (Gibson et al., 1999). Deep, semi-conformable altered zones are commonly spatially and temporally associated with subsurface synvolcanic intrusions (Galley, 1993, 2003). These intrusions can be individual granitic or porphyritic plutons or sheeted

dyke swarms (e.g. Gibson et al., 1983; de-Ronde et al., 1994; Brauhart et al., 1998). The tops of the subvolcanic intrusions and associated dykes may be included in the basal semiconformable altered zone (e.g. Galley, 1993; Brauhart et al., 1998). One of the best-documented examples of the spatial association between a subvolcanic intrusion, regional-scale semi-conformable altered zones and VHMS deposits comes from the Panorama district in Western Australia. Discordant chlorite + quartz zones directly beneath the VHMS deposits, are spatially associated with feldspar-destructive sericite + quartz zones in the top of the Strelley Granite pluton

FIGURE 6.6 | Geology and altered zones within part of the Strelley succession, Panorama district, Western Australia (modified after Brauhart et al., 1998).

1 4 4 | CHAPTER 6

1994). In the Snow Lake district Skirrow and Franklin (1994) estimated that approximately 1.1 x 107 metric tons of SiO2 was added by a minimum 12 km3 of hydrothermal fluid at 1—2 km depth. Interactions between large volumes of modified seawater and volcanic successions at depth are supported by geochemical and geophysical research at active ocean spreading ridges (e.g. Spooner and Fyfe, 1973; Bischoff andDickson, 1975). Deep, semi-conformable altered zones are characterised by mineral assemblages that reflect the reactions of glass and both primary and secondary minerals with seawater at temperatures up to 400°C (Galley, 1993).

Background K-Mg metasomatic zones

FIGURE 6.7 | Schematic section of the geology and altered zones in the Kangaroo Caves footwall succession, Panorama district, Western Australia (modified after Brauhart et al., 1998). See Figure 6.6 for legend.

(Figs 6.6 and 6.7; Brauhart et al., 1998). Faults bounding the discordant chlorite + quartz zone in the footwall of the Kangaroo Caves deposit (Fig. 6.7) controlled the distribution of the feldspar-destructive sericite + quartz zone in the Strelley Granite (Brauhart et al., 1998). The variations in alteration mineral assemblage down through the semi-conformable altered zones correspond to geochemical gradients in which there are gradual decreases in the Mg/Ca, Mg/Na and Na/Ca ratios of the altered rocks with increasing depths (Galley, 1993). Oxygen-isotope compositions suggest that the altered rocks are 18O enriched with respect to unaltered volcanic rocks (Munha and Kerrich, 1980; Barringa and Kerrich, 1984). The geochemical gradients and O-isotope data are consistent with metasomatic alteration resulting from the interaction of volcanic rocks with seawater (Muehlenbachs and Clayton, 1972; Lagerbald and Gorbatschev, 1985; Cathles, 1993). Although modified seawater is interpreted to be the main component, magmatic fluids may have also contributed to the hydrothermal fluid (Lagerbald and Gorbatschev, 1985). The spatial association between altered zones and subsurface intrusions suggests a genetic link where intrusions may have provided heat and or fluid to the hydrothermal system. The extent and intensity of deep, semi-conformable altered zones implies that very large volumes of fluid must have reacted with the host volcanic rocks (Skirrow and Franklin,

These zones are often described as the least-altered or diagenetically altered zones. At low temperatures (50-140°C) in the shallow subseafloor, the interaction of abundant seawater with the volcanic succession produces Mg-K-rich alteration assemblages (Seyfried and Bischoff, 1977; Galley, 1993). In felsic rocks these mineral assemblages include adularia and Mg-smectite, whereas in mafic rocks they are dominated by zeolites and Mg-smectite. Seawater becomes enriched in Si, Fe3+, Mn and lesser amounts of Ca, Mg and sulfur (Seyfried and Bischoff, 1977). Current mineral assemblages reflect the regional metamorphic grade. For example, at Snow Lake the diagenetically altered zone is characterised by quartz + biotite + garnet, Fe2O3, MgO and K2O gains and CaO, Na 2 O, Cu, Pb, Zn losses (Paradis et al., in press). These compositional changes are consistent with low-temperature seawater-dominated diagenesis of felsic volcanic facies to clays and zeolites (Section 5.3). The current mineral assemblage reflects the overprint of amphibolite facies metamorphism. In the Panorama district the background alteration mineral assemblage includes feldspar + calcite ± ankerite + quartz + pyrite ± sericite consistent with greenschist facies metamorphism of clays and zeolites in felsic volcanic rocks (Brauhart et al., 1998).

Transitional zone or Na-Mg metasomatic zones With increasing stratigraphic depth there is a transition from K-rich zones to Na-rich zones (Munha et al., 1980; Munha and Kerrich, 1980; Lagerbald and Gorbatschev, 1985; Schiffman and Smith, 1988; Brauhart et al., 2001). The NaMg metasomatic zones are characterised by the occurrence of feldspar, usually albite. Greenschist facies assemblages typically include chlorite, sericite, albite, epidote and quartz in mafic rocks, and albite, quartz, sericite ± chlorite ± carbonate (calcite or dolomite) in felsic rocks (Gibson et al., 2000). In the Panorama district, felsic rocks in the feldspar zone have the assemblage K-feldspar or albite + sericite + quartz + ankerite + leucoxene ± pyrite (Brauhart et al., 1998). The transition to Na-rich zones reflects the behaviour of Na and K in seawater at elevated temperatures. Between 140° and 200°C there is a transition between K- and Na-metasomatism (Seyfried and Bischoff, 1977). Munha et al. (1980) suggested that at lower temperatures (<150°C), Na in glass is exchanged for K in seawater, resulting in precipitation of K-rich zeolites

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 4 5

and possibly K-feldspar. At higher temperatures, K in the rock is exchanged for Na in seawater, resulting in the formation of albite. Thus at moderate temperatures (140—300°C), metasomatic reactions between modified seawater and the volcanic succession result in Na-Mg alteration assemblages (Seyfried et al., 1988). Regardless of the rock type, alteration mineral assemblages include Mg-smectite + chlorite + quartz + albite, and compositional changes are Na 2 O and MgO gains, and CaO, Zn and Cu losses (Gibson et al., 2000). The removal of Mg from seawater lowers the pH of the fluid and seawater evolves from a moderately alkali, Mg-K-Na-SO4rich fluid to a hot acidic Si-Na-Ca-rich hydrothermal fluid (Bischoff and Seyfried, 1978; Seyfvied et al., 1988). Silicified zones In greenschist facies felsic and some mafic rocks, the central altered zone is typically silicified, with assemblages of quartz

A. Central silicified zone Moderate, patchy quartz alteration in this andesite from the central silicified zone resulted in a fine-grained, pale rock, which resembles a rhyolite. Amulet Formation, Noranda district, Buttercup Hill, Canada.

B. Epidote + quartz zone This approximately one metre-wide patch of epidote + quartz alteration facies in the upper Amulet andesite has an irregular shape typical of patchy alteration in the basal episite + quartz zone. The groundmass has been pervasively epidote + quartz altered. Amulet Formation, Noranda district, Canada.

C. Epidote + quartz zone Amygdales in this patch of epidote + quartz-altered andesite from the basal epidote + quartz zone have amoeboid shapes and were lined with Fe-oxides and filled with epidote + quartz. Amulet Formation, Noranda district, Canada.

FIGURE 6.8 | Photographs from deep semi-conformable alteration zones in the Noranda district, Canada.

+ plagioclase or albite (Skirrow and Franklin, 1994; Gibson et al., 2000). In some mafic rocks, the central zone is sericitic, dominated by sericite + quartz ± chlorite (Gibson et al., 2000). Central silicified or sericite zones overprint regional albite zones (Galley, 1993). At Snow Lake the silicified zone is spatially and temporally associated with VHMS deposits and is zoned laterally from a silica zone to epidote and FeMg-metasomatic zones (amphibolite grade; garnet + chlorite ± biotite ± staurolite) (Skirrow and Franklin, 1994). Silicified zones are typically spatially associated with synvolcanic intrusions and the intensity and pervasiveness of alteration increases with proximity to the intrusions (Skirrow and Franklin, 1994; Paradis et al., in press). Silicified zones commonly contain patches of quartz + feldspar-altered rock, quartz-altered clasts in volcaniclastic facies, and quartz veins (e.g. Fig. 6.8A: Gibson et al., 1983; Skirrow and Franklin, 1994). The patches of quartz + feldsparaltered rock are restricted to flow-top breccias, and flow-

1 4 6 | CHAPTER 6

banded and vesicular lavas (Gibson et al., 1983). In mafic volcanic rock the originally glassy groundmass, elsewhere typically altered to chlorite, is altered to quartz in this zone (Gibson et al., 2000). This led to intensely silicified andesites in the Noranda sequence being misinterpreted as rhyolite: the Amulet rhyolite (Gibson et al., 1983). With increasing depth, seawater carries larger amounts of Si as Si solubility increases with temperature and pressure, and is enhanced in NaCl solutions or where the fluid is in contact with free Si or glass (Kennedy, 1950; Fournier, 1985). The Si-rich hydrothermal fluid is rapidly heated beyond the temperature of the quartz solubility maximum: 340-400°C at pressures below 900 bars (Fig. 6.9: Kennedy, 1950; MacGeehan, 1978; Skirrow and Franklin, 1994). The result is gains in SiO2 and Na 2 O, due to the precipitation of silica within pore spaces and albitisation, and losses in FeO, MgO, CaO, K 2 O, MnO and other metals from the volcanic rocks (Gibson et al., 1983, 2000; Lagerbald and Gorbatschev, 1985; Galley, 1993; Skirrow and Franklin, 1994). The Fe3+, Mg and possibly Zn leached from the silicified zone may have been transported laterally away from this environment, thereby producing semi-conformable Fe-Mg-metasomatised zones (Skirrow and Franklin, 1994). A second silicified zone is common directly beneath the seafloor in the Snow Lake, Noranda and Matagami Lake districts, where it is directly overlain by exhalites. This nearseafloor, silicified zone is related to low-temperature silicification during the hydrothermal alteration and devitrification of glass in cooling pillow basalts and andesites (Skirrow and Franklin, 1994; Galley et al., 2002).

FIGURE 6.9 | Calculated solubilities for quartz in water up to 900°C at pressures between 200 and 1000 bars (after Fournier, 1985). The shaded area outlines the conditions for retrograde solubility.

Basal Ca-Fe metasomatic zones Mineral assemblages in basal semi-conformable altered zones are dependent on the host-rock composition and porosity. In felsic rocks, basal zones are typically sericite or chlorite zones, whereas in mafic rocks they are clinozoisite or epidote + quartz zones (Gibson et al., 2000; Hannington et al., 2003a). Sericite zones are characterised by sericite + quartz ± Mgchlorite assemblages (Gibson et al., 2000). Chlorite zones are characterised by chloritoid + Fe-chlorite ± Fe-carbonate assemblages (Gibson et al., 2000). Epidote + quartz zones are characterised by epidote + quartz + calcite + actinolite + chlorite assemblages (Galley, 1993). Two alteration textures are persistent in epidote + quartz zones: pervasive and patchy. Pervasive epidote + quartz occurs as selective pervasive replacement of plagioclase phenocrysts or the groundmass by epidote, crystallisation of fine quartz patches in the groundmass (e.g. Fig. 6.8B), and replacement of Fe-Ti-oxide grain rims by sphene (Skirrow and Franklin, 1994). Patchy epidote + quartz occurs as less than 1 cm to 2 m irregular ovoids or amoeboid patches that infill vesicles and gas cavities within mafic lavas (e.g. Fig. 6.8C: Gibson et al., 1983; Skirrow and Franklin, 1994). These patchy textures are similar to the epidote + quartz metadomains described in spilites by Smith (1968, 1974, 1977; Smith et al., 1982). These basal zones may grade into the discordant (discharge) altered zones that cut through the overlying semiconformable and background altered zones (Brauhart et al., 1998). Epidote + quartz zones are enriched in CaO and Sr and depleted of MgO, Na 2 O, K 2 O, FeO ± MnO, Ba and base metals (MacGeehan, 1978; Gibson et al., 1983; Richardson et al., 1987; Schiffman and Smith, 1988; Skirrow and Franklin, 1994). Unlike the Canadian examples, the epidote + quartz zone in the Panorama district does not appear to have been the source of leached base metals (Brauhart et al., 2001). At high temperatures (300-500°C), Ca-Fe-S-base metalrich hydrothermal fluid reacts with the volcanic succession and possibly also with parts of the subsurface intrusion forming mineral assemblages typical of the basal Ca-Femetasomatic zones. Experimental work suggests that epidote + quartz alteration involved modified seawater (Mg-depleted, Ca-Na-K-Cl fluid) at temperatures of 35O-5OO°C and low water-rock ratios of less than three (Gibson et al., 2000). This zone is interpreted to represent the high-temperature interaction between modified seawater and the host volcanic facies to form metal-rich hydrothermal fluid at the deepest part of the hydrothermal convection system. Hence, it represents the roots of up-welling fluid discharge zones (Galley, 1993). Alternatively, Smith (1968, 1977) interpreted these district-scale zones of albitised basalt with Ca-rich epidote + quartz and pumpellyite + quartz metadomains to have formed during heterogenous burial metamorphism where local fluid flow promoted redistribution of elements. In some cases, he noted that the alteration was focussed suggesting that it was related to local hydrothermal systems and subseafloor fluid circulation.

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 4 7

Altered zones as part of a regional hydrothermal system Deep, semi-conformable altered zones superficially resemble regional diagenetic or metamorphic facies. This is because they are regionally extensive, vertically stacked altered zones with mineral assemblages similar to those formed during high-temperature diagenesis, regional greenschist facies metamorphism and hydrothermal seafloor alteration (Galley, 1993; Paradis et al., 1993). Discriminating between these processes and their products is difficult in submarine volcanic successions. The differences are essentially related to timing, temperatures, and fluid-rock ratios. In reality, there is a progression from diagenesis to isochemical metamorphism with increasing temperature and depth during burial (Fig. 6.10A). Porosity, permeability and fluid-rock ratios decrease with depth in diagenetic-metamorphic systems, thereby inhibiting the degree and pervasiveness of metasomatism at temperatures above 150°C. Typically once the temperature and pressure realm of metamorphism has been reached, the porosity and permeability of the host succession has been dramatically reduced, fluid flow inhibited and metasomatic reactions ceased. Gibson et al. (2000) suggested seawaterdominated diagenesis might also progress to deep regional hydrothermal alteration with increasing temperature and depth in shallow subseafloor hydrothermal systems (Fig. 6.1 OB). Deep regional hydrothermal alteration is interpreted to involve metasomatic reactions between seawater and the volcanic succession at temperatures transitional with diagenesis and greenschist facies metamorphism (i.e. 150— 400°C) (Galley, 1993). Although the processes of diagenesis and deep regional hydrothermal alteration are very similar, and both involve reactions between seawater (or modified seawater) and volcanic successions at increasing temperatures and depths, deep, semi-conformable altered zones are inconsistent with the diagenetic-metamorphic system. They

have anomalous mineral assemblages and alkali contents for igneous rocks (e.g. Fig. 6.11; Hughes, 1973), which suggest metasomatic rather than metamorphic origins (Gibson et al., 1983; Galley, 1993). Gibson et al. (1983) documented a vertically stacked sequence of altered zones in the Noranda sequence, from top to bottom: albite zone (spilites), silicified zone, and epidote + quartz zone (Fig. 3.16). This is consistent with a progression from low-temperature diagenesis to moderate- and hightemperature metasomatism with depth in the stratigraphy. Munha and Kerrich (1980) referred to this process of temperature and hence depth dependent metasomatism as 'hydrothermal metamorphism', a term that reflects the combined processes that operated in the subseafloor. In areas of volcanic and hydrothermal activity, it is probable that there is a spectrum of alteration between diagenesis and hydrothermal alteration where these processes operate in combination. Hydrothermal activity in the depositional basin would accelerate and intensify the process of diagenesis by contributing additional fluid and heat, and by promoting convection (Iijima, 1974; Marsaglia and Tazaki, 1992). Deep, semi-conformable altered zones are assumed to be the products of hydrothermal alteration within regional subseafloor hydrothermal systems (Gibson et al., 1983; Galley, 1993). These hydrothermal systems involve the largescale convection of modified seawater through the permeable volcanic successions (Spooner and Fyfe, 1973; Galley, 1993). The distribution of altered zones and spatial association with subsurface intrusions suggests that subsurface intrusions, augmented by heat from the cooling volcanic succession, may be the driving force for hydrothermal convection (Campbell et al., 1981; Lesher et al., 1986; Cathles, 1993; Galley, 1993; Skirrow and Franklin, 1994). Where the upper contacts of subsurface intrusions are sub-parallel to the volcanic-strata, the overlying isotherms are also semi-conformable with the strata and progressive temperature-dependent seawater-rock

FIGURE 6.10 j The relationships between fluid convection, diagenesis, metamorphism and regional hydrothermal alteration in submarine volcanic successions that host VHMS deposits. (A) Diagenetic-metamorphic system, where there is a progression from diagenesis to isochemical metamorphism with increasing temperature and depth in the subseafloor. This transition reflects the maximum depth to which seawater circulates and reacts with the host rocks. (B) Diagenetic-hydrothermal system, where a subsurface intrusion promotes deep circulation of fluid via the recharge, reservoir and discharge zones. The depth progression from diagenesis to regional hydrothermal alteration (deep, semi-conformable alteration) is dependent on temperature and fluid circulation.

1 4 8 | CHAPTER 6

FIGURE 6.11 | Alkali ratios for altered andesite samples from the Amulet rhyolite, Noranda district, Canada (after Gibson et al., 1983). Fields for the primary and metasomatised (albite-altered) andesites and basalts are from Hughes (1973).

reactions form a series of semi-conformable altered zones (Galley, 1993). In axial mid ocean ridge hydrothermal systems, downwelling seawaters traverse extremely steep temperature gradients in the upper crust, from less than 50°C near the seafloor to more than 250°C at 1-2 km depth (Mottl, 1983). Thus, vertically stacked deep, semi-conformable altered zones result from metasomatic reactions that take place at progressively higher temperatures with depth in the succession (Galley, 1993). The decrease in pervasiveness of alteration, from widespread nearly uniform diagenesis to more restricted and patchy silicification and Ca-Fe metasomatism, may reflect decreasing permeability with depth. The distribution, relative timing, and spatial association with VHMS deposits suggest a genetic link between regional hydrothermal alteration and mineralisation. Some authors have proposed that the deep, semi-conformable altered zones acted as reservoirs from which metals and sulfur were leached (e.g. Gibson et al., 1983; Lagerbald and Gorbatschev, 1985; Galley, 1993; Skirrow and Franklin, 1994). As such they represent much larger exploration targets than the discordant altered footwall zones (Galley, 1993).

6.3 | ALTERED ZONES WITHIN INTRUSIONS Intrusions are commonly modified by deuteric or hydrothermal alteration associated with emplacement and may subsequently undergo diagenesis, regional metamorphism or hydrothermal alteration.

Deuteric alteration Deuteric alteration, also referred to as autohydration or autometamorphism, is the alteration of recently crystallised magma by trapped magmatic fluid exsolved from the same cooling magma (Honnorez et al., 1979; Destrigneville et al., 1991). It has been recorded in intrusions and lavas in both

subaerial and submarine successions (Honnorez et al., 1979; Bohlke et al., 1980; McConnell et al., 1995). Sederholm (1929) originally defined deuteric alteration as the alteration that takes place 'in direct continuation of the consolidation of the magma' and thus it is considered a magmatic alteration process. It is the earliest alteration style and is a shortlived process, typically occurring as intrusions cool from temperatures of several hundred degrees centigrade (Ade-Hall etal., 1968; Honnorez etal., 1979). Small volume synvolcanic intrusions typically cool too rapidly to experience deuteric alteration (cf. Gromme et al., 1969). In contrast, granitoids and large-volume sills may have altered zones in their upper parts resulting from reactions between rising magmatic fluids and the cooling intrusions (e.g. Fig. 6.1A). Deuteric textural changes are minimal (Wilshire, 1959). Phenocrysts, particularly feldspars and mafic minerals, such as pyroxene or olivine, may be pseudomorphed by amphibole, chlorite or smectite (Fuller, 1938; Bohlke et al., 1980; Destrigneville et al., 1991). Open spaces, such as vesicles, mariolitic voids, and quench fractures, are lined or filled with smectite, zeolites, carbonate, biotite, chlorite and oxides (Wilshire, 1959; Furbish and Schrader, 1980; Destrigneville et al., 1991). High-Ti minerals, such as titanomagnetite, are oxidised and altered to low-Ti minerals, such as ilmenite ± hematite (Butler and Burbank, 1929; Ade-Hall et al., 1968; Surdam, 1968; Sherwood, 1988). Deuteric alteration does not involve major chemical composition changes; some components may be locally redistributed at sub-millimetre scales or undergo oxidation state changes (e.g. Fe2+/Fe3+ ratio, Scott and Hajash, 1976) that may alter rock thermomagnetic properties (Ade-Hall et al., 1968; Sherwood, 1988). The changes are quantitatively unimportant when compared to the products of long-lived diagenesis and hydrothermal alteration and may be difficult to distinguish from those of other alteration styles.

Hydrothermal alteration Altered zones within synvolcanic intrusions may also result from reactions with seawater or modified seawater circulating through the intrusion, either during the prograde hightemperature stage of hydrothermal activity or during cooling. If the hydrostatic pressure is high enough (at sufficient depths) seawater will be forced into thermal contraction fractures in the cooling intrusion (Burnham, 1979). The time interval for seawater-intrusion interaction may be limited by the rapid development of a local intrusion-related hydrothermal system in the host succession, which would result in the lithification and filling of primary pore space inhibiting fluid flow. Thereafter, episodic seawater-intrusion interaction would occur only if the fluid pressure exceeds the tensile strength causing the altered rock adjacent to the contact to fracture (Secor, 1965, in Fournier 1985; Phillips, 1973; Henley and McNabb, 1978). Alteration follows the advancing front of brittle fracturing to deeper and deeper levels within the intrusion (Burnham, 1979; Giggenbach, 1997). The resulting altered zones may be pervasive, occur along cooling fronts or more commonly as selective-pervasive alteration adjacent to fractures or veins. Alteration minerals fill vesicles, miarolitic voids and fractures, cement hydraulic

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 4 9

breccias, and pseudomorph primary magmatic minerals (Mevel and Cannat, 1991; Gillis et al., 1993; Kelley and Gillis, 1993; Nehlig, 1993; Davidson, 1998; Galley, 2003). Polya et al. (1986) and Davidson (1998) described proximal zones of hydrothermal alteration within the Cambrian Murchison Granite, western Tasmania. They described narrow, texturally destructive intense K-feldspar zones, associated with calcite veins, irregularly distributed in the margins of the granite and patchy selective-pervasive chlorite zones with chlorite pseudomorphs after biotite and hornblende or patches of chlorite ± pyrite ± sphene. The significant mineralogical, compositional and isotopic changes associated with proximal hydrothermal alteration within intrusions are consistent with seawater-rock interaction (Gregory and Taylor, 1981; Stakes and Taylor, 1992; Cathles, 1993; Galley, 2003). Galley (2003) identified three types of early hydrothermal and magmatic alteration within subvolcanic intrusions in the Snow Lake, Noranda and Sturgeon Lake districts. The earliest, a greenschist alteration facies in quartz diorite intrusions, is manifest as pervasive epidote + quartz, and epidote + actinolite + quartz + albite + magnetite ± sulfides, which replaced primary minerals and infilled miarolitic cavities, vesicles and fractures. Mass change calculations suggest CaO, Sr, Pb and CO 2 were gained and Fe2O3, MgO, Cu, Zn, Mo, Na 2 O, K2O and Ba lost. Galley (2003) interpreted this facies to be the product of hightemperature hydrothermal-magmatic alteration resulting from emplacement of quartz-diorite intrusions into a seawatersaturated succession. The second chlorite-rich alteration facies is characterised by quartz + chlorite + sericite and chlorite + sulfide-filled fractures and vein selvages. It is most intense near the margins of intrusions and directly beneath VHMS deposits. The chlorite-rich zones gained Fe2O3, MgO, Cu ± Pb, K2O, Ba, and Zn, and lost CaO, Na 2 O, Sr ± SiO2, Ba and CO 2 and are interpreted to result from hydrothermal alteration (Galley, 2003). Overprinting both of these zones is a biotite-rich alteration facies associated with silicification, and Cu-Mo-rich veins and breccia, which is interpreted as a magmatic alteration facies associated with late stage dykes (Galley, 2003). Hydrothermal alteration may also result from the absorption of fluid from and assimilation with sediment inclusions incorporated into the magma as it was emplaced into wet unconsolidated sediment. Wilshire and Hobbs (1962) described hydrothermal alteration in the margin of a peperitic latite intrusion in a submarine volcaniclastic succession, near Port Kembla in New South Wales. Alkali feldspar + chlorite + carbonate-rich zones coincide with abundant sediment inclusions and quench fractures in the margin of the intrusion, and the sedimentary inclusions have been chlorite ± zeolites ± carbonate altered. The altered latite gained Na 2 O and volatiles, and lost SiO2, A12O3, Fe2O3, K 2 O, MgO and CaO, whereas the sediment inclusions lost Na2O and volatiles, and gained CaO and MgO ± K 2 O.

6.4 | CONTACT ALTERED HALOS AROUND INTRUSIONS All magmatic intrusions transfer heat; they have thermal impacts on the enclosing rocks or sediments and may induce compositional changes. Contact alteration is used here as a non-genetic term referring collectively to the processes of contact metamorphism and contact hydrothermal alteration. Contact or thermal metamorphism involves changes in rock texture and mineralogy of the immediate host rock as a result of temperature increase (Yardley, 1989). The increased temperature drives dehydration and decarbonation reactions, and fluid migration away from the intrusion (Blatt et al., 1972; Manning and Bird, 1991). Only small volumes of H 2 O- and CO2-rich fluids are generated from these reactions and thus the metasomatic effect of contact metamorphism is negligible (Rose and Burt, 1979). Contact metamorphism typically results in only local remobilisation but extensive static recrystallisation of existing minerals or components. Contact hydrothermal alteration involves a substantial volume of heated fluid, typically comprising trapped seawater and pore fluid with or without magmatic fluid derived from the intrusion, which circulates through and reacts with the host facies (MacGeehan, 1978; Taylor and Forester, 1979; Polya et al., 1986; Galley, 2003). This promotes textural, mineralogical and compositional changes in the host facies. In submarine volcanic successions, abundant trapped seawater means that isochemical thermal metamorphism is rare. In addition, interaction between hot magma and wet unconsolidated sediment can result in: peperitic contacts, fluidisation of sediment (e.g. Schmincke, 1967; Kokelaar, 1982), fluid expulsion, induration (e.g. Einsele et al., 1980), secondary welding (e.g. Ito et al., 1984, in Kano, 1989), brecciation of host rock, local or regional hydrothermal alteration, quenching of the intrusion and magma-host rock assimilation (e.g. Wilshire and Hobbs, 1962; Puffer and Benimoff, 1997; WoldeGabriel et al., 1999).

Contact altered zones Contact altered zones are spatially associated with intrusion margins and may surround the intrusion as halos or aureoles. Successive contact altered zones reflect progressive changes in temperature or temperature and chemical conditions in the host succession away from the intrusion (Rose and Burt, 1979; Einsele et al., 1980; Yardley, 1989). Contact altered halos may vary in thickness from a few millimetres at the margins of thin, shallow synvolcanic sills (e.g. Einsele, 1985; Boulter, 1993; Skirrow and Franklin, 1994) to several kilometres around large subvolcanic plutons or intrusive complexes (e.g. Boulter, 1993; Schweitzer and Hatton, 1995; Galley, 2003). They may comprise one lowgrade altered zone, a sequence of roughly concentric altered zones, a series of asymmetric altered zones or overprinting altered zones (Fig. 6.12). Grades and mineral assemblages of contact altered zones vary considerably, reflecting: temperature and compositional differences between the host succession, the intrusion, and any fluid; duration of the alteration system; emplacement depth;

150 | CHAPTER 6

FIGURE 6.12 | Cartoons of the variety of contact altered zones around intrusions. (A) Cross-section showing two sills that were emplaced into wet unconsolidated sediment at shallow levels beneath the seafloor. The sills both have single low-grade indurated zones, which have a lower porosity than the surrounding host turbidites (after Einsele et al., 1980). (B) Schematic crosssection through roughly concentric zeolite and clay zones around a granite emplaced into felsic volcanic fades in the Green Tuff Belt, Japan. These zones are, from the intrusion outwards: a zeolite zone, devitrified zone, and leastaltered zone (after Utada, 1991). (C) Schematic section of the asymmetric altered halo around the Rustenberg Layered Suite intrusions in the Rooiberg felsic volcanic rocks of the Bushveld Complex, Africa (after Schweitzer and Hatton, 1995). Above the intrusion is a thick (>1.4 km) halo comprising biotite hornfeis and quartz + sericite + albite zones, which are enriched in K2O, MgO and base metals. Beneath the intrusion is a thinner (<400 m) granoblastic zone, in which primary volcanic textures are overprinted by metamorphic textures without significant compositional changes. Schweitzer and Hatton (1995) postulated that the reason for the asymmetrical zonation was that heated fluid convected freely above the intrusion and hydrothermal alteration dominated, whereas buoyant convection was inhibited beneath the intrusion acting as a seal, and thermal metamorphism dominated. (D) Map view of the prograde olivine, pyroxene and actinolite + chlorite zones associated with emplacement of the Skaergaard intrusion into mafic volcanic rocks, east Greenland (after Manning and Bird, 1991,1995). In the outer pyroxene zone and adjacent to fractures in the pyroxene and olivine zones, high-temperature mineral assemblages are overprinted by actinolite + chlorite, suggesting retrograde metamorphism occurred as temperatures dropped and cooler hydrothermal fluids migrated inwards through fractures.

rate of cooling; and the subsequent regional metamorphic grade. Although it is difficult to generalise about the mineralogy of contact altered zones, there are a few indicator minerals that are almost exclusively generated by intrusionrelated hydrothermal alteration (i.e. minerals associated with magmatic systems such as biotite, diaspore, fluorite, kaolinite, magnetite, pyrophyllite, rutile, topaz and tourmaline). Typically altered halos associated with small-volume intrusions emplaced at shallow depths below the seafloor comprise low-grade altered zones adjacent to the intrusion, which grade into partially altered zones at the peripheries. Low-grade altered zones may be manifest as indurated sediment (e.g. Einsele et al., 1980; Kano, 1989), fused glass (e.g. Ross and Smith, 1960; Smith, I960; Christiansen and Lipman, 1966; Schmincke, 1967; McPhie and Hunns,

1995), devitrified glass (e.g. Schweitzer and Hatton, 1995; WoldeGabriel et al., 1999), palagonite or clay minerals in mafic volcanic rocks (e.g. Upton and Wadsworth, 1970; Jakobsson, 1972; 1978) or zeolite and clay minerals in felsic volcanic rocks (e.g. Iijima, 1978; Utada, 1991). In contrast, high-grade altered zones tend to be associated with large volume subvolcanic intrusions and include hightemperature (up to 1000°C) mineral assemblages. For example, Seki et al. (1969) reported five high-grade altered zones around a large intrusion in the Neogene Green Tuff Belt, Japan. From the contact to the margin they were: amphibole zone, actinolite zone, pumpellyite + prehnite + chlorite zone, laumonite + mixed-layer chlorite zone, and clinoptilolite + vermiculite zone.

SYNVOLCANIC INTRUSION-RELATED ALTERATION I 1 5 1

Indurated or fused zones Thin contact altered zones of fused or secondary welded volcanic glass are common adjacent to intrusions and lavas in subaerial volcanic successions (e.g. Ross and Smith, I960; Smith, 1960; Christiansen and Lipman, 1966; Schmincke, 1967; WoldeGabriel et al., 1999), and also occur around intrusions in ancient submarine volcanic successions (e.g. Ito etal., 1984, in Kano, 1989; McPhie and Hunns, 1995). Christiansen and Lipman (1966) used the term fused for the induration and deformation of glassy clasts resulting from heating by adjacent lava, but emphasised that the term should not be taken to imply that melting (fusion) had occurred. They described altered subaerial tuffs adjacent to the Combs Peak rhyolite lavas and domes near Fortymile Canyon, southern Nevada. Three altered zones were developed parallel to the lava contact: an outer red zone characterised by the oxidation of glass, a middle indurated or partially fused zone and an inner densely fused zone characterised by fiamme and eutaxitic texture (Figs 6.13 and 6.14). In this case, eutaxitic texture was interpreted to result from the re-heating and accompanying load compaction of originally glassy pumice clasts in tuffs beneath the lava as a result of its emplacement (Christiansen and Lipman, 1966). The minimum temperature required for this partial welding is 535°C (Smith, 1960). Typically, indurated or fused zones closely parallel lava or intrusion contacts and may be several millimetres to tens of metres thick (e.g. Christiansen and Lipman, 1966; Einsele, 1985; Keating and Geissman, 1998). They are commonly associated with thin (< 1—100 m) intrusions that have peperitic or irregular contacts indicating emplacement into wet unconsolidated sediments (Kokelaar, 1982; Branney and Suthren, 1988; McPhie and Hunns, 1995; Keating and Geissman, 1998). Induration of sediment adjacent to contacts and around juvenile clasts in peperite is typically accompanied by changes in colour associated with thin (cm scale) carbonate, quartz or Fe-oxide altered halos (Fig. 6.15A: Schmincke, 1967; Kokelaar, 1982; Kano, 1989; Hunns and McPhie, 1999). The most significant textural changes in this zone are contact-parallel fiamme and eutaxitic textures in pumice-rich facies (Fig. 6.15B: McPhie and Hunns, 1995).

Devitrified zones Devitrified zones are characterised by high-temperature devitrification textures such as spherulites, lithophysae and micropoikilitic texture (Christiansen and Lipman, 1966; McPhie and Hunns, 1995). It is important to note that devitrification textures generated from re-heating of glassy volcanic facies by intrusions are indistinguishable from those formed during first cooling (Lofgren, 1971a, 1971b). Narrow zones oriented parallel to intrusion contacts may be completely or partially devitrified (e.g. Keating and Geissman, 1998; WoldeGabriel et al., 1999). They may overprint fused zones. For example, Christiansen and Lipman (1966) described superposition of three devitrified zones on to three fused zones in bedded rhyolitic tuffs (Fig. 6.14): an outer porous glassy zone, a middle dense glassy zone (vitrophyre that is commonly perlitic), and an inner crystalline zone with microlites, spherulites and lithophysae.

FIGURE 6.13 | Distribution of the fused zone adjacent to the Combs Peak rhyolite, near Fortymile Canyon, southern Nevada (modified after Christiansen and Lipman, 1966).

FIGURE 6.14 | Idealised relationships between the Combs Peak rhyolite (Nevada), the three fused zones and overprinting devitrified zones (modified after Christiansen and Lipman, 1966).

Compositional changes associated with devitrification are usually negligible. WoldeGabriel et al. (1999) found that devitrification in felsic volcaniclastic rocks in a 10m thick contact zone around a basaltic intrusion at Grants Ridge, New Mexico, was associated with minor gains in K2O and losses in H 2 O, Na 2 O, F, Fe2O3. The margins of the intrusion were slightly enriched in SiO2, K2O and P 2 O 5 and depleted in Fe2O3. They concluded that the thermal effects of the intrusion induced devitrification, dehydration and vapourphase expulsion in the contact zone. Vapour-phase expulsion of fluoride, chloride, hydroxide, sulfide, and CO 2 from silicic glass may have been responsible for the subtle chemical variations during devitrification (cf. Weaver et al., 1990).

Zeolite, clay or palagonite zones Low-temperature altered zones characterised by palagonite, zeolite and clay minerals in mafic volcanic rocks (e.g. Upton and Wadsworth, 1970), and zeolite and clay minerals in felsic volcanic rocks (Utada, 1991) are common around shallow synvolcanic intrusions in submarine volcanic successions. Mineral assemblages in these zones typically reflect the host rock compositions. For example, altered zones around granitoids in the felsic volcanic rocks of the Green Tuff Belt contain calcic zeolites (Iijima, 1978; Utada, 1991). In contrast, palagonite dominates altered zones around dykes in

1 5 2 | CHAPTER 6

A. Indurated siltstone in peperite The irregular clasts of indurated and silicified siltstone (grey) are mixed with feldspar-phyric rhyolite (green) clasts in this peperitic contact between rhyolite and siltstone. Away from the rhyolite contact, the host siltstone is green-grey, but fades to cream or pale green silicified siltstone in a zone about 1-2 cm wide adjacent to the rhyolite clasts in the peperite. This local colour change and silicification result from the thermal metamorphism of the unconsolidated silt in contact with hot rhyolite. Early Permian Berserker beds, Mount Chalmers district, Queensland.

B. Fused pumice breccia Well-developed fiamme (F) and eutaxitic texture characterise the fused zone in this pumice breccia immediately adjacent to a rhyolitic sill. Away from the rhyolite, fiamme in the pumice breccia are indistinct and parallel to bedding, whereas in the fused zone they parallel the pumice breccia-rhyolite contact. The fiamme and eutaxitic texture result from the partial welding and compaction of glassy pumice clasts during heating associated with emplacement of the rhyolite (McPhie and Hunns, 1995). Early Permian Berserker beds, Mount Chalmers district, Queensland.

Figure 6.15 | Photographs of hand specimens from the indurated and fused zones adjacent to rhyolite sills near the IVIount Chalmers VHMS deposit, Queensland.

submarine basaltic hyaloclastite at Surtsey (Jakobsson, 1972, 1978; Jakobsson and Moore, 1986), and chabazite, analcime, thomsonite, mesolite, phillipsite and natrolite characterise the zeolite zone associated with a swarm of sills in basaltic lavas and breccias at Piton des Neiges volcano on Reunion Island (Lacroix, 1936; Upton and Wadsworth, 1970). Zeolites in these zones may be accompanied by chlorite, epidote, carbonate and clay minerals. Compositional changes in the zeolite, clay or palagonite zones include K2O and MgO gains, and SiO2 and CaO losses (Hart, 1969; Thompson, 1973; Honnorezetal., 1979). These are consistent with low-temperature (<150°C) reactions with seawater promoting oxidation, fixation of alkalis, and exchange of seawater-Mg for rock-Ca to form smectite (Alt, 1999).

Rustenburg Layered Suite in the felsic Rooiberg volcanic rocks of the Bushveld Complex (Fig. 6.12C). The asymmetric aureole contains a biotite hornfels zone (immediately above the Rustenburg Layered Suite), and an overlying quartz + sericite + albite zone, which grades up into least-altered, devitrified volcanic rocks. In the quartz + sericite + albite zone, hornblende or chlorite replaced mafic phenocrysts, and quartz + chlorite + epidote replaced the glassy groundmass. Primary compositions may be significantly modified in greenschist facies zones. They are commonly enriched in K2O and MgO and depleted in CaO, Fe2O3, Na 2 O and MnO (Schweitzer and Hatton, 1995; Large et al., 1996). The behaviour of SiO2 is variable. The mineralogical and compositional changes reflect high-temperature (300-450°C) seawater-rock interactions similar to some proximal altered zones associated with VHMS ore deposits (Galley, 2003).

Greenschist facies zones Silicified zones Synvolcanic granitoids, large composite intrusions and clusters of sills or dykes commonly have altered zones with epidote-, chlorite-, sericite-, biotite- or K-feldspar-bearing mineral assemblages characteristic of greenschist facies metamorphism (Polya et al., 1986; Boulter, 1993; Neuhoff et al., 1997; Galley, 2003). For example, Schweitzer and Hatton (1995) described a 1.4 km thick greenschist facies aureole above the mafic

Silicified zones are typically pale grey in colour and can be massive pervasive or patchy in texture, filling vesicles and fractures, or cementing breccias (Humphris and Thompson, 1978; Skirrow and Franklin, 1994; Gifkins and Allen, 2001). They comprise chalcedony, cristobalite, quartz, quartz + feldspar, or quartz + sericite dominated alteration mineral assemblages.

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 5 3

Skirrow and Franklin (1994) described 10 cm to 2 m thick silicified contact aureoles associated with unaltered plagioclase- and quartz + plagioclase-phyric porphyry dykes in the submarine volcanic rocks beneath the Chisel Lake VHMS deposit in the Snow Lake district. The weakly silicified mottled zones consist of irregular light grey patches of quartz + plagioclase + hornblende + magnetite ± biotite, which coalesce into massive quartz + plagioclase rock in intensely silicified zones. Compositional changes include gains in total mass and SiO2, which may be accompanied by gains in K2O or Na 2 O, and losses in Fe2O3, MgO, CaO and Zn (MacGeehan, 1978; Skirrow and Franklin, 1994; Gifkins and Allen, 2001). Silicification is a common feature of hydrothermal alteration and incorporates both the addition of Si (largely as vein infill) and the redistribution of Si that was originally in glass or cristobalite (Henley and Ellis, 1983). Circulating heated seawater can leach Si from the intrusion or felsic volcanic glass in the host succession, resulting in a solution supersaturated with Si. Silica precipitation from this solution can occur by several mechanisms: cooling by conduction or mixing, decompression associated with boiling, heating into the temperature range for retrograde Si solubility, or a pH change (Dickson and Potter, 1982; Fournier, 1985). The solubility of Si generally increases with increasing temperature (Fig. 6.9); however, if a supersaturated solution is heated at constant pressure (<900 bars) it will either boil or reach a solubility maximum and may precipitate quartz upon further heating (Fournier, 1985). A supersaturated saline fluid may precipitate quartz at temperatures between 300° and 55O°C (Fournier, 1985). Thus Si-saturated seawater would deposit quartz on encountering temperatures greater than 300°C in the intrusion or the immediate host rocks adjacent to the intrusion. MacGeehan (1978) proposed this process, of Si leaching from volcanic glass and heating of the fluid into the retrograde solubility temperature range, to explain silicification in pillow basalts adjacent to synvolcanic gabbro sills in the Matagami district.

Genesis of contact altered zones Contact altered zones may develop adjacent to an intrusion as heat is transferred from the cooling intrusion and heated modified seawater reacts with the host succession (Fig. 6.16). Vapour or fluid exsolved from the crystallising magma may contribute both heat and elements to the hydrothermal fluid (Norton, 1984). Hydrothermal fluid temperatures are partly determined by the depth of emplacement, volume of the intrusion and the temperature and volume of contributed magmatic fluid (Polyaetal., 1986; Eastoeetal., 1987; Cathles, 1993; Galley, 2003). For example, two active hydrothermal systems are recognised in the Guaymas Basin (Geiskes et al., 1982; Kastner, 1982). One is a low temperature (<300°C) hydrothermal system associated with the emplacement of shallow sills into unconsolidated sediments below the seafloor. The other is a deep high-temperature hydrothermal system associated with dykes or magma chambers that fed the overlying sill complexes.

FIGURE 6,16 | Development of a contact metamorphic-hydrothermal system in a submarine volcanic succession after the emplacement of a synvolcanic intrusion. (A) Initial fluid expulsion and migration away from the intrusion as heat from the intrusion drives dehydration and decarbonation reactions in the host succession. A combination of thermal metamorphism and hydrothermal alteration, by seawater and magmatic volatiles and fluid, may produce a contact altered zone. Seawater heated by the intrusion is buoyant and rises towards the seafloor either by diffuse flow or along fractures and faults. (B) In response, cold seawater is drawn down and heated in the vicinity of the intrusion, promoting hydrothermal convection and alteration between the intrusion and the seafloor. (C) Hydrothermal convection collapses as the intrusion cools. Cold seawater may be drawn down along fractures to produce proximal zones of hydrothermal alteration within the intrusion, and retrograde zones that overprint higher temperature contact altered zones adjacent to the intrusion.

The volume of an intrusion influences the temperature and longevity of the alteration system. Large volume intrusions (e.g. plutons and thick sills) influence the temperature of the host rocks and the alteration system for longer than smaller volume intrusions (e.g. synvolcanic sills, cryptodomes and dykes). A small volume sill (-30 m thick) may cause the temperature at the sill-sediment contact to rise as high as 400°C (Einsele et al., 1980). However, calculations suggest

1 5 4 | CHAPTER 6

that the temperature at the contact will drop below boiling within five years. This will significantly reduce convection and remaining heat will be lost mainly by conduction through the contact. Generally relatively small volume intrusions have thermal effects restricted to several metres or hundreds of metres from the contacts (Utada, 1973). Although the contact temperatures may be high, high-temperature altered zones are rare because the isotherms dip sharply away from small volume intrusions (Reyes, 1990). In contrast, large volume intrusions, such as the Skaergaard intrusion in east Greenland, which had an estimated volume of 180 km3, may take 500,000 years to cool to ambient temperatures (Norton and Taylor, 1979; Norton, 1984). They may result in thick, high-grade contact metamorphichydrothermal altered zones (Seki et al., 1969) and may drive regional convection of modified seawater.

6.5 | CONTACT ALTERED ZONES ASSOCIATED WITH THE DARWIN GRANITE Cambrian granites along the eastern margin of the Mount Read Volcanics (Fig. 1.5) are extensively altered and surrounded by concentric altered zones (Polya, 1981; Polya et al., 1986;Eastoeetal., 1987; Abbott, 1992; Large etal., 1996; Davidson, 1998; Wyman, 2001). Well-developed K-feldspar, chlorite and sericite zones have been mapped around the margin of the Darwin Granite and its northward extension in the southern Mount Read Volcanics (Fig. 6.17: Jones, 1993; Large et al., 1996; Wyman, 2001). The Darwin Granite alteration halo is of particular interest because of its close spatial and possibly temporal relationships with several small tonnage but high-grade CuAu prospects (Fig. 6.17: Jones, 1993; Large et al., 1996). These prospects occur in the Central Volcanic Complex along the western margin of the granite and above its northern subsurface projection from Mount Darwin towards Mount Lyell. The deposit styles vary systematically with increasing distance from the granite: from Fe-oxide veins, stockworks of pyrite + chalcopyrite ± hematite ± magnetite and quartz + pyrite + chalcopyrite veins, disseminated pyrite + chalcopyrite ± covellite, to veins containing quartz, bornite, chalcopyrite and hematite (Wyman, 2001). Large et al. (1996) suggested that the Darwin Granite provided heat, metals and magmatic fluid to form VHMS deposits in the southern Mount Read Volcanics, such as those in the Mount Lyell field. This section summarises the setting, altered zones and genesis of the Darwin Granite system and presents data sheets of typical alteration facies (DG1 to DG6).

FIGURE 6.17 | Geological map of the Jukes-Darwin area in the southern Mount Read Volcanics (western Tasmania), showing the limited surface extent of the Darwin granite and the thick hydrothermally altered halo around the granite (modified after Wyman, 2001). The locations of the six data sheets are shown.

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 5 5

Geological setting

Alteration facies and zonation

The Darwin Granite is an I-type magnetite series equigranular granitoid pluton dominated by pink granite intruded by subordinate white granite, microgranite and quartz porphyry phases (Wyman, 2001). The surface extent of the pluton is approximately 5 x 1 km (Fig. 6.17). However, modelling of gravity and aeromagnetic data along the eastern margin of the Mount Read Volcanics has suggested that a semicontinuous body of granite extends subsurface approximately 100 km northwards to the Murchison Gorge (Leaman and Richardson, 1989; Payne, 1991; Large et al., 1996). In the Darwin-Jukes area, the Central Volcanic Complex includes feldspar-phyric dacite, quartz + feldspar-phyric rhyolite (e.g. data sheet DG1), pumice breccia, tuffaceous sandstone, blocky rhyolite breccia, and minor sedimentary facies (Jones, 1993; Wyman, 2001). A thick, columnar jointed, micropoikilitic or spherulitic rhyolite hosts the Jukes Cu-Au Prospect and altered zones. The emplacement age of Darwin Granite is constrained to the Cambrian as it intruded the Middle Cambrian Central Volcanic Complex, and both the granite and Central Volcanic Complex are unconformably overlain by late Middle Cambrian Tyndall Group volcaniclastic rocks, which contain pebbles of granite near Mount Darwin (Corbett, 1979, 1981, 1992; Jones, 1993; Wyman, 2001).

The altered zones associated with the Darwin Granite cover a 1 5 x 3 km area that extends north to Jukes Prospect (Wyman, 2001). The altered zones at Mount Darwin and Jukes Prospect represent two parts of the same hydrothermal system: those adjacent to the granite at Mount Darwin record the lateral extent of the altered halo, whereas those at Jukes Prospect occur at least 1 km above the north-plunging pluton (Fig. 6.18). Immediately adjacent to the granite at Mount Darwin is a thin (10—20 m) K-feldspar ± biotite (now chlorite) hornfels zone (Table 6.1). This grades outward into a 400 m wide K-feldspar + chlorite zone, which locally hosts coarse breccias with magnetite ± tourmaline matrices (Jones, 1993; Wyman, 2001). The K-feldspar + chlorite zone grades into a 300 m thick chlorite + magnetite zone (Wyman, 2001). A discontinuous wedge-shaped silicified zone separates the chlorite + magnetite zone from the outer K-feldspar + quartz zone. At Mount Darwin the K-feldspar + quartz zone occurs between 800 and 1000 m from the granite contact. At Jukes Prospect the K-feldspar + quartz zone is the central altered zone and is enclosed in chlorite, sericite and regional diagenetic-metamorphic zones. It is associated with Cu-Au mineralised rocks (Doyle, 1990; Large et al., 1996). The peripheral sericite zone merges with the regional diageneticmetamorphic albite + sericite zone (Wyman, 2001).

K-feldspar ±

<20

biotite hornfels

K-feldspar ± chlorite

Pervasive, hornfels

(after biotite) ± quartz ± sulfides

K-feldspar +

400

chlorite

K-feldspar + quartz

Intense

+ chlorite ± sericite

Magnetite ± tourmaline fill in hydraulic breccias and veins

SiO2, AI2O3, K2O and net

DG2

mass gains

± hematite (after magnetite) Silicified

' Quartz ± sericite ± pyrite ± hematite

Na2O losses Strong to

Texturally destructive,

Large gains in SiO2 and

intense

cryptocrystalline and

net mass

DG3

microcrystalline K-feldspar +

200

quartz

K-feldspar + quartz

Strong to

+ sericite + chlorite +

intense

Moderately texturally

K2O, SiO2 and Fe2O3

destructive, pervasive,

gains

pyrite ± magnetite ±

cryptocrystalline ,

chalcopyrite

pseudomorphs plagioclase,

DG4

Na2O losses

veins and vein envelope Chlorite + magnetite

300

Chlorite + sericite +

Moderate

Moderate preservation, chlorite

Fe2O3, MgO and K2O

magnetite ± dolomite

to intense

pseudomorphs plagioclase,

gains

± apatite ± pyrite,

domainal replacement of

chalcopyrite veins

groundmass and matrix textures, infill in hydraulic

DG5

SiO2, AI2O3, Na2O and net mass losses

breccias and veins, vein envelope Sericite

Sericite ± chlorite ±

Weak to

Sericite partially to completely

pyrite

strong

replaces plagioclase and is disseminated in the groundmass and matrix

K2O gain and Na2O CaO losses

DG6

1 5 6 | CHAPTER 6

FIGURE 6.18 | Schematic cross-section of the Darwin Granite and altered zones, illustrating relationships between the Mount Darwin and Jukes Prospect alteration systems, and surface maps presented in the data sheets (Large et al., 1996; Wyman, 2001).

seawater in reaction zones around the hotter portions of the discharge zones to form K- and Fe-rich alteration assemblages above the granite (Wyman, 2001). The K-feldspar, chlorite and sericite zones in the Jukes area all show depletions in Na 2 O and CaO, which reflect the breakdown of plagioclase and mafic minerals during alteration by modified seawater. Large et al. (1996) suggested that the distribution, composition and zonation of alteration facies around the Darwin Granite, regional zonation of metals with respect to the granite, distribution of Cu-Au-rich VHMS deposits in the Mount Lyell field and pre-Tyndall Group timing of both the granite and mineralisation support a genetic link between the granite and these deposits. Furthermore, magnetite + apatite ± pyrite veins in the Prince Lyell deposit are similar to those adjacent to and within the Darwin Granite and are consistent with magmatic fluid contributing to their formation. These authors proposed a model for the genesis of the Mount Lyell Cu-Au and related Pb-Zn-Cu massive sulfide deposits (Fig. 6.19), which involves seawater convection deep into the Central Volcanic Complex where it mixed with Fe, Cu, Au and P-rich magmatic fluids exsolved from the granite. The mixed magmatic-seawater hydrothermal system produced altered zones, magnetite veins (e.g. Jukes Prospect) and subseafloor Cu-Au deposits (e.g. Prince Lyell) close to the granite, and contemporaneous Pb-Zn-Cu massive sulfide deposits on the seafloor (e.g. Lyell-Comstock).

Genesis of the alteration system The proximity of intense hydrothermal alteration mineral assemblages to the granite contact and the northward extending cupola region above the buried pluton support the interpretation that the hydrothermal system was driven by heat from the intrusion (Eastoe et al., 1987; Large et al., 1996; Wyman, 2001). Overprinting relationships between Kfeldspar, sericite and chlorite facies indicate multiple alteration stages in which low-temperature mineral assemblages overprinted initial high-temperature assemblages (Wyman, 2001). Initial sericite and chlorite alteration assemblages were associated with fracturing and vein formation around and above the granite. As the fracture system evolved, the Jukes hydrothermal system became part of the discharge zone. The well-defined zones of hydrothermal alteration are interpreted to have formed from diffuse circulation of hydrothermal fluid through the volcanic rocks. Fluid access was enhanced by hydrothermal brecciation, and intense K-feldspar alteration and silicification was confined to the fracture zone above the granite (Wyman, 2001). Magnetite and tourmaline veins and breccias in the Kfeldspar-rich zones immediately adjacent to the granite contact, and in the centre of the Jukes alteration system, demonstrate that magmatic-hydrothermal fluids were exsolved during crystallisation (Large et al., 1996). The mass changes (i.e. gains in K 2 O, Fe2O3, Ba, Sr, Cu, Mo, W, Th and U) in the altered zones adjacent to the granite are consistent with hydrothermal alteration of the volcanic rocks by magmatic fluid (Wyman, 2001). These fluids mixed with modified

FIGURE 6.19 | Model of the zones of hydrothermal alteration and ore deposits in the Mount Lyell field and their relationship to the Darwin Granite, southern Mount Read Volcanics, western Tasmania (after Large et al., 1996).

SYNVOLCANIC INTRUSION-RELATED ALTERATION

DG1

Weak, regional, pervasive albite + sericite alteration facies Sample no.

143401

Alteration fades

weak, regional, pervasive albite + sericite

Location

Jukes Road

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

massive, feldspar + quartz-phyric rhyolite

Relict minerals Relict textures

feldspar (3%, 2 mm), quartz ( 1 % , <1 mm) micropoikilitic, porphyritic

Primary composition rhyolite Lithofacies

columnar jointed, massive

Interpretation

sill

Alteration minerals

albite + sericite + chlorite + hematite

Alteration textures

recrystallised groundmass, sericite pseudomorphs after feldspar, pervasiveselective sericite > chlorite, disseminated hematite

Distribution

regional

Preservation

good

Alteration intensity

weak

Timing

early

Alteration style

regional diagenetic and metamorphic

Hand specimen photograph

|

Geochemistry SiO2

76.17

Na2O

2.29

Rb

129

Zr

280

TiO2

0.31

K2O

3.66

Sr

22

Nb

13

AI2O3

13.30

P2O5

0.04

Ba

822

Y

39

Fe2O3

3.35

S

0

Cu

4

MnO

0.10

Total

100.00

Pb

20

Al

64

Zn

39 CCPI

Th

21

MgO

o.67

CaO

0.11

(

voLfree

Photomicrograph (ppl)

>

Ti/Zr

38 6 .6

157

1 5 8 | CHAPTER 6

Intense, pervasive K-feldspar + chlorite alteration facies Sample no.

143278

Alteration facies

intense, pervasive K-feldspar + chlorite

Location

Mount Darwin

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

massive, feldspar + quartz-phyric rhyolite

Relict minerals

quartz + plagioclase (5%, 1-2 mm)

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

unknown

Alteration minerals

quartz + K-feldspar > chlorite + sericite > hematite + pyrite plagioclase replaced by sericite, recrystallised groundmass of quartz + K-feldspar > chlorite + sericite, disseminated hematite + pyrite, selective chlorite

Alteration textures

Distribution

local

Preservation

none

Alteration intensity

intense

Timing

syn- to post-intrusion

Alteration style

proximal intrusion-related hydrothermal alteration

Hand specimen photograph

DG2

Geochemistry SiO2

71.35

Na2O

0.30

Rb

259

TiO2

0.27

K2O

9.27

Sr

80

AI2O3

14.59

P2O5

0.04

Ba

2463

Fe2O3

2.96

S

0

Cu

8

MnO

0.01

Total

99.72

Pb

31

MgO

0 . 9 2 free) (vol.

CaO

0.01

Photomicrograph (ppl)

Zn

89 Th

Ti/Zr

25

^

223

Nb

1g

y

46

A]

g7

ccp|

2J 7 3

SYNVOLCANIC INTRUSION-RELATED ALTERATION | 1 5 9

Strong, foliated quartz + sericite + pyrite alteration facies Sample no.

143237

Alteration facies

strong, foliated quartz + sericite + pyrite

Location

Slate Spur

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

feldspar + quartz-phyric rhyolite schist

Relict minerals

quartz + plagioclase (15%, 1-5 mm)

Relict textures

porphyritic, microcrystalline, partly spherulitic and possibly perlitic

Primary composition

rhyolite

Lithofacies

foliated

Interpretation

unknown

Alteration minerals

quartz + sericite + pyrite + hematite

Alteration textures

plagioclase replaced by sericite, pervasive microcrystalline groundmass of quartz + sericite + pyrite, hematite stylolites, quartz overgrowths, schistosity

Distribution/zonation

local

Preservation

poor

Alteration intensity

strong

Timing

syn- to post-intrusion

Alteration style

proximal intrusion-related hydrothermal alteration

Hand specimen photograph

DG3

Geochemistry SiO2

81.77

Na2O

1.54

Rb

145

Zr

142

TiO2

0.17

K2O

4.51

Sr

35

Nb

11

AI2O3

10.23

P2O5

0.03

Ba

889

Y

33

Fe2O3

1.37

S

0

Cu

3

MnO

0.01

Total

99.92

Pb

16

Al

76

MgO

0.29

(vol.

Zn

22

CCPI

20

CaO

0.01

Th

19

Ti/Zr

7.2

Photomicrograph (xn)

free)

1 6 0 | CHAPTER 6

DG4

Intense, pervasive K-feldspar + sericite alteration facies Sample no.

143360

Alteration fades

intense, pervasive K-feldspar + sericite

Location

Jukes Road

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

massive, feldspar + quartz-phyric rhyolite

Relict minerals

plagioclase (2%, 2 mm), quartz (1 %, 1 mm)

Relict textures

micropoikolitic, porphyritic

Primary composition

rhyolite

Lithofacies

columnar jointed, massive

Interpretation

sill

Alteration minerals

K-feldspar + quartz + sericite + chlorite + pyrite + magnetite pervasive recrystallised groundmass of K-feldspar + quartz + sericite > chlorite, cleavage defined by sericite, plagioclase altered to K-feldspar > chlorite + sericite, K-feldspar overgrowths, chlorite pseudomorphs of feldspar microphenocrysts

Alteration textures

Distribution/zonation

local

Preservation

moderate

Alteration intensity

intense

Timing

syn- to post-intrusion

Style

intrusion-related hydrothermal alteration

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

73.74 0.27 12.10 5.51 0.03 0.53 0.01 0.13

K2O 7.60 P2O5 0.04 S 0 Total 99.97 (vol. free)

Photomicrograph (ppl)

Rb Sr Ba Cu Pb Zn Th

161 38 2449 215 17 57 17

Zr Nb Y

257 11 40

Al CCPI Ti/Zr

98 42 6.3

SYNVOLCANIC INTRUSION-RELATED ALTERATION

Strong, pervasive sericite alteration facies Sample no.

143366

Alteration fades

strong, pervasive sericite

Location

Jukes Road

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic fades

massive feldspar + quartz-phyric

DG5

rhyolite Relict minerals

quartz (2%, 1 mm), plagioclase (2%, 3 mm), hornblende (<1%)

Relict textures

spherulitic, glomeroporphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

sill

Alteration minerals

chlorite > K-feldspar + sericite > pyrite + magnetite

Alteration textures

pervasive K-feldspar + quartz + chlorite, chlorite + sericite + magnetite pseudomorphs after plagiociase, chlorite pseudomorphs after hornblende, recrystallised spherulites

Distribution

local

Preservation

moderate

Alteration intensity

intense

Timing

syn- to post-intrusion

Alteration style

intrusion-related hydrothermal alteration

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3

72.77 Na2O 0.37

K2O

13.23 P2O5

0.15

Rb

167

Zr

284

5.06

Sr

18

Nb

13

0.05

Ba

1323

Y

27

0 Cu

234 Al

97

Fe2O3

7.11 S

MnO

0.04 Total

MgO

1.15 (vol. free)

Zn

CaO

0.05

Th

Photomicrograph (ppl)

99.97

Pb

6

114 CCPI 19

Ti/Zr

59 7.8

|

161

1 6 2 | CHAPTER 6

Strong, pervasive sericite alteration facies Sample no.

143400

Alteration facies

strong, pervasive sericite

Location

Jukes Road

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

massive feldspar + quartz-phyric rhyolite

Relict minerals

plagioclase (2%, 3 mm), quartz (1%, 1 mm)

Relict textures

porphyritic, spherulitic?

Primary composition

rhyolite

Lithofacies

massive

Interpretation

unknown

Alteration minerals

sericite + K-feldspar + quartz + hematite + pyrite plagioclase replaced by sericite, recrystallised groundmass of sericite + K-feldspar + quartz, hematite stylolites, disseminated pyrite, weakly developed cleavage

Alteration textures

DG6

Geochemistry

Distribution

local

SiO 2

Preservation

poor

TiO 2

Alteration intensity

strong

AIA

Timing

syn- to post-intrusion

Fe2O3

3.67 S

Alteration style

intrusion-related hydrothermal alteration

MnO

0.02 Total

MgO

0.99

CaO

0.04

Hand specimen photograph

75.23

Na2O

0.99

0.30 K2O 13.52

5.18 Sr

P2O5

0.05

173

Zr

278

22 Nb

13

Ba

1232

0 Cu

4

100.00

(vol. free)

Photomicrograph (ppl)

Rb

Y

41

Al

86

Pb

3

Zn

35

CCPI

41

Th

21

Ti/Zr

6.5

I 163

7 | LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS

An understanding of the mineralogical and chemical zonation of hydrothermally altered rocks around submarine massive sulfide deposits is vitally important to both ore genesis studies and to assist and focus mineral exploration. It has led to a vast amount of literature on hydrothermal alteration around VHMS deposits, which various workers have summarised (e.g. Franklin et al., 1981; Barriga et al., 1983; Urabe et al., 1983;Lydon, 1984; 1988; Large, 1992; Madeisky and Stanley, 1994; Barrett and MacLean, 1994b; Galley, 1995; Carvalho et al., 1999; Sanchez-Espana et al., 2000, 2002; Gemmell and Herrmann, 2001; Large et al., 2001c). Since the pioneering work in the 1950s and 1960s, when geologists first recognised the critical link between volcanicmagmatic processes and massive sulfide genesis (e.g. Stanton, 1955, 1959; Oftedahl, 1958; Gilmour, 1965; Horikoshi, 1969), it became widely accepted that these deposits form on the seafloor from hydrothermal activity generated during periods of local quiescence between volcanic eruptive cycles (Sangster, 1972; Solomon, 1976; Franklin et al., 1983; Ohmoto and Skinner, 1983). The discovery of seafloor black smokers and related sulfide chimneys on the presentday seafloor has further stimulated research and contributed to an improved understanding of ore forming processes in the ancient deposits (Rona and Scott, 1993). Over the last 15 years, many researchers have questioned whether all massive sulfide deposits form by exhalation on the seafloor. Although it has been recognised for some time that stringer zones and the lower parts of some massive sulfides formed by replacement (e.g. Large, 1997), several authors now consider that replacement of particular volcanic units below the seafloor maybe be a key process for massive sulfide formation (e.g. Barriga and Fyfe, 1988; Khin Zaw and Large, 1992; Allen, 1994b; Bodon and Valenta, 1995; Hannington et al., 1999; Doyle and Allen, 2003). This chapter is not a summary of VHMS ore genesis, but it highlights the features of alteration halos associated with VHMS deposits, particularly in the Mount Read province, western Tasmania, and fits the deposits and their alteration halos in to the broad range of ore deposits that are found in volcanic and volcano-sedimentary successions. In the later part of this chapter we provide descriptions, including data sheets depicting typical alteration facies, of examples from the Mount Read province and Mount Windsor Subprovince in

eastern Australia that illustrate the range in alteration styles and zonation associated with the spectrum of submarine volcanic-hosted base metal ores.

7.1

COMMON FEATURES OF VHMS DEPOSITS

VHMS deposits display the following features: • They are hosted by submarine volcanic or volcanosedimentary successions. • They are the same age as the host volcanic succession (i.e. the deposits are approximately synvolcanic and/or synsedimentary). • The host rocks vary from coherent to clastic volcanic or sedimentary facies and range in composition from basalt through andesite and dacite to rhyolite. • Most deposits are hosted in thin volcaniclastic units (<100 m thick) between major volcanic formations. • The economic parts of the deposits typically comprise massive sulfide, principally pyrite, subordinate sphalerite, chalcopyrite and galena. The term massive implies greater than 80 wt% sulfides (Sangster, 1972). • Massive sulfide lenses are commonly, but not always, aligned parallel to volcanic strata. • Stringer (or stockwork) sulfide zones commonly underlie the massive sulfides and may contain economic Cu grades. • Metal contents and metal ratios vary considerably. Deposits include Cu-rich, Au-rich, Cu-Zn, and polymetallic (CuZn-Pb-Ag-Au) types, but all contain more Zn than Pb. • Ore metals are typically vertically zoned within sulfide deposits from Cu at the stratigraphic base to Zn, Pb, Ag, Au and Ba in general order towards the top. Nevertheless, there are many exceptions to this zonation pattern and some deposits have no Ba. • Intense hydrothermal alteration of the footwall volcanic rocks stratigraphically below the massive sulfide, to chlorite, sericite and quartz is common. By comparison, the hanging wall rocks are weakly altered or unaltered. Over 700 VHMS deposits have been recorded around the world. They range in size from less than 100,000 tonnes to

1 6 4 | CHAPTER 7

over 510,000,000 tonnes (RioTinto, Iberian pyrite belt). The top 50 deposits, in terms of tonnes of contained Cu + Zn + Pb metal, are listed in Figure 7.1. Most of these deposits are from seven major VHMS provinces or districts (Fig. 7.2), from oldest to youngest: Abitibi belt in Canada, Skellefte district in Sweden, Mount Read province in Australia, Bathurst mining camp in Canada, Southern Urals in Russia, Iberian pyrite belt in Spain and Portugal, and Hokuroku district in Japan.

Cu+Zn+Pb metal content (million tonnes) 10

40 FIGURE 7.2 | Locations of the major VHMS provinces around the world.

7.2 | HYDROTHERMAL ALTERATION HALOS ASSOCIATED WITH VHMS DEPOSITS Hydrothermally altered zones proximal to VHMS deposits may include footwall alteration pipes, stratabound altered footwall zones and altered hanging wall zones. Previous studies (e.g. Franklin et al., 1983; Lydon, 1988) emphasised the pipe-like hydrothermal altered zones in the footwalls of many massive sulfide deposits. They are common in Archaean deposits in the Abitibi belt in Canada (e.g. Sangster, 1972) and in the Miocene Kuroko deposits of the Hokuroku district in Japan (e.g. Urabe et al., 1983). However, in other districts such as the Mount Read province, Mount Windsor Subprovince and Lachlan Fold Belt in eastern Australia, the Iberian pyrite belt, and the Bathurst mining camp, well-defined alteration pipes are less common, and stratabound altered footwall zones dominate (Large, 1992). These two styles of altered footwall zones are described below, with emphasis on the footwall alteration pipes, because they have received considerable attention from researchers and their alteration mineral zonation and genesis are better understood.

Footwali alteration pipes

FIGURE 7.1 Pb tonnes.

The 50 largest VHMS deposits in terms of contained Cu + Zn +

Figure 7.3 is a schematic cross-section of the geology, sulfide and alteration zonation related to a typical VHMS deposit. It is based on our understanding of the Hellyer deposit in the Mount Read province (Gemmell and Large, 1992), but also incorporates information on other Australian (Large, 1992), Canadian (Franklin et al., 1981; Lydon, 1988; Lentz and Goodfellow, 1996), Japanese (Date et al., 1983; Urabe et al., 1983) and Spanish-Portuguese deposits (Barriga et al., 1983; Leistel et al., 1998; Sanchez-Espana et al., 2000). Immediately below the thickest part of the massive sulfide ore the footwall alteration pipe, which may be oval in plan but is more commonly elongate along a synvolcanic fault, contains a concentric series of altered zones. These are, from the centre of the pipe outwards: siliceous core zone, chlorite zone, sericite zone, and albite zone, which grades into leastaltered volcanic rocks.

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 6 5 These zones are typically fine grained, dark and massive, preserving no volcanic or sedimentary textures. In many deposits these zones host pyrite + chalcopyrite stringer veins (e.g. Woodlawn deposit in eastern Australia and most Noranda district and Iberian pyrite belt deposits). Chlorite zones are commonly deformed during tectonic events, resulting in chlorite schist zones, which may be strung-out or dislocated from the massive sulfide ore (Sangster, 1972). Studies of chlorite composition from Canadian, Australian and Japanese deposits indicate that the inner chlorite zones are dominated by Mg-rich chlorite, with a general increase in Fe/Mg ratio passing from the inner chlorite zone to the outer edge of the sericite zone (e.g. Riverin and Hodgson, 1980; Urabe et al., 1983; Paulick et al., 2001). Nevertheless, reverse trends have been recorded, where chlorite becomes more Fe-rich towards the core (e.g. Eastoe et al., 1987; Lentz et al., 1997). Variations in chlorite composition are discussed in more detail in Section 4.2.

Sericite zones

FIGURE 7.3 | Cross-section of idealised mineralisation and alteration zonation patterns in a footwall alteration pipe beneath a typical VHMS deposit (modified after Gemmell and Large, 1992; Lydon, 1997). (A) Sulfide mineral zones and geology. (B) Hydrothermally altered zones.

Siliceous core zones The siliceous core zones are composed of quartz + pyrite and quartz + pyrite + sericite ± chlorite assemblages (e.g. data sheets HE6, RB4, TH4, WT7, HR8, HN6 and HN7). They may not always be present, and have only been described from a few deposits (e.g. Hellyer, Gemmell and Large, 1992; Brunswick No. 12 and other deposits in the Bathurst mining camp, Zhang et al., 2003). Siliceous core zones are the most intensely altered rocks in the centre of the pipes, and are commonly intersected by networks of pyrite + chalcopyrite stringer veins (Fig. 7.3). All primary rock textures within these zones have been completely destroyed due to the intensity of alteration. In some cases, quartz-rich alteration assemblages have overprinted earlier chlorite-rich alteration assemblages creating pseudobreccia textures. Mass-change calculations indicate that gains of 50-100 g/100 g, mainly due to Si addition, are common within siliceous core zones (Gemmell and Large, 1992, Fig. 11). Lentz and Goodfellow (1996) reported SiO2 gains of up to 300% in the siliceous core zone in the centre of the alteration system below the Brunswick No. 12 massive sulfide deposit. The siliceous core represents the zone of maximum hydro thermal fluid flow and highest temperatures.

Chlorite zones Chlorite zones are dominated by chlorite (>50 wt% and commonly >80 wt%), with subordinate quartz + pyrite + sericite ± carbonate (e.g. data sheets HE4, RB5 and HR7).

Sericite zones surround the inner chlorite zones and are characterised by assemblages of sericite + chlorite + quartz + carbonate + pyrite (e.g. data sheets HE3, RB3, WT3, WT6, HR6 and HN5). At Hellyer, rocks in the sericite zone are strongly to intensely altered, with up to 70 wt% sericite and sparse relict primary textures. In other deposits, the alteration intensity in the sericite zone decreases towards the outer margin, where altered rock grades into the least-altered footwall rocks. In many cases, the sericite zones are laterally extensive and merge with stratabound altered zones away from the central pipe (e.g. Mount Chalmers, Large and Both, 1980). Minor disseminated sphalerite or stockwork Zn may occur in the sericite zone, whereas Cu-enrichment is more common in the chlorite zone.

Albite zones Some authors have described weakly altered zones of albite + chlorite ± sericite that surround the main sericite zones (e.g. Iijima, 1974; Green et al., 1981; Urabe et al., 1983; Relvas et al., 1997; Goodfellow and McCutcheon, 2003). Although albite zones have not been widely described or accepted in all districts they are discussed here because of their significance to mineral exploration. It may be difficult to distinguish hydrothermal albite facies from the background diagenetic alteration facies, as within albite zones primary volcanic textures are commonly preserved. In the Hokuroku district this zone has an albite + sericite + chlorite assemblage (Iijima, 1974; Urabe et al., 1983). In some deposits of the Iberian pyrite belt an outermost halo of Na-sericite has been recognised (Relvas et al., 1997). In the Bathurst mining camp, Goodfellow and McCutcheon (2003) described an outermost altered zone of albite + Mg-rich chlorite surrounding the footwall alteration pipe. They described an increase in patchy albite, which replaced K-feldspar phenocrysts, in proximity to the pipe. Similar albite-altered rocks have been recognised adjacent to the Hercules footwall alteration pipe in the Mount Read province (Large et al., 1996). Further research is required

1 6 6 | CHAPTER 7

to characterise the features of these outermost albite + chlorite + sericite altered zones and to distinguish them from regional diagenetic albite alteration facies.

Variations in alteration zonation The three main hydro thermally altered zones (siliceous core, chlorite and sericite) are not present in all footwall alteration pipes beneath VHMS deposits, and in some cases additional zones, such as carbonate or talc zones, exist. Some variations from the idealised alteration pipe model outlined in Figures 7.3 and 7.4A are:

The Kuroko deposits do not have an inner chlorite zone. Shirozu (1974) described intensely altered volcanic rocks below the massive sulflde ore and surrounding the siliceous Cu-stockwork ore, as strongly silicified with abundant sericite and very little chlorite (Fig. 7.4B). Deposits in the Noranda district and the Iberian pyrite belt commonly have intense chlorite core zones surrounded by sericite zones (Fig. 7.4C), but without siliceous core zones (Franklin et al., 1983; Carvalho et al, 1999; SanchezEspana et al., 2000). In the Iberian pyrite belt, Relvas et al. (1997) recognised a Na-bearing sericite (paragonite) zone extending beyond and above the main sericite zone at both the Neves Corvo and Aljustrel deposits. In the Mattagami district in Canada, the intensely altered central core of the footwall alteration pipe is talc-rich (Fig. 7.4D) and surrounded by chlorite and sericite zones (e.g. Large, 1977; Roberts and Reardon, 1978). In several eastern Australian deposits, Mg- and/or Febearing carbonates are common in the alteration mineral assemblages (Large et al., 2001c). At Hellyer a chlorite + dolomite zone occurs below the massive sulflde near the top of the chlorite zone (Fig. 7.4A, data sheet HE5, Gemmell and Large, 1992). A more massive dolomite zone is developed at the western margin of the altered stringer pipe at Mount Chalmers (Large and Both, 1980). Carbonate-rich zones are common in the footwall of many Iberian pyrite belt deposits, either marginal to the massive sulflde or distributed throughout the footwall alteration systems (e.g. Rio Tinto, Williams et al., 1975, Solomon et al., 1980; La Zarza, Strauss et al., 1981; Tharsis, Tornos et al., 1998). Lydon (1988, Fig. 8) included an Fe-oxide zone at the top of the pipe below the massive sulflde in his footwall alteration pipe model, compiled from a number of deposits, but there are few examples of this facies.

Stratabound altered footwall zones

FIGURE 7.4 | Different patterns of mineral zonation in footwall alteration pipes. (A) Generalised model. (B) Hokuroku district model. (C) Noranda district and Iberian pyrite belt model. (D) Mattagami district model.

Many massive sulfide deposits, possibly half, do not have footwall alteration pipes, but are underlain by stratabound or semi-conformable altered zones, which extend laterally for up to several kilometres away from the deposits (Figs 3.16B, C and 7.5). These stratabound zones may extend for between 30 and several hundred metres below the massive sulfide. They are typically developed around sheet-like deposits, and are mainly associated with Zn-Pb-rich deposits (e.g. Rosebery, Scuddles and Bathurst mining camp deposits). Sangster (1972) and Goodfellow and McCutcheon (2003) consider stratabound altered footwall zones to originally have been pipes that were sheared and transposed parallel to stratigraphy during tectonic deformation. However, this does not explain stratabound footwall zones in weakly to moderately deformed volcanic successions, such as the Mount Read province and Iberian pyrite belt. In the Iberian pyrite belt, Tornos (in press) notes that although many previous studies have claimed pipelike morphologies to the footwall stockworks and altered zones (e.g. Costa et al., 1995; Carvalho et al., 1999; Saez et al., 1999), recent studies have defined irregular to stratabound morphologies for the footwall alteration associated with most

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 6 7

deposits (e.g. Tharsis, Tornos et al., 1998; Neves Corvo, Relvas et al., 2000). Stratabound altered footwall zones have similar alteration mineral assemblages to footwall alteration pipes, but the zones are distributed parallel to stratigraphy, rather than at right angles. In some cases, the siliceous core and chlorite zones are confined to the immediate footwall of the thickest Cu-rich part of the massive sulfide lens (e.g. Rosebery and Thalanga). In contrast to the pipes, the sericite zones of stratabound altered footwall zones are the volumetrically dominant zones being both laterally and vertically extensive. Massive carbonate zones are more common in stratabound than pipe-like alteration systems (Large et al., 2001c), typically occurring immediately along strike from the massive sulfide ore lenses (e.g. Rosebery, Thalanga and Mount Chalmers). From an exploration perspective stratabound altered footwall zones typically present broader targets than footwall alteration pipes; however, they tend to be more diffuse and thus create challenges when searching for the associated VHMS deposits.

Altered hanging wall zones Compared to footwall alteration, hanging wall alteration is typically less intense and therefore has not received much attention in the ore-deposit literature. Visible hanging wall alteration mineral assemblages are commonly sericite-rich and restricted to a few metres above the massive sulfide ore. However, detailed petrographic and geochemical studies have extended some altered zones to several tens of metres into the hanging wall. There are a number of exceptions to the generally limited altered hanging wall zones. Copper-Au-rich VHMS deposits that form by replacement below the seafloor may exhibit extensive sericite-rich altered hanging wall zones (Fig. 7.6A). Mount Lyell in the Mount Read province and Highway-Reward in the Mount Windsor Subprovince are good examples of subseafloor sericite zones and are described in detail in Sections 7.7 and 7.10. In stacked ore systems, such as Millenbach and Amulet deposits in the Noranda district, and Que River in the Mount Read province, the lower ore body in the stack has an intense altered hanging wall zone, similar to that in the footwall (Fig. 7.6B). This is due to hydrothermal fluids moving through the lower ore lens and hanging wall on their way to depositing the upper ore lenses. Significant altered hanging wall zones may have developed in situations where the hanging wall volcanic or sedimentary units were deposited while the massive sulfide was still forming on the seafloor. This is the case for the Hellyer deposit, which has a hanging wall alteration plume that comprises a core of fuchsite + carbonate (e.g. data sheet HE10), surrounded by successive halos of chlorite + carbonate, quartz + albite, and finally patchy sericite (Fig. 7.6C, Gemmell and Fulton, 2001). The interpretation that the hanging wall alteration zones formed after the seafloor massive sulfide is not in doubt because of sulfide and barite clasts in the directly overlying volcaniclastic debris-flow unit (McArthur and Dronseika, 1990;Sharpe, 1991). Although primary plagioclase destruction is a key process in the hydrothermal alteration associated with VHMS

FIGURE 7.5 | Examples of stratabound altered footwall zones (modified after Ashley et al., 1988; Large, 1992; Large et al., 2001c). (A) Scuddles, Western Australia, in plan view. (B) Teutonic Bore, Western Australia, in cross-section. (C) K lens at Rosebery, western Tasmania, in cross-section. (D) Thalanga, Queensland, in cross-section. Abbreviations are: FW = footwall and HW = hanging wall.

deposits, there are a few examples where albite zones exist in the hanging wall of the deposit. At the Henty gold deposit albite + quartz is a common hanging wall alteration mineral assemblage (see Section 7.8, data sheet HN3), and at Hellyer albite forms one of the altered zones in the hanging wall alteration plume (Fig. 7.6C, data sheet HE8). In addition to these examples of obvious altered hanging wall zones, some recent detailed lithogeochemical

1 6 8 | CHAPTER 7

Chemical reactions and mass changes Footwall alteration results from the reaction of hydrothermal fluid (principally composed of heated seawater) with volcanic rocks. The temperature of the fluids that form VHMS deposits are estimated to vary from about 200° to 350°C based on fluid inclusion evidence (Pisutha-Arnond and Ohmoto, 1983; Khin Zaw et al., 1996), the study of present day black smoker systems (Goldfarb et al., 1983) and thermodynamic calculations of mineral stabilities (e.g. Sato, 1973; Large, 1977; Ohmoto et al., 1983). Fluid salinities approximate that of seawater, although values of up to four times seawater have been recorded (de Ronde, 1995; Solomon et al., 2002). The pH varies from about 3 to 7 based on mineral assemblage, thermodynamic considerations and measurements at black smoker vents (Huston and Large, 1989; Scott, 1997). The principal result of interaction of this hot, mildly acidic to neutral fluid with volcanic rocks as it ascends towards the seafloor is the breakdown of feldspars and volcanic glass, and their replacement by sericite, quartz, chlorite and carbonate. Petrographic evidence of these reactions is shown in Figure 2.5 which depicts a series of progressively hydrothermally altered pumice-rich rocks from the footwall to the Hercules deposit in the Mount Read province. Reactions that describe these footwall alteration processes may include reaction R7.1 (from Sanchez-Espana et al., 2000) in the sericite zone: 3NaAlSi3O8 + K+ + 2H+ albite -» KAl3Si3O10(0H)2 + 6SiO2 + 3Na+ sericite quartz

FIGURE 7.6 | Examples of cross-sections through altered hanging wall zones associated with VHMS deposits. (A) Cu-Au deposit (modified after Large et al., 2001c). (B) A stacked ore system (after the Millenbach deposit, Knuckey etal., 1982). (C) Hellyer hanging wall alteration plume (modified after Gemmell and Fulton, 2001).

studies have defined subtle geochemical halos in otherwise least-altered hanging wall volcanic rocks, which may extend several hundreds of metres above the massive sulfide ore (Large et al., 2001b). For example, at Rosebery a hanging wall alteration halo can be defined using three geochemical parameters that may be applied during exploration: (1) wholerock Ba/Sr ratio, which outlines a halo extending about 100 m into the hanging wall; (2) Mn content of carbonate, which is anomalous over the same interval; and (3) whole-rock Tl content, which forms a halo that extends over 200 m into the hanging wall volcanic rocks. More detail of the Rosebery alteration system is provided in Section 7.6.

(R7.1)

This reaction is typical of sericite replacing albite in the outer part of the alteration system. The reaction involves a gain in K from hydrothermal fluid and loss of Na in the rock as albite is replaced. Silica is conserved by the deposition of quartz. Overall, the reaction leads to an increase in fluid pH due to the consumption of H+. Reactions in the sericite zone also involved sericitisation of K-feldspar and plagioclase in addition to albite. In the chlorite zone, there are two potential reactions (Pisutha-Arnond and Ohmoto, 1983):

Note, these reactions can be written with H4SiO4(aq) or SiO2 (quartz) + 2H 2 O. Both of these reactions involve the addition of Mg and 2+ Fe from the fluid and the loss of alkalies (Na or K) and H from the rock. However, the replacement of albite by chlorite

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 1 6 9

involves Si loss as H 4 SiO 4 , compared to the replacement of sericite by chlorite, which involves Si gain. Mass transfer calculations in chlorite zones invariably indicate significant loss of Si (e.g. Gemmell and Large, 1992; Barrett and MacLean, 1994b) suggesting that reaction R7.2 is the key reaction. Also both of these chlorite replacement reactions (R7.2 and R7.3) involve a release of H+ to the fluid and will cause an increase in fluid acidity. This means that continued chloritisation of volcanic rocks over an extensive area might produce moderately acidic fluids, which may subsequently cause silicic (quartz + sericite), or in the extreme case, argillic (pyrophyllite ± kaolinite ± sericite) alteration assemblages upflow from the chlorite zone. Examples of this occur at the Neves Corvo and Lagoa Salgada deposits in the Iberian pyrite belt where Relvas et al. (1994, 1997) have reported dombassite and pyrophyllite in the chlorite-rich central stockwork zones below the massive sulflde ores. Other reactions in the chlorite zone, not listed here, include chloritisation of mafic minerals such as biotite and amphibole. In the siliceous core zone at the centre of the hydro thermal system, three reactions are proposed:

Quartz is the main alteration mineral in this zone and is associated with minor sericite and chlorite. Mass balance calculations indicate that considerable SiO, gains. Aluminium is commonly immobile (Gemmell and Large, 1992; Barrett and MacLean, 1994b), but is significantly diluted by large Si gains. Quartz is probably deposited in the siliceous core zone according to reaction R7.4, as this involves mass gain of Si without loss of Al. Enrichment of Si in the hydrothermal fluid may be due to leaching of Si from the footwall volcanic rocks during chloritisation (reaction R7.2). Leached Si is then deposited in the hydrothermal vent immediately below the seafloor. Silica deposition may be caused by rapid conductive cooling (Fournier and Potter, 1982), mixing with seawater, or intense fluid-rock interaction at high-fluid-rock ratios. Replacement of both chlorite and albite by quartz (reactions R7.5 and R7.6) requires a strongly acidic fluid and results in loss of Al as Al(OH)3(aq). Aluminium mobility of this type is rare in VHMS systems, but may occur in intensely silicified zones associated with acid alteration. Mass balance calculations suggest this was the case in the siliceous core zone of the Henty volcanogenic gold deposit (Callaghan, 2001, see also Section 7.8). In summary, by writing simple chemical reactions to describe replacements in the major altered zones, it is possible to gain some idea of the chemical processes, elemental gains and losses, and variations in pH of the hydrothermal fluid.

Most reactions in the altered footwall zone involve the breakdown of feldspar and loss of Na (and usually Ca) to the fluid. Major gains include K in the sericite zone, Mgand Fe in the chlorite zone and Si in the siliceous core zone. In addition to Na and Ca, other losses include Si in the chlorite zone, and very rarely Al in the siliceous core. The fluid pH does not show any systematic unidirectional change. Initially mildly acidic fluids will become less acidic during sericitisation, but more acidic during chloritisation. Intense siliceous core zones may be related to rapid cooling, fluid mixing or intense fluidrock interaction of Si-saturated fluids during the peak of the hydrothermal activity.

Alteration box plot trends in altered footwall zones The AI-CCPI Alteration box plot (Section 2.5) is a simple way of tracking whole-rock compositional changes and relating them to alteration mineralogy and position in the altered footwall zones. In Figure 7.7, the fields of the major altered zones are shown on the Alteration box plot in relation to two alteration indices, Al and CCPI. Line AD represents an array of altered felsic volcanic rock samples passing from the outer edge of the altered footwall zone into the core zone proximal to massive sulfide ore. Line ED represents a similar sample array for mafic volcanic rocks. Let us first consider line AD. Point A represents a rhyolitic rock outside the altered zone. Sericite and weakly chlorite altered rocks in the margins of the altered zone increase the Al due to Na and Ca depletion, whereas the CCPI remains relatively constant. Altered samples plot progressively along the AB segment of the trend toward the plotted position of sericite (phengite) on the perimeter of the Alteration box plot. In the inner part of the sericite zone, the Al is commonly greater than 90 and the trend becomes vertical due to a strong increase in the CCPI caused by gains in Fe and Mg related to

FIGURE 7.7 | The AI-CCPI Alteration box plot showing trends for altered footwall zones. These data are based on case studies presented in Large et al. (2001a).

1 7 0 | CHAPTER 7

increasing pyrite and chlorite in the rock (segment BC). The final segment CD represents the chlorite zone, where CCPI reaches its maximum (80-100), due to the abundance of chlorite and pyrite proximal to massive sulfide ore. Although quartz is not plotted on the Alteration box plot, due to the absence of SiO2 from the two alteration indices, samples from the siliceous core zone commonly plot in the CD segment due to the presence of minor chlorite and pyrite. If carbonate is present in the chlorite zone, as at Hellyer and Thalanga, the samples typically plot between F and D. In cases where the footwall comprises mafic volcanic rocks, ED is the common trend from the edge to centre of the footwall alteration pipe. This difference is due to the fact that mafic rocks generally have higher Fe and Mg contents, and thus greater initial values of CCPI compared with felsic rocks (Fig. 2.10). Consequently, chlorite is generally more abundant in the outer sericite zone and the combination of increasing chlorite and pyrite gives a trend along ED toward the core of the footwall alteration pipe (e.g. Hellyer, Section 7.5).

The presence of tightly constrained altered zones, in a circular or more commonly elongate pipe, suggests that hydrothermal fluids were focussed along synvolcanic faults or fault intersections, and massive sulfide deposition occurred where the faults intersected the seafloor. The zonation in the footwall alteration pipe is commonly interpreted to reflect a decreasing thermal gradient away from the fluid conduit (e.g. Large, 1977; Riverin and Hodgson, 1980).

up, and becomes more acidic due to fluid-rock interaction, metals are leached from the volcanic succession (e.g. Kajiwara, 1973; Spooner and Fyfe, 1973; Solomon, 1976; Large, 1977; Ohmoto, 1996). Alternatively, metal-rich magmatic fluid may be derived from the crystallisation of a magma, which is also the source of volcanism (e.g. Urabe and Sato, 1978; Henley and Thornley, 1979; Sawkins and Kowalik, 1981; Stanton, 1985, 1990). For supporting evidence and relative merits of these two models the reader is referred to recent discussions by Lydon (1996) and Ohmoto (1996). Recent research suggests that both fluids and metals are probably derived from magmatic and seawater sources (e.g. Fig. 7.8C). Distinguishing criteria for the source includes the deposit style, proximity to volcanic centres, alteration mineralogy, metal ratios of the deposits and the salinity and composition of primary fluid inclusions. For example, Large (1992) suggested that relatively soluble chloride-metal complexes, such as Zn, Pb and Ag, are probably derived principally from seawater leaching of the volcanic succession, whereas the less soluble metals, such as Cu, Bi and Sn, may be sourced directly from the magma chamber. Gold could be derived either by seawater leaching of volcanic rocks as a bisulfide-Au complex, or directly from magma as a chloride complex or in a volatile phase (Fig. 7.8C). Goodfellow and McCutcheon (2003) proposed a similar dual metal source for the massive sulfide deposits of the Bathurst mining camp, with the largest deposits having a major magmatic-metal component. Recently Solomon et al. (2004) have compared the salinity and composition of fluid inclusions in the stringer zones of the Hellyer VHMS deposit with those of porphyry Cu deposits, and concluded that the metals at Hellyer had a magmatic source.

Source considerations

Fluid-rock interaction in the alteration pipe

There are two competing models for the source of fluids and metals that form VHMS deposits (Fig. 7.8A and B): (1) evolved seawater, and (2) magmatic fluid. The first model involves seawater convecting through a volcanic succession above an intrusion or magma chamber. As the seawater heats

The concept that the alteration zonation in footwall alteration pipes is a function of decreasing temperatures and fluid-rock ratios has recently been tested by Schardt et al. (2001) using a thermodynamic model of fluid-rock interaction between heated evolved seawater and an andesitic precursor (Fig.

The genesis of footwall alteration pipes

FIGURE 7.8 | Models for fluid flow and metal source in VHMS hydrothermal systems (after Large, 1992). (A) Metals are derived from deep seawater leaching of the volcanic succession and basement. (B) Metals are sourced directly from the magmatic vapour plume, with no significant leaching of volcanic rocks. (C) A mixture of volcanic and magmatic sources, with low-solubility metals (i.e. Cu and Au) provided from magma and high-solubility metals (i.e. Zn, Pb and Ag) from seawater leaching of volcanic rocks.

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 7 1 7.9). This modelling was based on geochemical data from the Hellyer deposit. The classical sequence of altered footwall zones observed in many VHMS deposits (from the core to the margin of the pipe: quartz —* chlorite —> sericite) was reproduced by simulating the reaction between a 250—350cC fluid, with a pH of 4.5—5.0, and andesite under conditions of decreasing fluid-rock ratio and temperature. Simulated cooling from 350° to 100°C reproduced the full range of footwall alteration mineral assemblages. The pH of the fluid showed little variation, from 4.5 to 4.0 (Schardt et al., 2001). Mg-rich chlorite formed in the inner chlorite zone, and Ferich chlorite developed in the outermost part of the sericite zone, similar to the pattern observed at many massive sulfide deposits. This modelling was carried out using a Mg-bearing

fluid, with the assumption that a component of seawaterderived Mg was incorporated into the fluid at depth. The modelling has shown that sericite zones form at temperatures below 250°C from the reaction of andesitic rocks with mildly acidic solutions (pH = 4.0—4.5). Extensive Mg-chlorite zones are favoured by higher temperatures (250— 300°C) and less acidic fluids (pH = 4.5-5.5). At lower pH, kaolinite and pyrophyllite are likely to develop in the sericite zone. At higher pH and lower temperatures (<200°C), Kfeldspar is developed at the outer margin of the sericite zone and in least-altered andesitic rocks (Schardt et al., 2001). Although carbonate alteration was not taken into account by this modelling, it is likely that chlorite + carbonate assemblages, such as those developed adjacent to massive sulfides or at the

FIGURE 7.9 | Thermodynamic model of fluid-rock interaction between heated evolved seawater and Hellyer andesite (modified after Schardt et al., 2001). (A) Modelled mineralogical variations resulting from fluid-rock interaction with decreasing temperatures. (B) Schematic representation of simulated water-rock interaction as a function of temperature.

1 7 2 | CHAPTER 7

periphery of many footwall alteration systems, are indicative of more alkaline conditions. These may develop where hot, near-neutral hydrothermal fluids have mixed with and heated seawater, leading to saturation of carbonate at the margins of the hydrothermal up-flow zones (Large et al., 2001c).

A model for the development of footwall alteration pipes Using the results of the thermodynamic modelling of Schardt et al. (2001), numerical fluid-flow modelling by Yang and Large (2001), and previous thermodynamic modelling of metal-sulfide growth by Huston and Large (1989), it is possible to speculate on the progressive development of subseafioor zoned alteration pipes (Fig. 7.10). Stage 1 (Fig. 7.10): Initial hydrothermal fluid flow is upwards along a sub-vertical permeable fault zone towards the seafloor. As the rising hydrothermal plume approaches within 1 km of the seafloor, secondary near-surface seawater convection above the plume head may enhance normal diagenetic reactions in volcanic rocks adjacent to the fault, causing increased formation of zeolites, smectites and Mg-rich chlorite, and albite replacement of primary feldspars. During ongoing diagenesis and subsequent metamorphism this will produce an outer albite zone (albite + sericite + chlorite), which is commonly difficult to distinguish from regional diagenetic and metamorphic mineral assemblages.

Stage 2 (Fig. 7.10): As low-temperature and mildly acidic hydrothermal fluids, (T<250°C, pH = 4.0-4.5) continue to move upwards to the seafloor, sericite-rich alteration overprints the early albite zone and expands out from the fluid conduit to form a pipe-like sericite zone. During this stage sphalerite + galena + pyrite massive sulfides deposit on the seafloor above the sericite zone. Minor pyrite + sphalerite + galena stringer mineralisation may also occur in the core of the sericite zone at these temperatures (Eldridge et al., 1983). Convective near surface reflux of seawater leads to an albite + chlorite zone extending laterally into the least-altered volcanic rocks surrounding the sericite zone. Stage 3 (Fig. 7.10): As the hydrothermal system intensifies and the temperature of the discharging fluid rises above 250°C, Mg-chlorite is stabilised adjacent to the main conduit and a chlorite zone develops in the core of the footwall alteration pipe. Under these high-temperature conditions, the fluid is capable of carrying significantly more Cu (e.g. Huston and Large, 1989), which is deposited in stringer veins in the central chlorite zone and in the base of the massive sulfide. Within the massive sulfide mound, Cu progressively displaces Zn upwards. Eldridge et al. (1983), Huston and Large (1989) and Hannington and Scott (1989) described this zone refining process. As the alteration system evolves, the pH of the fluid initially increases during formation of the sericite zone (from 4 to 5.5) due to consumption of H+ (reaction R7.1), and subsequently falls during the formation of the chlorite zone (reactions R7.2 and R7.3).

FIGURE 7.10 | Model for the evolution of the footwall alteration pipe in a mound-style massive sulfide deposit. Stage 1: an initial lowtemperature hydrothermal system produces an albite zone. Stage 2: increasing temperature results in the development of the sericite and Zn + Pb-rich sulfide zones. Stage 3: higher temperatures produce the chlorite and Cu + Zn +Pb-rich sulfide zones. Stage 4: maximum temperatures and low pH result in the siliceous core and Cu + Pb + Zn-rich sulfide zones. This model does not apply to all VHMS deposits, some of which may form in brine pools (Solomon and Groves, 1994; Solomon and Quesada, 2003).

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 173 Stage 4 (Fig. 7.10): With increasing temperature (300° to 35OCC) during chloritisation, Si is continually leached from the volcanic rocks (reaction R7.2) and the fluid becomes supersaturated in Si. Consequently, quartz is precipitated (reaction R7.4) in the upper-central part of the alteration pipe, forming a siliceous core zone. Maximum metal precipitation in both the stringer zone and massive sulfide is commonly associated with this stage. Subsequently, the hydrothermal system wanes and collapses with an influx of heated near-surface seawater that leads to overprinting by lower temperature mineral assemblages, which are commonly dominated by carbonates or barite, depending on the oxidation level of the overlying water column. Some previous workers have suggested that the zonation in alteration pipes, from Mg-Fe chlorite in the core to sericite at the margins, relates to the entrainment and mixing of Mg-bearing seawater with a Mg-poor hydrothermal fluid below the massive sulfide (e.g. Roberts and Reardon, 1978; Lydon and Galley, 1986). However, Riverin and Hodgson (1980) suggested that the presence of Mg-rich chlorite in the central and most intensely altered zone of the alteration pipe, and the abundance of Mg-chlorite in pyrite + chalcopyrite veins in the stringer zone, is consistent with Mg derived from hydrothermal fluid rather than seawater. In the model outlined in Figure 7.10, we have assumed the hydrothermal fluid is Mg-bearing, possibly either due to entrainment of seawater at considerable depth (>1 km) below the seafloor or due to the leaching of Mg from mafic volcanic rocks deep in the volcanic succession. Although near-surface entrainment of seawater is considered to be important in stages 1 and 2 of our model, it is likely to result in an increase in the rate and consequent grade of diagenetic alteration. This would lead to Na-Mg metasomatism and albite + chlorite formation at the margins of the alteration pipe, rather than Mg metasomatism and chlorite development within the core of the pipe, as previously proposed (cf. Franklin et al., 1981). In the thermodynamic modelling of Schardt et al. (2001), temperature and pH were shown to be the principal factors controlling the balance between chlorite and sericite zones in the footwall alteration pipe. However, two other factors also need consideration: (1) the composition of the immediate footwall volcanic rocks, and (2) the initial chemistry of the modified seawater as it rises up the conduit (e.g. Large, 1977). In the first case, particularly at low fluid-rock ratios, chlorite alteration is favoured in mafic host rocks and sericite alteration in felsic host rocks. However, at high fluid-rock ratios typical of the central parts of the hydrothermal pipe, the alteration mineral assemblage is controlled by fluid chemistry rather than rock chemistry. In the second case, seawater-rock interactions in deep rhyolite-dominated footwall volcanic successions, similar to those in the Hokuroku district and southern Mount Read province, will generate modified seawater hydrothermal solutions enriched in K and Si, but generally depleted in Mg, Fe and Ca. These fluids will result in sericite zones as they approach the seafloor. In contrast, footwall volcanic successions dominated by andesite and basalt, similar to those in the northern Mount Read province and the Abitibi belt, will generate evolved seawater fluids enriched in Mg, Fe and Ca, with lesser K and Si, and are more likely to develop zoned chlorite-sericite alteration pipes.

In summary, the footwall alteration mineral assemblages in VHMS systems are probably controlled by three factors: (1) the initial composition of the convective seawaterdominated ore fluid, which is constrained by the relative abundance of mafic versus felsic volcanic rocks deep in the succession; (2) the temperature and pH regime during fluidrock interactions in the footwall discharge zone (i.e. lower temperature, acidic conditions favour sericite development, whereas higher temperature and/or more neutral pH conditions favour chlorite formation); and (3) the composition of the immediate footwall host rocks.

Genesis of stratabound altered footwall zones Stratabound altered footwall zones (e.g. Rosebery, Scuddles and Teutonic Bore; Fig. 7.5) are interpreted to result from hydrothermal-fluid flow parallel to volcanic strata (Fig. 7.11), rather than at right angles to the stratigraphy as in the case for footwall alteration pipes. Alteration pipes are commonly developed in relatively impermeable footwall volcanic rocks (e.g. the coherent or clastic facies of lavas and synvolcanic intrusions) where fluids are focussed along sub-vertical synvolcanic faults (Fig. 7.10). In contrast, stratabound altered footwall zones are more commonly developed in volcanic rocks with moderate- to high-stratal permeability (e.g. volcaniclastic facies such as pumice breccia and volcanic

FIGURE 7.11 | Genetic models for the formation of stratabound altered footwall zones related to VHMS mineralisation. Fluid flow below and parallel to the seafloor and stratigraphy is controlled by the distribution of permeable volcanic facies (e.g. volcaniclastic units), or impermeable cap-rocks (e.g. sills or lavas). (A) Stratabound subseafloor replacement mineralised and altered zones (e.g. Mount Lyell deposit, Mount Read province and TAG deep Cu zone, Middle Valley, Juan de Fuca Ridge). (B) Stratabound ore lens and altered zones confined below an impermeable volcanic unit such as a sill (e.g. K lens at Rosebery, Mount Read province).

1 7 4 | CHAPTER 7

sandstone). In high-permeability rocks, hydrothermal fluids move laterally along the strata, sub-parallel to the seafloor, and metals are deposited due to the mixing of hydrothermal fluids with seawater and/or cooling. In these cases, fluids are poorly focussed and alteration tends to be of lower intensity and greater in lateral extent (Fig. 7.11). Sericite-altered rocks dominate stratabound altered footwall zones, and chloritic and siliceous zones are restricted to the immediate proximity of massive sulfides, where temperatures and fluid-rock ratios were at a maximum. At K lens in the Rosebery deposit, Allen (1994b) proposed that rising hydrothermal fluids were constrained to flow laterally below an impermeable quartz porphyry sill. As a result, stratabound massive sulfide and associated stratabound altered zones developed beneath the sill by replacement of more permeable and chemically reactive felsic pumice breccias (Fig. 7.1 IB).

There are two ways that acidic fluids may be generated to stabilise kaolinite or pyrophyllite in VHMS systems. The first is by acid-producing fluid-rock reactions, such as the replacement of albite (and volcanic glass) by Mg-Fe chlorite (reaction R7.2). This may result in a reduction in fluid pH by about 1 unit (from 5.5 to 4.5); however, buffering by sericite + chlorite assemblages in the rock will generally prevent the pH dropping below 3.5, which is needed for kaolinite and pyrophyllite formation. These acidic alteration minerals can only form by this method in volcanic rocks that contain negligible K (e.g. tholeiitic basalts) and thus contain no sericite to buffer the pH. The second method is by the introduction of magmatic volatiles at some stage during the life of the hydrothermal system (Ohmoto, 1996). Cooling of a magmatic gas containing SO2 at temperatures below 400°C will increase fluid acidity to levels below pH = 3.5 by a reaction similar to R7.7 (Burnham and Ohmoto, 1980):

Significance of pyrophyllite and kaolinite in VHMS systems

4SO2(g) + 4H2O(1) -» H2S(aq) + 3H+ + 3HSO4~

Pyrophyllite and kaolinite are generally rare in VHMS altered zones; however, because they are only stable under relatively acidic conditions their presence warrants some discussion. Kaolinite has been reported from the sericite and montmorillonite zones of some of the Japanese Kuroko deposits (e.g. Iijima, 1974; Ohmoto, 1996). It has also been reported in the altered footwall zone of the Mount Chalmers Cu-Au VHMS deposit in Queensland (McLeod, 1987). Pyrophyllite exists in the sericite zone of the Western Tharsis VHMS deposit in the Mount Lyell field, Mount Read province (e.g. data sheets WT4 and WT5: Huston and Kamprad, 2001), and is also reported in the stockwork zones of several VHMS deposits in the Iberian pyrite belt (Relvas etal., 1997). Figure 7.12 shows that the stability relationship between kaolinite and pyrophyllite is temperature dependent: pyrophyllite being stable above 280°C, whereas the stability between muscovite and pyrophyllite is controlled by the 3.K+/3.H+ ratio. For a fluid with 3.K+ varying from 0.1 to 0.01, which is considered the range for VHMS-related fluid, then pyrophyllite is stable at pH values of less than 3—4 units.

(R7.7)

Researchers have recently argued that the presence of pyrophyllite or kaolinite in VHMS altered zones supports the theory that these deposits are not simply the products of seawater convection, but that their genesis involves input of a magmatic-derived, low-pH fluid (e.g. Sillitoe et al., 1996; Huston and Kamprad, 2001).

Metamorphism of altered zones Few detailed studies have been published on the effects of contact and regional metamorphism of VHMS-related altered zones. Medium- to high-grade metamorphism of chlorite and sericite zones leads to assemblages containing cordierite, anthophyllite, garnet, biotite, andalusite, staurolite, gahnite, hornblende and plagioclase, depending on the bulk composition of the altered zones (Franklin et al., 1981). Cordierite + anthophyllite assemblages commonly result from metamorphism of chlorite-rich altered zones, with a spotted texture due to cordierite porphyroblasts, leading to the term dalmatianite (Fig. 3.7A: Riverin, 1977). Some examples of mineral assemblages from metamorphosed hydrothermally altered zones are provided in Table 7.1.

7.3 | THE SPECTRUM OF VOLCANICHOSTED DEPOSITS AND ASSOCIATED ALTERATION PATTERNS

FIGURE 7.12 | Stability relations among selected silicate minerals at 500 bars pressure (modified after Beane and Titley, 1981).

Our research into the Palaeozoic VHMS deposits of eastern Australia (Large, 1992; Gemmell et al., 1998; Gemmell and Herrmann, 2001; Large et al., 2001c) has revealed considerable variation in terms of volcanic environment, ore body shape, metal ratios, metal zonation, alteration mineral assembalges and zonation, and ratio of massive to stringer and disseminated styles of ore. Large (1992) identified 10 major styles of deposits in Australia (Fig. 7.13), only one of which was the classic mound style and associated footwall alteration pipe depicted in Figure 7.3. A number of factors

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 7 5 TABLE 7.1 | Mineral assemblages recorded in medium- to high-grade metamorphosed hydrothermally altered zones local to VHMS deposits.

Mattabi

Chloritoid + siderite + andalusite

Chlorite

Franklin et al. (1977)

Mattabi

Quartz + chloritoid + andalusite + kyanite

Silica core

Franklin et al. (1977)

Coronation

Cordierite + anthophyllite

Chlorite

Whitmore (1969)

Anderson Lake

Mg-chlorite + biotite + kyanite

Chlorite

Franklin etal. (1981)

Amulet A

Anthophyllite + cordierite

Chlorite

Beaty and Taylor (1979)

Geco

Bi

te r te + muscovite + almandine + cordierite + °Janthophylnte ! I f ? ++. chl0 Chlorite |! staurolite

Stanton (1984)

x

'

Balcooma

Chlorite + quartz + staurolite + biotite + cordierite + garnet

Chlorite

Huston etal. (1992)

Balcooma

Quartz + muscovite + biotite

Sericite

Huston etal. (1992)

Dry River South

Quartz + muscovite + biotite + staurolite + andalusite

Sericite

Huston et al. (1992)

Skellefte district

Chlorite + cordierite + andalusite

Chlorite

Weihed et al. (2000)

Boliden

Sericite + quartz + andalusite + corundum

Central zone

Nilsson (1968)

including temperature and salinity of the hydrothermal fluid; oxidation and H 2 S/SO 4 characteristics of the hydrothermal fluid and seafloor environment; composition of volcanic and sedimentary rocks deep in the succession; permeability of the footwall volcanic rocks; and depth of seawater control the metal carrying capacity of the fluid, and the chemical reactions that occur beneath and at the seafloor, leading to a broad range of deposit styles and associated local alteration halos. The spectrum of VHMS deposits found in submarine volcanic successions indicates that there may be continuum of deposit styles between the end members that form the basis for the current deposit classification in volcanic arc and rift environments: VHMS Cu-Zn-Pb, porphyry Cu-Au, epithermal Au-Ag and SEDEX Zn-Pb-Ag deposits. Recent workers have emphasised the possible continuum between VHMS and epithermal deposits (e.g. Lydon, 1996; Sillitoe et al., 1996). Large (2000, 2004) and Large et al. (2001c) extended this approach to include porphyry and SEDEX deposits. Sillitoe et al. (1996) introduced high-sulfidation and low-sulfidation VHMS deposits based on mineralisation style, hypogene alteration and sulfide mineralogy, seawater depth, and volcano-magmatic setting. High-sulfidation VHMS deposits develop in shallow-water environments, proximal to volcanic centres, and tend to be associated with zones containing argillic alteration minerals (e.g. pyrophyllite, allunite, kaolinite, diaspore) and high-sulfidation sulfide minerals (e.g. enargite, luzonite, bornite, tennantite). For most geologists the high-sulfidation—low-sulfidation classification scheme has proven difficult to embrace because these terms were originally based on the chemistry of the ore fluid and environment of mineral deposition, rather than a series of geological criteria that could be measured and applied in the field or in drill core. For this reason we do not endorse adoption of the high-sulfidation—low-sulfidation terminology for VHMS deposits. An alternative approach, suggested by Large et al. (2001c) and expanded here, is to place individual deposits within a range of features defined for ores in volcanic arcs and rifts. A diamond-shaped diagram (Fig. 7.14) shows deposits plotted

in terms of their attributes relative to end-member deposit models for VHMS, epithermal Au, porphyry Cu and SEDEX Zn-Pb deposits. The main eastern Australian deposits described in this chapter, and the Bathurst 12 deposit from the Bathurst mining camp, are plotted on the diagram. Hellyer plots very close to the ideal VHMS deposit end member. This is because the deposit exhibits most of the features of the idealised VHMS alteration-mineralisation system outlined in Figure 7.3. Mount Lyell (Western Tharsis, Section 7.7) and HighwayReward (Section 7. 10) plot toward the porphyry end of the spectrum with a significant magmatic component. This is because these deposits are Cu-Au-rich subsurface replacement ores that formed in proximal volcanic environments dominated by synvolcanic rhyolitic intrusions. Mount Lyell also contains alteration minerals typical of an acid fluid or high-sulfidation epithermal environment (pyrophyllite and zunyite), which may suggest that magmatic fluid was involved (Huston and Kamprad, 2001). These Cu-Au-rich massive sulfide deposits and others like them (e.g. Mount Morgan, Boliden, Bousquet) are considered to be hybrid VHMS-epithermal-porphyry deposits that are not easily classified as end members in the VHMS-epithermal-porphyry spectrum (Large, 2004). Henty is a gold-rich, base-metal-poor volcanic-hosted deposit within the Mount Read province (Section 7.8). The gold ore occurs in a stratabound subseafloor replacement zone surrounded by concentric altered zones dominated by quartz, sericite, carbonate and albite. The deposit is neither a typical VHMS nor an epithermal deposit, but has some features of both, and is best described as a hybrid VHMS-epithermal deposit. Rosebery is a Zn-Pb-Ag-Au massive sulfide deposit (Section 7.6). It differs from the classic VHMS deposit in its low Cu content, sheet-like stratiform nature with no stringer zone, and lack of a well-defined footwall alteration pipe. It is hosted in proximal and distal volcanic facies dominated by pumice breccia, volcanic sandstone and siltstone, and black shales. Rosebery and other sheet-like Zn-Pb-rich deposits, such as those in the Bathurst mining camp, have many features similar to SEDEX deposits even though they are in volcanic

1 7 6 | CHAPTER 7

FIGURE 7.13 | Schematic representation of the various shapes and alteration zonation pattens associated with VHMS deposits (modified after Large, 1992).

successions. Recognising their hybrid natures and range in features we have plotted the Rosebery and Brunswick No. 12 deposits on the boundary between VHMS and SEDEX deposits (Fig. 7.14).

Hydrothermal alteration related to the spectrum of deposits Deposits in the VHMS spectrum exhibit a continuum of alteration zonation patterns that are depicted in Figure 7.15. The shapes of the alteration halos, their mineral assemblages and zonation, change progressively along the spectrum. Porphyry Cu (-Au) deposits (Fig. 7.15A) exhibit a series of very extensive roughly concentric altered zones. These include potassic zones in the cores (K-feldspar and/or biotite),

enveloped by phyllic zones (quartz + sericite + pyrite), and finally propylitic zones (carbonate + chlorite + epidote) at the margins, which merge with the regional diagenetic or metamorphic facies. Hybrid massive sulfide Cu-Au deposits (Fig. 7.15B) exhibit similarly shaped, but less extensive, altered hanging wall zones compared to porphyry Cu deposits. Although they lack potassic zones, they comprise siliceous and/or chlorite core zones, which contain Cu-Au ore, and are surrounded by sericite zones and propylitic halos. Alteration minerals characteristic of highly acidic alteration (e.g. pyrophyllite, kaolinite, zunyite, topaz) may occur in the sericite zone in a similar pattern to porphyry systems. Classic mound or lens-shaped Cu + Zn ± Pb-rich VHMS deposits have both massive sulfide and footwall stringer zones with well-zoned footwall alteration pipes (Fig. 7.15Q and

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 7 7

KEY ALTERATION MINERALS chlorite sericite carbonate

subordinate hanging wall altered zones, unlike the previous members of the spectrum where altered rocks dominate the hanging wall. Mg-bearing altered zones (Mg-chlorite or talc) are common features of these deposits dependent on the chemistry and temperature of the fluids. Sheet style VHMS deposits are Zn-rich, strata parallel and have extensive stratabound altered zones (Fig. 7.15D). Stringer zones are less common, but where present they are stratabound rather than pipe-shaped. Carbonates are common alteration minerals, particularly around the margins of the massive sulfide ore. SEDEX Zn-Pb-Ag deposits are at the distal end of the massive sulflde spectrum in terms of proximity to volcanic centres, ratio of sedimentary to volcanic host rocks, and temperature of formation (Fig. 7.15E). Alteration halos are

FIGURE 7.15 | Variations in alteration halos for the spectrum of deposits from porphyry Cu-Au to SEDEX Zn-Pb-Ag. (A) Classical porphyry Cu-Au deposit (e.g. El Salvador, after Gustafson and Hunt, 1975; McMillan and Panteleyev, 1998). (B) Hybrid Cu-Au massive sulfide deposit (e.g. Mount Morgan, Mount Lyell, or Highway-Reward, after Large et al., 2001c). (C) Classic mound-style VHMS Cu-Zn or Cu-Zn-Pb deposit (e.g. Hellyer, after Gemmell and Fulton, 2001). (D) Sheet-style VHMS deposit (e.g. Rosebery, after Large et al., 2001b). (E) Classic SEDEX deposit (e.g. HYC, after Large et al., 2000).

1 7 8 | CHAPTER 7

commonly stratigraphically controlled and dominated by carbonate minerals (ferroan-dolomite, ankerite and siderite, Large et al., 2001c). Chlorite and siliceous zones are rare, and a sericite zone is usually restricted or entirely absent. Manganese and Tl are common alteration halo indicators in SEDEX and sheet style VHMS deposits, but are less common in the mound-style VHMS, hybrid Cu-Au deposits and porphyry Cu deposits (Large et al., 2001c).

7.4 | COMPARISONS BETWEEN ARCHAEAN, PALAEOZOIC AND CAINOZOIC VHMS ALTERATION SYSTEMS Australian Palaeozoic VHMS alteration halos The altered zones around eastern Australian Palaeozoic VHMS deposits have diverse morphologies and mineral assemblages related to variations in their volcanic settings and modes of formation (Large, 1992; Large et al., 2001c). Welldefined footwall alteration pipes are relatively uncommon, or perhaps unrecognised because of subsequent deformation. In Australian deposits, footwall alteration pipes are mainly associated with synvolcanic faults (e.g. Mount Morgan, Taube, 1986) and with relatively impermeable footwall rocks (e.g. Hellyer and Highway-Reward, Gemmell and Large, 1992; Large, 1992; Doyle, 2001). Laterally extensive stratabound altered footwall zones are more typical, especially beneath sheet-like Zn-rich polymetallic deposits, such as Rosebery, Hercules and Thalanga. Stratabound altered footwall zones are attributed to non-focussed discharge and lateral migration of hydrothermal fluids in permeable volcaniclastic units in the footwall (e.g. Rosebery, Green et al., 1981). Quartz + sericite + pyrite assemblages are volumetrically dominant in all types of alteration halos. The proximal altered footwall zones are typically quartz rich, containing less than 20% phyllosilicates but greater than 5% pyrite in disseminations and veins. Much broader footwall feldspardestructive altered zones, with mineral assemblages dominated by sericite or sericite + chlorite, a lower proportion of quartz, and a few percent of disseminated pyrite, envelop them. Chlorite-rich assemblages tend to be restricted to small zones immediately beneath ore lenses (e.g. Rosebery and Thalanga, Green et al., 1981; Paulick et al., 2001), and in several cases are associated with carbonate (e.g. Hellyer, Thalanga and Woodlawn, Davis, 1990; Herrmann and Hill, 2001) . Where there are footwall alteration pipes, chlorite exists in the medial zones, usually between a quartz-rich core and a surrounding sericite zone (e.g. Hellyer and Highway-Reward, Gemmell and Large, 1992; Doyle, 2001). However, the chlorite-rich footwall alteration pipes that are characteristic of many Canadian Archaean deposits do not seem to be present in the Australian Palaeozoic deposits. Several of the Tasmanian examples are virtually devoid of chlorite, such as Henty, Boco and Chester (Boda, 1991; Green and Taheri, 1992; Callaghan, 2001). These are base-metal-poor systems; some

of them contain aluminous minerals such as pyrophyllite and kaolinite. Recent interpretations suggest they are analogous to seafloor acid-sulfate epithermal systems, and possibly involve significant magmatic fluid in their formation (Large et al., 2001c). The sheet-like Zn-rich deposits do not have extensive visually recognisable hydrothermally altered hanging wall zones, although there may be subtle hanging wall geochemical halos (e.g. Rosebery, Large et al., 2001b). However, some deposits with vertical pipe-like footwall alteration and/or mineralised zones exhibit altered hanging wall zones that extend for several tens to hundreds of metres above the deposit. For example, the 200 m thick basalt unit overlying the Hellyer deposit contains an upward flaring plume of distinctive green, Cr-bearing muscovite (fuchsite), which is more or less concentrically enclosed by discontinuous chlorite + carbonate and quartz + albite and sericite zones (Gemmell and Fulton, 2001). A stratabound quartz + albite (± chlorite) altered zone up to 100 m thick exists above the pyritic zones at Henty (Halley and Roberts, 1997). The altered hanging wall zones at Hellyer and Henty contain only traces of pyrite, which suggests that they formed from fluids of very different composition to those in the footwall. In contrast, the quartz + sericite + pyrite zones extending above the Highway and Reward massive sulfide bodies are essentially similar to the footwall stringer zones. This supports the interpretation that these deposits formed entirely below the seafloor (Doyle and Huston, 1999). Carbonate alteration facies are common features of the Australian Palaeozoic polymetallic deposits. They are typically thin stratabound zones that enclose, lie immediately above, or are laterally equivalent to, the sulfide lenses (e.g. Rosebery, Henty and Thalanga, respectively). Except at Henty, the hydrothermal carbonates are generally not calcic; they have various Ca-Mg-Fe-Mn compositions, which in some cases vary systematically towards ore (e.g. Rosebery, Large et al., 2001b).

FIGURE 7.16 | Idealised cross-section of the four altered zones around Kurokotype massive sulfide deposits, Japan (modified from Franklin et al., 1981).

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS |

Japanese Cainozoic VHMS alteration halos The Miocene deposits of the Hokuroku district in northern Japan generally have similar hydrothermal alteration facies to the Australian deposits, although the transitions to diagenetic facies in the host succession is clearer because they are not deformed or metamorphosed. The idealised Kuroko deposit model has four altered zones surrounding the sulfide deposit (e.g. Fig. 7.16: Franklin et al., 1981). These include an inner quartz + sericite zone immediately beneath the sulfide deposit. The inner quartz + sericite altered zone typically encloses a quartz + pyrite stockwork zone (Keiko ore) that underlies the massive sulfide ore (Oko and Kuroko ores) (Ohmoto and Skinner, 1983) and is analogous to pyritic stringer zones beneath some Australian deposits. The inner quartz + sericite zone is laterally surrounded by a sericite + Mg-chlorite + montmorillonite zone that extends in a thin layer over the top of the deposit. It is succeeded outward by mixed-layer clay alteration facies (mineral assemblages of sericite + interlayered illite/smectite + chlorite + albite + K-feldspar), which may extend for up to several kilometres laterally and 200 m into the hanging wall. The outermost zeolite zone typically contains relict plagioclase in mineral assemblages progressing from analcime + montmorillonite + quartz ± calcite, through mordenite + montmorillonite + quartz ± inter-layered illite/smectite, to the background diagenetic clinoptilolite + mordenite alteration assemblage. There is some deposit-specific variation within that idealised Kuroko alteration zonation pattern. For example, at the Fukuzawa deposits the sericite + Mg-chlorite zone in the hanging wall contains relict plagioclase, and analcime seems to exist in more distal parts of the zeolite zone than mordenite (Date et al., 1983). Alteration mineral assemblages in the altered footwall zones around the Uwamuki deposits broadly conform to the idealised zonation but include peripheral zones of kaolinite in (presumably disequilibrium) mineral assemblages of sericite + chlorite + quartz + albite + pyrite (Urabe et al., 1983). These authors noted that kaolinite is not otherwise common around Kuroko deposits and, where present, usually occurs in the core zones with pyrophyllite + diaspore. Subsequent work around the Uwamuki deposit by Shikazono et al. (1998) indicated that the kaolinite zones are greater than 200 m from ore and extend into the hanging wall and therefore may not have been directly related ore deposition. Marumo (1989) also found kaolin minerals in the hanging wall of the small Inarizawa sulfide deposits and concluded that they formed during a low-temperature waning phase of the ore-related hydrothermal system. The existence of kaolinite ± pyrophyllite ± diaspore assemblages, characteristic of low pH, acid-sulfate systems, suggests that the Hokuroku district also contains a spectrum of volcanichosted deposits similar to those recently recognised in the early Palaeozoic belts of Tasmania and north Queensland (Large et al., 2001c). Carbonate-bearing assemblages have not been widely described in the Hokuroku district. Nevertheless, Shikazono et al. (1998) reported that magnesite, siderite, dolomite and calcite were common and characteristic in the ore horizon and hanging wall rocks. Based on isotopic and fluid inclusion data from Uwamuki they concluded that carbonates precipitated in a post-ore hydrothermal stage by interaction

of hydrothermal fluids with biogenic marine carbonates. The erratic distribution and the post-ore formation of carbonates, and indications of complex overprinting of different systems (VHMS and acid-sulfate) leaves some doubt about the genetic relationships between the massive sulfide deposits and carbonate assemblages.

Canadian and Australian Archaean VHMS alteration halos There are two major classes of altered zones associated with Canadian Precambrian massive sulfide deposits: (1) welldefined narrow footwall alteration pipes, and (2) broad irregular altered footwall zones that are transitional to deep semi-conformable alteration facies, with or without localised pipes (Morton and Franklin, 1987; Kerr and Gibson, 1993; Gibson et al., 1999). The former are commonly associated with small (<5 Mt) Cu-Zn deposits and are interpreted to have formed in deep water in dominantly coherent mafic volcanic rocks. The latter generally exist beneath larger deposits of variable metal associations formed in relatively shallow water (^500 m) and in dominantly felsic volcaniclastic rocks. Pipe-like altered footwall zones are epitomised by the CuZn deposits of the Noranda district in the Abitibi belt. These characteristically have upward flaring footwall alteration pipes that are roughly circular in plan view, generally with slightly smaller diameter but greater vertical extent than the overlying massive sulfide lenses. They are commonly recognisable for up to 1 km below the deposits (Franklin et al., 1981). The upper part of the footwall alteration pipe (Fig. 7.17) encloses a stringer zone or stockwork of pyrite ± chalcopyrite ± pyrrhotite veins in a core dominated by Fe-chlorite passing laterally and upward through Mg-chlorite to an outer zone dominated by sericite ± chlorite ± quartz (Lydon, 1984; Kerr and Gibson, 1993). Lydon (1996) noted the existence of talcbearing or aluminous assemblages in the upper parts of some footwall alteration pipes. Depletions of Si, Na, Ca and K and additions of Mg and Fe generally characterise the alteration of the chloritic core,

FIGURE 7.17 | Idealised cross-section of a typical zoned footwall alteration pipe beneath Noranda-type massive sulfide deposits, Abitibi belt, Canada (modified from Lydon, 1984; Kerr and Gibson, 1993).

179

1 8 0 | CHAPTER 7

whereas small additions of K and possibly Si occur in the sericitic shell (Barrett and MacLean, 1994b). This generally results in significant net loss of mass, due to Si loss, from the overall alteration pipe (Barrett and MacLean, 1991). In a few unusual cases there may be net mass gains (e.g. Norbec deposit: Barrett and MacLean, 1999). Footwall alteration pipes of this type may represent zones of hydrothermal discharge that were focussed by synvolcanic faults. Their vertical extent suggests that the hydrothermal fluid sources were very deep. Overprinting relationships indicate that the footwall alteration pipes were initially zones of sericite ± quartz altered rock. As the hydrothermal system intensified, sericite was replaced by chlorite concurrent with metal zone refining in the sulfide lenses (Kerr and Gibson, 1993). Non pipe-like, broad altered footwall zones have more variable morphologies and mineral assemblages as exemplified by the differences in the Home and Mattabi deposits. The Home deposit has a poorly defined altered footwall zone of quartz + sericite ± chlorite that is many times wider than the massive sulfide bodies (MacLean and Hoy, 1991; Kerr and Gibson, 1993). Calculations by MacLean and Hoy (1991) indicate that the Home footwall alteration was accompanied by significant net mass gains; mainly gains of Si, Fe and K, slightly offset by losses of Na, Ca and Mg. Beneath the Mattabi deposit are siderite + chloritoid ± andalusite, kyanite and pyrophyllite zones, which narrow with depth and are transitional downward and laterally into an extensive semiconformable ankerite + chlorite + sericite + quartz zone (Franklin et al., 1975; Morton and Franklin, 1987). Gibson et al. (1999) suggested that the aluminous assemblages at Mattabi (and several other deposits that are notably Au rich) were analogous to the advanced argillic assemblages formed by low pH fluids in acid-sulfate epithermal systems. The Archaean massive sulfide deposits of the Panorama district in the Pilbara of Western Australia occur near the top of a 2 km thick basaltic to rhyolitic volcanic succession above a large synvolcanic granite pluton (Brauhart et al., 2001). Large, semi-conformable altered zones of feldspar-destructive sericite + quartz and chlorite + quartz alteration assemblages occupy the lower and middle parts of the volcanic succession and extend almost the entire exposed strike length (20 km). Locally transgressive chlorite + quartz altered zones, bounded by synvolcanic faults, extend upwards from the semiconformable altered zones to beneath the massive sulfide prospects. Mass changes in the feldspar-destructive altered zones were modest: small gains of Si and losses Ca, Na, Fe and K in the lower sericite + quartz zones, and small gains of Mg, Fe, Si and losses of K, Na, Ca in the transgressive chlorite + quartz zones. In the Golden Grove district of the Archaean Yilgarn craton, Western Australia, the volcanic succession that hosts the Scuddles and Gossan Hill massive sulfide deposits also exhibits the effects of regional-scale, intense feldspar-destructive alteration. The entire footwall succession of altered andesitic to rhyolitic volcaniclastic units, although preserving primary volcanic textures, is composed essentially of quartz + chlorite (± minor sericite). The alteration process, interpreted as a syn-depositional or early hydrothermal regional metasomatic event, virtually removed all Ca, Na and K from the rocks and added substantial Si, Fe and Mg (Sharpe and Gemmell, 2001). At Gossan Hill the regional quartz + chlorite alteration

facies is overprinted by two alteration facies related to sulfide mineralisation. A narrow stratabound chlorite (± siderite, ankerite talc and andalusite) zone envelops the lower Cu-rich massive magnetite + sulfide ore lens. An intense quartz zone underlies the upper Zn-rich massive sulfide lens and encloses a discordant zone of sulfide stringer veins that connects the upper and lower lenses (Sharpe and Gemmell, 2001). There is considerable diversity among alteration facies aroundArchaean massive sulfide deposits. Features that appear to be common to most Archaean districts are large semiconformable altered zones and localised discordant altered footwall zones. Brauhart et al. (2001) highlighted some of the differences in mineral assemblages and mass changes between Panorama and the Canadian semi-conformable altered zones. However, the well-defined discordant footwall alteration pipes are typically chlorite rich (if not metamorphosed to higher grades) and characterised by significant net mass loss, which is attributable to major Si loss and only partly offset by Mg and Fe gains.

Comparisons Despite the many variations in mineral assemblage, morphology and extent of alteration facies associated with VHMS deposits, both within districts and across geologic time, there is one feature that is common to all: proximal altered footwall zones do not contain feldspar. Feldspar destruction is usually manifest in the presence of sericite, chlorite or smectite clays, or their higher grade metamorphic equivalents. One or more of these Al-bearing phyllosilicates is almost invariably present because, although VHMS-type hydrothermal fluids readily transport silica, alkalis and other cations, Al is generally immobile in these moderately acidic systems. One suspects that pyrite is an equally ubiquitous component of proximal altered footwall zones but it is frequently not mentioned in alteration mineral assemblages, due to the unnatural distinctions that many authors make between alteration and mineralisation. It is also becoming increasingly apparent that alumino-silicates such as kaolinite, pyrophyllite, andalusite and others exist locally in altered zones across the entire age spectrum of VHMS deposits, from Archaean to Cainozoic, and that there may always have been continua between moderate and low pH submarine hydrothermal systems. Because of the consistent feldspar destruction, alteration indices such as Al (Ishikawa et al., 1976), CCPI and S/Na2O (Large et al., 2001a) should be effective indicators of alteration intensity in all kinds of VHMS districts, at least at prospect scales. Footwall alteration facies of the Australian Palaeozoic and Japanese Cainozoic deposits are typically more sericitic than chloritic, and their proximal zones tend to be quartz rich. The limited mass change data available indicate they have typically undergone moderate to large net mass gains, which are dominantly attributable to gains in Si that significantly outweighed losses and gains in other components. This predominance of sericite is probably not due to low availability of Fe and Mg. Iron may be significantly added as pyrite and Mg commonly appears in carbonates or chlorite in upper or upper-peripheral altered footwall zones. In contrast, chlorite dominates the well-defined footwall alteration pipes that underlie many Canadian Archaean

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 8 1

deposits. These zones are characterised by significant net mass loss, in which the large loss of Si outweighed addition of Mg and Fe. Although there are some exceptions (e.g. Panorama), major Si and net mass losses are indicated wherever chlorite is dominant in an alteration mineral assemblage. This generalisation also applies to Australian Palaeozoic systems; for instance the small chlorite-rich zones in the Hellyer alteration pipe (Gemmell and Large, 1992) and Thalanga footwall (Herrmann and Hill, 2001). In terms of mass change, the major difference between Archaean deposits with chloride footwall alteration pipes, and Palaeozoic to Cainozoic deposits with quartz + sericitedominated altered footwall zones, is that the former lost mass and the latter gained mass. In addition, in all cases, the major contributor to net mass change was Si. This difference in the behaviour of Si is probably related to the evolution of the hydrothermal systems and particularly the compositions of hydrothermal fluids, which originated as seawater in both cases. Evidence of chlorite overprinting sericite ± quartz assemblages in the Canadian footwall alteration pipes suggests that fluid compositions changed as the hydrothermal system intensified. The initial fluid was probably over-saturated in Si and deposited quartz along the discharge zone to the seafloor as it cooled, whereas the later fluid, possibly of higher temperature and associated with Cu enrichment of the sulfide deposit, was undersaturated and leached Si from the core of the discharge zone. This change may be explained in terms of the regional deep semiconformable altered zones associated with Archaean deposits. The lower semi-conformable altered zone is typically a zone of silicification at temperatures greater than about 400°C (Kennedy, 1950; MacGeehan, 1978; Fournier, 1985; Galley, 1993; Skirrow and Franklin, 1994), attributed to down-going modified seawater being heated up to the range of retrograde Si solubility at 400-600°C (Fournier, 1985). If fluid deposited Si in the deep semi-conformable altered zone, it would then be undersaturated in Si as it ascended the discharge zone, even if it cooled as it ascended. The link between Archaean deposits and regional semiconformable altered zones, which are generally not recognised in the younger VHMS districts, suggests that Archaean crustal conditions (thin crust and large high-level plutonic intrusions) favoured large, intense and presumably long-lived systems. Palaeozoic and younger VHMS districts are not typically associated with large high-level plutons analogous to the Flavrian pluton of the Noranda district (Kerr and Gibson, 1993) or the Strelley granite at Panorama (Brauhart et al., 2001). Their absence may account for the less extensive, perhaps less evolved, altered footwall zones associated with Si and net mass gains that are most common beneath the Palaeozoic and younger massive sulfide deposits.

7.5 | HELLYER: A MASSIVE ELONGATE POLYMETALLIC LENS The Hellyer deposit is located in the northern part of the Cambrian Mount Read province, western Tasmania (Fig. 1.5). The pre-mining resource was 16.2 Mt of 13.9% Zn, 7.1% Pb, 0.4% Cu, 168 g/t Ag and 2.5 g/t Au (Gemmell and Large, 1992; McArthur, 1996). The deposit is a single elongate lens of massive sulfide about 800 m in length (north—south) by 200 m in width (east—west) and with an average vertical thickness of 45 m (Fig. 7.18). It occurs in the mainly calcalkaline, intermediate to mafic Que-Hellyer Volcanics at the base of the Mount Charter Group, which is equivalent to the western volcano-sedimentary sequences (Corbett and

Sericite + quartz zone Sericite zone Chlorite zone Quartz zone

FIGURE 7.18 | Hellyer plan showing the altered zones immediately below the massive sulfide ore (approximately 400 RL). The black line is the outline of the base of the massive sulfide. Modified after Gemmell and Large (1992).

182

|

CHAPTER 7

Komyshan, 1989). The massive sulflde lens is bisected and offset by a major north-south trending fault, the Jack fault (Figs 7.18 and 7.19). Beneath the massive sulflde lens is an elongate, carrot-shaped, zoned, footwall alteration pipe (described in detail by Gemmell and Large, 1992). Above the deposit is a moderately well developed altered hanging wall zone (Gemmell and Fulton, 2001). The massive sulflde lens exhibits classical metal zoning, with minor Cu concentrated in a pyritic core, followed by low grade Zn + Pb, and then high grade Zn + Pb + Ag in the upper parts of the massive sulflde lens (McArthur and Dronseika, 1990; Large, 1992; McArthur, 1996). The centre of the massive sulflde deposit is capped by a quartz + pyrite zone, which is flanked by thin irregular lenses of barite that

are directly over the high-grade massive sulflde (Sharpe, 1991). Barite and massive sulfide clasts occur in volcaniclastic mass-flow units flanking the deposit.

Geological setting The Hellyer ore body occurs above a footwall comprising feldspar-phyric andesitic and basaltic lavas and sills that consist of coherent and autoclastic facies, which are mainly hyaloclastite and peperite (Fig. 7.20: Waters and Wallace, 1992). Basalt (Hellyer basalt) and black mudstone (Que River Shale) dominate the hanging wall (Komyshan, 1986). The abundance of basalt-mudstone peperite indicates that most of the basalt units are sills that intruded the black mudstone (McPhie and Allen, 1992; Waters and Wallace, 1992). Very thick, graded quartz-bearing rhyolitic pumiceous and volcanic lithic breccias interbedded with turbidites and mudstones of the Southwell Subgroup occur in the upper parts of the hanging wall (Corbett, 1992; McPhie and Allen, 1992). The ore lens position is marked by a 0—40 m thick volcaniclastic unit, which mainly consists of coarse polymictic volcanic breccia, sandstone and laminated volcanic siltstone (Waters and Wallace, 1992). Trilobites in the Que River Shale, very thick sections of black mudstone and the abundance of graded mass-flow units collectively indicate that the Hellyer massive sulfide formed in a moderate to deep (>1000 m) submarine setting (Large et al., 2001a). The volcanic facies association indicates proximity to intrabasinal vents for effusive, basaltic and andesitic eruptions and synvolcanic intrusions.

Alteration facies and zonation Gemmell and Large (1992), and Gemmell and Fulton (2001) provided detailed description of both the footwall and hanging wall alteration facies, and zonation at Hellyer. The following section summarises their work.

Footwall alteration facies and zonation A zoned carrot-shaped footwall alteration pipe extends for at least 500 m beneath the Hellyer deposit (Figs 7.6C and 7.20). At the centre of the alteration pipe, immediately below the massive sulfide lens, is a siliceous core zone dominated by intense, pervasive quartz + sericite + pyrite alteration facies. This zone is progressively enclosed in chlorite, sericite and stringer envelope (or sericite + quartz) altered zones (Fig.

7.6C).

FIGURE 7.19 | Cross-section of the Hellyer deposit showing the distribution of rock types, mineralised zones, and altered footwall and hanging wall zones. (A) The ore lens and altered zones are offset along the Jack fault (modified after Gemmell and Large, 1993). (B) Reconstructed 10740 N/10870 N cross-section showing the massive sulfide and footwall alteration pipe prior to folding and faulting (modified after Downs, 1993).

The moderate, selective sericite + quartz alteration facies (e.g. data sheet HE2 in the stringer envelope zone) is the 10-50 m wide outermost part of the alteration pipe and grades outward into weak, selective-pervasive albite + chlorite alteration facies (least-altered footwall, data sheet HE1). Primary volcanic textures are preserved (although modified), lithic fragments exhibit sericite-altered margins, and feldspar phenocrysts are partly altered to sericite. The AI shows an increase from background values of around 30—55 to values of 60-70 (Fig. 2.8B).

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 8 3 Southwell Subgroup: crystal-rich volcaniclastic shale, greywacke and minor felsic lava Que River Shale: black shale and minor sandstone Hellyer basalt: massive to pillowed basalt, pillow breccia, hyaloclastite and andesitic lava Mixed sequence: polymict volcanic breccia, massive and auto-brecciated dacite and massive sulfide ore v

intense footwall alteration plume

" V V

V /

V

,i I V

n

Lower andesites and basalts: andesitic, dacitic and basaltic lavas, hyaloclastites and minor volcani-clastic facies

V

FIGURE 7.20 | Schematic stratigraphic section through the Que-Hellyer Volcanics showing the Lower basalt: massive to pillowed basalts, e x t e n t of a | tered z o n e s at t h e He || yer hyaloclastites and pillow breccias , ~ n. , . . . ..,. . ' and Que River deposits. Modified after Waters and Wallace (1992).

Sericite ± chlorite dominates the strong, selectivepervasive alteration facies in a 10—15 m wide zone marking the outer extent of an intense hydro thermal alteration system, recognised by obliteration of volcanic textures, presence of minor sulfides (mainly pyrite) and complete replacement of feldspar phenocrysts and feldspathic groundmass by sericite ± chlorite (e.g. data sheet HE3). The AI is typically between 70 and 85. On the Alteration box plot samples from this facies plot along a line from the least-altered footwall field toward the chlorite corner. In the intense, pervasive chlorite alteration facies, all primary minerals and glass in the footwall andesitic rocks are completely replaced by fine-grained chlorite with minor pyrite, sericite, quartz and carbonate (e.g. data sheet HE4). The AI is between 90 and 100 and the CCPI between 80 and 90. In the upper parts, immediately below the massive sulfide, this zone includes an intense, spheroidal chlorite + carbonate alteration facies (Fig. 7.6C), which has up to 50% dolomite in a fine-grained matrix of chlorite (e.g. data sheet HE5). This alteration facies has a lower AI (50—80) than the intense chlorite alteration facies due to the elevated whole-rock CaO related to the dolomite component in the rock. In the siliceous core zone, all volcanic textures are completely destroyed and the rock is composed of a fine intergrowth of quartz + sericite + pyrite + chlorite (intense, pervasive quartz + sericite + pyrite alteration facies, data sheet HE6). This zone also contains a series of sub-vertical pyrite + quartz + sphalerite + galena ± chalcopyrite ± carbonate ± barite veins, interpreted as hydrothermal feeders below the ore body. Alteration indices in the siliceous core zone are extremely high with values of both AI and CCPI exceeding 90. Hanging wall alteration facies and zonation

\

Hanging wall alteration facies at Hellyer are less well developed than the footwall alteration facies; however, recent detailed studies by Gemmell and Fulton (2001) have recognised an upward flaring zoned alteration system that is centred above the thickest part of the massive sulfide lens. The

altered hanging wall zone extends through the hanging wall pillow basalts up to the contact with the overlying Que River Shale (Fig. 7.19). Data sheet HE7 is an example of the leastaltered hanging wall andesite. The distribution and intensity of alteration facies in the altered hanging wall zone is patchy, with pillow margins more intensely and pervasively altered (e.g. data sheet HE9) than the pillow interiors. The outer margin of the altered hanging wall zone is defined by weak sericite alteration facies grading inwards to a pink-white, strong, pervasive albite alteration facies (e.g. data sheet HE8), moderate, pervasive chlorite + carbonate alteration facies (e.g. data sheet HE9), and in the centre of the system, a distinctive green, strong, pervasive fuchsite alteration facies (e.g. data sheet HE 10). There is no systematic trend in the alteration indices in the altered hanging wall zones. AI and CCPI values are commonly low in the albite alteration facies due to high Na 2 O, and low MgO and FeO values (e.g. data sheet HE9).

Ore genesis Based on geological, textural, and metal zonation studies, McArthur (1989, 1996), Large (1992) and Gemmell and Large (1992) concluded that the Hellyer massive sulfide deposit grew as a mound in a seafloor depression. The metal zonation from Fe -» Cu —» Pb-Zn —> Ba was interpreted to be an expression of hydrothermal zone refining (Large, 1992), which developed similarly to that described by Eldridge et al. (1983) for the Kuroko deposits. Solomon and Khin Zaw (1999), however, presented fluid inclusion data (indicating high ore fluid salinities: averaging 11 wt%) to propose that sulfide deposition occurred in a seafloor depression brinepool, directly above the footwall alteration pipe. Solomon and Gaspar (2001) provide textural evidence in support of sulfide accumulation in a brine pool. Solomon and Groves (1994) and Solomon et al. (2004) suggest that the abnormally high salinity and other chemical characteristics of the Hellyer fluid inclusions, are strongly suggestive of involvement of magmatic fluids in the hydrothermal system.

1 8 4 [ CHAPTER 7

Weak, selective-pervasive albite + chlorite alteration facies

HE1

Least-altered footwall Sample no.

Hfw-LAA(FPS-I)

Alteration facies

weak, selective-pervasive albite + chlorite

Location

footwall

Formation

Que-HellyerVolcanics

Succession

Mount Read Volcanics

Volcanic facies

monomictic mafic breccia

Relict mineralogy

feldspar

Relict texture

feldspar phenocrysts, perlitic fractures and curvi-angular cm-scale clasts, areas of jigsaw-fit clasts Primary composition andesite Lithofacies

massive, matrix supported and poorly sorted

Interpretation

andesitic hyaloclastite

Alteration minerals

albite + chlorite + sericite + calcite

Alteration textures

selective-pervasive chlorite + calcite + sericite in matrix and chlorite + sericite in clasts, chlorite infill in perlitic fractures, sericite + calcite-altered plagioclase

Distribution

regional

Preservation

good

Alteration intensity

weak

Timing

syn volcanic?

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 54.69 TiO2 0.64 AI2O3 17.92 Fe2O3 7.65 MnO 0.09 MgO 3.79 CaO 3.14 Na2O 6.45

K2O P2O5 S CO2 Total LOI

Photomicrograph (ppl)

1.55 0.12 1.92 97.96 3.96

Rb Sr Ba Cu Pb Zn Sb Tl

68 Zr 299 Nb 500 Y 0 0 Al 0 CCPI Ti/Zr

125 7.0 19 36 57 30.9

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 8 5

Moderate, selective sericite + quartz alteration facies

HE 2

Footwall

I

Sample no.

Hfw-SEZ

Alteration facies

moderate, selective sericite + quartz

Location

footwall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

polymictic mafic breccia

Relict minerals

feldspar

Relict textures

porphyritic and perlitic clasts, areas of jigsaw-fit clasts, clasts with curviplanar margins

Primary composition

andesite-basalt

Lithofacies

massive, clast supported, poorly sorted

Interpretation

resedimented polymictic hyaloclastite

Alteration minerals

sericite + chlorite + quartz + albite + calcite + pyrite

Alteration textures

selective domainal, calcite vein infill, disseminated pyrite, albite + sericite + calcite-altered feldspars

Distribution

alteration zone around pipe

Preservation

moderate

Alteration intensity

moderate

Timing

synmineralisation

Alteration style

peripheral hydrothermal

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

57.38 K2O 0.58 P2O5 15.33 S 7.76 CO2 0.10 Total 3.51 LOI 2.71 1.29

Photomicrograph (xn)

3.34 0.13 4.15 96.27 7.89

Rb Sr Ba Cu Pb Zn Sb Tl

127 97 6700 300 2700 4500

Zr Nb Y

123 7.0 26

Al CCPI Ti/Zr

63 69 28.3

1 8 6 | CHAPTER 7

Strong, selective-pervasive sericite + chlorite alteration facies

HE 3

Footwall Sample no.

Hfw-SZ

Alteration facies

strong, selective-pervasive sericite + chlorite

Location

footwall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

massive polymictic breccia

Relict minerals

rare feldspar

Relict textures

deformed feldspar phenocrysts and andesite clasts

Primary composition dacite-basalt Lithofacies

massive, matrix supported, poorly sorted

Interpretation

resedimented polymictic hyaloclastite

Alteration minerals

sericite + chlorite + quartz + pyrite + ankerite + (albite) selective pervasive, vein-halo (pyrite etc.), disseminated pyrite, and infill (carbonate)

Alteration textures Distribution

pipe

Preservation

poor

Alteration intensity

strong

Timing

synmineralisation

Alteration style

hydrothermal

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO

Na2O K ?°

Hand specimen photograph

54.50 0.52 14.51 10.93 0.21 5.16 1.47

P2O5 S CO2 Total LOI

0.90 ™ ' Ba

3 58

Photomicrograph (xn)

0.10 6.02

Cu Pb Zn 97.92 Sb 8.11 Tl Zr Nb

]]' Y 4100

500 5100 7800

94 6.0

18

Al CCPI Ti/Zr

79 77 33.4

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 8 7

Intense, pervasive chlorite alteration facies

HE 4

Footwall Sample no.

Hfw-CLZ

Alteration facies

intense, pervasive chlorite

Location

footwall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

coherent feldspar-phyric andesite

Relict minerals

nil

Relict textures

porphyritic, perlitic fractures

Primary composition andesite Lithofacies

massive

Interpretation

indeterminate

Alteration minerals

chlorite + pyrite + (quartz + sericite + calcite + galena)

Alteration textures

pervasive, chlorite infill in perlitic fractures, chlorite + quartz-altered feldspars

Distribution

pipe

Preservation

moderate

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2 37.69 TiO2 0.59 AI2O3 16.08 Fe2O3 18.86 MnO 0.41 MgO 11.38 CaO 0.64 Na2O 0.12

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

1.82 0.12 8.65

Rb Sr Ba Cu 96.35 Pb 12.29 Zn Sb Tl

79 39 2000 400 5200 9300

Zr Nb Y

140 8.0 25

Al CCPI Ti/Zr

95 94 25.2

CHAPTER 7

Intense, spheroidal chlorite + carbonate alteration facies

HE 5

Footwall Sample no.

135756

Alteration facies

intense, spheroidal chlorite +carbonate

Location

footwall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

coherent, feldspar-phyric andesite

Relict minerals

nil

Relict textures

porphyritic

Primary composition

andesite

Lithofacies

massive

Interpretation

indeterminate

Alteration minerals

chlorite + dolomite + (quartz + sericite)

Alteration textures

nodules-spheroids

Distribution

local

Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

Geochemistry

hydrothermal

SiO2 34.88 P2O5 TiO2 0.60 S AI2O3 15.63 CO2 Fe2O3 12.89 Total MnO 0.80 LOI MgO 13.07 CaO 4.86 Rb Na2O 0.01 Sr K2O 1.73 B a

Alteration style

Hand specimen photograph

Photomicrograph (xn)

0.13 4.13

Cu Pb Zn 88.72 Sb 15.40 Tl Zr 66 N b 58 1 1200

300 7200 11500

124 90

28

Al CCPI Ti/Zr

75 93 29.0

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 1 8 9

HE 6

Intense, pervasive quartz + sericite + pyrite alteration facies Footwall Sample no.

Hfw-SCZ

Alteration facies

intense, pervasive quartz + sericite + pyrite

Location

footwall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

feldspar-phyric andesite breccia

Relict minerals

nil

Relict textures

porphyritic, periitic fractures, jigsaw-fit clasts

Primary composition andesite Lithofacies

massive

Interpretation

in situ hyaloclastite

Alteration minerals

quartz + sericite + pyrite + (chlorite)

Alteration textures

Distribution

pervasive, pseudomorphs of feldspar ± pyroxene?, quartz and sericite veins disseminated pyrite, feldspar overgrowths on phenocrysts core of pipe

Preservation

moderate

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2 67.42 TiO2 0.30 AI2O3 8.13 Fe2O3 12.34 MnO 0.07 MgO 1.42 CaO 0.32 Na2O 0.03

K2O P2O5 S CO2 Total LOI Rb

Photomicrograph (xn)

2.17 Sr 0.05 Ba 8.87 Cu Pb 101.12 Zn 7.74 Sb Tl 79 Zr

15 Nb 14800 Y 2200 8000 Al 9800 CCPI Ti/Zr 74

4.0 28 91 85 24.7

1 9 0 | CHAPTER 7

Subtle, pervasive aibite + chlorite + calcite alteration facies

HE 7

Hanging wall

Sample no.

142562

Alteration facies

subtle, pervasive aibite + chlorite + calcite

Location

hanging wall

Formation

Que-HellyerVolcanics

Succession

Mount Read Volcanics

Volcanic facies

massive, feldspar-phyric amygdaloidal andesite

Relict minerals

plagioclase

Relict textures

massive, porphyritic, weakly amygdaloidal

Primary composition andesite Lithofacies

massive

Interpretation

lava

Alteration minerals

chlorite + aibite + calcite + quartz + (chalcopyrite) selective-pervasive, chlorite or quartz infill in amygdales, quartz + calcite veins

Alteration textures Distribution

regional

Preservation

good

Geochemistry

Alteration intensity

weak

Timing

synmineralisation

Alteration style

diagenetic to metamorphic

SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na,0

Hand specimen photograph

51.14 0.55 14.77 8.82 0.21 5.56 5.44 4.19

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

0.42 0.60 0.06 4.76 99.91 8.17

Rb Sr Ba Cu Pb Zn Sb TI

21 Zr 125 Nb 226 Y 745 6 Al 147 CCPI 1.1 Ti/Zr <0.5

151 7.1 21 38 75 21.8

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 9 1

Strong, pervasive albite alteration facies

HE 8

Hanging wall Sample no.

142622

Alteration facies

strong, pervasive albite

Location

hanging wall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

feldspar-phyric basalt breccia

Relict minerals

altered plagioclase

Relict textures

porphyritic

Primary composition basalt Lithofacies

massive and jigsaw-fit breccia

Interpretation

lava or sill

Alteration minerals

albite + chlorite + calcite + (pyrite)

Alteration textures

pervasive, albite + chlorite, massive chlorite + pyrite veins, patchy calcite domains possibly irregular pseudomorphs

Distribution

local, plume?

Preservation

poor

Alteration intensity

strong

Timing

post mineralisation

Alteration style

diagenetic-hydrothermal?

Hand specimen photograph

Photomicrograph (xn)

1 9 2 I CHAPTER 7

Moderate, pervasive chlorite + carbonate alteration facies

HE 9

Hanging wall Sample no.

142593

Alteration facies

moderate, pervasive chlorite + carbonate

Location

hanging wall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

monomictic basaltic andesite breccia

Relict minerals

altered feldspars

Relict textures

porphyritic, jigsaw-fit clasts

Primary composition

basaltic andesite

Lithofacies

massive

Interpretation

pillow lava

Alteration minerals

chlorite + caicite + sericite + (albite + quartz)

Alteration textures

pervasive, spheroidal and rhombic caicite, caicite veins with chlorite vein-halo alteration

Distribution

local, plume?

Preservation

good

Alteration intensity

moderate

Timing

post mineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2 42.34 TiO2 0.61 AI2O3 11.78 Fe2O3 5.41 MnO 0.12 MgO 4.34 CaO 15.44 Na2O 1.06

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

1.92 0.35 0.07 13.00 99.46 15.99

Rb Sr Ba Cu Pb Zn Sb Tl

55 Zr 281 Nb 737 Y 109 4 Al 51 CCPI 0.6 Ti/Zr 0.7

140 8.8 21 28 76 26.2

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 9 3

Strong, pervasive fuchsite alteration facies

HE 10

Hanging wall Sample no.

142643

Alteration facies

strong, pervasive fuchsite

Location

hanging wall

Formation

Que-Hellyer Volcanics

Succession

Mount Read Volcanics

Volcanic facies

massive feldspar-phyric basalt

Relict minerals

feldspar

Relict textures

porphyritic, minor amygdales

Primary composition basaltic andesite Lithofacies

massive

Interpretation

lava or sill

Alteration minerals

sericite (fuchsite) + calcite + ankerite + chlorite + pyrite

Alteration textures

pervasive, disseminated pyrite, sericite cleavage, calcite ± sericite pseudomorphs after feldspar

Distribution

plume

Preservation

poor

Geochemistry

Alteration intensity

strong

SiO 2

26.32

TiO2

0.54

AIA

17.35

Timing

post mineralisation

Alteration style

hydrothermal

Hand specimen photograph

K2O S

5.00

Rb

143 Zr

74

0.21

Sr

176

Nb

4.9

0.08

Ba

4473

Y

18

19.48 Cu

162

Fe2O3

3.38

CO2

MnO

0.25

Total

99.15 Pb

2

LOI

21.13 Zn

43

CCPI Ti/Zr

MgO

1.88

CaO

22.67

Sb

8.1

Na2O

0.00

Tl

18.3

Photomicrograph (ppl)

Al

23 50 43.9

1 9 4 | CHAPTER 7

7.6 | ROSEBERY: A POLYMETALLIC SHEET-STYLE DEPOSIT The Rosebery massive sulfide deposit is a sheet-style polymetallic Zn-Pb-Cu-Ag-Au VHMS deposit in the northern Central Volcanic Complex of the Mount Read Volcanics, western Tasmania (Fig. 1.5: Green et al., 1981; Large, 1992). The mining resource is 32 Mt at 14.7% Zn, 4.5% Pb, 0.6% Cu, 146 g/t Ag and 2.3 g/t Au (data from Pasminco Mining and Exploration). Compared with the Hellyer deposit, which comprises a single ore lens (described in the previous section), Rosebery is composed of at least 16 separate ore lenses (Fig. 7.21). These vary in size from 0.1 to 5 Mt. Unlike the carrotshaped footwall alteration pipe at Hellyer, the Rosebery ore lenses are enclosed in strata-parallel altered zones. The ore lenses are principally composed of massive and banded sulfides of sphalerite, galena, barite, pyrite and chalcopyrite, with minor tetrahedrite-tennantite, arsenopyrite, pyrrhotite, hematite and magnetite. In some sections of the mine (e.g. A and B lenses, Huston and Large, 1987) barite-rich lenses overlie the Zn-Pb-Cu ore lenses.

Geological setting The Rosebery, Hercules and South Hercules polymetallic ore bodies are hosted by the same stratigraphic sequence in the upper part of the Central Volcanic Complex, west of the Henty fault (Solomon, 1964; Green et al., 1981). The footwall comprises a thick (up to 500 m), poorly stratified rhyolitic-dacitic succession of weakly graded, feldspar-phyric pumice breccia, which is interpreted to be the product of large volume submarine caldera-forming eruptions, and

rhyolitic and dacitic sills (Hercules Pumice Formation: Lees, 1987; Allen, 1994b; Large et al., 2001b). The ore lenses occur in the 5 to 10 m thick, finely stratified pumiceous siltstone, sandstone, crystal-rich sandstone and claystone top (host rocks) of the footwall pumice breccias (Lees, 1987; Corbett and Solomon, 1989; Allen, 1994b; Large et al., 2001b). The host rocks are overlain by black mudstone, which represents a hiatus in volcanism marked by non-volcanic sedimentation. The hanging wall comprises a 5 to 400 m thick succession of massive to stratified, feldspar + quartz-phyric rhyodacitic volcaniclastic units of the White Spur Formation, interbedded with black mudstone (Lees, 1987; Allen, 1994b). The footwall rhyolitic pumice breccias haveTi/Zr of 7-9 (Fig. 4.5), whereas the host interval porphyry sill has Ti/Zr of 12-14, and the volcaniclastic facies 10-30 (Large et al., 2001b). Bedforms and textures within the footwall pumice breccias and host rocks are consistent with deposition from volcaniclastic turbidity currents, debris flows and suspension in a below-wave-base environment (McPhie and Allen, 1992). The footwall pumice breccias are interpreted to represent the submarine deposits from a large, felsic explosive eruption (Allen, 1994a). The host rocks may have been derived from water-settled suspension sedimentation or the influx of volcaniclastic turbidites from distal rhyolitic volcanic centres (Large et al., 2001b). The hanging wall volcaniclastic units probably comprise the medial to distal facies from an extrabasinal felsic volcanic centre (Allen, 1994a). A below-wave-base submarine setting for the RoseberyHercules succession is indicated by the presence of sedimentary structures in the footwall pumice breccias consistent with deposition from cold water-supported gravity flows and water-settled fall, rare intercalated black pyritic mudstone and the associated VHMS deposits (Gifkins and Allen, 2002).

FIGURE 7.21 | Long-section of the Rosebery mine, western Tasmania, showing the drives and main ore lenses, labelled alphabetically (provided by Zinifex Rosebery mine, 2004). K lens is at depth at the north end of the mine.

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 1 9 5

Alteration facies and zonation

Genesis of the ore lenses and alteration system

Four strata-parallel altered zones enclose the Rosebery ore lenses. From the periphery to the core of the alteration system these are: sericite zone, chlorite zone, Mn-carbonate zone, and quartz + sericite zone (Fig. 7.5C: Large et al., 2001b). Outside the altered footwall zone, least-altered rhyolitic volcanic rocks contain plagioclase crystals in a pumice- and shard-rich matrix (e.g. data sheet RBI). This matrix is weakly sericite + chlorite + quartz altered, commonly enhancing the shard and pumice textures. Plagioclase crystals are weakly altered with disseminated fine-grained sericite and albite overgrowths. These rocks have AI = 30—60 and CCPI = 15— 40 and plot in the least-altered box of the Alteration box plot. Data sheet RB2 is an example of the least-altered hanging wall rocks. The outer part of the hydrothermal alteration system is a broad sericite zone with scattered Mn-carbonate blebs, which extends up to 300 m into the footwall, but less than 25 m into the hanging wall rocks (Fig. 7.5C). It extends along the upper contact of footwall pumice breccia for at least 1000 m beyond the ore lenses (Large et al., 2001b). One texturalcompositional variant of this enveloping zone is represented in alteration facies data sheet RB3. The white mica contents of the facies varies from about 20 to 60% as plagioclase crystals are increasingly replaced by carbonate and sericite, and the glass shard-rich matrix by fine sericite, with proximity to massive sulfide. The AI increases from 60 to 95 as the sericite proportions increase. Intense, schistose chlorite alteration facies (e.g. data sheet RB5) is concentrated in the immediate footwall of the ore lenses forming a thin (typically less than 5 m thick) chloritic zone, which is commonly thickest (5-10 m) beneath the Curich sulfide lenses at the south end of the mine. This alteration facies has variable chlorite (15—50 wt%) and sulfide (10— 30 wt%) contents. The AI is between 95 and 100, and the CCPI between 70 and 90. Commonly overlying the ore lenses is a zone of intense, proximal Mn-carbonate alteration facies up to 10 m thick (e.g. data sheet RB6), which is closely associated with massive sulfide, but locally extends several tens of metres beyond the limits of known sulfide lenses. The intense, proximal Mncarbonate alteration facies typically has a spotty texture, with 25—60% Mn-carbonate spots in a sericitic, or locally chloritic, matrix with low sulfide content. Carbonate composition in this facies varies from rhodochrosite (MnCO 3 ), to manganosiderite ((Mn,Fe)CO3) and kutnahorite (CaMn(CO3)2) (Braithwaite, 1974; Large etak, 2001b). Typically the massive and semi-massive sulfides occur in strata-parallel zones of intense quartz + sericite alteration facies. This alteration facies has a bleached appearance with textures that vary from massive-pervasive to spotty and augenschist textured (e.g. data sheet RB4). The latter comprises an anastomosing fabric of strongly foliated sericite dominated domains wrapping around siliceous knots of quartz with minor sericite. The intense quartz + sericite alteration facies continues laterally beyond the margins of the ore lenses, where it contains 1-10% disseminated sulfides (pyrite, sphalerite, galena).

Most previous workers have interpreted the Rosebery deposit to be synvolcanic exhalative in origin (e.g. Braithwaite, 1974; Green et al., 1981; Huston and Large, 1987; Green and Iliff, 1989; Large, 1992; Khin Zaw et al., 1999). Solomon and Groves (1994) consider that the sheet-like form, stratiform sulfide banding, large size and high Zn-Pb metal content indicate that Rosebery formed within a brine pool from relatively high-salinity fluids, similar to the genesis of many SEDEX deposits. However, the ore lenses are not associated with well-developed stringer sulfide zones or alteration pipes typical of seafloor systems. Instead, there are footwall zones of disseminated sulfides, with altered zones that are aligned parallel to the ore lenses and the volcanic strata. These features, combined with textures in the massive sulfides indicative of replacement, suggest that the ore lenses did not form immediately above hydrothermal vents, but may have formed from lateral fluid flow, either on the seafloor, or below the seafloor by replacement of the fine-grained tops of permeable pumice breccia units (Fig. 7.11: Allen, 1994a; Doyle and Allen, 2003; Martin, 2004).

1 9 6 | CHAPTER 7

Weak, selective-pervasive albite + quartz * sericite alteration facies

RB1

Least-altered footwal! Sample no.

139602

Alteration facies

weak, selective-pervasive albite + quartz + sericite

Location

footwall

Formation

Kershaw Pumice Formation (CVC)

Succession

Mount Read Volcamics

Volcanic facies

feldspar-phyric pumice breccia

Relict minerals

plagioclase

Relict textures

plagioclase crystals and cm-sized plagioclase porphyritic tube pumice clasts

Primary composition rhyolite Lithofacies

massive, clast supported, poorly sorted

Interpretation

subaqueous mass flow deposit

Alteration minerals

albite + sericite + quartz > chlorite + pyrite + hematite

Alteration textures

selective-pervasive, disseminated, foliated, sericite ± chlorite fiamme, sericite + hematite stylolites, albite + sericite altered plagioclase, feldspar overgrowths on plagioclase

SiO2

73.26

Distribution

regional

TiO2

0.32

Preservation

good

AI2O3

14.13

Alteration intensity

weak

Fe2O3

Timing

synvolcanic to burial

2.11 0.03

Total

0.97

LOI

Alteration style

diagenetic

Hand specimen photograph

Geochemistry

MnO MgO CaO

0.99

Na2O

3.87

K2O

2.38

P2O5

0.04

S

0.01

CO2

Photomicrograph (ppl)

98.20

Rb Sr Ba

110 Zr 257 Nb 866 Y

Cu Pb Zn

1 3 Al 26 CCPI

Sb Tl

0.0 Ti/Zr 0.0

257 14.2

33 41 31 7.5

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 197

Subtle, selective albite * quartz + sericite alteration fades

RB2

Least-altered hanging wall Sample no.

139586

Alteration facies

subtle, selective albite + quartz + sericite

Location

hanging wall

Formation

White Spur Formation

Succession

Mount Read Volcamics

Volcanic facies

feldspar > quartz crystal-rich pumiceous sandstone

Relict minerals

plagioclase and quartz

Relict textures

clastic (feldspar and quartz crystals, and pumice shards)

Primary composition rhyolite-dacite Lithofacies

massive

Interpretation

subaqueous mass-flow deposit

Alteration minerals

albite + sericite + quartz + chlorite

Alteration textures

selective clast alteration, disseminated sericite, albite + sericite-altered plagioclase

Distribution

regional

Preservation

good

Alteration intensity

subtle

Timing

synvolcanic

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 74.40 TiO2 0.35 AI2O3 14.34 Fe2O3 2.06 MnO 0.04 MgO 0.57 CaO 0.58 Na2O 4.41

K2O P2O5 S CO2 Total LOI Rb

Photomicrograph (xn)

1.99 0.05 0.03 0.27 99.26

75

Sr Ba Cu Pb Zn Sb Tl Zr

261 1472 3 62 142 0.0 0.0 197

Nb Y

10.8 33 :

Al CCPI Ti/Zr

34 27 10.6

1 9 8 I CHAPTER 7

KB 3

Moderate, foliated sericite alteration facies Footwall Sample no.

139747

Alteration facies

moderate, foliated sericite

Location

footwall

Formation

Kershaw Pumice Formation

Succession

Mount Read Volcanics

Volcanic facies

feldspar-phyric pumice breccia

Relict minerals

plagioclase

Relict textures

porphyritic, fibrous tube pumice clasts

Primary composition rhyolite Lithofacies

massive, normally graded

Interpretation

syneruptive, mass-flow-emplaced deposit

Alteration minerals

sericite + albite + quartz + carbonate

Alteration textures

foliated, schistose, stylolites, fiamme?, fractured and albite-altered plagioclase, quartz veinlets

Distribution

local

Preservation

poor

Alteration intensity

moderate

Timing

synmineralisation

Alteration style

hydrothermal and metamorphic

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

Hand specimen photograph

71.46 K2O 0.32 P2O5 12.93 S 2.44 CO2 0.11 Total 1.65 LOI 1.49 1.36

Photomicrograph (xn)

4.16 Rb 0.06 Sr 0.06 Ba 1.87 Cu 98.02 Pb 3.89 Zn Sb Tl

182 Zr 56 Nb 1042.5 Y 3 5 Al 29 CCPI 2.7 Ti/Zr 0.9

229.9 12.6 32 67 41 8.3

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 1 9 9

RB4

Intense, augen schistose quartz + seriate alteration facies Footwal! Sample no.

139778

Alteration facies

intense, augen schistose quartz + sericite

Location

footwall

Formation

Kershaw Pumice Formation (CVC)

Succession

Mount Read Volcamics

Volcanic facies

feldspar-phyric pumice breccia

Relict minerals

nil

Relict textures

foliated clasts (pumice)

Primary composition rhyolite Lithofacies

massive

Interpretation

indeterminate

Alteration minerals

sericite + quartz + sulfides

Alteration textures Distribution

augen schistose, sericite + sulfide cleavage local

Preservation

poor

Alteration intensity

intense

Geochemistry

Timing

synmineralisation

Alteration style

hydrothermal and metamorphic

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O

Hand specimen photograph

72.95 K2O 0.27 P 2 O 5 12.86 S 2.58 co 2 0.17 Total 1.14 LOI 0.27 0.01

Photomicrograph (xn)

4.24 Rb 0.04 Sr 1.94 Ba 0.40 98.19 3.57

Cu Pb Zn Sb Tl

202 Zr 18 Nb 1767 Y 239 4300 Al 6900 CCPI 7.5 Ti/Zr 4.5

224 13.2 35 95 45 7.2

2 0 0 I CHAPTER 7

Intense, schistose chlorite alteration facies

•RB5

Footwall Sample no.

139743

Alteration facies

intense, schistose chlorite

Location

footwall

Formation

Kershaw Pumice Formation (CVC)

Succession

Mount Read Volcanics

Volcanic facies

feldspar-phyric breccia

Relict minerals

nil

Relict textures

porphyritic?

Primary composition

rhyolite

Lithofacies

massive

Interpretation

indeterminate

Alteration minerals

chlorite + pyrite + sphalerite + quartz + sericite

Alteration textures

foliated, schistose

Distribution

local

Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

SiO2

44.22

K2O

3.37

Alteration style

hydrothermal and metamorphic

TiO2

0.33

0.06

MA

14.48

P2O5 S

Fe2O3

16.58 1.49

co2 Total

0.95 91.03

2.22

LOI

7.55

Geochemistry

MnO MgO CaO Na2O

Hand specimen photograph

0.10 0.05

Photomicrograph (ppl)

7.18

Rb Sr Ba Cu Pb Zn Sb Tl

174 Zr 14 Nb 948 Y

247 12.0

35

1678

604 Al 68200

CCPI

3.9 Ti/Zr 7.1

97 83 8.0

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 2 0 1

RB6

Intense, proximal In-carbonate alteration facies Hanging waff Sample no.

139740

Alteration facies

intense, proximal Mn-carbonate

Location

hanging wall

Formation

White Spur Formation

Succession

Mount Read Volcamics

Volcanic facies

feldspar-phyric pumice breccia?

Relict minerals

nil

Relict textures

rare feldspar crystals

Primary composition dacite Lithofacies

massive

Interpretation

indeterminate

Alteration minerals

rhodochrosite + sericite + pyrite

Alteration textures

nodular-spheroidal rhodochrosite, disseminated pyrite, sericite-altered feldspar

Distribution

local, immediate hanging wall ofsulfide lens

Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Geochemistry SiO2

16.79

K2O

3.64

TiO2

0.39

P2O5

0.13

AI2O3

10.73

Fe2O3

7.50

co2

MnO MgO

29.36

Total

98.64

1.97

LOI

23.35

CaO

2.04 0.05

Rb

183

Na2O

Hand specimen photograph

S

Photomicrograph (ppl)

3.24 22.20

Sr Ba

27 4297

Cu Pb Zn Sb

20 674 995 8.8

TI Zr

24.6

142

Nb Y

7.2 20

Al

73 70

CCPI Ti/Zr

16.5

I CHAPTER 7

7.7 | WESTERN THARSIS: A HYBRID CuAu VHMS DEPOSIT Western Tharsis is a stratabound disseminated pyrite + chalcopyrite deposit in the northwestern part of the Mount Lyell mining field, western Tasmania (Fig. 1.5). It contains around 12.4 Mt at 1.3% Cu and 0.3 g/t Au (Huston and Kamprad, 2001). Although discovered in 1897, it is the only non-exploited deposit of at least 22 deposits in the Mount Lyell field, which together produced a total of 113 Mt of ore at average grades of 1.36% Cu, 6.8 g/t Ag and 0.4 g/t Au (Corbett, 2001). The mineralised zone is a sub-vertical stratabound lens, up to 150 m thick and narrowing towards the surface, with a down dip extent of greater than 1000 m and strike extent of about 300 m. Most of the deposit consists of disseminated pyrite + chalcopyrite in a gangue of quartz + sericite ± chlorite and, locally, magnetite. Smaller bornite-rich mineralised zones similar to the North Lyell ores exist in the upper parts, particularly associated with quartz and quartz + pyrophyllite alteration assemblages. The bornite zones also contain minor chalcocite, chalcopyrite, mawsonite, digenite, enargite, molybdenite, woodhouseite and barite.

Geological setting The Lyell ore bodies and their altered halos are focussed along the Great Lyell fault and occur at a variety of stratigraphic intervals in the Central Volcanic Complex and in the overlying Mount Julia Member of the Comstock Formation, Tyndall Group (Corbett, 2001). Western Tharsis is situated in a steeply west-dipping, overturned, east-facing succession of altered intermediate to felsic volcanic rocks assigned to the Central Volcanic Complex of the Middle Cambrian Mount Read Volcanics. In the Mount Lyell area, these volcanic rocks are reverse-faulted to the east against the late Cambrian to Early Ordovician siliciclastic conglomerate and sandstone of the Owen Group. All the Mount Lyell field deposits lie within a 6 x 1 km pyritic altered zone adjacent to the complex fault contact (Corbett, 2001). Two units of rhyolitic volcaniclastic rocks with subordinate interbedded volcanogenic sandstone and siltstone comprise the immediate stratigraphic footwall and host rocks at Western Tharsis (Huston and Kamprad, 2001). These units, each several hundred metres thick, contain some ash- to lapilli-sized clasts but primary volcanic textures are typically obscured by alteration and their volcaniclastic origin is largely interpretative. They are separated by a 10-50 m thick group of andesitic volcaniclastic rocks and locally amygdaloidal coherent lavas or sills. The stratigraphic hanging wall consists of a 200—300 m thick complex of intercalated felsic and intermediate volcaniclastic rocks. A thin unit of felsic quartz porphyry, possibly a correlate of the lower Tyndall Group, occurs between the altered Central Volcanic Complex and the faulted contact with the Owen Group. In this part of the field, the (North Lyell) fault contact dips at 70° to the southwest. Deep drilling indicates that the Western Tharsis mineralised zone may intersect the fault at around 1500 m below surface (Corbett, 2001). The presence of thick, graded beds in the Central Volcanic Complex at Lyell indicates a subaqueous environment of

deposition for the host succession. The occurrence of exhalative massive sulfide bodies, limestone with shallow marine fauna, and welded ignimbrite in the Tyndall Group at Comstock suggest a shallow marine setting for mineralisation (Jago et al., 1972; Corbett et al., 1974; White and McPhie, 1997; Corbett, 2001).

Alteration facies and zonation Corbett (2001) showed that the deposit is enclosed by a 400-500 m wide zone of quartz + sericite + pyrite schist adjacent to the North Lyell fault. This is part of a pyritic core zone extending 4 km from the Lyell Highway to the Lyell Comstock mine. This proximal, strong to intense, feldspardestructive altered zone grades outwards to less intense sericite + chlorite alteration facies in felsic and intermediate volcanic rocks. At surface above Western Tharsis and around the upper mineralised zone, the proximal quartz + sericite + pyrite alteration facies includes numerous bodies up to 20 m across, of microcrystalline quartz ± pyrite, termed silica heads. Huston and Kamprad (2001) subdivided the Western Tharsis system into five main alteration facies. An intense, pervasive, proximal quartz + chlorite + pyrite ± sericite alteration facies (e.g. data sheet WT8) exists in the chalcopyrite + pyrite mineralised zone at depths greater than 350 m below surface. An intense, proximal quartz + pyrophyllite + pyrite alteration facies (e.g. data sheets WT4 and 5) occurs in a 150 m wide zone associated with the bornite + chalcopyrite mineralised zone between 100 and 400 m below surface and in a 50 m thick zone extending along the stratigraphic footwall to 750 m below surface. This facies includes narrow zones of quartz + topaz assemblages (particularly in the upper parts, which may correspond to Corbett's silica heads) and locally minor phases including fluorite, barite, zunyite and woodhouseite. A strong, pervasive, medial quartz + sericite + pyrite alteration facies (e.g. data sheets WT3, 6 and 7), which encloses the two proximal facies above, and extends up to 150 m outwards into the stratigraphic footwall and through most of the hanging wall. The outer margins, adjacent to the weak, medial chlorite + sericite ± carbonate alteration facies and in upper part of hanging wall succession, contain minor disseminated carbonate. A weak, pervasive, medial chlorite + sericite ± carbonate alteration facies (e.g. data sheet WT2) occurs in the peripheral zones. It exists in both footwall and hanging wall, about 150200 m outwards from the mineralised zone, particularly in andesitic volcanic rocks with minor pyrite or hematite. A weak, selective quartz + chlorite + sericite + carbonate alteration facies (e.g. data sheet WTl) grades westward into least-altered rocks composed of quartz + albite + chlorite (± sericite and carbonate). Sericite compositions in the outer zones are slightly phengitic (up to 0.5 Fe + Mg atoms per formula unit) grading to essentially non-phengitic and slightly sodic (molecular Na/ Na+K <0.15) in the proximal to medial zones. This variation has potential as a deposit-scale exploration vector, which can be effectively measured by short wavelength infrared (SWIR) spectral analysis (Herrmann et al., 2001). SWIRspectrometry

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 0 3

is similarly effective in delineating pyrophyllite, topaz and zunyite-bearing zones (see Section 8.2 for more detail). Chlorite compositions are moderately Fe-rich (molecular Mg/Mg + Fe ratios = 27-48) with a subtle trend to Fe enrichment towards the mineralised zone. Carbonates in the distal to medial alteration facies are ankeritic to sideritic in composition. They show a distinct trend of Fe enrichment from the footwall towards the mineralised zone (Huston and Kamprad, 2001). Carbonates in the hanging wall are moderately manganiferous (up to 0.2 Mn atoms per formula unit).

Ore genesis Metallogenic interpretations of the Mount Lyell deposits have fuelled geological debate for over a century and remain controversial today (Corbett, 2001; Huston and Kamprad, 2001). Early models that related mineralisation to Devonian or Cambrian intrusions were succeeded, during the 1960s, by acceptance of Cambrian synvolcanic origins. In the 1980s and early 1990s, the North Lyell type bornite ores were popularly attributed to re-mobilisation during Devonian deformation. Large et al. (1996) revived the magmatic connection, interpreting Cambrian granites to be the source of hydrothermal fluids and metals. In recent years, a magmatic connection has been further supported by wider recognition of advanced argillic type assemblages, which are consistent with the involvement of magmatic volatiles and acidic fluids. Nevertheless, there is still disagreement over the timing of mineralisation. Huston and Kamprad (2001) pointed to an apparent (Pb-isotopic) 40 Ma age difference between stratiform synvolcanic Pb + Zn + Cu sulfide lenses at Lyell Comstock and the disseminated Cu + Au deposit at Prince Lyell. They suggested a two event history: Middle Cambrian

syngenetic stratiform Pb + Zn + Cu mineralisation followed by a 40 Ma period of tectonism that culminated in highsulfidation type Cu + Au mineralised zones derived from deep Ordovician granites. However, the more extensive field evidence gathered by Corbett (2001) indicates that all the alteration and mineralisation was restricted to Middle Cambrian, and ceased during deposition of the lower part of the Tyndall Group. Rather than a temporal overprinting of different styles, Corbett (2001) envisaged a single, vertically extensive, submarine volcanic, hybrid magmatic-seawater hydrothermal system. It produced disseminated chalcopyrite + pyrite (and locally magnetite + apatite) mineralised zones in the deeper parts, high-sulfidation type bornite mineralised zones and intense siliceous altered zones in the upper subseafloor zones, and deposited exhalative Pb + Zn + Cu massive sulfide lenses at the seafloor. The Western Tharsis zone encompasses the transition between deep chalcopyrite + pyrite and upper highsulfidation types of mineralisation (Fig. 7.22). Corbett's diagrammatic representation shows the system as sub-vertical, cutting through sub-horizontal volcanic strata and focussed along or adjacent to the Great Lyell fault. However, the Western Tharsis deposit appears to be subvertical and stratabound. This is possibly a misinterpretation; primary volcanic textures and facies associations are largely obscured in the intensely altered zones. Furthermore, Corbett's (2001) model suggests diapir-like upward movement of the phyllosilicate-rich altered volcanic rocks on the hanging wall side of the fault zone, which may have disrupted the volcanic sequence. The arguments about Mount Lyell are not yet settled. Nevertheless, the emerging recognition of high-sulfidation ore deposits may renew interest in exploration in western Tasmania.

pyntic core zone (senate + chlorite + pyrite silica schists)

disseminated chalcopyrite-pyrite bodie

FIGURE 7.22 | Cross-section model of the Mount Lyell, vertically extensive, submarine, hybrid magmatic-seawater hydrothermal, alteration and mineralisation system, western Tasmania (modified after Corbett, 2001).

2 0 4 | CHAPTER 7

Weak, selective chlorite + sericite + quartz + carbonate alteration facies

WT 1

Least-altered footwall Sample no.

113084

Alteration facies

weak, selective chlorite + sericite + quartz + carbonate >250 m stratigraphicaliy below mineralised zone

Location Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

massive plagioclase + quartz-phyric rhyolite phenocrysts of albitised plagioclase > quartz

Relict minerals Relict textures

porphyritic

Primary composition rhyolite Lithofacies

massive

Interpretation

rhyolite lava

Alteration minerals

sericite + chlotite + carbonate > quartz

Alteration textures

selective-pervasive, matrix altered to 2040 |jm chlorite, aligned sericite cleavage, albite + sericite-altered plagioclase

Distribution

regional?

Preservation

moderate

Alteration intensity

weak

Timing

synvolcanic and syndeformation

Alteration style

diagenetic and tectonic-metamorphic

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

75.69 0.26 13.00 2.45 0.04 0.78 1.04 1.15

K20 P2O5 S CO2 Total LOI Rb

Photomicrograph (xn)

3.33 0.04 0.37 1.17 99.32 3.12 112

Sr Ba Cu Pb Zn Sb Tl Zr

37 243 3 14 28 0.7 0.9 291

Nb Y

13 40

Al CCPI Ti/Zr

65 40 5.4

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 0 5

Weak, pervasive, medial chlorite + sericite ± carbonate alteration facies Sample no.

113086

Alteration facies

weak, pervasive, medial chlorite + sericite ± carbonate -200 m stratigraphically below mineralised zone

Location Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

altered andesite

Relict minerals

nil

Relict textures

nil

y\/T 2

Primary composition andesite Lithofacies

massive

Interpretation

indeterminate

Alteration minerals

chlorite + carbonate > sericite

Alteration textures

pervasive, foliated, carbonate-altered plagioclase, carbonate veins

Distribution Preservation

poor

Geochemistry

Alteration intensity

weak

Timing

synvolcanic and subsequent syndeformation diagenetic and tectonic-metamorphic

SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

Alteration style

Hand specimen photograph

47.78 0.51 14.56 12.59 0.19 4.63 5.53 2.07

K20 P2O5 S CO2 Total LOI

Photomicrograph (xn)

1.47 0.09 0.04 7.84 97.30 10.71

Rb Sr Ba Cu Pb Zn Sb Tl

46 83 367 10 6 199 0.8 0.5

Zr Nb Y Al CCPI Ti/Zr

72 3 17 45 82 42.5

2 0 6 | CHAPTER 7

Strong, pervasive, medial quartz + sericite + pyrite alteration facies Sample no.

113092

Alteration facies

strong, pervasive, medial quartz + sericite + pyrite -100 m stratigraphically below mineralised zone

Location Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

volcaniclastic rhyolite

Relict minerals

nil

Relict textures

relict granular texture, possibly clastic

WT 3

Primary composition rhyolite Lithofacies Interpretation

indeterminate

Alteration minerals

quartz + sericite + siderite

Alteration textures

pervasive, mosaic of sutured 20-600 |j fractured quartz grains with interstitial sericite and patches coarse siderite

Distribution Preservation

Geochemistry poor

Alteration intensity

strong

Timing

synmineralisation

Alteration style

Hand specimen photograph

hydrothermal

SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na,0

70.58 0.23 12.12 8.03 0.11 0.68 0.16 0.21

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

3.63 0.04 0.02 3.11 98.92 4.44

Rb Sr Ba Cu Pb Zn Sb TI

126

Zr

246

19

Nb

13

Y

42

Al

92

758 18 5

CCPI

67

2.5 Ti/Zr

63

5.6

0.9

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 0 7

Intense, proximal quartz + pyrophyllite + pyrite alteration facies Sample no.

113102

Alteration fades

intense, proximal quartz + pyrophyllite + pyrite

Location

proximal to upper mineralised zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic fades

altered rhyolite

Relict minerals

nil

Relict textures

nil

Primary composition

rhyolite

yyi 4

Lithofacies Interpretation

indeterminate

Alteration minerals

quartz + pyrophyllite > sericite + pyrite

Alteration textures ,

pervasive, mosaic of sutured 40-400 prn quartz grains with ragged patches of semialigned pyrophyllite, minor sericite and disseminated, fractured pyrite

Geochemistry

Distribution Preservation

poor

Alteration intensity

intense

Timing Alteration style

Hand specimen photograph

hydrothermal

SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O K2O

80.78 0.24 12.25 2.21 0.01 0.12 0.03 0.36 1.43

P2O5 S CO2 Total LOI Rb Sr Ba

Photomicrograph (xn)

0.05 Cu 1.61 Pb 0.40 Zn 99.49 Sb 3.08 Tl Zr 31 Nb 117 Y 1170

72 7 6 0.6 <0.5 249 13 5

Al CCPI Ti/Zr

80 54 5.8

2 0 8 | CHAPTER 7

Intense, proximal quartz + pyrophyllite + pyrite alteration facies Sample no.

113105

Alteration facies

intense, proximal quartz + pyrophyllite + pyrite

Location

-200 m straigraphicaily above mineralised zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

altered rhyolite

WT 5

Relict minerals Relict textures Primary composition

rhyolite

Lithofacies Interpretation

indeterminate

Alteration minerals

quartz + topaz + pyrite > carbonate

Alteration textures

pervasive, domainal 50-100 pm microcrystalline quartz and granular topaz, minor disseminated pyrite, irregular mmscale patches > carbonate veinlets

Distribution Preservation

poor

Alteration intensity

intense

Timing Alteration style

Hand specimen photograph

hydrothermal

Geochemistry SiO2 68.25 P2O5 TiO2 0.28 S AI2O3 18.13 CO2 Fe2O3 1.35 Total MnO 0.30 LOl MgO 1.12 CaO 2.05 Rb Na2O 0.05 Sr K2O 0.06 Ba

Photomicrograph (xn)

0.07 0.37 3.11 95.14 8.19

Cu Pb Zn Sb Tl Zr 1 Nb 33 Y 171

10 21 14 0.1 <0.5 364 14 10

Al CCPI Ti/Zr

36 96 4.6

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 0 9

Strong, pervasive, medial quartz + sericite + pyrite alteration facies Sample no.

113110

Alteration facies

strong, pervasive, medial quartz + sericite + pyrite -100 m stratigraphically above mineralised zone

Location Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

altered dacite

yyy g

Relict minerals Relict textures Primary composition

dacite

Lithofacies Interpretation

indeterminate

Alteration minerals

quartz + sericite + pyrite > chlorite and carbonate pervasive, 50-100 |jm microcrystalline quartz with interstitial shreds and seams of aligned sericite, disseminated euhedral pyrite, some highly deformed and recrystallised quartz + carbonate > chlorite veins/patches

Alteration textures

Distribution Preservation

poor

Alteration intensity

strong

Timing Alteration style

Hand specimen photograph

hydrothermal

Geochemistry SiO2 TiO2 AI,0 '2^3 Fe2O3 MnO MgO CaO Na2O

60.29 0.35 12.84 10.77 0.52 1.08 1.26 0.13

K2O S

co 2 Total LOI

Photomicrograph (xn)

3.97 0.08 6.28 2.64 100.21 8.09

Rb Sr Ba Cu Pb Zn Sb

120 25 1737 88 214 138 1.0

Tl Zr Nb Y

1.0 182 9 23

Al CCPI Ti/Zr

78 72 11.5

2 1 0 I CHAPTER 7

Strong, pervasive, medial quartz + sericite + pyrite alteration facies Sample no.

113284

Alteration facies

strong, pervasive, medial quartz + sericite + pyrite

Location

deep mineralised zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

altered rhyolite

Relict minerals

nil

Relict textures

nil

WT 7

Primary composition rhyolite Lithofacies Interpretation

indeterminate

Alteration minerals

quartz + sericite + pyrite + chalcopyrite

Alteration textures

pervasive: 50-100 pm microcrystalline quartz, interstitial shreds and seams of sericite, disseminated euhedral pyrite, > chalcopyrite

Distribution

Geochemistry

Preservation

poor

Alteration intensity

strong

Timing Alteration style

Hand specimen photograph

hydrothermal

SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

72.75 0.22 10.11 6.11 0.01 0.36 0.33 0.05

K2O

S CO2 Total LOI

Photomicrograph (xn)

2.80 Rb 0.35 Sr 4.17 Ba 0.02 Cu 97.28 Pb 4.22 Zn Sb Tl

64 111 1330 13800 21 32 99.2 <0.5

Zr Nb Y

200 11 12

Al CCPI Ti/Zr

67 6.6

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 1 1

Intense, pervasive, proximal quartz + chlorite + pyrite ± sericite alteration facies Sample no.

113264

Alteration facies

intense, pervasive, proximal quartz +

WT 8

chlorite + pyrite ± sericite Location

deep mineralised zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

altered dacite

Relict minerals Relict textures Primary composition dacite Lithofacies Interpretation

indeterminate

Alteration minerals

quartz + sericite + chlorite + pyrite chalcopyrite pervasive: mosaic of 40-200 pm dusty quartz, interstitial and anastomosing sericite cleavage, irregular domains of chlorite + sulfides

Alteration textures

Distribution Preservation

poor

Alteration intensity

intense

Hand specimen photograph

58.74 0.49 AI2O3 13.65 Fe2O3 12.66 MnO 0.23 MgO 1.75 CaO 0.26 Na2O 0.03 SiO 2

Timing Alteration style

Geochemistry

hydrothermal

TiO 2

3.45 P2O5 0.27 5.38 S co 2 0.17 Total 97.08 LOI 6.73 K2O

Photomicrograph (xn)

Rb Sr Ba Cu Pb Zn Sb Tl

92 36 2097 12000 15 99 0.9 <0.5

Zr Nb Y

167 13 25

Al 95 CCPI 79 Ti/Zr 17.6

212 | CHAPTER 7

7.8 | HENTY: A VOLCANOGENIC GOLD DEPOSIT The Henty volcanogenic gold mine and nearby Mount Julia gold prospect are hosted by Cambrian Mount Read Volcanics, near the junctions of the North and South Henty faults and the Great Lyell fault (Fig. 1.5). The deposit comprises at least six steeply dipping, thin, stratabound, disconnected siliceous lenses of up to a few hundred metres vertical extent, distributed over about 2.5 km of strike length (Callaghan, 2001). The estimated total geological resource in December 2001 was 2,154,000 tonnes at 12.1 g/t Au (838,800 oz.). This included production of 820,000 tonnes at 17.5 g/t Au (462,000 oz.).

Geological setting The ore lenses lie in a laterally extensive but narrow stratabound altered zone (A-zone of Callaghan, 2001) at the stratigraphic boundary between the Central Volcanic Complex and the base of the Tyndall Group. In the Henty area these units trend NNW to NNE and face east, with steep easterly to slightly overturned steep westerly dips. The Henty fault zone, trending about 015° and dipping at 70° to the west, obliquely truncates the volcanic succession. The stratabound altered and mineralised zones occur in the immediate footwall of the fault zone, extending about 200 m down-dip from the fault (Halley and Roberts, 1997; Callaghan, 2001). The intersection of the fault and the favourable stratigraphic horizon plunges at a low angle to the south. This is a consequence of the gradual change in trend and slight overturning of the host sequence, from NNW with steep easterly dip in the south, to NNE and steep westerly dip in the northern part of the mine area (Halley and Roberts, 1997). As in many Au deposits, grade cut-offs rather than lithological differences define the ore zones (Callaghan, 1998). Most of the high-grade ore exists in thin lenses or sheet-like bodies up to 7 m thick in the intense, massive quartz (MQ) alteration facies (Halley and Roberts, 1997) but this facies is not uniformly auriferous (Callaghan, 2001). The stratigraphic upper part of the mineralised A-zone typically has a high disseminated base-metal-sulfide content or is spatially associated with lenses of massive pyrite or massive to banded sphalerite + galena (Penney, 1998). Discontinuous massive pyrite lenses up to 2 m thick exist at this stratigraphic level for 600 m of strike but extend less than 150 m down-dip from the Henty fault. Feldspar-phyric to aphyric dacitic lavas and rare basaltic lavas intercalated with dacitic to basaltic hyaloclastite and polymictic volcanic breccias of the Central Volcanic Complex dominate the stratigraphic footwall between the Henty fault and A-zone (e.g. data sheet HN1: Callaghan, 1998). The footwall succession includes discontinuous calcareous volcaniclastic units and hematitic fossiliferous limestone. Immobile element ratios indicate that the protoliths of the altered and mineralised zone were compositionally uniform dacitic volcanic units. The stratigraphic hanging wall, immediately east of the A-zone, comprises massive, andesitic, feldspar crystal-rich volcanic sandstone (e.g. data sheet HN2), dacitic volcanic breccia, lavas and polymictic volcano-sedimentary breccia,

intercalated with calcareous volcanic sandstone, hematitic fossiliferous limestone and minor mudstone. This lithologically diverse part of the hanging wall succession, up to 200 m thick, is recognised as the Lynchford Member: the lowermost unit of the Tyndall Group (Callaghan, 2001). It is succeeded eastwards by the Mount Julia Member consisting of graded rhyolitic breccia, quartz + feldspar-rich volcanic sandstone and minor siltstone intruded by southward thickening quartz + feldspar porphyritic rhyolite sills, cryptodomes and associated hyaloclastites. Overlying this is a thick succession of quartzrich epiclastic sandstone and volcanolithic conglomerate (Zig Zag Hill Formation), which passes conformably eastwards into siliciclastic and micaceous sandstone and conglomerate. White and McPhie (1996) interpreted the massive crystal-rich volcanic sandstone of the Lynchford and Mount Julia Members as originating from large subaerial or shallow marine explosive eruptions that produced pyroclastic flows, which transgressed into a shallow marine environment. The fossil assemblage in the limestone units (Jago et al., 1972) and local welded ignimbrite units in the Mount Julia Member (White and McPhie, 1996) also indicate that the Henty host rocks, or at least those immediately overlying the mineralised zone, were deposited in a near-shore, shallow-marine setting.

Alteration facies and zonation The distribution of alteration facies in this elongate and stratabound alteration system reflects decreasing alteration intensity down-dip away from its intersection with the Henty fault zone, and differing thermo-chemical conditions from footwall to hanging wall. Table 7.2 summarises the features of the Henty—Mount Julia alteration facies. A moderate, footwall sericite + quartz ± carbonate alteration facies (MA) occupies the wedge of stratigraphic footwall between the Au-bearing A-zone and the Henty fault. The A-zone has a discontinuous sheet-like inner zone of intense, massive quartz alteration facies (MQ, e.g. data sheet 7), which is composed of microcrystalline quartz with multiple generations of fine veinlets of quartz + calcite ± sulfides (pyrite, chalcopyrite and galena, which contain most of the Au). It grades outwards (stratigraphically up and down as well as down-dip away from the fault) through intense, proximal, domainal quartz + sericite alteration facies (MV, e.g. data sheet HN6) to intense, foliated sericite + quartz + chlorite + pyrite alteration facies (MZ, e.g. data sheet HN5). These enveloping alteration facies have progressively less quartz and greater phyllosilicate contents and exist in variable proportions in different parts of the deposit. They typically have sericitic groundmasses with foliated to schistose fabrics anastomosing around small siliceous domains that possibly represent silicified lithic clasts (Callaghan, 2001). Massive pyrite lenses at the stratigraphic top of the A-zone grade laterally and down-dip to massive carbonate alteration facies (CB, e.g. data sheet HN4) in peripheral areas. This facies includes massive rocks composed of carbonate + chlorite, and thin bands and lenses of carbonate that are difficult to distinguish from fossiliferous limestone. Some of the carbonate bands contain fragments of red jasper. A zone of strong, selective, hanging wall albite + quartz alteration facies (e.g. data sheet HN3) lies either immediately

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 1 3

base-metal sulfides disseminated in the upper altered footwall zone were due to the shallow-marine, near-shore setting where input of meteoric water produced low-salinity hydrothermal fluids, which boiled and cooled at some depth below the seafloor. Although the Au in the intense, massive quartz alteration facies exists largely in late-stage veinlets related to Devonian brittle deformation and remobilisation, there is no evidence for addition of metals during this event. Callaghan (2001) proposed synvolcanic inputs of magmatic volatiles, fluids and metals to account for the Au + Cu + Bi + Ag + Te metal association and Al mobility in the intense, massive quartz alteration facies, which are atypical of seawater dominated VHMS systems. He envisaged a low pH, high salinity, submarine, subseafloor type of pulsed magmatic plus seawater, high-sulfidation epithermal system. This model invokes the proto-Henty fault as a magmatic volatile and fluid conduit that reactivated during Devonian deformation to dislocate the eastern and western halves of the hydrothermal system.

above (Callaghan, 2001) or 20-40 m stratigraphically above the mineralised zone (Halley and Roberts, 1997). In proximal areas, the hanging wall albite + quartz alteration fades is up to 100 m thick, possibly extending up to the base of the Zig Zag Hill Formation (Callaghan, 1998). It is considered to be of hydrothermal origin, distinct from regional-scale diagenetic albite alteration facies.

Ore genesis Halley and Roberts (1997) interpreted Henty as a Au-rich volcanogenic massive sulfide deposit because of its association with conformable pyrite and carbonate lenses, colloform textures in pyrite, the presence of red jasper clasts that resemble siliceous exhalites, and C-, O- and Pb-isotopic data that indicate a Cambrian synvolcanic origin for the stratabound alteration system. They suggested that its unusual high Au/Ag ratios, extent of footwall silicification and high proportion of

Table 7.2 | The Henty-Mount Julia alteration facies and their defining characteristics (Callaghan, 1998). Alteration facies

Code

Mineral assemblage

Sulfides

Gold

(%)

(g/t)

Distribution

Intense, massive quartz

MQ

Quartz ± (carbonate, sericite, pyrite, chalcopyrite, galena, gold)

~2

Variable; average 36

Thin lenses in core of A-zone.

Intense, proximal quartz + sericite

MV

Quartz + sericite ± (carbonate, pyrite, chalcopyrite, galena, sphalerite)

0.1 to 5

0.1 to 1

Enclosing and gradational to intense massive quartz alteration facies.

Intense, peripheral sericite + quartz + chlorite + pyrite

MZ

Sericite + quartz + pyrite + chlorite ± (carbonate, chalcopyrite, galena)

2-10

0.5 to 2

Peripheral, enveloping the intense proximal quartz + sericite alteration facies.

Massive carbonate

CB

Calcite ± chlorite

<10

?

Discontinuous stratiform lenses at stratigraphic top of A-zone in peripheral parts of the system; laterally equivalent to massive pyrite lenses.

Moderate, footwall sericite + quartz ± carbonate

MA

Sericite + quartz ± (carbonate, pyrite)

<2

?

Stratigraphic footwall, in felsic volcanic rocks between Henty fault and A-zone.

Strong, hanging wall albite + quartz

AS

Albite + quartz ± (chlorite)

0

0

Directly adjacent to A-zone and extending up to 100 m into hanging wall succession.

2 1 4 | CHAPTER 7

Moderate albite + chiorite + calcite alteration facies

HN1

Least-altered host rock Sample no.

255005

Alteration Facies

moderate albite + chlorite + calcite

Location

down dipofA-zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

massive plagioclase-phyric dacite

Relict minerals

albitised plagioclase phenocrysts

Relict textures

porphyritic

Primary composition dacite Lithofacies

massive to brecciated

Interpretation

dacite lava

Alteration minerals

albite + chlorite + calcite + quartz

Alteration textures

selective-pervasive in irregular chlorite and calcite veinlets and blebs, albite ± sericitealtered plagioclase

Distribution

regional

Preservation

moderate

Alteration intensity

moderate

Timing

synvoicanic plus subsequent fault-related deformation

Alteration style

diagenetic and tectonic deformation

Hand specimen photograph

Geochemistry K2O

TiO 2

0.46

AI2O3

12.50

Fe 2 O 3

4.42

co 2

MnO

0.13

Total

MgO

1.51

LOI

0.97 0.17 0.01 6.82 99.15 7.91

CaO

8.71

Na2O

5.05

Au

0.005

SiO 2

58.40

S

Photomicrograph (xn)

Sr

114

Nb

Ba

327

Y

28

Cu

7

Al

15 48

Pb

8

Zn

139

CCPI

Sb

1.2

Ti/Zr

Tl

0.5

Zr

188

40.9

14.7

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 1 5

Weak, selective aibite + chlorite + caicite alteration facies

HN2

Least-aitered hanging wall Sample no.

255024

Alteration facies

weak, selective aibite + chlorite + caicite

Location

lower hanging wall unit + down dip of Azone

Formation

Lynchford Member (Tyndall Group)

Succession

Mount Read Volcanics

Volcanic facies

massive, feldspar crystal-rich volcanicastic sandstone

Relict minerals

albitised plagioclase crystals

Relict textures

abundant sand-sized crystals and sparse

Primary composition

andesite

Lithofacies

massive to crudely banded, moderately

Interpretation

volcaniclastic mass-flow deposit

Alteration minerals

aibite + chlorite + caicite > sericite

Alteration textures

selective-pervasive aibite + chlorite-altered

lithic clasts, subangular

well sorted, matrix supported

matrix, irregular discontinuous caicite veinlets, domainal microcrystalline quartz + chlorite + pyrite Distribution

regional

Preservation

moderate

Alteration intensity

weak

Timing

synvolcanic

Alteration style

diagenetic and tectonic deformation

Hand specimen photograph

Geochemistry SiO2 54.90 TiO2 0.64 AI2O3 14.40 Fe2O3 3.73 MnO 0.19 MgO 1.81 CaO 9.05 Na2O 6.74

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

0.29 Au 0.09 Sr 0.08 Ba 6.97 Cu 98.89 Pb 8.72 Zn Sb Tl

0.005 Zr 134 Nb 127 Y 7 5 Al 132 CCPI 0.8 Ti/Zr 0.5

130 29.6 22 12 42 29.5

2 1 6 | CHAPTER 7

Strong, selective, hanging-wall albite + quartz alteration facies

HN3

Hanging waff Sample no.

255038

Alteration facies

strong, selective, hanging wall albite + quartz (AS)

Location

hanging wall, + 70 m stratigraphically above A-zone

Formation

Mount Julia Member (Tyndall Group)

Succession

Mount Read Volcanics

Volcanic facies

massive volcanicastic sandstone

Relict minerals

quartz and albitised plagioclase crystals

Relict textures

sparse crystals and few lithic clasts in 10-20 urn matrix

Primary composition rhyolite Lithofacies

massive to crudely banded

Interpretation

volcaniclastic mass-flow deposit

Alteration minerals

albite + quartz > sericite > chlorite + calcite

Alteration textures

selective-pervasive altered matrix, microcrystalline albite + quartz matrix with minor interstitial chlorite, quartz + calcite infill irregular veinlets, sericite in later parallel veinlets

Distribution Preservation Alteration intensity Timing Alteration style

local and stratabound in hanging wall sequence

Geochemistry 0.29

Sr

47

Nb

TiO 2

77.70 K2O 0.16

0.01

Ba

100

Y

AI 2 O 3

12.10

0.00

Cu

4

SiO 2

S

moderate

Fe 2 O 3

1.13

co 2

strong

MnO

0.02

Total

albite + quartz probably syn volcanic, sericite veinlets syn deformation

MgO

0.36

LOI

CaO

0.75

hydrothermal?

Na2O

6.57

Hand specimen photograph

Au

Photomicrograph (xn)

14.6 26

0.99

Pb

2

100.09

Zn

100

CCPI

17

0.90

Sb

0.9

Ti/Zr

5.3

Tl

0.5

Zr

182

0.005

Al

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 217

Massive carbonate alteration facies

HN4

Host-rock equivalent? Sample no.

255050

Alteration facies

massive carbonate

Location

upper A-zone, laterally equivalent to massive pyrite

Formation

Lynchford Member (Tyndall Group) or Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

marine limestone with minor volcaniclastic component

Relict minerals

calcite + plagioclase + quartz

Relict textures

plagioclase and quartz crystal fragments

Primary composition Lithofacies

massive to thinly bedded

Interpretation

impure marine carbonate

Alteration minerals

calcite?

Alteration textures

10-50 |jm microcrystalline calcite, calcite veins, stylolites

Distribution

local and stratabound in peripheral upper part of A-zone

Preservation

moderate

Alteration intensity

weak

Timing

synvolcanic diagenesis, syndeformational dynamic recrystallisation

Alteration style

diagenetic and tectonic-metamorphic, doubtful hydrothermal carbonate component

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

22.80 0.18 4.16 2.33 0.26 0.85 36.40 1.38

K2O P2O5 S CO2 Total LOI Au

Photomicrograph (xn)

0.58 Sr 0.08 Ba 0.72 Cu 24.70 Pb 94.44 Zn 29.11 Sb Tl 0.017 Zr

263 476 21 76 100 10.0 0.5 49

Nb Y A! CCPI Ti/Zr

1.0 16 4 60 22.0

2 1 8 | CHAPTER 7

Intense, foliated sericite + quartz + chlorite + pyrite alteration facies

HNS

Footwali? Sample no.

255030

Alteration facies

intense, foliated sericite + quartz + chlorite + pyrite (MZ)

Location

peripheral altered zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

volcaniclastic breccia

Relict minerals

minor quartz

Relict textures

blocky clasts

Primary composition dacite Lithofacies

indeterminate

Interpretation

indeterminate

Alteration minerals

sericite + quartz + pyrite + chlorite

Alteration textures

foliated, semi-mylonitic, disseminated pyrite, non-foliated domains of quartz + calcite > sericite

Distribution

local, enclosing mineralised lens

Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal and tectonic-metamorphic

Hand specimen photograph

Geochemistry SiO2 65.30 TiO2 0.53 AI2O3 15.00 Fe2O3 6.18 MnO 0.03 MgO 1.28 CaO 0.97 Na2O 0.37

K2O P2O5 S CO2 Total LOI Au

Photomicrograph (xn)

5.35 Sr 0.12 Ba 3.31 Cu 0.88 Pb 99.32 Zn 4.85 Sb Tl 0.424 Zr

15 475 22 30 100 3.1 2.0 172

Nb Y

29.4 25

Al CCPI Ti/Zr

83 54 18.5

I

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 1 9

Intense, proximal, domainal quartz + sericite alteration facies

Hi 6

Footwali? Sample no.

255053

Alteration facies

intense, proximal, domainal quartz + sericite (MV) proximal altered zone enclosing and transitional to mineralised MQ

Location Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

indeterminate

Relict minerals Relict textures Primary composition dacite Lithofacies

indeterminate

Interpretation

indeterminate

Alteration minerals

quartz + sericite + pyrite

Alteration textures

10-40 |jm microcrystalline sutured mosaic of quartz, selective domainal sericite with cleavage, disseminated pyrite

Distribution

local, enclosing mineralised lens

Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry 88.10 K2O SiO 2 TiO2 0.44 P A 6.29 S MA Fe2O3 MnO MgO CaO Na2O

1.21 co 2 0.01 Total 0.31 LOI 0.31 0.08 Au

Photomicrograph (xn)

1.94 0.03 0.54 0.37 99.63 1.44 0.318

Sr Ba

9 Nb 100 Y

Cu

913 24 Al 100 CCPI 1.6 Ti/Zr 0.9 178

Pb Zn Sb Tl Zr

12.2 14 85 41 14.8

2 2 0 | CHAPTER 7

intense, massive quartz alteration facies Footwall? Sample no.

255044

Alteration facies

intense, massive quartz (MQ)

Location

central siliceous core of A-zone

Formation

Central Volcanic Complex

Succession

Mount Read Volcanics

Volcanic facies

indeterminate

Relict minerals

sparse polycrystalline quartz crystals?, pseudomorphs of feldspar?

Relict textures

porphyritic?

Primary composition dacite Lithofacies

indeterminate

Interpretation

indeterminate

Alteration minerals

quartz + calcite + pyrite + chalcopyrite

Alteration textures

2CM0 (jm microcrystalline in coarse and fine equigranuiar domains of quartz, disseminated calcite blebs, possible calcite pseudomorphs after feldspar crystals, pyrite + chlorite ± calcite veinlets

Distribution

local, ore lens

SiO 2

Preservation

poor

TiO 2

Alteration intensity

intense

AI2O3

Timing

synmineralisation

Fe 2 O 3

Alteration style

hydrothermal

Geochemistry

MnO MgO CaO Na2O

Hand specimen photograph

87.70 0.39 0.80 1.17 0.07 0.12 5.29 0.05

K2O P2O5 S

co 2 Total LOI

Photomicrograph (xn)

0.24 0.04 0.22 4.25 100.35 4.33

Au Sr Ba Cu Pb Zn Sb Tl

0.048 23 100 1690 47 100 1.6 0.5

Zr Nb Y

163 2.3 12

Al CCPI Ti/Zr

6 80 14.3

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 221

7.9 | THALANGA: A POLYMETALLIC SHEET-STYLE DEPOSIT The Thalanga deposit is located near the western end of the Early Ordovician Mount Windsor Subprovince (Fig. 1.7). It was the most economically significant deposit in the Mount Windsor Subprovince. The sulfide lenses were up to 25 m thick and distributed over about 3000 m strike and 400 m vertical extent. The pre-mining resource estimate was 6.6 Mt grading 1.8% Cu, 2.6% Pb, 8.4% Zn, 69 g/t Ag and 0.4 g/t Au.

Geological setting The Thalanga deposit consisted of several semi-connected, thin, stratabound and stratiform massive sulfide lenses hosted in a distinctive quartz crystal-rich volcanic unit, which is sandwiched between the underlying rhyolitic Mount Windsor Formation and the overlying mixed andesitic-dacitic Trooper Creek Formation (Fig. 7.23). The host unit (known as the Thalanga horizon or favourable unit) is composed of quartz + feldspar crystal-rich volcanic breccia, sandstone and siltstone, and co-magmatic, peperitic quartz + feldspar intrusions (Paulick and McPhie, 1999).

The ores are massive to semi-massive lenses dominated by pyrite and sphalerite with variable proportions of galena, chalcopyrite, pyrrhotite, magnetite and barite (Gregory et al., 1990). Barite-rich zones exist in the up-dip and western peripheries of the West and East Thalanga lenses. Chlorite + tremolite + carbonate rocks, interpreted as metamorphosed chlorite + carbonate alteration assemblages (Herrmann and Hill, 2001), are closely associated with the West Thalanga ore lenses. Magnetite-bearing quartzite bodies in the peripheral or upper parts of some of the sulfide lenses, and also intercalated with volcanic siltstone of the host unit to the west, are interpreted to be metamorphosed exhalative siliceous ironstones (Duhig et al., 1992). The stratigraphic footwall is a laterally extensive, 1200 m thick, submarine rhyolitic succession. It is dominated by sparsely quartz + feldspar-phyric coherent rhyolitic lavas (e.g. data sheet TH1) and domes that may have formed a low volcanic rise in the Thalanga area (Paulick and McPhie, 1999). Rhyolitic hyaloclastite breccias and volcanic sandstones are locally significant, particularly in the upper part of the footwall beneath the western sulfide lenses. The hanging wall succession is composed mainly of unaltered to weakly altered coherent lavas and sills of feldspar-phyric to aphyric dacite (e.g. data sheet TH2), and minor basalticandesite. At Thalanga, it includes minor volcaniclastic rocks

FIGURE 7.23 | Schematic facies architecture of the submarine volcanic succession, from the Mount Windsor Volcanics, through the Trooper Creek Formation, to the Rollston Range Formation, northwest of the Thalanga mine, Queensland. Modified after Hill (1996).

2 2 2 | CHAPTER 7

of mixed dacitic-rhyolitic derivation, including lithic massflow breccia, sandstone and massive to laminated cherty siltstone, which increase in proportion westward. Paulick and McPhie (1999) interpreted the volcanic facies assemblage to indicate that the deposit formed in a below-storm-wavebase environment on an elevated, lava-dominated, rhyolitic centre. The compositions of the footwall and hanging wall successions, respectively, indicate that they were rhyolitic magmas derived from crustal melting, and mixed-mafic-felsic magmas from subduction-modified mantle, in an extensional back-arc-basin setting (Stolz, 1995). Regional deformation and metamorphism, related to Mid-Late Ordovician granitoid intrusions produced upper greenschist facies metamorphic mineral assemblages and a near-vertical foliation, particularly in phyllosilicate-rich hydrothermally altered volcanic rocks.

Alteration facies and zonation Underlying the Thalanga deposit is an extensive zone of strong, pervasive quartz + sericite + pyrite ± chlorite alteration facies (e.g. data sheet TH3) characterised by 1—4% disseminated pyrite and an absence of primary feldspars. This alteration style was pervasive in both clastic and coherent rhyolites, and typically produced pseudoclastic breccia and mottled, domainal alteration textures (Paulick and McPhie, 1999). The zone extends beneath the entire strike length of the deposit and is at least 200 m thick in the Central area, pinching out to less than 50 m near the lateral and down-dip extremities (Herrmann and Hill, 2001). It has a broad, upward flaring shape and gradational boundaries with the surrounding leastaltered rhyolites. Within the broad zone of feldspar destruction there are semi-stratiform stringer zones of intense, pervasive quartz + pyrite alteration facies (e.g. data sheet TH4) up to 50 m thick. They extend obliquely up through the footwall at about 15° to the host unit and intersect it beneath the East, Central and "eastern edge of the West Thalanga ore lenses, suggesting that they were paths of maximum hydrothermal fluid flow. These zones are composed essentially of quartz and up to 20% pyrite in disseminated grains and anastomosing veins. They typically contain less than 20% phyllosilicates (sericite ± chlorite). Intense, macrocrystalline quartz + K-feldspar alteration facies (e.g. data sheet TH5) exist in the immediate footwall, lateral to the sulfide lens and stringer zone at East Thalanga, and also stratigraphically lower in the footwall succession at Central and West Thalanga. Stratabound chlorite + dolomite altered zones, subsequently metamorphosed to chlorite + tremolite ± carbonate assemblages, formed in permeable volcaniclastic footwall rocks close to the palaeo-seafloor and lateral to the West Thalanga ore lenses (stratabound alteration facies; data sheets TH6, 7 and 8). Local zones of non-pyritic, foliated altered rhyolite with relict plagioclase (moderate, foliated sericite + chlorite alteration facies; data sheet TH9), exist within the least-altered rhyolite, mainly around the peripheries of the feldspar-destructive, strong quartz + sericite + pyrite + chlorite alteration facies, and may represent low-temperature hydrothermal recharge zones.

Ore genesis There is consensus amongst researchers that Thalanga is a sheet-like, synvolcanic, deformed and metamorphosed VHMS deposit formed in a deep-marine back-arc rift. Isotopic data suggests that the hydrothermal fluid and sulfur were dominantly of seawater origin (Hill, 1996; Herrmann and Hill, 2001). The hydrothermal system pervasively altered a very broad zone in the mainly coherent rhyolitic footwall succession to quartz + sericite + pyrite + chlorite assemblages. The ore-forming fluids were focussed in low-angle quartz + pyrite stringer zones. These pyritic stringer zones have a semistratiform distribution, which suggests control by volcanic facies related permeability contrasts. However, some appear to cut through coherent rhyolite units and thus may represent deformed synvolcanic fault zones (Paulick and McPhie, 1999). A major proportion of the massive sulfide ore was deposited in thin, extensive, stratiform and stratabound lenses, either directly on the palaeo-seafloor or a few metres below it. The distribution of massive pyrite and Cu-rich zones suggests that the down-dip eastern parts of West and Central Thalanga ore bodies, and central part of East Thalanga, were sites of high-temperature hydrothermal discharge (Hill, 1996; Paulick et al., 2001). Subordinate stratabound semi-massive ore lenses were formed by subsurface replacement and/or infilling of coarse volcaniclastic units of the host unit, which were deposited by syneruptive, synhydrothermal mass flows on top of the accumulating seafloor massive sulfide lenses. Chlorite + carbonate (pre-metamorphic) alteration mineral assemblages intimately associated with the West Thalanga sulfide lenses, probably formed by mixing of hydrothermal fluid and cold seawater in permeable volcaniclastic units immediately below the palaeo-seafloor, in proximal to medial parts of the hydrothermal discharge system. Apart from disseminated and vein-type pyrite in the altered footwall zones, all the sulfides were deposited at the top of the rhyolitic Mount Windsor Formation, or in the quartz crystal-rich unit that immediately overlies it. The massive, coherent dacite lavas and sills of the hanging wall succession are essentially unaltered and unmineralised; their emplacement appears to have ended local hydrothermal circulation.

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 2 3

TH1

Weak, patchy quartz + sericite alteration facies Least-altered footwal! Sample no.

140802

Alteration facies

weak, patchy quartz + sericite

Location

Thalanga footwall

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz and albitised plagioclase phenocrysts

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

rhyolite lava

Alteration minerals

quartz + sericite + biotite + chlorite

Alteration textures

selective-pervasive (patchy) microcrystalline quartz matrix, weakly aligned sericite ± biotite in cleavage

Distribution

regional, broadly stratabound, footwall succession

Preservation

moderate

Geochemistry

Alteration intensity

weak

SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na,0

Timing

synvolcanic

Alteration style

diagenetic

Hand specimen photograph

76.40 0.11 11.90 1.64 0.04 0.67 1.42 2.27

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

4.04 0.02 0.01 <0.1 98.52 0.60

Rb Sr Ba Cu Pb Zn Sb

110 Tl 82 Zr 1056 Nb 5 Y 20 48 Al 0.2 CCPI Ti/Zr

0.5 146 13 40 56 25 4.5

2 2 4 | CHAPTER 7

Subtle, selective-pervasive quartz 4- albite alteration fa.cies

TH2

Least-altered hanging wall Sample no.

140799

Alteration fades

subtle, selective-pervasive quartz + albite

Location

Thalanga hanging wall

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive sparsely plagioclase-phyric

dacite Relict minerals

plagioclase phenocrysts

Relic textures

weakly porphyritic, faintly flow banded

Primary composition

dacite

Lithofacies

massive to flow banded

Interpretation

dacite lava or sill

Alteration minerals

quartz + albite ± (chlorite + actinolite ± epidote)

Alteration textures

selective-pervasive; mosaic of 20 pm quartz + albite, dissemiated chlorite and acicular prisms actinolite defining weak relict flow banding, quartz ± calcite veins

Distribution

regional, broadly stratabound, hanging wall succession

Preservation

good

Alteration intensity

subtle

Timing

synvolcanic

Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO2 74.00 TiO2 0.33 AI2O3 12.70 Fe2O3 1.47 MnO 0.03 MgO 0.52 CaO 1.47 Na2O 4.10

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

3.82 0.06 <0.01 98.50 0.47

Rb Sr Ba Cu Pb Zn Sb

64 76 1285 3 11 29 0.2

Tl Zr Nb Y

0.5 164 9 26

Al CCPI Ti/Zr

44 19 12.1

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 2 5

Strong, pervasive quartz + sericite + pyrite ± chlorite alteration fades

TO 3

Footwall Sample no.

140808

Alteration facies

strong, pervasive quartz + sericite + pyrite ± chlorite

Location

Thalanga footwall

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

rhyolite lava

Alteration minerals

quartz + sericite + chlorite + pyrite

Alteration textures

pervasive, microcrystalline quartz matrix with strongly aligned sericite (cleavage), scattered 1-2 cm elliptical chlorite-rich domains local, broadly stratabound in footwall beneath entire Thalanga system, >200 m thick, thinning laterally

Distribution

Preservation

poor

Alteration intensity

strong

Timing

synmineralisation

Alteration style

footwall hydrothermal

Hand specimen photograph

Geochemistry SiO2 75.70 K2O TiO2 0.07 P2O5 AI2O3 11.40 S Fe2O3 5.14 CO2 MnO 0.08 Total MgO 2.38 LOI CaO <0.01 Na2O 0.21

Photomicrograph (xn)

2.39 0.01 0.65 <0.1 98.03 2.75

Rb Sr Ba Cu Pb Zn Sb Tl

64 8 326 18 11 77 0.2 <0.5

Zr Nb Y

128 14 36

Al CCPI Ti/Zr

96 73 3.3

2 2 6 I CHAPTER 7

Intense, pervasive quartz + pyrite alteration facies

TH4

Footwall Sample no.

140724

Alteration facies

intense, pervasive quartz + pyrite

Location

Thalanga footwall

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz phenocrysts

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

rhyolite lava

Alteration minerals

quartz + sericite + pyrite > biotite ± chlorite

Preservation

pervasive, mosaic of 100-200 pm quartz and pyrite with interstitial shreds of semialigned white mica > biotite local, broadly stratabound in footwall, discrete sheet-like zones at -15° to host unit, intersecting it beneath suifide lenses poor

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Alteration textures

Distribution

Hand specimen photograph

Geochemistry SiO2 67.00 TiO2 0.05 AI2O3 6.60 Fe2O3 14.34 MnO 0.02 MgO 1.67 CaO 0.05 Na2O 0.06

K2O P2O5 S CO2 Total LOI Rb

Photomicrograph (xn)

1.99 0.01 10.61 <0.1 102.40 8.24 68

Sr Ba Cu Pb Zn Sb Tl Zr

16 2300 22 16 74 1.1 1.9 79

Nb Y

7 44

Al CCPI Ti/Zr

97 88 3.8

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 2 7

intense, microcrystalline quartz + K-feldspar alteration facies

TH5

Footwall Sample no.

140902

Alteration facies

intense, microcrystalline quartz + Kfeldspar

Location

Thalanga footwall

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz and albitised plagioclase phenocrysts

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

rhyolite lava

Alteration minerals

quartz + K-feldspar > trace pyrite

Alteration textures

selective-pervasive; matrix of 10-50 pm microcrystalline quartz + K-feldspar, albitised plagioclase, quartz veins with feldspar-alteration selvage

Distribution

local, lozenge shaped zones, broadly stratabound in footwall

Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2 84.20 K2O TiO2 0.06 P2O5 AI2O3 7.40 S Fe2O3 0.60 CO2 MnO <0.01 Total MgO 0.08 LOI CaO 0.11 Na2O 0.34 Rb

Photomicrograph (xn)

5.87 0.01 0.48 99.15 0.49 94

Sr Ba Cu Pb Zn Sb Tl Zr

42 1936 3 70 118 0.4 1.5 80

Nb Y

9 20

Al CCPI Ti/Zr

93 9 4.5

2 2 8 | CHAPTER 7

TH8

Intense, pervasive, stratabound chlorite + tremolite alteration facies Footwall Sample no.

145401

Alteration facies

intense, pervasive, stratabound chlorite + tremolite

Location

Thalanga footwall

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

quartz + plagioclase-phyric rhyolitic volcaniclastic breccia?

Relict minerals

nil

Relict textures

nil

Primary composition

rhyolite

Lithofacies

indeterminate

Interpretation Alteration minerals

phlogopite + chlorite + tremolilte > minor pyrite, sphalerite, chalcopyrite

Alteration textures

pervasive, foliated phlogopite ± chlorite, 30% coarse prisms and bands of randomly oriented tremolite

Distribution

local, stratabound, typically immediately below sulfide lenses

Preservation

nil

Alteration intensity

intense

Timing

synmineralisation

Alteration style

footwall hydrothermal

Hand specimen photograph

Geochemistry SiO2 44.57 TiO2 0.135 AI2O3 13.06 Fe2O3 3.88 MnO 0.15 MgO 17.67 CaO 6.76 Na2O 0.18

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

3.48 0.04 1.66 0.30 91.89 3.23

Rb Sr Ba Cu Pb Zn Sb Tl

38300 1500 200 5900

Zr Nb Y

171 10 32

Al CCPI Ti/Zr

75 85 4.7

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 2 2 9

Intense, stratabound. pervasive chlorite + tremolite + calcite alteration facies

TH7

Host rock Sample no.

145417

Alteration facies

intense, stratabound, pervasive chlorite + tremolite + calcite

Location

Thalanga host rock

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

quartz + plagioclase-phyric rhyolitic volcaniclastic breccia?

Relict minerals

nil

Relict textures

nil

Primary composition

rhyolite

Lithofacies

indeterminate ?

Interpretation Alteration minerals

chlorite + tremolite + calcite > pyrite + chalcopyrite + sphalerite + galena

Alteration textures

pervasive; coarse interlocking prisms of tremolite with 10% interstitial sulfides and ragged calcite patches

Distribution

local, stratabound, proximal to medial, closely associated with or lateral to West Thalanga sulfide lenses

Preservation

nil

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2

33.10

K2O

0.58

Rb

Zr

89

TiO2

0.060

P2O5

0.15

Sr

Nb

6

Y

13

AI2O3

6.48

S

4.82

Ba

8700

Fe2O3

4.82

CO2

7.10

Cu

7500

MnO

0.57

Total

92.41

Pb

4200

MgO

19.00

LOI

5.86

Zn

23800

CaO

15.60

Sb

Na2O

0.13

Tl

Photomicrograph (xn)

Al

55

CCPI

97

Ti/Zr

4.0

2 3 0 I CHAPTER 7

Intense, strataboynd tremolite + dolomite + calcite alteration facies Host rock Sample no.

145418

Alteration facies

intense, stratabound tremolite + dolomite + calcite

Location

Thalanga host rock

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies

quartz + plagioclase-phyric rhyolitic volcaniciastic breccia?

Relict minerals Relic textures

nil nil

Primary composition

rhyolite

Lithofacies

indeterminate

Interpretation

?

Alteration minerals

dolomite + calcite > minor tremolite + chlorite + pyrite + chalcopyrite + sphalerite + galena

Alteration textures

pervasive; mosaic of 1 mm sutured spheroidal dolomite and interstitial calcite disseminated sulfides and sparse ragged tremolite prisms

Distribution

local, stratabound, proximal to medial, closely associated with or lateral to West Thalanga sulfide lenses

Preservation

nil

Alteration intensity

intense

Timing

synmineralisation

Alteration style

footwall hydrothermal, seawater mixing?

Hand specimen photograph

Photomicrograph (xn)

TH8

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 2 3 1

TH9

Moderate, foliated sericite + chlorite alteration fades FootwaSI Sample no.

140727

Alteration facies

moderate, foliated sericite + chlorite

Location

Thalanga footwall

Formation

Mount Windsor Formation

Succession

Mount Windsor Subprovince

Volcanic facies Relict minerals

massive quartz + plagioclase-phyric rhyolite quartz and albitised plagioclase

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive, foliated

Interpretation

rhyolite lava

Alteration minerals

sericite + quartz + biotite + chlorite

Alteration textures

selective-pervasive 50-100 pm microcrystalline quartz matrix in <1 mm lenses wrapped by foliated sericite ± biotite (augen texture), broken grains, cleavage, sericite altered plagioclase local, stratabound in upper part of medial to distal footwall, particularly down dip of ore zones

Distribution

Preservation

moderate

Alteration intensity

moderate

Timing

synmineralisation

Alteration style

hydrothermal, tectonic-metamorphic

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO

70.90 0.10 14.50 1.69 0.04 3.94 0.12

Na2O K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

1.05 3.92 0.01 <0.01 0.10 96.37 2.67

Rb Sr Ba Cu Pb Zn Sb Tl

145 40 952 3 5 231 0.2 2.6

Zr Nb Y

162 19 54

Al CCPI Ti/Zr

87 52 3.7

2 3 2 | CHAPTER 7

7.10 | HIGHWAY-REWARD: A PIPE STYLE Cu-Au VHMS DEPOSIT The Highway-Reward Cu-Au deposit, in the central part of the Mount Windsor Subprovince (Fig. 1.8), represents a contrast in style of deposit and stratigraphic setting. It consists of two discordant, vertical pipe-like bodies of massive pyrite about 200 m apart, hosted in the proximal facies of a non-explosive, submarine felsic volcanic centre located near the top of the Trooper Creek Formation (Fig. 7.23).

Geological setting The lithofacies association at Highway-Reward represents a deep marine intrusion-dominated felsic volcanic centre. Doyle and McPhie (2000) recognised at least 13 coherent feldspar- and quartz + feldspar-phyric dacitic to rhyolitic synvolcanic sills, small cryptodomes (e.g. data sheets HR1 and 4) and lavas in the immediate area. The abundance and complex overlapping relationships of coherent intrusive units indicate a proximal volcanic setting. Thin volcanic sandstone and siltstone units and thicker units of crystal- and pumicerich sandstone and breccia separate the intrusions. The crystaland pumice-rich facies were mainly derived from explosive eruptions and deposited in the submarine basin from watersupported gravity flows. The succession is upright and dips at 20-30° to the southeast. Massive pyrite ± chalcopyrite exists in two vertical pipelike bodies about 150 m apart. Both are discordant to bedding, parallel to a locally developed northeast trending sub-vertical cleavage (S4) and have irregular-amoeboid outlines with plan dimensions of about 200 x 75-150 m. The western pipe (Highway) has a vertical extent of 250 m and the eastern pipe (Reward) of at least 350 m (Doyle and Huston, 1999). They are dominantly composed of fine-grained (<0.5 mm) pyrite with interstitial chalcopyrite, minor tennantite, sphalerite, quartz and sericite, and traces of chlorite, galena, barite, hematite and aikinite (PbCuBiS3). The massive sulfide pipes are intersected by chalcopyrite, barite, quartz + carbonate and anhydrite veins, and contain inclusions of quartz + sericite + pyrite altered volcanic rocks in their margins. The Highway and Reward massive sulfide pipes contain approximately 2 Mt and 5 Mt of pyrite, respectively. They include hypogene sulfide resources estimated at 1.2 Mt @ 5.5% Cu, 1.2 g/t Au and 6.5 g/t Ag in the Highway pipe and 0.2 Mt @ 3.5% Cu, 1 g/t Au and 13 g/t Ag in the Reward pipe. The pipes are enveloped by a broad 200 x 500 m halo of vein and disseminated low-grade Zn + Pb + Ba sulfides. Within that are several small zones of massive to laminated sphalerite + pyrite + galena + chalcopyrite + barite. A 20-30 m thick stratabound lens of sphalerite-rich massive sulfide exists in volcaniclastic rocks 50 m above and south of the Reward pipe. It has a pyrite-rich base that thickens northwards into a discordant lens of massive pyrite lying above the southern edge of the Reward pipe. Sphalerite-rich sulfides also exist locally in narrow discordant zones at the margins of the main Highway and Reward massive sulfide pipes.

Alteration facies and zonation A discordant zone of feldspar-destructive hydrothermal alteration envelopes the massive sulfide pipes. It has an elliptical area of 500 x 250 m in plan and extends from 60 m above to at least 150 m below the massive sulfide bodies (Doyle and Huston, 1999). Doyle and Huston's (1999) alteration zonation is here simplified down to six alteration facies. Intense, stringer quartz + sericite + pyrite alteration facies (e.g. data sheet HR8), locally flanked by intense, pervasive chlorite + pyrite alteration facies (e.g. data sheet HR7) occupy feeder zones which extend vertically beneath both pipes and possibly meet at depth. Zones of similar quartz + sericite + pyrite altered rocks extend into the hanging wall above the southern parts of both massive sulfide pipes (Doyle and Huston, 1999). These intensely altered zones pass laterally outwards to enveloping zones of intense sericite + quartz + pyrite (e.g. data sheet HR6) and strong, pervasive chlorite + sericite + quartz + pyrite alteration facies (e.g. data sheet HR5). These locally enclose, and in turn grade laterally and upwards in to, non-pyritic zones of moderate, pervasive chlorite alteration facies (data sheet HR3). The weak, regional, selective albite ± hematite alteration facies (e.g. data sheet HR2) exists at greater than 50—200 m from the massive sulfide pipes. It comprises two sub-facies with mineral assemblages of feldspar + carbonate ± quartz ± chlorite + sericite and hematite + quartz ± sericite ± chlorite ± albite, which are regionally distributed in the Trooper Creek Formation and are respectively attributed to alteration during diagenesis and synvolcanic low-temperature fluid convection.

Ore genesis The deposits were initially thought to have had a two-stage origin (Beams et al., 1998). The stratiform Zn-rich zone was interpreted as a syngenetic Cambro-Ordovician sulfide lens and the pyrite + chalcopyrite pipes as Siluro-Devonian syndeformational deposits, because of their discordance to host volcanic rocks, parallelism to the youngest cleavage (S4) and the observation that anhydrite overprinted the dominant S3 cleavage. However, Doyle and Huston (1999) refuted this microtextural relationship and argued for a syngenetic volcanic-associated, subseafloor replacement origin for the massive sulfide pipes. Lead isotopic ratios, the gradation from stratiform Zn-rich sulfides into discordant Cu-rich massive pyrite, relict framboidal sulfide textures, hydrothermal alteration facies and their relationships to primary volcanic facies, and the S3 tectonic overprint are all consistent with Early Ordovician synvolcanic formation of all the sulfide zones. The Highway-Reward massive sulfide pipes have some similarities with disseminated to massive Cu-Au deposits in the Mount Lyell field (Large et al., 2001c). The similarities include metal ratios, dominance of pyritic subseafloor replacement style mineralisation and low 534S values; mostly in the range 5 to 7.5%o at Highway-Reward and 5 to 10%o at Mount Lyell (Solomon etal., 1969; Doyle and Huston, 1999). Given the emerging evidence for involvement of magmatic fluids at Mount Lyell (Corbett, 2001; Huston and Kamprad, 2001) it is reasonable to similarly classify Highway-Reward as a hybrid seawater-magmatic hydrothermal system.

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS I 2 3 3

Weak, selective-pervasive quartz + sericite +albite alteration facies

HR1

Least-altered rhyolite Sample No.

137068

Alteration Facies

weak, selective-pervasive quartz + sericite + albite

Location

upper medial, 50 m east of Highway pipe (10075N)

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz, plagioclase

Relict textures

porphyritic

Primary composition

rhyolite

Lithofacies

massive

Interpretation

partly extrusive cryptodome

Alteration minerals

quartz + sericite + albite? > (calcite, pyrite)

Alteration textures

selective-pervasive, microcrystalline groundmass, albite + sericite or calcitealtered plagioclase

Distribution

regional

Preservation

moderate

Alteration intensity

weak

Timing Alteration style

diagenetic

Hand specimen photograph

Geochemistry SiO 2

75.08

TiO 2

0.30

AI 2 O 3

12.75

K2O S

2.22

Rb

50

0.06

Sr

46

0.51

Ba

1382

Fe 2 O 3

1.57

co 2

Cu

43

MnO

0.08

Total

97.37

Pb

14

LOI

2.88

Zr Nb

9

Y

22

Al

60

MgO

2.02

Zn

104

CCPI

CaO

0.61

Sb

2.2

Ti/Zr

Na2O

2.16

Tl

1.0

Photomicrograph (xn)

158

44 11.4

2 3 4 | CHAPTER 7

Weak, regional, selective albite + hematite alteration facies Sample no.

137105

Alteration facies

weak, regional, selective albite + hematite

Location

upper periphery, 250 m east of Reward pipe

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz, piagiociase

Relict textures

porphyritic, giomeroporphyritic piagiociase, microcrystalline groundmass

Primary composition

rhyolite

Lithofacies

massive

Interpretation

synvolcanic sill

Alteration minerals

quartz + albite + chlorite > (sericite + calcite + hematite)

Alteration textures

selective-pervasive in groundmass; 20-60 urn microcrystalline quartz + albite ± calcite-altered piagiociase, disseminated chlorite and hematite patches

HR2

Geochemistry 70.58 K2O SiO2 TiO2 0.33

0.34

Rb

8

0.06

Sr

142

<0.01

Ba

92

Cu

<2

Pb

5

Zn

Distribution

regional

Preservation

good

Alteration intensity

weak

Fe 2 O 3

2.41

CO2

Timing

synvolcanic

MnO

0.05

Total

98.07

Alteration style

diagenetic

MgO

2.26

LOI

2.42

CaO

1.78

Sb

Na2O

6.51

Tl

<0.5

Hand specimen photograph

13.75

S

Photomicrograph (xn)

Zr Nb

161 9

Y

21

Al

24

38

CCPI

39

0.8

Ti/Zr

12.3

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 3 5

Moderate, pervasive chlorite alteration facies Sample no.

137079

Alteration facies

moderate, pervasive chlorite

Location

HR3

upper proximal zone, between Highway and Reward sulfide pipes

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyolite

Relict minerals

quartz, mafic phenocrysts?

Relict textures

porphyritic, amygdaloidal?

Primary composition

rhyolite

Lithofacies

massive

Interpretation

partly extrusive cryptodome

Alteration minerals

quartz + sericite + chlorite

Alteration textures

pervasive, microcrystalline mosaic of quartz + chlorite + sericite, sericite pseudomorphs after plagioclase phenocrysts, anastomosing wispy sericite foliation, recrystallised overgrowths on quartz local; medial to proximal zones laterally equivalent to upper parts of sulfide pipes

Distribution Preservation

moderate to poor

Alteration intensity

moderate

Timing

synmineralisation

Alteration style

hydrothermal

Geochemistry SiO 2 TiO 2

0.30

AI 2 O 3

14.09 3.40 0.15 3.59 0.39 0.16

Fe2O3 MnO MgO CaO Na2O

Hand specimen photograph

71.21

K2O S

70

Tl

1.5

Sr

21

Zr

165

Ba

1791

3.27

Rb

0.05 0.01

Nb

9

Y

23

93 66

Cu

6

Total

96.61

Pb

7

LOI

3.61

Zn

120

Al

Sb

0.8

CCPI

CO2

Photomicrograph (xn)

Ti/Zr

10.9

2 3 6 I CHAPTER 7

Weak, pervasive albite + sericite alteration facies

HR4

Least-altered dacife Sample no.

136919

Alteration facies

weak, pervasive albite + sericite

Location

Highway, medial footwall

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive, sparsely plagioclase-phyric massive dacite

Relict minerals

plagioclase <1mm

Relict textures

porphyritic and micropoikilitic

Primary composition dacite Lithofacies

massive to weakly flow banded

Interpretation

cryptodome

Alteration minerals

albite + chlorite + sericite > (zeolite?, quartz)

Alteration textures

pervasive groundmass, microcrystalline partly preserving micropoikilitic texture, albite ± sericite-altered plagioclase, chlorite veinlets

Distribution

regional good

SiO 2 TiO2

66.74

Preservation Alteration intensity

weak

AI2O3

16.63

synvolcanic

Fe2O3

4.30

diagenetic

MnO

0.18

co 2 Total

MgO

2.48

LOI

CaO

0.19

Na2O

4.40

Timing Alteration style

Hand specimen photograph

Geochemistry 0.55

K2O

1.81

Rb

44

Zr

161

P2O5

0.10

Sr

68

Nb

S

0.00

Ba

515

Y

9 25

Cu

1

Photomicrograph (xn)

97.38 2.42

Pb

2

Al

48

Zn

162

CCPI

51

Sb

0.3

Ti/Zr

Tl

<0.5

20.5

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 3 7

Strong, perfasiwe chlorite + sericite + quartz + pyrite alteration facies Sample no.

137127

Alteration facies

strong, pervasive chlorite + sericite + quartz + pyrite

Location

Highway footwall, 100 m east of stringer zone

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive quartz + plagioclase-phyric rhyodacite

Relict minerals

quartz + altered plagioclase

Relict textures

porphyritic

HR 5

Primary composition rhyodacite Lithofacies

massive

Interpretation

cryptodome

Alteration minerals

quartz + sericite + pyrite + chlorite ± rutile

Alteration textures

pervasive in groundmass, millimetre patches of microcrystalline quartz and wispy domains of aligned sericite, some broken quartz phenocrysts

Geochemistry K2O

local; medial to proximal zones laterally equivalent to upper parts of sulfide pipes

SiO2

64.55

TiO2

0.44

poor

AI2O3

17.02

Alteration intensity

strong

Fe2O3

5.54

co 2

Timing

synmineralisation

MnO

0.04

Total

MgO

1.87

LOI

Alteration style

hydrothermal

CaO

0.20

Na2O

0.18

Distribution Preservation

Hand specimen photograph

S

Photomicrograph (xn)

4.53

Rb

92 Zr

0.08

Sr

28

Ba

3683

Cu

16

2.97

149

Nb

8

Y

25

32 Al

94

68

CCPI

59

Sb

2.0

Ti/Zr

Tl

2.0

97.43 Pb 4.93 Zn

17.8

2 3 8 | CHAPTER 7

Intense sericite + quartz + pyrite alteration facies Sample no.

137080

Alteration fades

intense sericite + quartz + pyrite

Location

upper proximal, 20 m east of Highway pipe

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive, sparsely plagioclase-phyric dacite

Relict minerals

altered plagioclase

Relict textures

porphyritic

HR§

Primary composition dacite Lithofacies

massive

Interpretation

cryptodome

Alteration minerals

quartz + sericite + pyrite

Alteration textures

pervasive, polycrystalline quartz pseudomorphs after plagioclase, microcrystalline matrix of quartz + sericite > pyrite local, proximal zone enveloping Highway sulfide pipe

Distribution Preservation

poor

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2 TiO2 AI2O3 Fe2O3 MnO MgO CaO Na2O

73.82 0.41 12.23 4.55 0.02 0.76 0.09 0.08

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

3.65 0.06 2.89

Rb Sr Ba Cu 98.56 Pb 4.09 Zn Sb Tl

78 14 2222 1385 10 90 1.3 5.3

Zr Nb Y

107 6 13

Al CCPI Ti/Zr

96 57 23.0

LOCAL HYDROTHERMAL ALTERATION RELATED TO VHMS DEPOSITS | 2 3 9

Intense, pervasive chlorite + pyrite alteration facies

HR7

Footwall Sample no.

137083

Alteration fades

intense, pervasive chlorite + pyrite

Location

Highway footwall, adjacent to stringer zone

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic fades

altered dacitic pumice breccia?

Relict minerals

nil

Relict textures

nil

Primary composition dacite Lithofacies

indeterminate

Interpretation

indeterminate

Alteration minerals

chlorite + pyrite > quartz

Alteration textures

pervasive cryptocrystalline groundmass or matrix of chlorite, cleavage, 5% disseminated euhedral pyrite with quartz pressure shadows

Distribution

local; narrow zones enveloping footwall quartz + pyrite stringer zone

Preservation

nil

Alteration intensity

intense

Timing

synmineralisation

Alteration style

hydrothermal

Hand specimen photograph

Geochemistry SiO2 24.90 TiO2 0.61 AI2O3 18.79 Fe2O3 16.73 MnO 0.35 MgO 22.70 CaO 0.21 Na2O 0.04

K2O P2O5 S CO2 Total LOI

Photomicrograph (xn)

0.01 Rb 0.12 Sr 6.37 Ba Cu 90.83 Pb 13.80 Zn Sb Tl

<1 9 21 102 12 522

Zr Nb Y

169 9 27

Al CCPI Ti/Zr

99 100 21.7

2 4 0 | CHAPTER 7

Intense, stringer quartz + sericite + pyrite alteration fades

HR8

Footwali Sample no.

137129

Alteration facies

intense, stringer quartz + sericite + pyrite

Location

Reward footwall stringer zone

Formation

Trooper Creek Formation

Succession

Mount Windsor Subprovince

Volcanic facies

massive

Relict minerals

nil

Relict textures

nil

Primary composition

rhyodacite

Lithofacies

massive

Interpretation

synvolcanic sill?

Alteration minerals

quartz + sericite + pyrite

Alteration textures

pervasive, irrregular sericite pseudomorphs after feldspar in <10 pm microcrystalline quartz, interstitial sericite groundmass, disseminated 5-10% euhedral pyrite

Distribution

local; footwall stringer zones beneath

Geochemistry SiO2

75.78

nil

TiO2

Alteration intensity

intense

AI2O3

Timing

synmineralisation

Fe2O3

9.89

CO2

hydrothermal

MnO

0.02

Total LOI

sulfide pipes Preservation

Alteration style

Hand specimen photograph

Stacked SWIR spectra (hull quotient)

K2O

1.80

Rb

0.18

P2O5

0.02

5.87

S

7.16 101.21

Zr

52

Sr

13

Nb

3

Ba

1404

Y

8

Cu

11

Pb

3

Al

90

MgO

0.26

Zn

12

CCPI

CaO

0.20

Sb

1.2

Ti/Zr

Na2O

0.03

Tl

1.2

Photomicrograph (xn)

TiO2-Zr immobile element plot

5.81

30

83 20.8

I 241

8 | FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS

Recognising alteration facies that may be genetically related to ore deposits is an important step in mineral exploration. Even more helpful is the ability to identify alteration facies that are likely to be associated with particular ore deposit types and thus prioritise exploration targets. The characteristics of alteration facies have the potential to be used as exploration vectors, guiding explorers to the most prospective altered zones in a system, and thereby enabling efficient and early testing of the best targets, avoiding expensive, protracted exploration programs, and improving the chance of success. The processes that alter volcanic facies and the range of textures and mineral assemblages they produce are complex and challenging. As described in previous chapters, there are a variety of alteration processes, which can produce a broad range of alteration mineral assemblages and textures. Ancient volcanic rocks commonly contain a complex assemblage of overprinting alteration minerals and textures, which reflect multiple episodes of alteration by a variety of processes: diagenesis, hydrothermal alteration, deformation, metamorphism or weathering. In early Palaeozoic volcanic regions, like the Mount Read Volcanics, western Tasmania and the Mount Windsor Subprovince, north Queensland, patience and experience are required to unravel the complexities of altered rocks and recognise those altered zones that are 'red-herrings' to mineral explorers. In fact, numerous geologists have initially doubted that the foliated, weathered and mosscovered rocks encountered in western Tasmania really were of volcanic origin. Several intensive, protracted and ultimately unsuccessful exploration programs have been conducted in the Mount Read Volcanics on unfavourable altered zones. On the other hand, there may be altered zones that remain under-explored because favourable alteration facies were not recognised. Recognising the occurrence of altered rocks and identifying favourable or prospective alteration facies and zones are important steps toward minimising risk and expenditure during exploration in these environments. This chapter draws together the descriptive and geochemical techniques described in Chapters 2, 3 and 4, and an understanding of the different alteration processes that modify submarine volcanic successions. It proposes methods for discriminating alteration facies associated with particular processes, identifying favourable altered zones for mineral

exploration, and guiding exploration within those zones toward potentially mineralised areas.

8.1 | PRINCIPLES OF DISCRIMINATING BETWEEN DIAGENETIC, HYDROTHERMAL AND METAMORPHIC ALTERATION FACIES Diagenetic facies As discussed in Chapter 5, the characteristics of diagenetic facies are: • They are typically widespread with district or regional-scale distribution. • At local scales, they display variable alteration intensity and patchy distribution. This is mainly controlled by distribution of coherent versus clastic volcanic facies, and variations in the primary composition, permeability, porosity and the proportion of glassy to crystalline facies. • They occur in vertically-stacked, extensive, sub-horizontal altered zones, which have mineral assemblages that reflect increasing temperature with depth. • They have undergone relatively minor (< 10 wt%) chemical changes that are predominantly in response to hydration or alkali-exchange reactions between the volcanic facies and modified seawater. Mass transfers are generally small, an order of magnitude less than those in intense hydrothermal alteration facies. The scale of migration of elements is also small (millimetres to tens of centimetres) and thus on a larger scale (i.e. basin scale) the changes are essentially isochemical. • Their mineralogical and textural changes vary from subtle to strong. Quartz phenocrysts, for example, are relatively stable and commonly well preserved, whereas mafic phenocrysts and volcanic glass are relatively unstable and typically completely altered. These changes are commonly overprinted or obscured

2 4 2 I CHAPTER!

by subsequent alteration as diagenesis is often the earliest preserved post-emplacement process.

Metamorphic facies Metamorphic facies share some characteristics with diagenetic facies, but also differ significantly in these ways: • Distribution varies in scale: contact metamorphic facies associated with intrusions may be as narrow as a few centimetres and up to several kilometres wide. Regional metamorphic facies (either burial metamorphism or metamorphism associated with deformation) can be tens or hundreds of kilometres wide and several kilometres thick. • Metamorphic facies are uniform and pervasive: they are not typically patchy at a scale of metres to tens of metres. Primary volcanic textures have virtually no influence on high-grade metamorphic facies, which are principally determined by whole-rock compositions and metamorphic conditions. • Chemical changes are minor; metamorphism is generally a process of phase-change in response to changing temperature and pressure at low water-rock ratios, which limits the redistribution of chemical components in and out of the system. The most common metamorphic reactions are dehydration and decarbonation reactions. The composition and mineralogy of metamorphic facies are generally strongly influenced by the primary composition of volcanic facies. • Mineralogical and textural changes vary from subtle to intense depending on the degree of metamorphism. Typically, primary volcanic quartz phenocrysts are well preserved up to about amphibolite grade, but fine-grained or glassy facies and some mafic phenocrysts are unlikely to survive even low grades of zeolite and greenschist facies or contact metamorphism. Metamorphic re-crystallisation produces a wide variety of distinctive textures, such as granoblastic, porphyroblastic, decussate, schistose, and gneissic, which are not easily confused with primary volcanic or diagenetic textures.

Hydrothermal alteration facies Hydrothermal alteration facies are unlike diagenetic and metamorphic facies in their potential for major compositional change. This is because hydrothermal alteration typically involves large volumes of fluid, which facilitate large-scale mass transfers into, out of, or around hydrothermal systems. Depending on the intensity of alteration, this characteristic determines or limits the other characteristics of hydrothermal alteration facies. • Hydrothermal alteration facies generally have local distribution, limited to tens or hundreds of metres and rarely exceeding a few kilometres. • Hydrothermally altered zones commonly have high aspect ratios (i.e. narrow lateral and great vertical extents) because convecting, typically ascending, fluids produce them. • Locally, on small-scales, the distribution of hydrothermal alteration facies is generally uniform, or pervasive. However,

the distribution is mainly dependent on permeability and porosity; therefore hydrothermal alteration facies may be restricted to fractures and vein selvedges in coherent or otherwise impermeable rocks. • The degrees of mineralogical and textural preservation, and chemical changes are extremely dependent on alteration intensity and pre-hydrothermal alteration composition and texture of the facies. Pre-existing textures and minerals are less likely to be preserved in proximal zones of hydrothermal systems, through which hot reactive fluids are flushed, than in peripheral zones with lower temperature, partly neutralised fluids and lower fluid-rock ratios. As in the other types of alteration, quartz crystals in felsic volcanic facies tend to survive intense alteration, except where major loss of silica is involved (e.g. in chlorite zones). Other primary crystal phases are commonly progressively replaced (e.g. feldspars altered to sericite) and may be useful as indicators of alteration intensity. Hydrothermal alteration facies rarely preserve primary textures in originally glassy facies. • Hydrothermal mineral assemblages are largely controlled by fluid composition and physicochemical conditions, and are not noticeably influenced by primary compositions; at least in the intensely altered zones, which had high fluidrock ratios. Thus, an intensely hydrothermally altered zone may cut across volcanic lithofacies of different primary compositions and textures (e.g. coherent andesite and rhyolitic breccia) and comprise only one alteration facies in which the protoliths are mineralogically and texturally indistinguishable. • Hydrothermal alteration commonly involves significant mass transfer of chemically mobile elements. Elements may be gained through precipitation or lost through dissolution. These mass transfers may produce large positive or negative net mass changes within particular alteration facies (generally with implications for volume change) or balance each other out to produce negligible net change. Significant mass changes are commonly evident in composition data and derivative alteration indices. For example, Na depletion typically accompanies hydrolysis and sericitisation of plagioclase. However, substantial mass changes in some major elements are commonly obscured by the constant sum effect; this applies especially to Si. • Major chemical modifications are frequently reflected in exotic mineral assemblages. For example, VHMS-related alteration facies commonly contain disseminated pyrite or base-metal sulfides, and several types of Zn deposits are associated with Mn-rich mineral assemblages. It is important for economic geologists to recognise hydrothermal alteration facies, which may indicate the largescale transport and deposition of economically valuable elements, and to discriminate these from alteration facies that result from other alteration processes that are unrelated to ore deposition. In some cases, examination of an individual altered sample can reveal important facts that help to identify the alteration process. For example, a rock with gneissic fabric is metamorphic; a rock composed essentially of quartz and pyrite is probably of hydrothermal origin. However, alteration textures and mineral assemblages may not easily distinguish some weak hydrothermal alteration facies, perhaps in peripheral zones, from diagenetic or metamorphic facies. One of the main criteria distinguishing hydrothermal

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS | 2 4 3

from other alteration facies is the distribution or extent of the altered zone. This cannot be determined by observation of an individual specimen or outcrop. It requires systematic prospect-scale mapping and knowledge of the district-scale geological context. Unfortunately, in the last decade of the 20 th century there has been a significant decline in in-field and on-ground geological data collection, particularly in the mineral exploration and mining industries. There is a trend towards using high technology remote sensing to rapidly explore large areas at continuously improving resolutions. However, to ensure meaningful interpretation of these data, it is imperative that this virtual geology is not disconnected from real rocks. The combination of a variety of criteria, and high-quality mapping, will lead to the best interpretation of alteration facies.

8.2 | EXPLORATION VECTORS AND PROXIMITY INDICATORS Mineral zonation Mapping of sulfide distribution, particularly pyrite, is an important exploration technique. Sulfide abundances are easily estimated by eye, even in weathered samples, and should be applied at an early stage of exploration wherever geological exposure permits. VHMS deposits commonly have extensive footwall zones of disseminated pyrite. For example, the Mount Lyell Cu-Au deposits (western Tasmania) all lie in a zone of greater than 1 % disseminated pyrite, which is 6 km long and 1 km wide at the surface (Corbett, 2001). Such pyritic zones provide very large exploration targets for initial area selection. They have the potential to be delineated into high-abundance zones, in order to reduce the size of the targets for intensive exploration and drill testing. Interpretation of sulfide vectors is straightforward: more is better, and sulfide proportions generally increase with proximity to sulfide deposits. Other components of alteration mineral assemblages that are easily recognisable in all sample types and may be spatially zoned around mineral deposits include silicates, carbonates and Fe-oxides. The ratios of quartz to phyllosilicates, sericite to chlorite, and carbonate to silicates are commonly systematically zoned around VHMS deposits, and recognition of the zonation patterns can provide useful exploration vectors, at least on a prospect scale. Unfortunately, interpretation of the patterns is rather complex. Australian VHMS deposits are typically associated with siliceous proximal zones (Section 7.4) but there are many variations even within mineral fields and districts. Therefore it is unwise to be too strictly empirical or model-driven in applying this approach. It is better to map out mineral distributions and relate them to alteration intensities, rather than rely on the recognition of specific zonation patterns, which may relate to an ore deposit model. Carbonate + chlorite assemblages, for example, are indicators of ore proximity in some VHMS deposits, such as Rosebery, Hellyer and Thalanga deposits, but only occur in the peripheral or least-altered zones of the Western Tharsis deposit (Mount Lyell field). Therefore, rigidly applying a carbonate + chlorite

vector could be misleading and potentially guide exploration away from some Cu-Au deposits in the Mount Lyell field. Mapping of mineral zonation is effective where large systematic datasets are available (i.e. where there are plenty of outcrops or drill cores) and mineral assemblages are visually distinctive or can be determined by simple field tests (e.g. effervescence in acid for carbonate or sodium-cobaltinitrite staining for K-feldspar). However some mineral assemblages that are not readily identified visually, are discretely zoned and may be diagnostic of a deposit style. New field-based mineralogical tools, such as portable SWTR spectrometers (Section 2.4), will facilitate major improvements, which will not only aid exploration, but also contribute to understanding these deposit systems (Thompson et al., 1999). SWIR spectral studies have recently shown some spectacular examples of mineral zonation, particularly in acid-sulfate type systems (e.g. case studies in Thompson et al., 1999, and Huston and Kamprad, 2001).

Major element lithogeochemistry Although intense hydrothermal alteration frequently produces simple alteration mineral assemblages, the minerals are commonly fine grained. These minerals may be difficult to recognise visually, and it can also be difficult to estimate their abundances. In these cases, lithogeochemistry can frequently help to identify minerals and quantify compositional changes even in less intensely altered rocks that contain incipient, overprinting or domainal alteration minerals. Analysis of whole-rock samples to determine major element abundance is a way of supporting and augmenting estimates of mineral proportions and alteration intensity, which have been determined visually or by other methods (e.g. Section 2.4). Quantitative lithogeochemical data can be used in two ways: (1) to indicate alteration intensity, and (2) to estimate mineral proportions in mineral assemblages where the mineral species and their individual compositions are known. Interpreting exploration vectors based on compositional data seems straightforward. The data can be plotted as contour maps or cross-sections (e.g. Figs 2.7 and 2.12), down-hole line graphs (e.g. Fig. 2.14) and so on, and the vectors inferred based on expected variations in mineral abundance or composition. Decreases in Na 2 O contents of volcanic rocks, for instance, are usually related to increasing sericite or chlorite at the expense of plagioclase. Na 2 O depletion is a popular and reliable vector used in VHMS exploration (e.g. Na 2 O halo maps of the Fukuzawa area in Date et al., 1983). Variations in carbonate content, both increases and decreases, are typically evident in CO 2 data and in CaO, MgO or Fe2O3, depending on the carbonate species. Sulfide content can be quantified by sulfur analyses. Weight percent sulfur is generally nearly equivalent to volume percent of pyrite in felsic rocks, if pyrite is the only sulfur-bearing phase. This is due to pyrite having a density of just under twice the density of felsic rock, and sulfur comprising just over half the mass in pyrite. However, major element data are subject to distortion by closure, otherwise known as the constant sum effect. This phenomenon is more fully explained in Section 4.1. It particularly affects the dominant chemical components (e.g. SiO2, A12O3 and Fe2O3) and can be significant in

2 4 4 | CHAPTER!

hydrothermally altered rocks with large net mass gains. Although additions of exotic hydrothermal components such as sulfur and CO 2 , and depletions of Na 2 O, are relatively immune to the effects of closure, it seriously compromises the use of some other major components as exploration vectors. For example, SiO2 may not provide effective vectors in hydrothermal systems where mineralisation was associated with silicification. If closure in major element data is likely to obscure the effects of alteration on the compositional data and exploration vectors, then it is preferable to estimate the individual component mass changes (by immobile element techniques, Section 4.1) and use those as exploration vectors. The alternative approach is to convert quantitative major element lithogeochemical data to modal mineral proportions using a method such as MINSQ (Herrmann and Berry, 2002) or GENMIX (Le Maitre, 1981). This does not remove the effects of closure, but is a way of quantifying mineral proportions, which can then be used as vectors in mineral exploration. This approach was used by Large et al. (2001b, Fig. 6) to demonstrate systematic variations in proportions of alteration minerals around the Rosebery K-lens sulfide deposit (Fig. 2.14).

Alteration indices Alteration indices formulated from two or more components of major element analyses (Section 2.4) enhance the compositional contrast between variably altered samples and thus are frequently more effective as exploration vectors than single component lithogeochemical data. For example, sulfur and Na 2 O proportions in the footwall of the Rosebery K-lens deposit range from 0.01% (limit of detection) up to about 7.2% and 5.6%, respectively (Large and Allen, 1997). However, the ratio S/Na2O ranges from 0.002 to 194, because those components increase and decrease respectively in response to increasing alteration intensity (Large et al., 2001b). Both components vary over two to three

orders of magnitude, whereas S/Na2O varies across about five orders of magnitude. Carefully formulated indices can in this way amplify compositional changes and reflect variations in more than one mineral composition or abundance. Where systematic lithogeochemical data are available, plotting and contouring of alteration indices on plans and crosssections provides numerical indications of alteration intensity (e.g. Fig. 2.7). Datasets of alteration indices are of assistance in guiding exploration towards potentially mineralised altered zones, especially when used in combination with alteration facies or mineral zonation maps. Mineral explorers have increasingly applied these techniques to VHMS exploration over the last two decades; however, few results or case studies have been published. Exploration data are commonly limited to a few samples or drill holes and are not suitable for contouring. Nevertheless, useful vectors can be inferred from sparse but strategically or fortuitously located data. This is exemplified in lithogeochemical data from a few drill holes near the northern end of the Rosebery deposit (65R, 109R, 113R and 128R; Table 8.1). If, in a VHMS exploration scenario, the first two holes were drilled in sequence (65R followed by 109R), then the lithogeochemical vectors would suggest that exploration was heading away from the most favourable zone. The intermediate third hole, 113R, would then be superfluous, merely reinforcing interpretation of vectors in the first two holes. The anomalous values in the near-miss hole (65R) would encourage further persistence. If, on the other hand, the first hole in a greenfields exploration program was 109R, the major element or alteration indices data would not justify continuing exploration in that vicinity, even if the favourable stratigraphic setting was recognised. In this case, success would depend on the explorer recognising other vectors or indicators of proximity, such as the distal trace element Tl and Sb halos identified by Large et al., (2001b). In favourable geologic settings, limited lithogeochemical data, even from a single drill hole, may yield useful vectors. For example, samples from a single hole, such as HL6 or

TABLE 8.1 | Selected major element data and alteration indices for samples of pumice breccia from the footwall to the Rosebery K-lens massive sulfide deposit, western Tasmania. The values tabulated are (A) averages of three samples from the top 30 m of the footwall unit and (B) the uppermost sample of the footwall intersected in eachdriii hole. The alteration indices, S/Na2O and Al, generally show greater increases with proximity to ore than the changes in Na2O, S and Zn. Averaging the uppermost three samples smoothes the gradients towards ore, but diminishes the anomalies in the medial intersection, 113R. Data from Large and Allen (1997). Locations of the drill holes are shown in Figure 2.7 of this volume and Figure 2 of Large et al. (2001b).

(A) Averages of three samples from the top 30 m of the footwall unit 109R

450

2.99

0.18

0.00

0

44

113R

250

1.66

0.37

0.13

2

61

65R

75

0.28

0.79

0.18

44

89

128R

0

0.01

0.49

0.49

49

89

(B) Uppermost sample of footwall unit 109R

450

1.53

0.29

0.01

0

51

113R

250

0.21

1.09

0.38

5

89

65R

75

0.08

1.02

0.02

13

95

128R

0

0.01

0.37

0.19

37

89

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS | 2 4 5

HL14 drilled through the footwall zones beneath the Hellyer massive sulfide deposit, generally exhibit gradually increasing alteration indices (Figs 9 and 20 of Gemmell and Large, 1992). Recognition of these variations, in combination with the visible alteration facies, confirms that an altered zone exists and indicates the direction of increasing alteration intensity, guiding further exploration. Drill hole NC4, which intersected the Tyndall GroupCentral Volcanic Complex contact south of Henty is another interesting example. In this hole, an abrupt down-hole increase in the alteration indices is associated with a change of lithotype (Fig. 4.4). The lithogeochemical data support the recognition of an extensive stratabound altered zone of which the upper boundary is most favourable for VHMS exploration. Bivariate (x-y) plots of two alteration indices, such as the AI-CCPI Alteration box plot (Large et al., 2001a), are useful in identifying compositional trends and different alteration facies. This graphic approach simplifies the recognition of rock compositions that lie outside the normal range of primary volcanic compositions (i.e. those that have been modified by chemical or depositional processes; Fig. 2.9). It also assists classification of different alteration facies and identifying the zones of greatest prospectivity (Fig. 2.11). In recent CODES research projects, box plots of customdesigned alteration indices have been effective in several other types of hydrothermal systems, including low- and high-sulfidation epithermal Au-Ag deposits (Williams, 2000) and Broken Hill type Pb-Zn-Ag deposits (Large, 2004). The Ishikawa et al. (1976) alteration index (AI) has been successfully applied to many plagioclase-destructive and/ or K-feldspar-bearing alteration styles, but there is scope for more experimentation with new indices. As outlined in Section 4.1, the formulae for alteration indices typically have chemical components that were increased by alteration in their numerators, and components that were decreased in the denominators. The gained or lost components can often be inferred from the differences in alteration mineral assemblages. However, immobile-element-based mass change calculations provide a more rigorous method of selecting components for formulating alteration indices. Section 4.1 summarises several techniques of estimating mass changes by comparing compositions of alteration facies to their leastaltered precursor compositions and their potential application to exploration vectors is discussed below.

Mass change vectors Hydrothermal alteration commonly involves major changes in chemical composition; in fact these changes are one of the characteristic features of hydrothermally altered rocks. Significant masses of mobile chemical components may have been gained or removed from an altered zone. However, closure in composition data will obscure or distort the amounts of these changes, except in the special cases where the mass gains and loses balance exactly, so that there is no net mass change. It is unsound to assume zero net mass change in an alteration facies, and in these cases the unquantified effect of closure on raw major element data limits their usefulness, or that of alteration indices based on them, as indicators of alteration intensity.

The solution to the closure problem is to estimate the mass changes of all mobile major-element components, using an immobile-element-based method of the type described in Section 4.1. Spatially located mass change data can then be used as direct indicators of alteration intensity or as multiple component alteration indices, in the same way as major element lithogeochemical data. This approach has the potential to target favourable areas during exploration. It provides closure-free quantification of compositional changes, which help delineate hydrothermal fluid pathways, zones of greatest alteration intensity and prospective areas. Furthermore, the quantification of absolute mass changes is a means of estimating the 'quality' of an altered zone. For example, let us consider a hypothetical program of lithogeochemical sampling over two altered zones of similar dimensions in a VHMS district. Mass change estimates might show that the first altered zone involved negligible mass transfers and the second had significant mass gains, of the order of 20—30 g/lOOg and equating to tens of millions of tonnes of altered rock (cf. Thalanga footwall zone, Herrmann and Hill, 2001). In this case, we would conclude that the second altered zone has greater mineral potential. Substantial mass changes demonstrate that a hydrothermal system had the intensity, and perhaps duration, to move large amounts of SiO2, CO 2 , S and other components into the alteration facies. Therefore, it probably also had the capacity, if fluid compositions were suitable, to move large amounts of base and precious metals and potentially, if a favourable site and process for deposition is available, form an ore deposit. The first altered zone in our hypothetical example was produced by near-isochemical alteration and resulted in negligible changes to the whole-rock composition, suggesting that alteration involved less reactive or smaller volumes of fluid, perhaps over a short duration. The differences may be semi-evident in alteration mineral assemblages and intensities, and possibly in the composition data despite distortion by closure, but the only way to quantify the difference for objective exploration decisions is by mass transfer techniques. The major difficulty in this method is in determining precursor compositions to compare with the altered compositions. Poor exposures, limited lithogeochemical data, lateral variation in the primary composition of volcanic facies or structural complexity make the pairing of alteration facies and unaltered (or least-altered) precursors problematic, and frequently impossible, in practical application. There are no published examples where mass change calculations have led to a mineral discovery, probably because of the leastaltered precursor problem and the only recent development of easy mass change calculation techniques. Nevertheless, the mass change approach will contribute to a higher level of lithogeochemical interpretation and exploration targeting where host volcanic successions are compositionally uniform and sufficiently understood to enable its confident application.

Mineral chemistry vectors As noted in Section 4.2, the main limitations to the wide use of mineral chemistry in exploration have been that the analytical

2 4 6 | CHAPTER 8

tools — electron microprobe and X-ray diffraction — are complex laboratory-based instruments requiring considerable expertise in operation and data interpretation. That has made mineral analysis slow and expensive relative to geochemical analyses of rocks, soils and sediments, and consequently explorers have largely ignored the mineral chemistry vector possibilities. Researchers at CODES are currently developing laser ablation ICP-MS techniques for micro-analysis of trace elements in sulfides. These are likely to provide exploration vectors but, for similar reasons, they may not ultimately be widely applied by mineral explorers. However, the advent during the last decade of portable short wavelength infrared (SWIR) spectrometers, which can indirectly measure compositional variations in micas, clays and carbonates, could establish mineral composition mapping as a viable exploration technique (Sections 3.1 and 4.2, and references therein). SWIR spectrometers such as PIMA are relatively inexpensive at about US$21,000 to purchase or US$70 per day for hire. They can analyse up to a few hundred samples per day of all types of geological materials, which require no preparation apart from drying. SWIR spectrometers are simple to operate and the PC-based spectral recognition software now available has simplified spectral interpretation and data manipulation, so that an operator can quickly become an expert interpreter. White micas, chlorites and clays in altered zones around mineral deposits frequently show spatial compositional variations that could be exploration vectors (Section 4.2). The ease of SWIR spectral analysis now enables explorers to rapidly test for the existence of mineral composition vectors in a large enough set of orientation samples. If the results are promising, the technique can be inexpensively applied on a routine basis to assist exploration targeting. If, on the other hand, SWIR spectral features are invariant or spatially erratic, then little time and money will have been expended. There are not yet many published mineral exploration case studies involving portable SWIR spectral analysis because it is a relatively new technique (e.g. Denniss et al., 1999; Huston et al., 1999; Merry and Pontual, 1999; Herrmann et al., 2001; Jones et al., in prep.). Nevertheless, recent and current research at CODES shows great potential for SWIRdetermined white mica composition vectors, on scales of tens to hundreds of metres, in a variety of volcanic-hosted gold and base-metal deposits. Further work is required on spatial SWIR spectral variations in chlorites and clay minerals. It is likely that mineral explorers will rapidly adopt this technique over the next few years. Part of the stimulus comes from very recent developments in airborne high-resolution visible-to-SWIR spectral scanning systems, such as HyMap", which offer great promise for districtscale mineral mapping in exploration of well-exposed bedrock areas (Taranik, 2001). For example, mineral maps from a trial HyMap* airborne spectral survey of the Panorama VHMS district, Western Australia, apparently 'show the complete hydro thermal convective system' (Cudahy et al., 2000). At Panorama, these authors consider that spectrally interpreted distributions of white mica, pyrophyllite and topaz define altered zones that formed at the boundary between magmatic fluid and seawater convection, in addition to seawater recharge zones, and hydrothermal discharge zones. The discharge zones are prospective for massive sulfides. A similar HyMap* survey

of the Mount Lyell area in western Tasmania has produced mineral distribution and pyrophyllite abundance maps (e.g. Fig. 8.1). These illustrate the high spatial resolution now available from airborne spectral surveys, and their enormous potential for alteration mapping and using vectors during exploration in well-exposed, thinly vegetated areas. Remote sensing spectral systems are also finding applications in regolith mapping (Craig, 2001) and exploration of partly covered areas. Bierwirth et al. (2002) used HyMap data to map distributions of a range of minerals — including pyrophyllite, white mica, Mg- and Fe-chlorite, calcite, dolomite, kaolinite, tourmaline, hematite and goethite — in altered zones associated with epithermal and lode Au deposits in the poorly exposed, largely alluvium- and calcrete-covered Indee District of the Central Pilbara. These demonstrations of district-scale mineral and mineral compositional mapping by remote sensing tools should certainly encourage explorers to use spectral data in prospect-scale investigations. In addition, high-output, multipurpose visible-SWIR and thermal infrared spectral, and laser instruments such as CSIRO's HyLogging and HyChips systems (Syddell, 2004) and the OARS prototype (CSIRO, 2002), are being developed for routine logging of drill core, cuttings, soil and other geological sample materials.

Isotopic vectors Section 4.3 introduces the potential for isotope geochemistry to yield interpretations of hydrothermal fluid sources, temperatures, water-rock ratios, and broad halos for exploration targeting. Oxygen isotopes are particularly useful in exploration because oxygen is a major component of hydrothermal fluids, and it readily exchanges isotopes with silicate minerals at fractionation factors that are mineral specific and temperature dependent. Furthermore, the 518O-depletion halos observed around several deposit types typically extend further from ore than most other geochemical anomalies and may provide direct vectors to ore zones. For example, the 518O-depletion zone around the Fukuzawa deposits in the Hokuroku district, Japan, extends up to 1 km beyond the Na2O-depletion anomaly (Green et al., 1983). Waring et al. (1998) found 618O-depletion anomalies in dolomitic shale at Mount Isa (Queensland), which extend up to 2 km beyond Cu ore zones, with low and uniform isotopic gradients (<2%o per 100 m) that allow estimates of the distance to ore. Most importantly, the O-isotopic anomalies produced in hydrothermally altered zones appear to survive subsequent deformation and metamorphism. For instance, Cartwright (1999) argued convincingly that a hydrothermally related regionalscale 618O depletion zone in Proterozoic metapelites in the Broken Hill district, NSW, had survived high-grade regional metamorphism up to granulite facies. The final section of this chapter summarises several VHMS-related alteration studies and exploration programs, which have applied whole-rock Oisotope geochemistry. Sulfur-isotope geochemistry has been widely applied to interpretations of sulfur (and hence fluid) sources, and hydrothermal temperatures, which have been used in the development of VHMS genetic models. For example,

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS I 2 4 7

Map projection: UTM zone 55, AGD66

FIGURE 8.1 | Mineral maps of the Mount Lyell mine area, western Tasmania, interpreted from HyMap® airborne hyperspectrai data. Map A shows the zonal distributions of eight important alteration minerals. Map B shows relative abundance of pyrophyllite (warm colours = high abundance), and discriminates the pyrophyllite-rich facies at North Lyell, Western Tharsis and Glen Lyell from weaker responses in the Owen Group exposed on Mount Lyell. The spatial resolution (pixel size) is about 5 m. Mineral spectral responses are partly restricted by vegetated areas, which appear as dark grey tones on the HyMap band (greyscale) background airphoto images. These maps were created by K. Yang, M.A. Quigley and J.F. Huntington as part of the 2003 HyMap® mineral mapping project for Copper Mines of Tasmania and Mineral Resources Tasmania, carried out through the C-Vista strategic alliance between CSIRO and HyVista Corporation.

2 4 8 | CHAPTER 8

S-isotope compositions constrained some of the genetic interpretations for formation of the Hellyer deposit (e.g. Gemmell and Large, 1992; Solomon and Khin Zaw, 1997). It also has exploration potential for discriminating different types of deposits and hydrothermally altered zones, which may have different economic potential. The regional study of sulfide deposits in the Mount Read province by Solomon et al. (1988) found considerable variation in S-isotope compositions consistent with different geologic settings and metal associations, and which contributed to interpretations of hydro thermal geochemistry. Green and Taheri (1992) took both a genetic and discriminatory approach to the interpretation of low 634S values of pyrite (-1.2 to +4.7%o) at the Boco prospect, western Tasmania. They suggested that altered zones at Boco formed in a seawater-hydrothermal system, which leached sulfur from volcanic host rocks at temperatures that were too low to inorganically reduce seawater sulfate, and transported base metals to form an ore deposit. Subsequent recognition of advanced argillic alteration mineral assemblages at Boco and several other Tasmanian and Victorian prospects indicate possible involvement of magmatic fluids, and hence a magmatic source of sulfur (Herrmann et al., 2004). Regardless of the genetic uncertainties, sulfide 634S values of less than 5%o could distinguish barren pyritic altered zones from more prospective base and precious metalrich VHMS systems in the Mount Read province. There are few published accounts of S-isotope compositions as direct vectors in mineral exploration. However, existing data for Rosebery (Davidson et al., 2000) and Hellyer (Gemmell and Large, 1992) suggest broad halos of 634S enrichment in disseminated pyrite in footwall zones lateral to the main upflow zones, which could be used to increase exploration target sizes and zero-in on Zn-rich VHMS deposits, particularly in permeable volcaniclastic successions (Large et al., 2001c). At regional scales, two recent studies of deeply covered areas have promoted S-isotope compositions of sulfates in groundwater as potential indicators of buried oxidising PbZn-Ag sulfide deposits, in the Broken Hill region of New South Wales (Waring et al., 1998) and Gawler Craton in South Australia (Kirste et al., 2003). In the latter case, sulfates from oxidising sulfide deposits with low 534S signatures (-2.5 to +5.6%o) appear to have contributed to anomalous low 534S values in groundwater sulfates, detectable several hundred metres downstream from the Menninnie Dam prospect. Background 634S values of sulfates in ground waters are 16 to 18%o in the Gawler region and -13.5%o in the Broken Hill region. The concept is probably less applicable to exploration for sulfide deposits with higher 634S signatures (e.g. Tasmanian VHMS deposits, 8 to 17%o, Solomon et al., 1988), which would provide less contrast against background groundwater compositions. Furthermore there are many, typically difficult to determine, hydrological and geochemical factors that complicate interpretations of local groundwater isotopic anomalies. This new application of S-isotope geochemistry is one that will probably appeal only to the most persistent of under-cover mineral explorers. Carbon isotopes, like sulfur, are used for interpreting fluid sources and hydrothermal conditions but have not been widely applied as exploration vectors. Huston's (1999) review of stable isotopes in VHMS systems found carbonate 613C values in most deposits occupy a narrow range of-5 to

0%o, consistent with seawater dissolved bicarbonate sources. Low fractionation factors, and the limited occurrences of carbonates in massive sulfide deposits (Ohmoto and Goldhaber, 1997) restrict the applications of C isotopes, except in conjunction with O isotopes. For example, Khin Zaw and Large (1992) interpreted a coupled positive trend of 613C and 618O data in Mn-rich carbonates at South Hercules, Tasmania, as temperature-related, and then, with additional fluid inclusion temperature data, estimated the isotopic compositions of the hydrothermal fluid. Although their paper did not describe spatial zonation of isotopic data, the genetic discussion speculated that mineralised and altered facies were zoned according to variations in temperature and hydrothermal fluid-seawater mixing ratios, controlled by permeabilities in the volcaniclastic succession. In these types of deposits associated with lateral carbonate facies, isotopic data could provide prospect-scale exploration vectors if the hydrothermal temperature gradients were consistent. Callaghan's (2001) study of the Henty-Mount Julia gold deposit, Tasmania, used carbonate 613C and 618O data in a boomerang shaped trend for intensive modelling of fluid compositions and genetic concepts. The data, crudely divided into proximal and distal carbonates, lie on two trends joined at an abrupt inflection. Both of the fluid mixing or fluidlimestone interaction models proposed by Callaghan (2001) to account for the trends offer potential for prospect to districtscale isotopic vectors, or at least methods of discriminating hydrothermal and sedimentary carbonates.

Whole-rock O-isotope vectors in VHMS exploration In most cases, the proximal altered zones of VHMS systems show significant 618O depletion, partly attributable to highfluid temperatures and low-fractionation factors of some minerals (e.g. chlorite) in seafloor hydrothermal discharge zones, and partly due to the contrast with 618O enrichment caused by low-temperature seawater-rock reactions in normal submarine volcanic successions. A classic semi-regional study by Cathles (1993) in the Noranda district, Canada, discovered a low 618O anomaly (<6%o) in volcanic rocks around the Flavrian felsic pluton. The pluton is surrounded by a discontinuous halo of high wholerock 618O anomalies (>9%o) 10-15 km from the intrusion. Several narrow finger-like zones of low 618O values extend radially from the inner 18O-depleted zone through the high 618O halo, in the directions of most of the known VHMS deposits in the district (Fig. 8.2). These low 518O zones record areas of high hydrothermal-fluid flow and concentrated discharge, which are favourable for mineral deposits. The concentric zones of 18O depletion and enrichment around the pluton closely match the isotopic zonation patterns of Cathles' (1983) numerical model. He concluded that wholerock 818O sampling, at 0.5 km intervals along traverses adjacent to the margins of plutons, could identify plutons with sufficient energy to drive long-lived hydrothermal systems, and favourable settings for detailed massive sulfide exploration. Another district-scale study, in the Panorama area of Western Australia, showed a similar pattern of low whole-rock 618O around the perimeter of a large subvolcanic intrusion

FINDING ORE DEPOSITS IN ALTERED VOLCANIC ROCKS | 2 4 9

FIGURE 8.2 | Map showing spatial relationships between felsic plutons, whole-rock 618O anomalies, and massive sulfide deposits in the Noranda area, Abitibi belt, Canada (modified after Cathles, 1993).

(Brauhart et al., 2000). The granitoid pluton underlies a 1.5 km thick mixed mafic to felsic volcanic succession that hosts several small polymetallic massive sulfide deposits and prospects along a single favourable horizon at the stratigraphic top of the sequence (Fig. 6.6). Narrow radial zones of low 618O point to most of the known deposits and prospects (Fig. 8.3). These low 618O zones coincide with intense feldspar-destructive sericite + quartz and chlorite + quartz zones. Brauhart et al. (2000) calculated hydrothermal temperatures from the 618O data. They used fractionation factors calculated to suit the specific modal mineralogy of each sample, an initial fluid 618O value of+2%o and assumed high water-rock ratios. The resulting calculated temperature distribution closely matched the O-isotopic pattern, the low 618O zones coinciding with temperatures greater than about 300°C (Fig. 8.3). This indicates that temperature was the main control on low whole-rock 818O. It is consistent with increased temperature with depth in the volcanic succession, and in the transgressive discharge or feeder zones beneath the sulfide deposits. The authors concluded that whole-rock O-isotope mapping could be used as a regional exploration vector, and pointed to additional favourable targets in the Panorama district. Green and Taheri (1992) followed up the Hokuroku work of Green et al. (1983) with several empirical isotopic studies of alteration systems in the Mount Read province. The altered footwall zones beneath the Hellyer deposit exhibit a subtle whole-rock 618O anomaly with values ranging from 8.3 ± 1.3%o in the central stringer zone, through 9.8 ± 1.7%o in the enclosing sericitic zone, to background values around 11.3±0.9%o in adjacent least-altered footwall andesites. There is also a subtle depletion anomaly of 10.6 ± 1.2%o in the basalts immediately above the deposit, compared to background values of 11.8±2.2%o. However, the 518Odepletion zone is narrow, reflecting the strong fault or fracture control on hydrothermal-fluid flow. This limits its utility in exploration. The Hercules alteration system also shows a range of whole-rock 6 !8 O values from 6.8%o in footwall zones to background values of 14.0 to 15.5%o. There are some unexpectedly high values (around 15%o) in relatively proximal parts of the footwall and lowvalues (down to 6.8%o)

FIGURE 8.3 | Distribution of whole-rock 618O values and estimated hydrothermal temperatures in the Panorama district of the Pilbra region, Western Australia (modified after Brauhart et al., 2000).

2 5 0 | CHAPTER!

in the apparent hanging wall rocks to the east. These may be partly due to fault displacements that dismembered the alteration system. They highlight the difficulty of applying broad-scale geochemical exploration techniques in deformed and structurally complex terrains. In contrast to Hellyer and Hercules, the apparently barren Boco altered zone has whole-rock 618O values that are not significantly different to background values (9.9 ±1.0 and 10.5 ± l.l%o, respectively). This is consistent with Green and Taheri's (1992) interpretation that the Boco alteration facies formed in a low-temperature (<200cC) seawater hydrothermal system, incapable of transporting base metals and reducing seawater sulfate. Alternatively, the 618O values could indicate a higher temperature, isotopically heavier fluid (>280°C, --5%o), representing either evolved seawater or mixed seawater and magmatic water. The presence of advanced argillic assemblages in parts of the Boco system implies highly acidic fluid conditions, which supports a magmatic fluid input (Herrmann et al., 2004). The least-altered volcanic rocks in VHMS-hosting successions typically have anomalously high background whole-rock 618O values (>9 or 10%o), which are attributable to low-temperature diagenetic alteration. The curves in Figure 4.17 indicate that re-equilibration with quite small proportions of cold seawater can produce large positive shifts in volcanic rock 618O values. On the other hand, zones of low 618O reflect high-temperature hydrothermal alteration at high water-rock ratios. The empirical data from Noranda and Panorama show that low 818O zones may be regionally

extensive at depths of greater than 1 km below favourable horizons, and in narrow finger-like zones that point toward favourable sites for hydrothermal discharge. They may form relatively broad halos around massive sulfide deposits. Despite these promising research results, VHMS explorers have been less than enthusiastic about O-isotope vectors and there are few examples of successful application in Australia. This may be largely attributable to the expense of isotopic analysis (currently around US$150 per sample) and the recognition that interpretation of isotopic data is not straightforward. A notable exception is the case of the Thalanga West 45 deposit, documented by Miller et al. (2001). These authors took a similar approach to Brauhart et al. (2000), using estimates of modal mineralogy to determine tailor-made fractionation factors for each sample, to calculate isotopic equilibration temperatures from whole-rock 618O data, at assumed high water-rock ratios and fluid isotopic composition. They found that zones of apparent high temperatures (>230°C) coincided with the known Central, East and Orient massive sulfide lenses. The existence of an additional isotopic-temperature anomaly, about 1 km west of the known resources, stimulated further exploration that turned up a favourable REE geochemical anomaly in the same sector. Subsequent exploratory drilling discovered a 0.23 Mt polymetallic massive sulfide lens. It remains sub-economic, but may represent the first successful VHMS exploration application of O-isotope geochemistry in Australia.

251

I REFERENCES

Abbott, P. D. B., 1992. Geology of a barite-galena occurrence exposed in the Anthony Power Development Tunnel, western Tasmania: Unpublished Honours thesis, University of Tasmania, 62 p. Ade-Hall, J. ML, Khan, A., Dagley, P., and Wilson, R. L., 1968. A detailed opaque petrological and magnetic investigation of a single Tertiary lava flow from Skye, Scotland - I irontitanium oxide petrology: Geophysical Journal of the Royal Astronomical Society, v. 16, p. 375—88. Alabaster, T, and Pearce, J. A., 1985. The inter-relationship between magmatic and ore-forming hydrothermal processes in the Oman Ophiolite: Economic Geology, v. 80, p. 1-16. Allen, R. L., 1988. False pyroclastic textures in altered silicic lavas, with implications for volcanic-associated mineralisation: Economic Geology, v. 83, p. 1424—46. Allen, R. L., 1990. Subaqueous welding, or alteration, diagenetic compaction and tectonic dissolution? [abs.]: IAVCEI International Volcanological Congress, Mainz, Germany, September 3-8, 1990, Abstracts volume, p. 2. Allen, R. L., 1991. Structure, stratigraphy and volcanology of the Rosebery-Hercules Zn-Pb-Cu-Au massive sulfide district, Tasmania: Results 1988-1990: Unpublished Report to Pasminco Exploration and Mining, 3 volumes, 54 p. Allen, R. L., 1994a. Volcanic facies analysis indicates large pyroclastic eruptions, sill complexes, synvolcanic grabens and subtle thrusts in the Cambrian 'Central Volcanic Complex' volcanic centre, western Tasmania [abs.], in Cooke, D. R., and Kitto, P. A., eds., Contentious issues in Tasmanian geology, a symposium: Hobart, Australia, November 3-4, 1994, Geological Society of Australia Abstracts, v. 39, p. 41—3. Allen, R. L., 1994b. Synvolcanic, subseafloor replacement model for Rosebery and other massive sulfide ores [abs.], in Cooke, D. R., and Kitto, P. A., eds., Contentious issues in Tasmanian geology, a symposium: Hobart, Australia, November 3—4, 1994, Geological Society of Australia Abstracts, v. 39, p. 107-8. Allen, R. L., 1997. Rosebery alteration study and regional alteration studies in the Mount Read Volcanics - the record of diagenetic alteration in the strongly deformed, felsic volcaniclastic succession enclosing the Rosebery and Hercules massive sulfide deposits: AMIRA/ARC Project P439, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 5, October 1997, p. 135-73. Allen, R. L., and Cas, R. A. E, 1990. The Rosebery controversy: distinguishing prospective submarine ignimbrite-like units from true subaerial ignimbrites in the Rosebery-Hercules Zn-Cu-Pb massive sulfide district, Tasmania [abs.]: 10th

Australian Geological Convention, Gondwana: terranes and resources, Hobart, Australia, February 4-9, 1990, Geological Society of Australia Abstracts, v. 25, p. 31-2. Allen, R. L., and Large, R., 1996. Rosebery alteration study: AMIRA/ARC Project P439, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 3, October 1996, p. 143-52. Alt, J., C, 1995a. Sulfur isotopic profile through the oceanic crust sulfur mobility and seawater-crustal sulfur exchange during hydrothermal alteration: Geology, v. 23, p. 585—88. Alt, J. C, 1995b. Subseafloor processes in mid-ocean ridge hydrothermal systems: AGU Geophysical Monograph, v. 91, p. 85-114. Alt, J. C, 1999. Hydrothermal alteration and mineralisation of oceanic crust - mineralogy, geochemistry and processes: Reviews in Economic Geology, v. 8, p. 133-55. Alt, J. C, and Honnorez, J., 1984. Alteration of the upper oceanic crust, Deep Sea Drilling Project site 417 — mineralogy and chemistry: Contributions to Mineralogy and Petrology, v. 87, p. 149-69. Alt, J. C, Honnorez, J., Laverne, C, and Emmermann, R., 1986. Hydrothermal alteration of a 1-km section through the upper oceanic crust, Deep Sea Drilling Project hole 504B - mineralogy, chemistry, and evolution of seawater-basalt interactions: Journal of Geophysical Research, v. 91, p.10309-35. Alt-Epping, P., and Smith, L., 1997. Geochemical patterns associated with the onset of convective circulation at a sedimented ridge crest [abs.]: Geological Society of America Abstracts with Programs, v. 29, p. 74. Amstutz, G. C, and Park, W. C, 1967. Stylolites of diagenetic age and their role in the interpretation of the southern Illinois fluorspar deposits: Mineralium Deposita, v. 2, p. 44—53. Anderson, J. E. J., 1969. Development of snowflake texture in a welded tuff, Davis Mountain, Texas: Geological Society of America Bulletin, v. 80, p. 2075-80. Anonymous, 1997. Mineral analysis guide, ANALABS, Bentley, Western Australia, 41 p. Ashley, P. M., Dudley, R. J., Lesh, R. H., Marr, J. M., and Ryall, A. W, 1988. The Scuddles Cu-Zn prospect, an Archaean volcanogenic massive sulfide deposit, Golden Grove District, Western Australia: Economic Geology, v. 83, p. 918—51. Bailes, A. H., and Galley, A. G., 1999. Evolution of the Palaeoproterozoic Snow Lake arc assemblage and geodynamic setting for associated volcanic-hosted massive sulfide deposits, Flin Flon Belt, Manitoba, Canada: Canadian Journal of Earth Sciences, v. 36, p. 1789-805.

2 5 2 | REFERENCES Baillie, P. W., 1989. Stratigraphy, sedimentology and structural setting of the Cambrian Sticht Range Formation, western Tasmania: Tasmania Geological Survey Bulletin, v. 65, p. 1— 34. Baker, J. H., 1985. The petrology and geochemistry of 1.8—1.9 Ga granitic magmatism and related subseafloor hydrothermal alteration and ore-forming processes, W. Bergslagen, Sweden: GUA Papers of Geology, v. 1, 203 p. Banwell, C. J., Cooper, E. R., Thompson, G. E. K., and McCree, K. J., 1957. Physics of the New Zealand thermal area: Department of Scientific and Industrial Research Report, no. 123, 109 p. Baragar, W. R. A., Plant, A. G., Pringle, G. J., and Schau, M., 1977. Petrology and alteration of selected units of Mid-Atlantic Ridge basalts sampled from sites 332 and 335, Deep Sea Drilling Project: Canadian Journal of Earth Sciences, v. 14, p. 837-74. Baragar, W. R. A., Plant, A. G., Pringle, G. J., and Schau, M., 1979. Diagenetic and post-diagenetic changes in the composition of an Archaean pillow: Canadian Journal of Earth Sciences, v. 16, p. 2102-21. Barker, A. J., 1990. Introduction to metamorphic textures and microstructures: Glasgow, Blakie and Sons Ltd, 162 p. Barnes, H. L., ed., 1979. Geochemistry of hydrothermal ore deposits, 3rd edition: New York, USA, John Wiley and Sons, 798 p. Barrett, T. J., and MacLean, W. H., 1991. Chemical, mass, and oxygen isotope changes during extreme hydrothermal alteration of an Archaean rhyolite, Noranda, Quebec: Economic Geology, v. 86, p. 406-14. Barrett, T. J., and MacLean, W. H., 1994a. Chemostratigraphy and hydrothermal alteration in exploration for VHMS deposits in greenstones and younger volcanic rocks, in Lentz, D. R., ed., Alteration and alteration processes associated with ore forming systems: Geological Society of Canada Short Course Notes, v. 11, p. 433-67. Barrett, T. J., and MacLean, W. H., 1994b. Mass changes in hydrothermal alteration zones associated with VMS deposits of the Noranda area: Exploration and Mining Geology, v. 3, p. 131-60. Barrett, T. J., and MacLean, W. H., 1999. Volcanic sequences, lithogeochemistry and hydrothermal alteration in some bimodal volcanic-associated massive sulfide systems: Reviews in Economic Geology, v. 8, p. 101—31. Barrett, T. J., Jarvis, I., and Jarvis, K. E., 1990. Rare earth element geochemistry of massive sulfides-sulfates and gossans on the Southern Explorer Ridge: Geology, v. 18, p. 583-6. Barrett, T. J., Cattalani, S., and MacLean, W H., 1993. Volcanic lithogeochemistry and alteration at the Delbridge massive sulfide deposit, Noranda, Quebec: Journal of Geochemical Exploration, v. 48, p. 135-73. Barrett, T. J., MacLean, W. H., andTennant, S. C, 2001. Volcanic sequence and alteration at Parys Mountain volcanichosted massive sulfide deposit, Wales, United Kingdom — applications of immobile element lithogeochemistry: Economic Geology, v. 96, p. 1279—305. Barriga, F. J. A. S., and Fyfe, W. S., 1988. Giant pyritic base-metal deposits - the example of Feitais (Aljustrel, Portugal): Chemical Geology, v. 69, p. 331-43. Barriga, E J. A. S., and Kerrich, R., 1984. 18O-enriched volcanics and 18O-evolved marine water, Aljustrel, Iberian pyrite belt - transition from high to low Raleigh number convective regimes: Geochimica et Cosmochimica Acta, v. 48, p. 1021— 32. Barriga, F. J. A. S., Carvalho, D., and Oliveira, J. T, 1983. Carboniferous volcanogenic sulfide mineralisations in South Portugal, Iberian pyrite belt: Memorias dos Servicos Geologicos de Portugal, v. 29, p. 99—113.

Bateman, A. M., 1951. Economic mineral deposits: New York, USA, John Wiley and Sons, 916 p. Beams, S. D., Dronseika, E. V, and Doyle, M. G., 1998. The exploration history, geology and geochemistry of the polymetallic Highway-Reward deposit, Mt Windsor Subprovince [abs.], in Beams, S. D., ed. Economic geology of northeast Queensland, the 1998 perspective: Townsville, Australia, July 6-10, 1998, Geological Society of Australia Abstracts, v. 49 , p. 189-206. Beane, R. E., 1982. Hydrothermal alteration in silicate rocks southwestern North America, in Titley, S. R., ed., Advances in geology of porphyry copper deposits - southwestern North America: Tucson, USA, University of Arizona Press, p. 117-37. Beane, R. E., and Titley, S. R., 1981. Porphyry copper deposits - Part II, hydrothermal alteration and mineralisation: Economic Geology, 75 th anniversary volume, p. 235—69. Beaty, D. W, and Taylor, H. P., Jr., 1979. Oxygen-isotope geochemistry of the Abitibi greenstone belt, Ontario; evidence for seawater-rock interaction and implications regarding the isotopic composition and evolution of the ocean and oceanic crust [abs.]: Geological Society of America, Abstracts with Programs, v. 11, p. 386. Beaudoin, G., and Therrien, P., 1999. Isotopic fractionation calculator, Laboratoie de geochimie isotopique, Universite Laval, Quebec <www.ggl.ulaval.ca/personnel/beaudoin/ labo.html>. Bell, T. H., and Etheridge, M. A., 1973. Microstructure of mylonites and their descriptive terminology: Lithos, v. 6, p. 337-48. Berry, R. E, and Crawford, A. J., 1988. The tectonic significance of Cambrian allochthonous mafic-ultramafic complexes in Tasmania: Australian Journal of Earth Sciences, v. 35, p. 523-33. Berry, L. G., Mason, B., and Dietrich, R. V, 1983. Mineralogyconcepts, descriptions, determinations: San Francisco, USA, W H . Freeman and Co., 561 p. Berry, R. E, Huston, D. L, Stolz, A. J., Hill, A. P., Beams, S. D., Kuronen, U., and Taube, A., 1992. Stratigraphy, structure, and volcanic-hosted mineralisation of the Mount Windsor Subprovince, north Queensland, Australia: Economic Geology, v. 87, p. 739-63. Bettison-Varga, L., Varga, R. J., and Schiffman, P., 1992. Relation between ore-forming hydrothermal systems and extensional deformation in the Solea Graben spreading centre, Troodos Ophiolite, Cyprus: Geology, v. 20, p. 987-90. Bevins, R. E., and Robinson, D., 1992. Low-grade metamorphism: Geology Today, January—February, p. 23—7Bierlein, F. P., Arne, D. C, McKnight, S., Lu, J., Reeves, S., Besanko, J., Marek, J., and Cooke, D., 2000. Wall-rock petrology and geochemistry in alteration halos associated with mesothermal gold mineralisation, central Victoria, Australia: Economic Geology, v. 95, p. 283-311. Bierwirth, P., Huston, D., and Blewett, R. S., 2002. Hyperspectral mapping of mineral assemblages associated with gold mineralisation in the central Pilbara, Western Australia: Economic Geology, v. 97, p. 819—26. Bigger, S. E., and Hanson, R. E., 1992. Devitrification textures and related features in the Carlton Rhyolite in the Blue Creek Canyon area, Wichita Mountains, southwestern Okalahoma: Okalahoma Geological Notes, v. 52, p. 124—42. Bischoff J. L., and Dickson, F. W, 1975. Seawater-basalt interaction at 200cC and 500 bars: implications for origin of seafloor heavy metal deposits and regulation of seawater chemistry: Earth and Planetary Science Letters, v. 25, p. 385—97. Bischoff, J. L, and Seyfried, W E., Jr., 1978. Hydrothermal chemistry of seawater from 25°C to 350°C: American Journal of Science, v. 278, p. 838-60.

REFERENCES | 2 5 3

Bish, D. L., and Aronson, J. L., 1993. Paleogeothermal and paleohydrologic conditions in silicic tuff from Yucca mountains, Nevada: Clays and Clay Minerals, v. 41, p. 148— 61. Blatt, H., Middleton, G., and Murray, R., 1972. Origin of sedimentary rocks: New Jersey, USA, Prentice-Hall Inc., 634 p. Boda, S. P., 1991. The geology, structural setting and genesis of the Chester Mine, northwest Tasmania: Unpublished Honours thesis, Australian National University, 111 p. Bodon, S. B., and Cooke, D. R., 1998. Geochemical modelling of low temperature (5° to 150°C) seawater-andesite interaction — implications for regional alteration assemblages in VHMS districts: AMIRA/ARC Project P439, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 3, October 1996, p. 121-33. Bodon, S. B., and Valenta, R. K., 1995. Primary and tectonic features of the Currawong Zn-Cu-Pb(-Au) massive sulfide deposit, Benambra, Victoria — implications for ore genesis: Economic Geology, v. 90, p. 1694-721. Bohlke, J. K., Honnorez, J., and Honnorez-Guerstein, B. M., 1980. Alteration of basalts from Site 396B, Deep Sea Drilling Project — petrographic and mineralogical studies: Contributions to Mineralogy and Petrology, v. 73, p. 341-64. Boles, J. R., 1974. Structure, stratigraphy and petrology of mainly Triassic rocks, Hokonui Hills, Southland, New Zealand: New Zealand Journal of Geology and Geophysics, v. 17, p.337-74. Boles, J. R., 1977. Zeolites in low-grade metamorphic rocks, in Mumpton, F. A., ed., Mineralogy and geology of natural zeolites: Mineral Society of America Short Course Notes, v. 4, p. 103-35. Boles, J. R., 1982. Active albitization of plagioclase, Gulf Coast Tertiary: American Journal of Science, v. 282, p. 165-80. Boles, J. R., and Coombs, D. S., 1975. Mineral reactions in zeolitic Triassic tuff, Hokonui Hills, New Zealand: Geological Society of America Bulletin, v. 86, p. 163-73. Boles, J. R., and Coombs, D. S., 1977. Zeolite facies alteration of sandstone in the Southland Syncline, New Zealand: American Journal of Science, v. 277, p. 982—1012. Bonatti, E., 1965. Palagonite, hyaloclastites and alteration of volcanic glass in the ocean: Bulletin Volcanologique, v. 28, p. 257-70. Boulter, C. A., 1993. Comparison of RioTinto, Spain, and Guaymas Basin, Gulf of California - an explanation of a supergiant massive sulfide deposit in an ancient sill-sediment complex: Geology, v. 21, p. 801-4. Braithwaite, R. L., 1974. The geology and origin of the Rosebery ore deposit, Tasmania: Economic Geology, v. 69, p. 1086-101. Branney, M. J., and Sparks, R. S. J., 1990. Fiamme formed by diagenesis and burial compaction in soils and subaqueous sediments: Journal of the Geological Society of London, v. 147, p. 919-22. Branney, M. J., and Suthren, R. J., 1988. High-level peperitic sills in the English Lake District — distinction from block lavas and implications for Borrowdale Volcanic Group stratigraphy: Geological Journal, v. 23, p. 171-87. Brauhart, C. W., Groves, D. I., and Morant, P., 1998. Regional alteration systems associated with volcanogenic massive sulfide mineralisation at Panorama, Pilbara, Western Australia: Economic Geology, v. 93, p. 292-302. Brauhart, C. W., Huston, D. L., and Andrew, A. S., 2000. Oxygen isotope mapping in the Panorama VMS district, Pilbara Craton, Western Australia — applications to estimating temperatures of alteration and to exploration: Mineralium Deposita, v. 35, p. 727-40. Brauhart, C. W, Huston, D. L., Groves, D. I., Mikucki, E. J., and

Gardoll, S. J., 2001. Geochemical mass-transfer patterns as indicators of the architecture of a complete volcanic-hosted massive sulfide hydrothermal alteration system, Panorama District, Pilbara, western Australia: Economic Geology, v. 96, p. 1263-78. Brey, G., and Schmincke, H-U., 1980. Origin and diagenesis of the Roque Nublo Breccia, Gran Canaria (Canary Islands) - petrology of Roque Nublo volcanics, II: Bulletin Volcanologique, v. 43, p. 15—33. Browne, P. R. L., 1978. Hydrothermal alteration in active geothermal fields: Annual Review of Earth and Planetary Sciences, v. 6, p. 229-50. Burnham, C. W, 1979. Magmas and hydrothermal fluids, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits, 3 rd edition: New York, USA, John Wiley and Sons, p. 71—136. Burnham, C. W, and Ohmoto, H., 1980. Late-stage processes of felsic magmatism: Mining Geology, v. 8, p. 1-11. Busby-Spera, C. J., and White, J. D. L., 1987. Variation in peperite textures associated with differing host-sediment properties: Bulletin of Volcanology, v. 49, p. 765-75. Butler, B. S., and Burbank, W S., 1929. The copper deposits of Michigan: United States Geological Survey Professional Paper 144, 238 p. Butterfield, D. A., 2000. Deep ocean hydrothermal vents, in Sigurdsson, H., ed., Encyclopaedia of volcanoes, San Diego, California, USA, Academic Press, p. 857-75. Cagatay, M. N., 1993. Hydrothermal alteration associated with volcanogenic massive sulfide deposits — examples from Turkey: Economic Geology, v. 88, p. 606-21. Callaghan, T, 1998. Geology and alteration of the Mt Julia deposit, Henty gold mine, Tasmania: Unpublished MSc thesis, University of Tasmania, 78 p. Callaghan, T, 2001. Geology and host-rock alteration of the Henty and Mt Julia gold deposits, western Tasmania: Economic Geology, v. 96, p. 1073-88. Campana, B., and King, D., 1963. Palaeozoic tectonism, sedimentation and mineralisation in west Tasmania: Journal of Geological Society of Australia, v. 10, p. 1-54. Campbell, I. H., Franklin, J. M., Gorton, M. P., Hart, T. R., and Scott, S. D., 1981. The role of subvolcanic sills in the generation of massive sulfide deposits: Economic Geology, v. 76, p. 2248-53. Cann, J. R., 1969. Spilites from the Carlsberg Ridge, Indian Ocean: Journal of Petrology, v. 10, p. 1—19. Carmichael, I. S. E., 1979. Glass and glassy rocks, in Yoder, H. S. J., ed., The evolution of igneous rocks: Princeton, USA, Princeton University Press, p. 233-44. Carozzi, A. V, 1960. Microscopic sedimentary petrography: New York, USA, John Wiley and Sons, 485 p. Carson, D. J. T, 1973. Petrography, chemistry, age and emplacement of the plutonic rocks of Vancouver Island: Canada Geological Survey Paper, v. 72—44, 69 p. Cartwright, I., 1999. Regional oxygen isotope zonation at Broken Hill, New South Wales, Australia - large-scale fluid flow and implications for Pb-Zn-Ag mineralisation: Economic Geology, v. 94, p. 357-73. Carvalho, C. M. N., and Barriga, E J. A. S., 2000. Preliminary report on the detailed mineralogical study at the alteration zones surrounding the Feitais VMS ore body (Aljustrel, Portugal, IPB) [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits, Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 19-21. Carvalho, D., Barriga, E J. A. S., and Munha, J., 1999. The Iberian pyrite belt of Portugal and Spain — examples of bimodal siliciclastic systems: Reviews in Economic Geology, v. 8, p. 385-418.

2 5 4 | REFERENCES

Cas, R. A. E, and Wright, J. V., 1987. Volcanic successions-modern and ancient - a geological approach to processes, products and successions: Oxford, UK, Unwin Hyman Inc., 528 p. Casey, W. H., and Bunker, B., 1990. Leaching of mineral and glass surfaces during dissolution: Reviews in Mineralogy, v. 23, p. 397-426. Cathelineau, M., and Nieva, D., 1985. A chlorite solid solution geothermometer — the Los Azufres (Mexico) geothermal system: Contributions to Mineralogy and Petrology, v. 91, p. 235-44. Cathles, L. M., 1977. An analysis of the cooling of intrusives by ground-water convection, which includes boiling: Economic Geology, v. 72, p. 804-26. Cathles, L. M., 1981. Fluid flow and genesis of hydrothermal ore deposits: Economic Geology, 75 th anniversary volume, p. 424-57. Cathles, L. M., 1983. An analysis of the hydrothermal system responsible for massive sulfide deposition in the Hokuroku Basin of Japan: Economic Geology Monographs, v. 5, p. 439-87. Cathles, L. M., 1993. Oxygen isotope alteration in the Noranda mining district, Abitibi greenstone belt, Quebec: Economic Geology, v. 88, p. 1483-511. Cerling, T. E., Brown, E H., and Bowman, J. R., 1985- Lowtemperature alteration of volcanic glass — hydration, Na, K, 18 O and Ar mobility: Chemical Geology, v. 52, p. 281-93. Christiansen, R. L., and Lipman, P. W., 1966. Emplacement and thermal history of a rhyolite lava flow near Fortymile Canyon, southern Nevada: Geological Society of America Bulletin, v. 77, p. 671-84. Coombs, D. S., 1954. The nature and alteration of some Triassic sediments from Southland, New Zealand: Transactions of the Royal Society of New Zealand, v. 82, p. 65-109. Coombs, D. S., 1974. On the mineral facies of spilitic rocks and their genesis, in Amstutz, G. C, ed., Spilites and spilitic rocks: New York, USA, Springer-Verlag, p. 373-86. Coombs, D. S., Ellis, A. J., Fyfe, W. S., and Taylor, A. M., 1959. The zeolite facies, with comments on the interpretation of hydrothermal syntheses: Geochimica et Cosmochimica Acta, v. 17, p. 53-107. Corbett, K. D., 1979. Stratigraphy, correlation and evolution of the Mt Read Volcanics in the Queenstown, Jukes-Darwin and Mt Sedgwick areas: Geological Survey Tasmania Bulletin, v. 58, 74 p. Corbett, K. D., 1981. Stratigraphy and mineralisation in the Mt Read Volcanics, western Tasmania: Economic Geology, v. 76, p. 209-30. Corbett, K. D., 1986. The geological setting of mineralisation in the Mt Read Volcanics, in Large, R. R., ed., The Mount Read Volcanics and associated ore deposits: Burnie, Australia, November 20-21, 1986, Geological Society of Australia, Tasmanian Division, Proceedings, p. 1—10. Corbett, K. D., 1989. Stratigraphy, palaeogeography and geochemistry of the Mt Read Volcanics, in Burnett, C. E, and Martin, E. L., eds., Geology and mineral resources of Tasmania: Geological Society of Australia Special Publication, v. 15, p. 86-106. Corbett, K. D., 1992. Stratigraphic-volcanic setting of massive sulfide deposits in the Cambrian Mount Read Volcanics, Tasmania: Economic Geology, v. 87, p. 564—86. Corbett, K. D., 1994. Stratigraphic mapping, Tyennan connections, Cambrian orogenies, the Arthur Lineament, and the tectonic context of the Mount Read Volcanics - a fresh look at western Tasmania [abs.], in Cooke, D. R., and Kitto, P. A., eds., Contentious issues in Tasmanian geology, a symposium: Hobart, Australia, November 3-4, 1994, Geological Society of Australia Abstracts, v. 39, p. 35-7.

Corbett, K. D., 2001. New mapping and interpretations of the Mount Lyell mining district, Tasmania - a large hybrid CuAu system with an exhalative Pb-Zn top: Economic Geology, v. 96, p. 1089-122. Corbett, K. D., 2002. Western Tasmanian regional minerals program, Mount Read Volcanics compilation - updating the geology of the Mount Read Volcanics belt: Tasmanian Geological Survey Record, no. 19, 31 p. Corbett, K. D., and Komyshan, P., 1989. Geology of the Hellyer-Mt Charter area: Mt Read Volcanics Project Geological Report, Tasmania Department of Mines, no. 1, 48 p. Corbett, K. D., and Lees, T. C, 1987. Stratigraphic and structural relationships and evidence for Cambrian deformation at the western margin of the Mt Read Volcanics, Tasmania, Australia: Australian Journal of Earth Sciences, v. 34, p. 4 5 67. Corbett, K. D., and Solomon, M., 1989. Cambrian Mt Read Volcanics and associated mineral deposits, in Burrett, C. E, and Martin, E. L., eds., Geology and mineral resources of Tasmania: Geological Society of Australia Special Publication, v. 15, p. 84-153. Corbett, K. D., Reid, K. O., Corbett, E. B., Green, G. R., Wells, K., and Sheppard, N. W, 1974. The Mount Read Volcanics and Cambro-Ordovician relationships at Queenstown, Tasmania: Journal of Geological Society of Australia, v. 21, p. 173-86. Costa, U. R., Barnett, R. L., and Kerrich, R., 1983. The Mattagami Lake mine Archaean Zn-Cu sulfide deposit, Quebec - hydrothermal co-precipitation of talc and sulfides in a seafloor brine pool — evidence from geochemistry, lsO/i6O and mineral chemistry: Economic Geology, v. 78, p. 1144— 203. Costa, I. A., Barriga, E J. A. S., and Fouquet, Y., 1995. Active hydrothermal sites and ancient massive sulfides - Lucky Strike versus Iberian pyrite belt [abs.]: Terra Abstracts, v. 7, p. 212. Coutts, B. P. C, 1990. The geology, geochemistry and hydrothermal alteration of the Hollway Andesite, western Tasmania: Unpublished Honours thesis, University of Tasmania, 76 p. Cox, S. F., 1981. The stratigraphic and structural setting of the Mt Lyell volcanic-hosted sulfide deposits: Economic Geology, v. 76, p. 231-45. Craig, M. A., 2001. Regolith mapping for geochemical exploration in the Yilgarn Craton, Western Australia: Geochemistry - Exploration, Environment, Analysis, v. 1, p. 383-90. Craig, J. R., and Vaughan, D. J., 1981. Ore microscopy and ore petrology: New York, USA, John Wiley and Sons, 406 p. Crawford, A. J., and Berry, R. F., 1992. Tectonic implications of Late Proterozoic-Early Palaeozoic igneous rock associations in western Tasmania: Tectonophysics, v. 214, p. 37-56. Crawford, A. J., Corbett, K. D., and Everard, J. L, 1992. Geochemistry of the Cambrian volcanic-hosted massive sulfide-rich Mount Read Volcanics, Tasmania, and some tectonic implications: Economic Geology, v. 87, p. 597619. Creasey, S. C, 1959. Some phase relations in hydro thermally altered rocks of porphyry copper deposits: Economic Geology, v. 54, p. 351-73. CSIRO, 2002. Automated mineralogical logging of drill core, chips and powders: CSIRO/AMIRA Project P685, <www.syd.dem. csiro.au/research/MMTG/P685%20Core%20Logging/ P685.htm>. Cudahy, T. J., Okada, K., and Brauhart, C, 2000. Targeting VMSstyle Zn mineralisation at Panorama, Australia, using airborne hyperspectral VNIR-SW1R Hymap data: 14th International Conference on Applied Geologic Remote Sensing: Ann Arbor, Las Vegas, USA, November 6-8, 2000,

REFERENCES | 2 5 5

Proceedings, v. 14, p. 395-402. Dana, E. S., 1957. A textbook of mineralogy, with an extended treatise on crystallography and physical mineralogy, (4th edition, revised and enlarged by W. E. Ford): New York, USA, Wiley, 851 p. Date, J., Watanabe, Y., Iwaya, S., and Horiuchi, M., 1979. A consideration on the alteration of dacite below the Fukazawa ore deposits, Fukazawa Mine, Akita Prefecture: Mining Geology, v. 29, p. 187-96. Date, J., Watanabe, Y., and Saeki, Y., 1983. Zonal alteration around the Fukazawa Kuroko deposits, Akita Prefecture, northern Japan: Economic Geology Monographs, v. 5, p. 365-86. Davidson, P., 1998. The Murchison Granite: Unpublished Honours thesis, University of Tasmania, 128 p. Davidson, G. J., and Kitto, P., 1997. The sulfur isotope signature of Cambrian faults in the Mt Read Volcanic Belt and their implications for sulfur enrichment in submarine volcanosedimentary rift basins: AMIRA/ARC Project P291A, Centre for Ore Deposit Research, University of Tasmania, Unpublished Final Report, March 1997, p. 121-69. Davidson, G. J., Garven, G., Kitto, P. A., and Berry, R., 2000. Geochemically discrete fluid bodies formed by convection at the heated edge of porous seafloor aquifers [abs.], in Woodhead, J. D., Hergt, J. M., and Noble, W P., eds., Beyond 2000, new frontiers in isotope geoscience: Lome, Australia, January 30-February 4, 2000, Abstracts and Proceedings, p. 39-40. Davies, J. F., Whitehead, R. E., Huang, J., and Nawaratne, S., 1990. A comparison of progressive hydrothermal carbonate alteration in Archaean metabasalts and metaperidotites: Mineralium Deposita, v. 25, p. 65-72. Davis, L. W, 1990. Silver-lead-zinc-copper mineralisation in the Captains Flat-Goulburn synclinorial zone and the Hill End synclinorial zone: Australasian Institute of Mining and Metallurgy Monograph, v. 14, p. 1375-84. Davis, B. K., andMcPhie, J., 1996. Spherulites, quench fractures and relict perlite in a Late Devonian rhyolite dyke, Queensland, Australia: Journal of Volcanology and Geothermal Research, v. 71, p. 1-11. deRonde, C.E.J., 1995. Fluid chemistry and isotopic characteristics of seafloor hydrothermal systems and associated VMS deposits: potential for magmatic contributions, in Thompson, J. F. H., ed., Magmas, fluids and ore deposits: Victoria, Canada, Mineralogical Association of Canada, v. 23, p. 479-509. de Ronde, C. E. J., de Wit, M. J., and Spooner, E. T. C, 1994. Early Archaean (>3.2 Ga) Fe-oxide-rich, hydrothermal discharge vents in the Barberton greenstone belt, South Africa: Geological Social of America Bulletin, v. 106, p. 86-104. de Rosen-Spence, A. F., Provost, G., Dimroth, E., Gochnauer, K., and Owen, V, 1980. Archaean subaqueous felsic flows, Rouyn-Noranda, Quebec, Canada, and their Quaternary equivalents: Precambrian Research, v. 12, p. 43-77. Deelman, J. C, 1975. 'Pressure solution' re-examined: Neues Jahrbuch fuer Geologie und Palaeontologie Abhandlungen, v. 149, p. 384-97. Deelman, J. C, 1976. Lithification analysis-experimental observations: Geologishe Rundschau, v. 65, p. 1055—78. Deer, W. A., Howie, R. A., and Zussman, J., 1966. An introduction to the rock forming minerals: London, UK, Longman Scientific and Technical, 528 p. Denniss, A. M., Colman, T. B., Cooper, D. C, Hatton, W. A., and Shaw, M. H., 1999. The combined use of PIMA and Vulcan technology for mineral deposit evaluation at the Parys Mountain Mine, Anglesea, UK: 13* International Conference on Applied Geologic Remote Sensing, Vancouver, Canada, March 1-3, 1999, Proceedings, p. 125-32. Destrigneville, C, Schott, J., Caristan, Y, and Agrinier, P., 1991.

Evidence of an early alteration process driven by magmatic fluid in Mururoa Volcano: Earth and Planetary Science Letters, v. 104, p. 119-39. Dickinson, W. R., 1962. Petrology and diagenesis of Jurassic andesitic strata in central Oregon: American Journal of Science, v. 260, p. 481-500. Dickson, F. W, and Potter, J. M., 1982. Rock-brine chemical interactions: Electric Power Research Institute Project 6532, Final Report, AP-2258, 89 p. Dill, H. G., Bosse, H. R., Henning, K. H., Fricke, A., and Ahrendt, H., 1997. Mineralogical and chemical variations in hypogene and supergene kaolin deposits in a mobile fold belt in the Central Andes of northwestern Peru: Mineralium Deposita, v. 32, p. 149-63. Dimroth, E., and Lichtblau, A. P., 1979. Metamorphic evolution of Archaean hyaloclastites, Noranda area - Part 1, comparison of Archaean and Cenozoic seafloor metamorphism: Canadian Journal of Earth Science, v. 16, p. 1315—40. Downs, R. C, 1993. Syndepositional fault controls on the Hellyer volcanic-hosted massive sulfide deposit: Unpublished MSc thesis, University of Tasmania, 63 p. Doyle, M. G., 1990. The geology, mineralisation and alteration of the Jukes Proprietary prospect, Tasmania: Unpublished Honours thesis, University of Tasmania, 114 p. Doyle, M. G., 1997. A Cambro-Ordovician submarine volcanic succession hosting massive sulfide mineralisation: Mount Windsor Subprovince, Queensland: Unpublished PhD thesis, University of Tasmania, 264 p. Doyle, M. G., 2001. Volcanic influences on hydrothermal and diagenetic alteration - evidence from Highway-Reward, Mount Windsor Subprovince, Australia: Economic Geology, v. 96, p. 1133-48. Doyle, M. G., and Allen, R. L., 2003. Subseafloor replacement in volcanic-hosted massive sulfide deposits: Ore Geology Reviews, v. 23, p. 183-222. Doyle, M. G., and Huston, D. L., 1999. The subseafloor replacement origin of the Ordovician Highway-Reward volcanic-associated massive sulfide deposit, Mount Windsor Subprovince, Australia: Economic Geology, v. 94, p. 82543. Doyle, M. G., and McPhie, J., 2000. Facies architecture of a silicic intrusion-dominated volcanic centre at HighwayReward, Queensland, Australia: Journal of Volcanology and Geothermal Research, v. 99, p. 79-96. Drysdale, D. J., 1991. Perlitic texture and other fracture patterns produced by hydration of glassy rocks: Department of Geology, University of Queensland Scientific Paper, v. 12, no. 3, p. 278-85. Dudas, F. O., Campbell, I. H., and Gorton, M. P., 1983. Geochemistry of igneous rocks in the Hokuroku District, northern Japan: Economic Geology Monographs, v. 5, p. 115-33. Dugdale, J. S., 1992. Lithostratigraphy of the White Spur area, western Tasmania: Unpublished Honours thesis, University of Tasmania, 93 p. Duhig, N. C, Stolz, J., Davidson, G. J., and Large, R. R., 1992. Cambrian microbial and silica gel textures in silica-iron exhalites from the Mount Windsor volcanic belt, Australia — their petrography, chemistry, and origin: Economic Geology, v. 87, p. 764-84. Eastoe, C. J., Solomon, M., and Walshe, J. L., 1987. Districtscale alteration associated with massive sulfide deposits in the Mount Read Volcanics, western Tasmania: Economic Geology, v. 82, p. 1239-58. Eggleton, R. A., and Keller, J., 1982. The palagonitization of limburgite glass - a TEM study: Neues Jahrbuch fur Mineralogie Monatshefte, p. 321-36. Eilu, P., Mikucki, E. J., and Groves, D. I., 1997. Wallrock alteration

2 5 6 | REFERENCES and primary geochemical dispersion in lode-gold exploration: 4 th biennial meeting of the Society for Geology Applied to Mineral Deposits, Turku, Finland, August 11-13, 1997, Short course notes, p. 65. Einsele, G., 1985. Basaltic sill-sediment complexes in young spreading centres — genesis and significance: Geology, v. 13, p. 249-52. Einsele, G., Gieskes, J. M., Curray, J., Moore, D. M., Aguayo, E., Aubry, M., Fornari, D., Guerrero, J., Kastner, M., Kelts, K., Lyle, M., Matoba, Y., Molina-Cruz, A., Niemitz, J., Rueda, J., Saunders, A., Schrader, H., Simoneit, B., and Vacquier, V., 1980. Intrusion of basaltic sill into highly porous sediments and resulting hydrothermal activity: Nature, v. 283, p. 441—

5. Eldridge, C. S., Barton, P. B. J., and Ohmoto, H., 1983. Mineral textures and their bearing on formation of the Kuroko ore bodies: Economic Geology Monographs, v. 5, p. 241-81. Elliott-Meadows, S. R., and Appleyard, E. C., 1991. The alteration geochemistry and petrology of the Lar Cu-Zn deposit, Lynn Lake area, Manitoba, Canada: Economic Geology, v. 86, p.486-505. Ellis, A. J., and Mahon, W. A. J., 1964. Natural hydrothermal systems and experimental hot-water/rock interactions: Geochimica et Cosmochimica Acta, v. 28, p. 1323—57. Ewart, A., 1979. A review of the mineralogy and chemistry of Tertiary-Recent dacitic, latitic, rhyolitic, and related salic volcanic rocks, in Barker, E, ed., Trondhjemites, dacites, and related rocks: Amsterdam, Netherlands, Elsevier Scientific Publication, p. 13—121. Faure, G., 1986. Principles of isotope geology: New York, USA, John Wiley and Sons, 589 p. Finlow-Bates, T, and Strumpfl, E. E, 1981. The behaviour of socalled immobile elements in hydrothermally altered rocks associated with volcanogenic submarine-exhalative ore deposits: Mineralium Deposita, v. 16, p. 319-28. Fisher, R. V, I960. Classification of volcanic breccias: Geological Society of America Bulletin, v. 71, p. 973-82. Fisher, R. V., 1961. Proposed classification of volcaniclastic sediments and rocks: Geological Society of America Bulletin, v. 72, p.1409-14. Fisher, R. V., and Schmincke, H-U., 1984. Pyroclastic rocks: New York, USA, Springer-Verlag, 472 p. Fiske, R. S., 1969. Recognition and significance of pumice in marine pyroclastic rocks: Geological Society of America Bulletin, v. 80, p. 1-8. Fitzgerald, F. G., 1974. The Primrose Pyroclastics in the HerculesWhite Spur area, western Tasmania: Unpublished Honours thesis, University of Tasmania, 188 p. Floyd, P. A., and Winchester, J. A., 1978. Identification and discrimination of altered and metamorphosed volcanic rocks using immobile elements: Chemical Geology, v. 21, p. 2 9 1 306. Folk, R. L., 1965. Some aspects of recrystallisation in ancient limestones: Society of Economic Palaeontologists and Mineralogists Special Publication, v. 13, p. 14—48. Fournier, R. O., 1985. The behaviour of silica in hydrothermal solutions: Reviews in Economic Geology, v. 2, p. 45-61. Fournier, R. O., and Potter, R. W, II, 1982. An equation correlating the solubility of quartz in water from 20° to 900°C at pressures up to 10,000 bars: Geochimica et Cosmochimica Acta, v. 46, p. 1969-73. Fowler, A. D., Jensen, L. S., and Peloquin, S. A., 1987. Varioles in Archaean basalts — products of spherulitic crystallization: The Canadian Mineralogist, v. 25, p. 275-89. Franklin, J. M., Kasarda, J., and Poulsen, K. H., 1975. Petrology and chemistry of the alteration zone of the Mattabi massive sulfide deposit: Economic Geology, v. 70, p. 63—79.

Franklin, J. M., Gibb, W, Poulsen, K. H., and Severin, P., 1977. Archean metallogeny and stratigraphy of the South Sturgeon Lake area: Ontario, Canada, Institute Lake Superior Geology, 73 p. Franklin, J. M., Sangster, D. M., and Lydon, J. W, 1981. Volcanicassociated massive sulfide deposits: Economic Geology, 75 th anniversary volume, p. 485—627. Franklin, J. M., Sangster, D. E, Roscoe, S. M., and Loveridge, W D., 1983. Lead isotope studies in Superior and Southern provinces: Geological Survey of Canada Bulletin, v. 351, 60 p. Friedman, I., and Long, W, 1984. Volcanic glasses — their origins and alteration processes: Journal of Non-Crystalline Solids, v. 67, p. 127-33. Friedman, I., and Smith, R. L., 1958. The deuterium content of water in some volcanic glasses: Geochimica et Cosmochimica Acta, v. 15, p. 218-28. Friedman, I., Smith, R. L., and Long, W. D., 1966. Hydration of natural glass and formation of perlite: Geological Society of America Bulletin, v. 77, p. 323-8. Fritsch, K. V, and Reiss, W, 1868. Geologische Beschreibung der Insel Teneriffa: Winterthur, Verlag von Wurster and Co., 494 p. Fuller, R. E., 1938. Deuteric alteration controlled by the jointing of lavas: American Journal of Science, v. 35, p. 161—71. Furbish, W. J., and Schrader, E. L., 1980. Secondary minerals of deuteric and diagenetic origin filling voids and encrusting surfaces on basalts from Deep Sea Drilling Project Leg 54: Proceedings of the Ocean Drilling Program, Initial Reports, v. 54, p. 807-18. Furnes, H., 1975. Experimental palagonitization of basaltic glasses of varied composition: Contributions to Mineralogy and Petrology, v. 50, p. 105-13. Fyfe, W. S., Turner, F. J., and Verhoogen, J., 1958. Metamorphic reactions and metamorphic facies: Geological Society of America Memoir, v. 73, 259 p. Galley, A. G., 1993. Characteristics of semi-conformable alteration zones associated with volcanogenic massive sulfide districts: Journal of Geochemical Exploration, v. 48, p. 175—99. Galley, A. G., 1995. Target vectoring using lithogeochemistry applications to the exploration for volcanic-hosted massive sulfide deposits: CIM Bulletin, v. 88, p. 15-27. Galley, A. G., 2003. Composite synvolcanic intrusions associated with Precambrian VMS-related hydrothermal systems: Mineralium Deposita, v. 38, p. 443-73. Galley, A. G., Bailes, A., Hannington, M., Hoik, G., Katsube, J., Paquette, F., Paradis, S. J., Santaguida, F., Taylor, B., and Hillary, B., 2002. Database for CAMIRO Project 94E07 - inter-relationships between subvolcanic intrusions, largescale alteration zones and VMS deposits: Geological Survey of Canada Open File 4431 (CD). Galloway, J. N., 1979. Alteration of trace metal geochemical cycles due to the marine discharge of wastewater: Geochimica et Cosmochimica Acta, v. 38, p. 207—18. Garwin, S., 2002. The geologic setting of intrusion-related hydrothermal systems near the Batu Hijau porphyry coppergold deposit, Sumbawa, Indonesia: SEG Special Publication, v. 9, p. 333-66. Gary, M., McAfee, R., Jr., and Wolf, C. L, eds., 1974. Glossary of geology: Washington, D.C., American Geological Institution, 805 p. Geiskes, J. M., Kastner, M., Einsele, G., Kelts, K., and Niemitz, J., 1982. Hydrothermal activity in the Guaymas Basin, Gulf of California - a synthesis: Proceedings of the Ocean Drilling Program, Initial Reports, v. 64, p. 1159-68. Gemmell, J. B., and Fulton, R., 2001. Geology, genesis, and exploration implications of the footwall and hanging-

REFERENCES | 2 5 7 wall alteration associated with the Hellyer volcanic-hosted massive sulfide deposit, Tasmania, Australia: Economic Geology, v. 96, p. 1003-35. Gemmell, J. B., and Herrmann, W, 2001. Preface to a special issue devoted to alteration associated with volcanic-hosted massive sulfide deposits, and its exploration significance: Economic Geology, v. 96, p. 909-12. Gemmell, J. B., and Large, R. R., 1992. Stringer system and alteration zones underlying the Hellyer volcanogenic massive sulfide deposit, Tasmania, Australia: Economic Geology, v. 87, p. 620-49. Gemmell, J. B., and Large, R. R., 1993. Evolution of a VHMS hydrothermal system, Hellyer deposit, Tasmania, Australia — sulphur isotope evidence, in Ishihara, S., Urabe, T, and Ohmoto, H., eds., Mineral resources symposia: Tokyo, Japan, August 24, 1992, Society of Resource Geologists of Japan Special Issue, v. 17, p. 108-19. Gemmell, J. B., Large, R. R., and Khin Zaw, 1998. Palaeozoic volcanic-hosted massive sulfide deposits: AGSO Journal of Australian Geology and Geophysics, v. 17, p. 129—37. Ghiara, M. R., Franco, E., Petti, C, Stanzione, D., and Valentino, G. M., 1993. Hydrothermal interaction between basaltic glass, de-ionized water and seawater: Chemical Geology, v. 104, p. 125-38. Gibson, H. L., 1989. The mine sequence of the Central Noranda Volcanic Complex - geology alteration, massive sulfide deposits and volcanological reconstruction: Unpublished PhD thesis, Carleton University, 715 p. Gibson, H. L., and Watkinson, D. H., 1990. Volcanogenic massive sulfide deposits of the Noranda cauldron and shield volcano, Quebec: Canadian Institute of Mining and Metallurgy Special Issue, v. 43, p. 119-32. Gibson, H. L, Watkinson, D. H., and Comba, C. D. A., 1983. Silicification — hydrothermal alteration in an Archaean geothermal system within the Amulet Rhyolite Formation, Noranda, Quebec: Economic Geology, v. 78, p. 954-71. Gibson, H. L., Morton, R. L., Hudak, G. J., and Harper, G. D., 1999. Submarine volcanic processes, deposits, and environments favourable for the location of volcanic-associated massive sulfide deposits: Reviews in Economic Geology, v. 8, p. 1351. Gibson, H. L., Santaguida, F., Paquette-Mihalasky, F. I., and Watkinson, D. H., 2000. Evolution of regional semiconformable alteration assemblages within an Archaean subseafloor hydrothermal system, and relationship to VMS deposits, at Noranda, Quebec, Canada [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits: Hobart, Australia, November 1619, 2000, CODES Special Publication, v. 3, p. 61-3. Gifkins, C. C, 2001. Submarine volcanism and alteration in the Cambrian, northern Central Volcanic Complex, western Tasmania: Unpublished PhD thesis, University of Tasmania, 239 p. Gifkins, C. C, and Allen, R. L., 2001. Textural and chemical characteristics of diagenetic and hydrothermal alteration in glassy volcanic rocks - examples from the Mount Read Volcanics, Tasmania: Economic Geology, v. 96, p. 9731002. Gifkins, C. C, and Allen, R. L., 2002. Voluminous, submarine, intracaldera pumice breccia generated by explosive eruptions - the Cambrian Mount Black and Kershaw Pumice Formations, western Tasmania [abs.], in Chapman Conference — Explosive subaqueous volcanism: Dunedin, New Zealand, January 21-25, 2002, Abstracts Volume, p. 26. Gifkins, C. C, and Kimber, B., 2004. Revolutionising our understanding of an ancient volcanic succession — an

interactive map of the Mount Read Volcanics, western Tasmania [abs.], in McPhie, J., and McGoldrick, P. J., eds., Dynamic earth — past, present and future: Hobart, Australia, February 8-13, 2004, Geological Society of Australia Abstracts, v. 73: p. 81. Gifkins, C. C, McPhie, J., and Allen, R. L, 2002. Pumiceous rhyolitic peperite in ancient submarine volcanic successions: Journal of Volcanology and Geothermal Research, v. 114, p. 181-203. Gifkins, C. C, Allen, R. L., and McPhie, }., in press. Apparent welding textures in altered pumice-rich rocks: Journal of Volcanology and Geothermal Research. Giggenbach, W. E, 1997. The origin and evolution of fluids in magmatic hydrothermal systems, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits: New York, USA, John Wiley and Sons, p. 737-96. Gillis, K. M., and Robinson, P. T, 1990. Patterns and processes of alteration in the lavas and dykes of the Troodos Ophiolite, Cyprus: Journal of Geophysical Research, v. 95, p. 2152348. Gillis, K. M., Thompson, G., and Kelley, D. S., 1993. A view of the lower crustal component of hydrothermal systems at the Mid-Atlantic Ridge: Journal of Geophysical Research, v. 98, p. 19597-619. Gilmour, P., 1965. The origin of the massive sulfide mineralisation in the Noranda district, northwestern Quebec: Proceedings of the Geological Association of Canada, v. 16, p. 63-81. Goldfarb, M. S., Converse, D. R., Holland, H. D., and Edmond, J. M., 1983. The genesis of hot spring deposits on the East Pacific Rise, 21° N: Economic Geology Monographs, v. 5, p. 184-97. Goodfellow, W D., and Franklin, J. M., 1993. Geology, mineralogy, and chemistry of sediment-hosted clastic massive sulfides in shallow cores, Middle Valley, northern Juan de Fuca Ridge: Economic Geology, v. 88, p. 2033-64. Goodfellow, W. D., and McCutcheon, S. R., 2003. Geologic and genetic attributes of volcanic sediment-hosted massive sulfide deposits of the Bathurst mining camp, northern New Brunswick - a synthesis: Economic Geology Monographs, v. 11, p. 245-301. Grant, J. A., 1986. The isocon diagram — a simple solution to Gresens' equation for metasomatic alteration: Economic Geology, v. 81, p. 1976-82. Grapes, R. H., 1976. Low-grade alteration of basic rocks in Hokkaido, Japan: Journal of the Faculty of Science, Hokkaido University, v. 17, p. 373-85. Green, G. R., and Iliff, G. D., 1989. Rosebery, in Burrett, C. E, and Martin, E. L., eds., Geology and mineral resources of Tasmania: Geological Society of Australia Special Publication, v. 15, p. 132-7. Green, G. R., andTaheri, J., 1992. Stable isotopes and geochemistry as exploration indicators, in Baillie, P., Dix, M., and Richardson, R., eds., Tasmania — an island of potential: Geological Survey Bulletin, v. 70, p. 84-91. Green, G. R., Solomon, M., andWalshe, J. L., 1981. The formation of the volcanic-hosted massive sulfide deposit at Rosebery, Tasmania: Economic Geology, v. 76, p. 304-38. Green, G. R., Ohmoto, H., Date, J., and Takahashi, T, 1983. Whole-rock oxygen isotope distribution in the FukuzawaKosaka area, Hokuroku District, Japan, and its potential application to mineral exploration: Economic Geology Monographs, v. 5, p. 395-411. Gregory, R. T, and Taylor, H. P., Jr., 1981. An oxygen isotope profile in a section of Cretaceous oceanic crust, Samail Ophiolite, Oman - evidence for d18O buffering of the oceans by deep (>5 km) seawater-hydrothermal circulation at mid-ocean ridges: Journal of Geophysical Research, v. 86, p. 2737—55.

2 5 8 | REFERENCES Gregory, P. W., Hartley, J. S., and Wills, K. J. A., 1990. Thalanga zinc-lead-copper-silver deposit, in Hughes, F. E., ed., Geology of the mineral deposits of Australia and Papua New Guinea: Australasian Institute of Mining and Metallurgy Monograph, v. 14, p. 1527-37. Gresens, R. L., 1967. Composition-volume relationships of metasomatism: Chemical Geology, v. 2, p. 47-65. Gramme, C. S., Wright, T. L, and Peck, D. L., 1969. Magnetic properties and oxidation of iron-titanium oxide minerals in Alae and Makaoputi lava lakes, Hawaii: Journal of Geophysical Research, v. 74, p. 5277-93. Guilbert, J. M., and Park, C. F. J., 1986. The geology of ore deposits: New York, USA, W.H. Freeman and Co., 985 p. Gulson, B. L., and Porritt, P. M., 1987. Base metal exploration of the Mount Read Volcanics, western Tasmania — Part II, lead isotope signatures and genetic implications: Economic Geology, v. 82, p. 291-307. Gustafson, L. B., and Hunt, J. P., 1975. The porphyry copper deposit at El Salvador, Chile: Economic Geology, v. 70, p. 857-912. Hajash, A., 1975. Hydrothermal processes along mid-ocean ridges an experimental investigation: Contributions to Mineralogy and Petrology, v. 53, p. 205—26. Hajash, A., 1977. Experimental seawater/basalt interactions effect of water/rock ratio and temperature gradient [abs.]: Geological Society of America Abstracts with Programs, v. 9, p. 1002. Hajash, A., and Archer, P., 1980. Experimental seawater/basalt interactions - effects of cooling: Contributions to Mineralogy and Petrology, v. 75, p. 1-13. Hajash, A., and Chandler, G. W, 1981. An experimental investigation of high-temperature interactions between seawater and rhyolite, andesite, basalt and peridotite: Contributions to Mineralogy and Petrology, v. 78, p. 240-54. Halley, S. W, and Roberts, R. H., 1997. Henty - a shallow water gold-rich volcanogenic massive sulfide deposit in western Tasmania: Economic Geology, v. 92, p. 438—47. Hannington, M. D., and Herzig, P. M., 2000. Submarine epithermal deposits and the VMS-epithermal transition — a new exploration target [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits: Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 75-7. Hannington, M. D., and Scott, S. D., 1989. Gold mineralisation in volcanogenic massive sulfides — implications of data from active hydrothermal vents on the modern seafloor: Economic Geology Monographs, v. 6, p. 491-507. Hannington, M. D., Poulsen, K. H., and Thompson, J. F. H., 1999. Volcanogenic gold in the massive sulfide environment: Reviews in Economic Geology, v. 8, p. 325-56. Hannington, M. D., Santaguida, E, Kjarsgaard, I. M., and Cathles, L. M., 2003a. Regional-scale hydrothermal alteration in the Central Blake River Group, western Abitibi subprovince, Canada - implications for VMS prospectivity: Mineralium Deposita, v. 38, p. 393-422. Hannington, M. D., Kjarsgaard, I. M., Galley, A. G., and Taylor, B. E., 2003b. Mineral-chemical studies of metamorphosed hydrothermal alteration in the Kristineberg volcanogenic massive sulfide district, Sweden: Mineralium Deposita, v. 38, p. 423-43. Hanor, J. S., 1979. The sedimentary genesis of hydrothermal fluids, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits, 3 rd edition: New York, USA, John Wiley and Sons, p. 137-72. Hanson, R. E., 1991. Quenching and hydroclastic disruption of andesitic to rhyolitic intrusions in a submarine island-arc sequence, northern Sierra Nevada, California: Geological

Society of America Bulletin, v. 103, p. 804-16. Hart, S. R., 1969. K, Rb, Cs contents and K/Rb, K/Cs ratios of fresh and altered submarine basalts: Earth and Planetary Science Letters, v. 6, p. 295-303. Hashiguchi, H., Yamada, R., and Inoue, T, 1983. Practical application of low Na 2 O anomalies in footwall acid lava for delimiting promising areas around the Kosaka and Fukazawa Kuroko deposits, Akita Prefecture, Japan: Economic Geology Monographs, v. 5, p. 387-94. Hawkins, D. B., 1981. Kinetics of glass dissolution and zeolite formation under hydrothermal conditions: Clays and Clay Minerals, v. 29, p. 331-40. Hay, R. L., 1963. Stratigraphy and zeolitic diagenesis of the John Day Formation of Oregon: University of California Publications in Geological Sciences, v. 42, p. 199—262. Hay, R. L., 1966. Zeolites and zeolitic reactions in sedimentary rocks: Geological Society of America Special Paper, v. 85, 130 p. Hay, R. L., 1977. Geology of zeolites in sedimentary rocks in Mumpton, F. A., ed., Mineralogy and geology of natural zeolites: Mineralogical Society of America Short Course Notes 4, p. 53-63 Hay, R. L., 1978. Geological occurrence of zeolites, in Sand, L. B., and Mumpton, F. A., eds., Natural zeolites — occurrence, properties, use: Oxford, UK, Peregamon Press, p. 135-43. Hay, R. L., and Guldman, S. G., 1987. Diagenetic alteration of silicic ash in Searles Lake, California: Clays and Clay Minerals, v. 35, p. 449-57. Hay, R. L., and Iijima, A., 1968a. Nature and origin of palagonite tuffs of the Honolulu Group on Oahu, Hawaii: Geological Society of America Memoir, v. 116, p. 338-76. Hay, R. L., and Iijima, A., 1968b. Petrology of palagonite tuffs of KOKO Craters, Oahu, Hawaii: Contributions to Mineralogy and Petrology, v. 17, p. 141-54. Heaton, T. H. E., and Sheppard, S. M. F, 1977. Hydrogen and oxygen isotope evidence for seawater-hydrothermal alteration and ore deposition, Troodos Complex, Cyprus [abs.]: London, UK, January, 1976, Institute of Mining and Metallurgy Monograph - Volcanic processes in ore genesis, p. 21-22. Hein, J. R., and Scholl, D. W, 1978. Diagenesis and distribution of late Cenozoic volcanic sediment in the southern Bering Sea: Geological Society of America Bulletin, v. 89, p. 197-210. Hellman, P. L., Smith, R. E., and Henderson, P., 1977. Rare earth element investigation of the Cliefden outcrop, NSW, Australia: Contributions to Mineralogy and Petrology, v. 65, p. 155-64. Helmold, K. P., and van de Kamp, P C , 1984. Diagenetic mineralogy and control on albitization and laumontite formation in Palaeogene arkoses, Santa Ynez Mountains, California, USA: AAPG Memoir, v. 37, p. 239-76. Hemley, J. J., and Jones, W. R., 1964. Chemical aspects of hydrothermal alteration with emphasis on hydrogen metasomatism: Economic Geology, v. 59, p. 538-67. Henderson, R. A., 1986. Geology of the Mt Windsor Subprovince — a Lower Palaeozoic volcano-sedimentary terrane in the northern Tasman Orogenic Zone: Australian Journal of Earth Sciences, v. 33, p. 343-64. Henley, R. W, and Brown, K. L., 1985. A practical guide to the thermodynamics of geothermal fluids and hydrothermal ore deposits: Reviews in Economic Geology, v. 2, p. 25—44. Henley, H. R., and Ellis, A. J., 1983. Geothermal systems ancient and modern - a geochemical review: Earth Science Reviews, v. 19, p. 1-50. Henley, R. W., and McNabb, A., 1978. Magmatic vapour plumes and ground-water interaction in porphyry copper emplacement: Economic Geology, v. 73, p. 1-19.

REFERENCES | 2 5 9

Henley, R. W., andThornley, P., 1979. Some geothermal aspects of polymetallic massive sulfide formation: Economic Geology, v. 74, p. 1600-12. Henneberger, R. C, and Browne, P. R. L., 1988. Hydrothermal alteration and evolution of the Ohakuri Hydrothermal System, Taupo Volcanic Zone, New Zealand: Journal of Volcanology and Geothermal Research, v. 34, p. 211—31. Herrmann, W., and Berry, R. E, 2002. MINSQ - a least-squares spreadsheet method for calculating mineral proportions from whole rock major element analyses: Geochemistry: Exploration, Environment, Analysis, v. 2, p. 361—8. Herrmann, W., and Hill, A. P., 2001. The origin of chloritetremolite-carbonate rocks associated with the Thalanga volcanic-hosted massive sulfide deposit, north Queensland, Australia: Economic Geology, v. 96, p. 1149—73. Herrmann, W., and MacDonald, G., 1996. Volcanic facies, alteration and exploration targets in EL 8/96 South Henty, Tasmania: Resolute Ltd, Unpublished Report, November 1996, 43 p. Herrmann, W., Blake, M. D., Doyle, M. G., Huston, D. L., Kamprad, J., Merry, N., and Pontual, S., 2001. Short wavelength infrared (SWIR) spectral analysis of hydrothermal alteration zones associated with base metal sulfide deposits at Rosebery and Western Tharsis, Tasmania and Highway-Reward, Queensland: Economic Geology, v. 96, p. 939—55. Herrmann, W., Green, G. R., and Barton, M. D., 2004. Boco prospect - mineralogic and isotopic implications for VHMS exploration [abs.], in McPhie, J., and McGoldrick, P. J., eds., Dynamic earth - past, present and future: Hobart, Australia, February 8—13, 2004, Geological Society of Australia Abstracts, v. 73, p. 87. Herzig, P. M., Petersen, S., Hannington, M. D., and Jonasson, I. R., 2000. Conical Seamount - a submarine analogue of the Ladolam epithermal gold deposits on Lihir Island, Papua New Guinea [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits: Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 83-84. Hill, A. P., 1996. Structure, volcanic setting, hydrothermal alteration and genesis of the Thalanga massive sulfide deposit: Unpublished PhD thesis, University of Tasmania, 404 p. Hill, A. P., and Orth, K., 1994. Textures and origin of carbonate associated with some Australian VHMS deposits [abs.]: AMF Symposium: Australian Research on Ore Genesis, Adelaide, Australia, 1994. Hill, A. P., and Orth, K., 1995. Textures and origin of carbonate associated with the volcanic-hosted massive sulfide deposit at Rosebery, Tasmania: AMIRA/ARC Project, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 1, November 1995, p. 129-41. Hodges, D. J., and Manojlovic, P. M., 1993. Application of lithogeochemistry to exploration for deep VMS deposits in high-grade metamorphic rocks, Snow Lake, Manitoba: Journal of Geochemical Exploration, v. 48, p. 201-24. Hoefs, J., 1973. Stable isotope geochemistry: Berlin, Germany, Springer-Verlag, 140 p. Honnorez, J., 1969. La formation actuelle d'un gisement sousmarin de sulfures fumerolliens a Vulcano (mer tyrrhenienne) - Partie 1, les Mineraux sulfures des tufs immerges a faible profondeur, Translated Title: Present-day formation of a submarine deposit of fumarolic sulfides on Vulcano (Tyrrhenian sea) - Part 1, sulfide minerals of submerged tuffs at shallow depths: Mineralium Deposita, v. 4, p. 114-31. Honnorez, J., 1978. Generation of phillipsites by palagonitization of basaltic glass in seawater and the origin of K-rich deepsea sediments, in Sand, L. B., and Mumpton, E A., eds., Natural zeolites: occurrence, properties and use: New York, USA, Pergamon Press, p. 245-58.

Honnorez, J., Bohlke, J. K., and Honnorez-Guerstein, B. M., 1979. Petrographical and geochemical study of the lowtemperature submarine alteration of basalt from Hole 396B, Leg 46: Proceedings of the Ocean Drilling Program, Initial Reports, v. 46, p. 299-318. Horikoshi, E., 1969. Volcanic activity related to the formation of the Kuroko-type deposits in the Kosaka district, Japan: Mineralium Deposita, v. 4, p. 321-45. Hudak, G. J., Morton, R. L., Peterson, D. M., and Franklin, J. M., 2000. The Sturgeon Lake caldera complex, northwestern Ontario - volcanological evolution of an Archaean shallow water VHMS belt [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits, Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 89-91. Hughes, C. J., 1973. Spilites, keratophyres, and the igneous spectrum: Geological Magazine, v. 6, p. 513-27. Humphris, S. E., and Thompson, G., 1978. Hydrothermal alteration of oceanic basalts by seawater: Geochimica et Cosmochimica Acta, v. 42, p. 107-25. Hunns, S. R., and McPhie, J., 1999. Pumiceous peperite in a submarine volcanic succession at Mount Chalmers, Queensland, Australia: Journal of Volcanology and Geothermal Research, v. 88, p. 239-54. Huston, D. L., 1993. The effect of alteration and metamorphism on wall rocks to the Balcooma and Dry River South volcanichosted massive sulfide deposits, Queensland, Australia: Journal of Geochemical Exploration, v. 48, p. 277—307. Huston, D. L., 1999. Stable isotopes and their significance for understanding the genesis of volcanic hosted massive sulfide deposits — a review: Reviews in Economic Geology, v. 8, p. 157-79. Huston, D. L., and Kamprad, J., 2000. The Western Tharsis deposit, a high sulfidation Cu-Au deposit in the Mt Lyell field, of possible Ordovician age: AGSO Research Newsletter, v. 32, p. 2-6. Huston, D. L., and Kamprad, J., 2001. Zonation of alteration facies at Western Tharsis — implications for the genesis of CuAu deposits, Mt Lyell field, western Tasmania: Economic Geology, v. 96, p. 1123-32. Huston, D. L., and Large, R. R., 1987. Genetic and exploration significance of the zinc ratio (100 Zn/(Zn + Pb)) in massive sulfide systems: Economic Geology, v. 82, p. 1521—39. Huston, D. L., and Large, R. R., 1989. A chemical model for the concentration of gold in volcanogenic massive sulfide deposits: Ore Geology Reviews, v. 4, p. 171—200. Huston, D. L., Taylor, T, Fabray, J., and Patterson, D. J., 1992. A comparison of the geology and mineralization of the Balcooma and Dry River South volcanogenic massive sulfide deposits, northern Queensland: Economic Geology, v. 87, p. 785-811. Huston, D. L., Kamprad, J., and Brauhart, C, 1999. Definition of high-temperature alteration zones with PIMA - an example from the Panorama VHMS district, central Pilbara Craton: AGSO Research Newsletter, v. 30, p. 10-12. Iijima, A., 1974. Clay and zeolitic alteration zones surrounding Kuroko deposits in the Hokuroku District, Northern Akita, as submarine hydrothermal-diagenetic alteration products: Mining Geology Special Issue, v. 6, p. 267—89. Iijima, A., 1978. Geological occurrences of zeolites in marine environments, in Sand, L. B., and Mumpton, F. A., eds., Natural zeolites — occurrence, properties, use: Oxford, UK, Peregamon Press, p. 175—98. Iijima, A., and Hay, R. L., 1968. Analcime composition in tuffs of the Green River Formation of Wyoming: American Mineralogist, v. 53, p. 184-200. Iijima, A., and Utada, M., 1971. Present-day zeolitic diagenesis of

2 6 0 | REFERENCES the Neogene geosynclinal deposits in the Niigata oil field, Japan: American Chemical Society Advances in Chemistry Series, v. 101, p. 342-9. Iijima, A., and Utada, M., 1972. A critical review on the occurrence of zeolites in sedimentary rocks in Japan: Japan Journal of Geology and Geography, v. 42, p. 61-84. Ineson, P. R., 1989. Introduction to practical ore microscopy: New York, USA, John Wiley and Sons, 181 p. Ishikawa, Y., 1983. Volcanic activities and mineralisation of the Fukazawa-Ezuri Kuroko deposits area: Akita University Mining College Scientific and Technological Report, no. 4, p. 23-32. Ishikawa, Y., Sawaguchi, T., Iwaya, S., and Horiuchi, M., 1976. Delineation of prospecting targets for Kuroko deposits based on modes of volcanism of underlying dacite and alteration halos: Mining Geology, v. 26, p. 105-17. Ito,T., Matsumoto, R., Kano, K., and Sakuyama, M., 1984. Welding of acidic tuff and altered rhyolite fragments caused by the intrusion of magma - examples in the Neogene Shirahama Group in the southern part of Izu Peninsula: Journal of the Geological Society of Japan, v. 90, p. 191-205. Jack, D. J., 1989. Hellyer host rock alteration: Unpublished MSc thesis, University of Tasmania, 182 p. Jago, J. B., Reid, K. O., Quilty, P. G., Green, G. R., and Daily, B., 1972. Fossiliferous Cambrian limestone from within the Mt Read Volcanics, Mt Lyell mine area, Tasmania: Journal of Geological Society of Australia, v. 19, p. 379—82. Jakobsson, S. P., 1972. On the consolidation and palagonitization of the tephra of the Surtsey volcanic island, Iceland: Surtsey Research Progress Report, v. 6, p. 121-8. Jakobsson, S. P., 1978. Environmental factors controlling the palagonitization of the Surtsey tephra, Iceland: Geological Society of Denmark Bulletin, v. 27, p. 95-105. Jakobsson, S. P., and Moore, J. G., 1986. Hydrothermal minerals and alteration rates at Surtsey Volcano, Iceland: Geological Society of America Bulletin, v. 97, p. 648-59. Jenner, G. A., 1998. Trace element geochemistry of volcanic rocks — a practical guide to geochemical nomenclature and analytical geochemistry, Centre for Ore Deposit Research, University of Tasmania, Master of Economic Geology course work manual, v. 17, p. 11—48. Jensen, M. L., and Bateman, A. M., 1981. Economic mineral deposits: New York, USA, John Wiley and Sons, 593 p. Jezek, P. A., and Noble, D. C, 1978. Natural hydration and ion exchange of obsidian: American Mineralogist, v. 63, p. 266— 73. Jiang, S., Palmer, M. R., Li, Y, and Xue, C, 1996. Ba-rich micas from the Yindongzi-Daxigou Pb-Zn-Ag and Fe deposits, Qinling, northwestern China: Mineralogical Magazine, v. 60, p. 433-45. Jolly, W. T., and Smith, R. E., 1972. Degradation and metamorphic differentiation of the Keweenawan tholeiitic lavas of northern Michigan, USA: Journal of Petrology, v. 13, p. 273-309. Jones, A. T, 1993. The geology, geochemistry and structure of the Mount Darwin-South Darwin Peak area, western Tasmania: , Unpublished Honours thesis, University of Tasmania, 106 p. Jones, A. T, 1999. Volcanology and geochemistry of the Cambrian Mount Read Volcanics in the Basin Lake area, western Tasmania: Unpublished MSc thesis, University of Tasmania, 190 p. Jones, S., Herrmann, W, and Gemmell, J. B., in prep. SWIR spectral characteristics of the HW horizon: implications for exploration at Myra Falls VHMS camp, Vancouver Island, British Columbia, Canada: Economic Geology. Kajiwara, Y, 1973. Significance of cyclic seawater as a possible determinant of rock alteration facies in the earth's crust:

Geochemical Journal, v. 7, p. 23-36. Kano, K-L, 1989. Interactions between andesitic magma and poorly consolidated sediments — examples in the Neogene Shirahama Group, South Izu, Japan: Journal of Volcanology and Geothermal Research, v. 37, p. 59-75. Karkhanis, S. N., Bancroft, G. M., Fyfe, W S., and Brown, J. D., 1980. Leaching behaviour of rhyolite glass: Nature, v. 284, p. 435. Kastner, M., 1982. Evidence for two distinct hydrothermal systems in the Guaymas Basin: Proceedings of the Ocean Drilling Program, Initial Reports, v. 64, p. 1143-57. Kastner, M., and Gieskes, J. M., 1976. Interstitial water profiles and sites of diagenetic reactions, Leg 35, Bellingshausen abyssal plain: Earth and Planetary Science Letters, v. 33, p. 11-20. Keating, G. N., and Geissman, J. W, 1998. Cooling history of shallow-level mafic intrusions and host tuffs, Paiute Ridge, Nevada — field paleomagnetic, and thermal modelling results [abs.]: Geological Society of America Abstracts with Programs, v. 30, p. 107. Keene, J. B., Clague, A., and Nishimori, R. K., 1976. Experimental hydrothermal alteration of tholeiitic basalt - resultant mineralogy and textures: Journal of Sedimentary Research, v. 46, p. 647-53. Kelley, D. S., and Gillis, K. M., 1993. Fluid evolution in submarine magma-hydrothermal systems at the Mid-Atlantic Ridge: Journal of Geophysical Research, v. 98, p. 19579-96. Kelley, D. S., Robinson, P. T, and Malpas, J. G., 1992. Processes of brine generation and circulation in the oceanic crust - fluid inclusion evidence from the Troodos Ophiolite, Cyprus: Journal of Geophysical Research, v. 97, p. 9307-22. Kennedy, G. C, 1950. A portion of the system silica-water: Economic Geology, v. 45, p. 629—53. Kerr, D. J., and Gibson, H. L., 1993. A comparison of the Home volcanogenic massive sulfide deposit and intracauldron deposits of the Mine Sequence, Noranda, Quebec: Economic Geology, v. 88, p. 1419-42. Khin Zaw, and Large, R. R., 1992. The precious metal-rich, South Hercules mineralisation, western Tasmania - a possible subseafloor replacement volcanic-hosted massive sulfide deposit: Economic Geology, v. 87, p. 931—52. Khin Zaw, Gemmell, J. B., Large, R. R., Mernagh, T. P., and Ryan, C. G., 1996. Evolution and source of ore fluids in the stringer system, Hellyer VHMS deposit, Tasmania, Australia — evidence from fluid inclusion microthermometry and geochemistry: Ore Geology Reviews, v. 10, p. 251-78. Khin Zaw, Huston, D. L., and Large, R. R., 1999. A chemical model for the Devonian remobilization process in the Cambrian volcanic-hosted massive sulfide Rosebery deposit, western Tasmania: Economic Geology, v. 94, p. 529-46. Kirste, D., de Caritat, P., and Dann, R., 2003. The application of the stable isotopes of sulfur and oxygen in groundwater sulfate to mineral exploration in the Broken Hill region of Australia: Journal of Geochemical Exploration, v. 78-79, p. 81-4. Klein, G. D. V, and Lee, Y. I., 1984. A preliminary assessment of geodynamic controls on depositional systems and sandstone diagenesis in back-arc basins, western Pacific Ocean: Tectonophysics, v. 102, p. 119—52. Knuckey, M. J., Comba, C. D. A., and Riverin, G., 1982. Structure, metal zoning and alteration at the Millenbach deposit, Noranda, Quebec, in Hutchinson, R. W, Spence, C. D., and Franklin, J. M., eds., Precambrian sulfide deposits: Geological Association of Canada Special Paper, v. 25, p. 255-317. Kokelaar, B. P., 1982. Fluidization of wet sediments during the emplacement and cooling of various igneous bodies: Journal of the Geological Society of London, v. 139, p. 21-33. Komyshan, P., 1986. Geology of the Hellyer-Mt Charter area, in

REFERENCES | 2 6 1 Large, R. R., ed., The Mount Read Volcanics and associated ore deposits: Burnie, Australia, November 20—21, 1986, Geological Society of Australia, Tasmanian Division, Proceedings, p. 53-5. Kristmannsdottir, H., 1976. Types ofclay minerals in hydro thermally altered basaltic rocks, Reykjanes, Iceland: Jokull, v. 26, p. 30-9. Kuniyoshi, S., and Liou, J. G., 1976. Burial metamorphism of the Karmutsen volcanic rocks, northeastern Vancouver Island, British Columbia: American Journal of Science, v. 276, p. 1096-119. Kutina, J., Park, C. E, Jr., and Smirnov, V. I., 1965. On the definition of zoning and on the relation between zoning and paragenesis: Symposium - Problems of post-magmatic ore deposition with special reference to the geochemistry of ore veins, Prague, Czech Republic, 1965, Proceedings, v. 2, p. 589-95. Lacroix, A., 1936. Le volcan actif de l'ile de la Reunion et ses produits: Paris, France, Gauthier-Villars, 297 p. Lagerbald, B., and Gorbatschev, R., 1985. Hydrothermal alteration as a control of regional geochemistry and ore formation in the central Baltic Shield: Geologische Rundschau, v. 74, p. 33-49. Landis, C. A., and Coombs, D. S., 1967. Metamorphic belts and orogenesis in southern New Zealand: Tectonophysics, v. 4, p. 501-18. Large, R. R., 1977. Chemical evolution and zonation of massive sulfide deposits in volcanic terrains: Economic Geology, v. 72, p. 549-72. Large, R. R., 1992. Australian volcanic-hosted massive sulfide deposits - features, styles, and genetic models: Economic Geology, v. 87, p. 471-510. Large, R. R., 1997. Variability is the key to understanding the genesis of Australian Palaeozoic volcanic hosted massive sulfide deposits [abs.], in Barriga, F. J. A. S., ed. SEG Neves Corvo field conference: Lisbon, Portugal, May 11-14, 1997, SEG Abstracts and Program, p. 54. Large, R. R., 2000. The significance of hybrid deposits in the continuum of VHMS-epithermal and porphyry systems - an example from the Mt Read Volcanics [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits: Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 113. Large, R. R., 2004. Development of an alteration vector diagram for use in the exploration for BHT systems [abs.], in McPhie, J., and McGoldrick, P. J., eds., Dynamic earth — past, present and future: Hobart, Australia, February 8-13, 2004, Geological Society of Australia Abstracts, v. 73, p. 91. Large, R. R., and Allen, R. L., 1997. Preliminary report on the Rosebery lithochemical halo study: AMIRA/ARC Project P439, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 4, May 1997, p. 259-330. Large, R. R., and Both, R. A., 1980. The volcanogenic sulfide ores at Mount Chalmers, eastern Queensland: Economic Geology, v. 75, p. 992-1009. Large, R. R., Huston, D. L., McGoldrick, P. J., Ruxton, P. A., and McArthur, G., 1989. Gold distribution and genesis in Australian volcanogenic massive sulfide deposits and their significance for gold transport models: Economic Geology Monographs, v. 6, p. 520-36. Large, R. R., Doyle, M.G., Raymond, O. L, Cooke, D. R., Jones, A. T, and Heasman, L., 1996. Evaluation of the role of Cambrian granites in the genesis of world-class VHMS deposits in Tasmania: Ore Geology Reviews, v. 10, p. 21530. Large, R. R., Bull, S. W., and McGoldrick, P. J., 2000.

Lithogeochemical halos and geochemical vectors to stratiform sediment hosted Zn-Pb-Ag deposits - Part 2, HYC deposit, McArthur River, Northern Territory: Journal of Geochemical Exploration, v. 68, p. 105-26. Large, R. R., Gemmell, J. B., and Paulick, H., 2001a. The Alteration box plot - a simple approach to understanding the relationship between alteration mineralogy and lithogeochemistry associated with volcanic-hosted massive sulfide deposits: Economic Geology, v. 96, p. 957—71. Large, R. R., Allen, R. L., Blake, M. D., and Herrmann, W., 2001b. Hydrothermal alteration and volatile element halos for the Rosebery K lens volcanic-hosted massive sulfide deposit, western Tasmania: Economic Geology, v. 96, p. 1055-72. Large, R. R., McPhie, J., Gemmell, J. B., Herrmann, W., and Davidson, G. J., 2001c. The spectrum of ore deposit types, volcanic environments, alteration halos, and related exploration vectors in submarine volcanic successions — some examples from Australia: Economic Geology, v. 96, p. 913-38. Larsen, G., and Chilingar, G. V, 1979. Introduction - diagenesis of sediments and rocks, in Larsen, G., and Chilingar, G. V, eds., Diagenesis in sediments and sedimentary rocks, New York, USA, Elsevier Scientific Publishing Co., v. 1, p. 1-29. Le Maitre, R. W., 1981. GENMIX - a generalised petrological mixing model program: Computers and Geosciences, v. 7, p. 229-47. Leaman, D. E., and Richardson, R. G., 1989. The granites of west and northwest Tasmania — a geophysical interpretation: Tasmania Department of Mines, Geological Survey Bulletin, v. 66, 146 p. Lee, Y. I., and Klein, G. D., 1986. Diagenesis of sandstones in the back-arc basins of the western Pacific Ocean: Sedimentology, v. 33, p. 651-75. Lees, T. C, 1987. Geology and mineralisation of the RoseberyHercules area, Tasmania: Unpublished MSc thesis, University of Tasmania, 160 p. Lees, T. C, Khin Zaw, Large, R. R., and Huston, D. L, 1990. Rosebery and Hercules copper-lead-zinc deposits, in Hughes, F. E., ed., Geology of the mineral deposits of Australia and Papua New Guinea: Australasian Institute of Mining and Metallurgy Monograph, v. 14, p. 1241-47. Leistel, J. M., Marcoux, E., Thieblemont, D., Quesada, C, Sanchez, A., Almodovar, G. R., Pascual, E., and Saez, R., 1998. The volcanic-hosted massive sulfide deposits of the Iberian pyrite belt: Mineralium Deposita, v. 33, p. 2—30. Lentz, D. R., and Goodfellow, W. D., 1996. Intense silicification of footwall sedimentary rocks in the stockwork alteration zone beneath the Brunswick No. 12 massive sulfide deposit, Bathurst, New Brunswick: Canadian Journal of Earth Sciences, v. 33, p. 284-302. Lentz, D. R., Hall, D. C, and Hoy, L. D., 1997. Chemostratigraphic, alteration, and oxygen isotopic trends in a profile through the stratigraphic sequence hosting the Health Steele B zone massive sulfide deposit, New Brunswick: The Canadian Mineralogist, v. 35, p. 841-74. Lesher, C. M., Goodwin, A. M., Campbell, I. H., and Gorton, M. P., 1986. Trace-element geochemistry of ore-associated and barren, felsic metavolcanic rocks in the Superior province, Canada: Canadian Journal of Earth Sciences, v. 23, p. 22237. Levi, B., 1970. Burial metamorphic episodes in the Andean geosyncline, Central Chile: Geologishe Rundschau, v. 59, p. 994-1013. Lindgren, W., 1933. Mineral deposits: New York, USA, McGrawHill, 936 p. Lipman, P. W., 1965. Chemical comparison of glassy and crystalline volcanic rocks: United States Geological Survey Bulletin,

2 6 2 | REFERENCES 1201-D, 24 p. Lipman, P. W., Christiansen, R. L., and van Alstine, R. E., 1969. Retention of alkalis by calc-alkalic rhyolites during crystallisation and hydration: American Mineralogist, v. 54, p. 286-91. Lofgren, G., 1970. Experimental devitrification rate of rhyolite glass: Geological Society of America Bulletin, v. 81, p. 553-60. Lofgren, G., 1971a. Experimentally produced devitrification textures in natural rhyolite glass: Geological Society of America Bulletin, v. 82, p. 111-24. Lofgren, G., 1971b. Spherulitic textures in glassy and crystalline rocks: Journal of Geophysical Research, v. 76, p. 5635-48. Lowell, J. D., and Guilbert, J. M., 1970. Lateral and vertical alteration-mineralisation zoning in porphyry ore deposits: Economic Geology, v. 65, p. 373—408. Lydon, J. W., 1984. Volcanogenic massive sulfide deposits - Part 1, a descriptive model: Geoscience Canada, v. 11, p. 195-202. Lydon, J. W., 1988. Volcanogenic massive sulfide deposits — Part 2, genetic models: in Sheahan, P. A., and Cherry, M. E., eds., Ore deposit models: Geoscience Canada Reprint Series, v. 3, p. 145-155. Lydon, J. W, 1996. Characteristics of volcanogenic massive sulfide deposits — interpretations in terms of hydrothermal convection systems and magmatic hydrothermal systems: Boletin Geologico y Minero, v. 107, p. 215-64. Lydon, J. W., 1997. Sedex deposits and global evolution [abs.]: Geological Society of America Abstracts with Programs, v. 29, p. 17-18. Lydon, J. W., and Galley, A., 1986. The chemical and mineralogical zonation of the Mathiati alteration pipe, Cyprus, and its genetic significance, in Gallagher, M. J., et al., ed., Metallurgy of basic and ultra-basic rocks: London, UK, Institution of Mining and Metallurgy, p. 49-68. Lydon, J. W, and Jamieson, H. E., 1984. The generation of oreforming hydrothermal solutions in the Troodos Ophiolite Complex — some hydrodynamic and mineralogical considerations: Geological Survey of Canada Paper, v. 841A, p. 617-25. MacGeehan, P. J., 1978. The geochemistry of altered volcanic rocks at Matagami, Quebec — a geothermal model for massive sulfide genesis: Canadian Journal of Earth Sciences, v. 15, p. 551-70. MacLean, W. H., and Barrett, T. J., 1993. Lithogeochemical techniques using immobile elements: Journal of Geochemical Exploration, v. 48, p. 109-33. MacLean, W. H., and Hoy, L. D., 1991. Geochemistry of hydrothermally altered rocks at the Home Mine, Noranda, Quebec: Economic Geology, v. 86, p. 506—28. MacLean, W. H., and Kranidiotis, P., 1987. Immobile elements as monitors of mass transfer in hydrothermal alteration Phelps Dodge massive sulfide deposit, Matagami, Quebec: Economic Geology, v. 82, p. 951—62. Madeisky, H. E., and Stanley, C. R., 1994. Metasomatism in felsic volcanic hosted massive sulfide footwall alteration zones - lithogeochemical patterns, fluid-rock reactions, and their t application to exploration in metamorphosed terranes [abs.]: Waterloo, Canada, May 16—18,1994, Geological Association of Canada Program with Abstracts, v. 19, p. 70. Manning, C. E., and Bird, K. K., 1991. Porosity, evolution and fluid flow in the basalts of the Skaergaard magma-hydrothermal system, east Greenland: American Journal of Science, v. 291, p. 201-57. Manning, C. E., and Bird, D. K., 1995. Porosity, permeability, and basalt metamorphism: Geological Society of America Special Paper, v. 296, p. 123-40. Marsaglia, K. M., and Tazaki, K., 1992. Diagenetic trends in Leg 126 sandstones: Proceedings of the Ocean Drilling Program,

Scientific Results, v. 126, p. 125-38. Marshall, C. E., 1961. Reactions of feldspars and micas with aqueous solutions: Economic Geology, v. 56, p. 1330. Martin, N. K., 2004. Genesis of the Rosebery massive sulfide deposit, western Tasmanian, Australia: Unpublished PhD thesis, University of Tasmania, 273 p. Marumo, K., 1989. Genesis of kaolin minerals and pyrophyllite in Kuroko deposits of Japan — implications for the origins of the hydrothermal fluids from mineralogical and stable isotope data: Geochimica et Cosmochimica Acta, v. 53, p.2915-24. Mason, B., 1966. Principles of geochemistry: New York, USA, John Wiley and Sons, 329 p. McArthur, G. J., 1989. Hellyer, in Burrett, C. E, and Martin, E. L, eds., Geology and mineral resources of Tasmania: Geological Society of Australia Special Publication, v. 15, p. 144—48. McArthur, G. J., 1996. Textural evolution of the Hellyer massive sulfide deposit (3 volumes): Unpublished PhD thesis, University of Tasmania, 272 p. McArthur, G. J., and Dronseika, E. V., 1990. Que River and Hellyer zinc-lead-silver deposits: Australasian Institute of Mining and Metallurgy, Monograph, v. 14, p. 1229-39. McArthur, A. N., Cas, R. A. E, and Orton, G. J., 1998. Distribution and significance of crystalline, perlitic and vesicular textures in the Ordovician Garth Tuff (Wales): Bulletin of Volcanology, v. 60, p. 260-85. McBride, E. E, 1978. Porosity loss in sandstone by ductile grain deformation during compaction: AAPG (American Association of Petroleum Geologists) Bulletin, v. 62, p. 1761. McConnell, V. S., Shearer, C. K., Eichelberger, J. C, Keskinen, M. J., Layer, P. W, and Papike, J. J., 1995. Rhyolite intrusions in the intracaldera Bishop Tuff, Long Valley Caldera, California: Journal of Volcanology and Geothermal Research, v. 67, p. 41-60. McKibben, J. A., 1993. The geology and geochemistry of the North Pinnacles Ridge, western Tasmania: Unpublished Honours thesis, University of Tasmania, 120 p. McLeod, R. L., 1987. Alteration associated with volcanogenic sulfide ores at Mount Chalmers, Queensland, Australia: Institution of Mining and Metallurgy, Transactions, Section B: Applied Earth Science, v. 96, p. 117-27. McLeod, R. L., and Stanton, R. L., 1984. Phyllosilicates and associated minerals in some Palaeozoic stratiform sulfide deposits of southeastern Australia: Economic Geology, v. 79, p. 1-22. McMillan, W J., and Panteleyev, A., 1998. Porphyry copper deposits, in Roberts, R. G., and Sheahan, P. A., eds., Ore deposit models, Geoscience Canada Reprint Series, v. 3, p. 45-58. McNeill, A. W, and Corbett, K. D., 1992. Geology and mineralisation of the Mt Murchison area (including Map 4): Mt Read Volcanics Project Geological Report, Tasmania Department of Mines, no. 3, 39 p. McPhie, J., 1996. Hall Rivulet Canal-Hercules-Mount Read-Read Hills-Anthony Dam: AMIRA/ARC Project P439, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 2, May 1996, p. 19—28. McPhie, J., and Allen, R. L., 1992. Facies architecture of mineralised submarine volcanic sequences — Cambrian Mount Read Volcanics, western Tasmania: Economic Geology, v. 87, p. 587-96. McPhie, J., and Allen, R. L., 2003. Submarine, silicic, syneruptive pyroclastic units in the Mount Read Volcanics, western Tasmania — influence of vent setting and proximity on lithofacies characteristics: AGU Geophysical Monograph, v. 140, p. 245-58.

REFERENCES | 2 6 3

McPhie.J., andHunns, S. R., 1995. Secondary welding of submarine, pumice-lithic breccia at Mount Chalmers, Queensland, Australia: Bulletin of Volcanology, v. 57, p. 170-8. McPhie, J., Doyle, M. G., and Allen, R. L., 1993. Volcanic textures — a guide to the interpretation of textures in volcanic rocks: Hobart, Australia, Centre for Ore Deposit and Exploration Studies, University of Tasmania, 197 p. Meinert, L. D., 1983. Variability of skarn deposits — guides to exploration, in Boardman Shelby, J., ed., Revolution in the earth sciences - advances in the past half-century: Dubuque, USA, Kendall/Hunt Publishing Co., p. 301-16. Meinert, L. D., 1993. Skarns and skarn deposits, in Sheahan, P. A., and Cherry, M. E., eds., Ore deposit models, 2, Geoscience Canada Reprint Series, v. 2, p. 117—34. Merino, E., 1975. Diagenesis in Tertiary sandstones from Kettleman North Dome, California — Part I, diagenetic mineralogy: Journal of Sedimentary Petrology, v. 45, p. 320—36. Merino, E., Ortoleva, P., and Strickholm, P., 1983. Generation of evenly spaced pressure-solution seams during (late) diagenesis — a kinetic theory: Contributions to Mineralogy and Petrology, v. 82, p. 360-70. Merry, N., and Pontual, S., 1999. Rapid alteration mapping using field portable infrared spectrometers: Bali, Indonesia, October 10-13, 1999, Proceedings of PACRIM '99 congress [modified], p. 693-8. Mevel, C, and Cannat, M., 1991. Lithospheric stretching and hydrothermal processes in oceanic gabbros from slowspreading ridges, in Peters, T. J. (ed.) Ophiolite genesis and evolution of the oceanic lithosphere, Ministry of Petroleum and Minerals, Muscat, Sultante of Oman, p. 293-312. Meyer, C, and Hemley, J. J., 1967. Wall-rock alteration, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits: New York, USA, Holt, Rinehart and Winston, p. 166-235. Miller, C. R., 1996. Geological and geochemical aspects of the Liontown VHMS deposit, northeastern Queensland: Unpublished MSc thesis, University of Tasmania, 60 p. Miller, C. R., Halley, S., Green, G. R., and Jones, M. T, 2001. Discovery of the West 45 volcanic-hosted massive sulfide deposit using oxygen isotopes and REE geochemistry: Economic Geology, v. 96, p. 1227—37. Miyashiro, A., and Shido, E, 1970. Progressive metamorphism in zeolite assemblages: Lithos, v. 3, p. 251-60. Moore, J. G., 1966. Rate of palagonitization of submarine basalt adjacent to Hawaii: United States Geological Survey Professional Paper 55OD, D163-71. Morrow, D. W, and Mcllreath, I. A., 1990. Diagenesis general introduction, in Mcllreath, I. A., and Morrow, D. W, eds., Diagenesis: Newfoundland, Canada, Geological Association of Canada, Geoscience Canada Reprint Series, v. 4, p. 1-8. Morton, R. L., and Franklin, J. M., 1987. Two-fold classification of Archaean volcanic-associated massive sulfide deposits: Economic Geology, v. 82, p. 1057—63. Mottl, M. J., 1983. Metabasalts, axial hot springs, and the structure of hydrothermal systems at mid-ocean ridges: Geological Society of America Bulletin, v. 94, p. 161-80. Mottl, M. J., and Seyfried, W. E., 1977. Experimental basaltseawater interaction - rock- vs. seawater-dominated systems and the origin of submarine hydrothermal deposits [abs.]: Geological Society of America Abstracts with Programs, v. 9, p. 1104-5. Muehlenbachs, K., and Clayton, R. N., 1972. Oxygen isotope studies of fresh and weathered submarine basalts: Canadian Journal of Earth Sciences, v. 9, p. 172-84. Mungall, J. E., and Martin, R. E, 1994. Severe leaching of trachytic glass without devitrification, Terceira, Azores: Geochimica et Cosmochimica Acta, v. 58, p. 75-83. Munha, J., and Kerrich, R., 1980. Seawater basalt interaction

in spilites from the Iberian pyrite belt: Contributions to Mineralogy and Petrology, v. 73, p. 191-200. Munha, J., Fyfe, W. S., and Kerrich, R., 1980. Adularia the characteristic mineral of felsic spilites: Contributions to Mineralogy and Petrology, v. 75, p. 15-19. Nakajima, T, 1988. Geology of the Hokuroku basin in Kuroko deposits and geothermal fields in northern Honshu: The Society of Mining Geologists of Japan, Guidebook 3. Nehlig, P., 1993. Interactions between magma chambers and hydrothermal systems - oceanic and ophiolitic constraints: Journal of Geophysical Research, v. 98, p. 19621-33. Nesbitt, B. E. 1996. Applications of oxygen and hydrogen isotopes to exploration for hydrothermal mineralisation: SEG Newsletter, v. 27, p. 1 and 8-13. Neuhoff, P. S., Watt, W S., Bird, D. K., and Pedersen, A. K., 1997. Timing and structural relations of regional zeolite zones in basalts of the east Greenland continental margin: Geology, v. 25, p. 803-6. Nielsen, T E D., Soper, N. J., Brooks, C. K., Faller, A. M., Higgens, A. C, and Matthews, D. W, 1981. The pre-basaltic sediments and the lower basalts at Kangerdlugssuaq, east Greenland — their stratigraphy, lithology, palaeomagnetism and petrology: Meddelelser om Gronland, Geoscience, v. 6, p. 25. Nilsson, C. A., 1968. Wall-rock alteration at the Boliden deposit, Sweden: Economic Geology, v. 63, p. 472—94. Noble, D. C, 1967. Sodium, potassium and ferrous iron contents of some secondarily hydrated natural silicic glasses: American Mineralogist, v. 52, p. 280-6. Noh, J. H., and Boles, J. R., 1989. Diagenetic alteration of perlite in the Guryongpo area, Republic of Korea: Clays and Clay Minerals, v. 37, p. 47-58. Norton, D. L., 1984. Theory of hydrothermal systems: Annual Review Earth Planetary Science, v. 12, p. 155—77. Norton, D. L., and Taylor, H. P., Jr., 1979. Quantitative simulation of the hydrothermal systems of crystallizing magmas on the basis of transport theory and oxygen isotope data - an analysis of the Skaergaard intrusion: Journal of Petrology, v. 20, p. 421-86. Norton, D. L., Taylor, H. P., Jr., and Bird, D. K., 1984. The geometry and high-temperature brittle deformation of the Skaergaard intrusion: Journal of Geophysical Research, v. 89, p. 10178-92. Offler, R., and Whitford, D. J., 1992. Wall-rock alteration and metamorphism of a volcanogenic massive sulfide deposit at Que River, Tasmania — petrology and mineralogy: Economic Geology, v. 87, p. 686-705. Oftedahl, C, 1958. A theory of exhalative-sedimentary ores: Geologiska Foereningen i Stockholm Foerhandlingar, v. 80, p. 1-19. Ogihara, S., 1996. Diagenetic transformation of clinoptilolite to analcime in silicic tuffs of Hokkaido, Japan: Mineralium Deposita, v. 31, p. 548-53. Ohmoto, H., 1978. Submarine caldera — a key for the formation of massive sulfide deposits: Mining Geology Special Issue, v. 28, p. 219-31. Ohmoto, H., 1986. Stable isotope geochemistry of ore deposits: Reviews in Mineralogy, v. 16, p. 491-559. Ohmoto, H., 1996. Formation of volcanogenic massive sulfide deposits — the Kuroko perspective: Ore Geology Reviews, v. 10, p. 135-77. Ohmoto, H., and Goldhaber, M. J., 1997. Sulfur and carbon isotopes, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits, 3 rd edition: New York, USA, John Wiley and Sons, p. 517-612. Ohmoto, H., and Rye, R. O., 1974. Hydrogen and oxygen isotopic compositions of fluid inclusions — an introduction and

2 6 4 | REFERENCES summary of new findings: Economic Geology Monographs, v. 5, p. 1-8. Ohmoto, H., and Skinner, B. J., 1983. The Kuroko and related volcanogenic massive sulfide deposits — introduction and summary of new findings: Economic Geology Monographs, v. 5, p. 1-8. Ohmoto, H., and Takahashi, T., 1983. Submarine calderas and Kuroko genesis - Part III: Economic Geology Monographs, v. 5, p. 39-54. Ohmoto, H., Mizukami, M., Drummond, S. E., Eldridge, C. S., Pisutha, A. V., and Lenagh, T. C., 1983. Chemical processes of Kuroko formation: Economic Geology Monographs, v. 5, p. 570-604. O'Keefe, J. A., 1984. Natural glass: Journal of Non-Crystalline Solids, v. 67, p. 1-17. O'Neil, J. R., and Silberman, M. L., 1974. Stable isotope relations in epithermal Au-Ag deposits: Economic Geology, v. 69, p. 902-9. Palmasson, G., Arnorsson, S., Fridleifsson, I. B., Kristmannsdottir, H., Saemundson, K., Stefansson, V., Steingrimsson, B., and Tomasson, J., 1979. The Iceland crust — evidence from drill hole data on structure and processes, in Talwani, M., et al., eds., Deep drilling results in the Atlantic Ocean: oceanic crust, AGU Maurice Ewing Series, v. 2, p. 43-65. Paradis, S., Taylor, B. E., Watkinson, D. H., and Jonasson, I. R., 1993. Oxygen isotope zonation and alteration in the northern Noranda District, Quebec - evidence for hydrothermal fluid flow: Economic Geology, v. 88, p. 1512-25. Paradis, S., Bailes, A. H., and Galley, A. G., in press. Regionalscale synvolcanic alteration zones associated with Snow Lake volcanogenic massive sulfide deposits, Flin Flon belt, Manitoba, Canada - Part II, mineralogical and chemical characteristics: Mineralium Deposita. Passaglia, E., Artioli, G., Gualtieri, A., and Carnevali, R., 1995. Diagenetic mordenite from Ponza, Italy: European Journal of Mineralogy, v. 7, p. 429-38. Paulick, H., 1999. The Thalanga sequence - facies architecture, geochemistry, alteration and metamorphism of felsic volcanics hosting the Thalanga massive sulfide deposit (Early Ordovician, northern Queensland, Australia): Unpublished PhD thesis, University of Tasmania, 240 p. Paulick, H., and McPhie, J., 1999. Facies architecture of the felsic lava-dominated host sequence to the Thalanga massive sulfide deposit, Lower Ordovician, northern Queensland: Australian Journal of Earth Sciences, v. 46, p. 391-405. Paulick, H., Herrmann, W., and Gemmell, J. B., 2001. Alteration of felsic volcanics hosting the Thalanga massive sulfide deposit (northern Queensland, Australia) and geochemical proximity indicators to ore: Economic Geology, v. 96, p. 1175-200. Payne, B., 1991. Geophysical interpretation of the Mt SedgewickRed Hills area, western Tasmania: Unpublished Honours thesis, University of Tasmania, 107 p. Peacock, M. A., 1926. The petrology of Iceland - Part I, the basic tuffs: Royal Society of Edinburgh Transactions, v. 55, p. 53— 76. Pearce, J. A., and Cann, J. R., 1973. Tectonic setting of basic volcanic rocks determined using trace element analyses: Earth and Planetary Science Letters, v. 19, p. 290-300. Pemberton, J., and Corbett, K. D., 1992. Stratigraphic-facies associations and their relationship to mineralisation in the Mount Read Volcanics: Bulletin Geological Survey Tasmania, v. 70, p. 167—76. Pemberton, J., Vicary, M. J., and Corbett, K. D., 1991. Geology of the Cradle Mountain Link Road-Mt Tor area (maps 7 and 8): Mt Read Volcanics Project Geological Report, Tasmania Department of Mines, v. 4, 74 p. Penney, A., 1998. Geology and genesis of the massive pyrite and

carbonate, Henty Gold Mine, Tasmania: Unpublished Honours thesis, University of Tasmania, 111 p. Peter, J. M., and Goodfellow, W. D., 1996. Mineralogy, bulk and rare earth element geochemistry of massive sulfide-associated hydrothermal sediments of the Brunswick horizon, Bathurst mining camp, New Brunswick: Canadian Journal of Earth Sciences, v. 33, p. 252—83. Peterson, D. W., 1979. Significance of the flattening of pumice fragments in ash-flow tuffs, in Chapin, C. E., and Elston, W. E., eds., Ash-flow tuffs, Geological Society of America Special Paper, no. 180, p. 195-204. Pettijohn, F. J., 1957. Sedimentary rocks: New York, USA, Harper and Row, 718 p. Phillips, W. J., 1973. Mechanical effects of retrograde boiling and its probable importance in the formation of some porphyry ore deposits: Institute Mining Metallurgy Transcripts, Section B, v. 82, p. B90-8. Pichler, H., 1965. Acid hyaloclastites: Bulletin of Volcanology, v. 28, p. 293-310. Pichler, H., 1981. Italienische Vulkan-Gebiete III: Lipari, Vulcano, Stromboli, Tyrrhenisches Meer: Translated Title: Italian volcanic zones III: Lipari, Volcano, Stromboli, Tyrrhenisches Sea: Sammlung Geologischer Fuhrer, v. 69, 270 p. Pichler, T, Giggenbach, W. E, Mclnnes, B. I. A., Buhl, D., and Duck, B., 1999. Fe-sulfide formation due to seawater-gassediment interaction in a shallow-water hydrothermal system at Lihir Island, Papua New Guinea: Economic Geology, v. 94, p. 281-8. Pisutha-Arnond, V., and Ohmoto, H., 1983. Thermal history, and chemical and isotopic compositions of the ore-forming fluids responsible for the Kuroko massive sulfide deposits in the Hokuroku District of Japan: Economic Geology Monographs, v. 5, p. 523—58. Plimer, I. R., and de Carvalho, D., 1982. The geochemistry of hydrothermal alteration at the Salgadinho copper deposit, Portugal: Mineralium Deposita, v. 17, p. 193-211. Polya, D. A., 1981. The geology of the Murchison Gorge: Unpublished Honours thesis, University of Tasmania, 170 p. Polya, D. A., Solomon, M., Eastoe, C. J., and Walshe, J. L, 1986. The Murchison Gorge, Tasmania — a possible cross-section through a Cambrian massive sulfide system: Economic Geology, v. 81, p. 1341-55. Pontual, S., Merry, N., and Gamson, P., 1997. Spectral interpretation field manual, G-Mex, volume 1: Melbourne, Ausspec International Pty. Ltd., 169 p. Potts, P. J., 1987. A handbook of silicate rock analysis: London, UK, Blackie and Son Ltd., 622 p. Puffer, J. H., and Benimoff, A. I., 1997. Fractionation, hydrothermal alteration, and wall-rock contamination of an Early Jurassic diabase intrusion, Laurel Hill, New Jersey: Journal of Geology, v. 105, p. 99-110. Ramdohr, P., 1980. The ore minerals and their intergrowths: Oxford, UK, Pergamon Press, 1205 p. Ratterman, N. G., and Surdam, R. C, 1981. Zeolite mineral reactions in a tuff in the Laney Member of the Green River Formation, Wyoming: Clays and Clay Minerals, v. 29, p. 365-77. Relvas, J. M. R. S., Barriga, E J. A. S., Bernardino, F. B. C. P., Oliveira, V. M. S., and Matos, J. X., 1994. Ore zone hydrothermal alteration in drill hole IGM-LS1, at Lagoa Salgada, Grandola, Portugal — a first report on pyrophyllite in a central stockwork, Boletin Sociedad Espanola de Mineralogia, v. 17, p. 157-8. Relvas, J. M. R. S., Barriga, F. J. A. S., Ferreira, A., Noiva, P. C, and Carvalho, P., 1997. Footwall alteration and stringer ores, Corvo ore body, Neves Corvo, Portugal [abs.], in Barriga,

REFERENCES | 2 6 5 F. J. A. S., ed. SEG Neves Corvo field conference: Lisbon, Portugal, May 11-14, 1997, SEG Abstracts and Program, p. 93. Relvas, J. M. R. S., Barriga, E J. A. S., Ferreira, A., and Noiva, P. C., 2000. Ore geology, hydrothermal alteration and replacement ore-forming mechanisms in the Corvo ore body, Neves Corvo, Portugal [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits, Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 167-8. Reyes, A. R., 1990. Petrology of Philippine geothermal systems and the application of alteration mineralogy to their assessment: Journal of Volcanology and Geothermal Research, v. 43, p.279-309. Richardson, C. J., Cann, J. R., Richards, H. G., and Cowan, J. G., 1987. Metal-depleted root zones of theTroodos ore-forming hydrothermal systems, Cyprus: Earth and Planetary Science Letters, v. 84, p. 243-53. Riech, V., 1979. Diagenesis of silica, zeolites and phyllosilicates at sites 397 and 398: Proceedings of the Ocean Drilling Program, Initial Reports, v. 47, p. 741-60. Rittmann, A., 1962. Volcanoes and their activity: New York, USA, John Wiley and Sons, 305 p. Riverin, G., 1977. Wall-rock alteration at the Millenbach Mine, Noranda, Quebec: Unpublished PhD thesis, Queen's University. Riverin, G., and Hodgson, C. J., 1980. Wall-rock alteration at the Millenbach Cu-Zn mine, Noranda, Quebec: Economic Geology, v. 75, p. 424-44. Roberts, R. G., and Reardon, E. J., 1978. Alteration and ore-forming processes at Mattagami Lake Mine, Quebec: Canadian Journal of Earth Sciences, v. 15, p. 1-21. Robinson, P., 2001. The place of XRF and ICP-MS in the analysis of silicate rocks, sulfide ores, ironstones and carbonates [abs.]: 2nd international workshop, Siberian Geoanalytical Seminar, Irkutsk, Russia, July 24-26, 2001, Abstracts volume, p. 56. Rollinson, H. R., 1993. Using geochemical data - evaluation, presentation, interpretation: Harlow, UK, Longman Scientific and Technical, 352 p. Rona, P. A., and Scott, S. D., 1993. A special issue on seafloor hydrothermal mineralisation - new perspectives: Economic Geology, v. 88, p. 1933-2245. Rona, P. A., Hannington, M. D., Raman, C. V, Thompson, G., Tivey, M. K., Humphris, S. E., Lalou, C, and Petersen, S., 1993. Active and relict seafloor hydrothermal mineralisation at the TAG hydrothermal field, Mid-Atlantic Ridge: Economic Geology, v. 88, p. 1987-2013. Rose, A. W, 1970. Zonal relations of wall-rock alteration and sulfide distribution at porphyry copper deposits: Economic Geology, v. 65, p. 920-36. Rose, A. W, and Burt, D. M., 1979. Hydrothermal alteration, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits, 3 rd edition: New York, USA, John Wiley and Sons, p. 173-235. Ross, C. S., and Smith, R. L., 1955. Water and other volatiles in volcanic glass: American Mineralogist, v. 40, p. 1071—89. Ross, C. S., and Smith, R. L., 1960. Ash-flow tuffs - their origin, geologic relations and identification: United States Geological Survey Professional Paper 366, 54 p. Saccocia, P. J., Ding, K., Berndt, M. E., Seewald, J. S., and Seyfried, W. E., Jr., 1994. Experimental and theoretical perspectives on crustal alteration at mid-ocean ridges, in Lentz, D. R., ed., Alteration and alteration processes associated with ore-forming systems, II: Waterloo, Canada, Geological Association of Canada short course notes, p. 403-31. Saeki, Y, and Date, J., 1980. Computer application to the alteration data of the footwall dacite lava at the Ezuri Kuroko deposits,

Akita Prefecture: Mining Geology, v. 30, p. 241—50. Saez, R., Pascual, E., Toscano, M., and Almodova, G. R., 1999. The Iberian type of volcano-sedimentary massive sulfide deposits: Mineralium Deposita, v. 34, p. 549-70. Sakai, H., Shinaki, R., Kishima, M., and Tazaki, K., 1978. Oxygen isotope exchange and chemical reactions in rhyolite-seawater system at 300°C and 1000 bars: EOS-Transactions, American Geophysical Union, v. 59, p. 1220-1. Sales, R. H., and Meyer, C, 1948. Wall-rock alteration at Butte, Montana: American Institute of Mining, Metallurgical and Petroleum Engineers Technical Publication, v. 12, no. 3, 25 p. Sanchez-Espana, J., Velasco, E, and Yusta, I., 2000. Hydrothermal alteration of felsic volcanic rocks associated with massive sulfide deposition in the northern Iberian pyrite belt (SW Spain): Applied Geochemistry, v. 15, p. 1265-90. Sanchez-Espana, J., Carcia de Cortazar, A., Gil, P., and Velasco, E, 2002. The discovery of the Borobia world-class stratiform magnesite deposit (Soria, Spain) — a preliminary report: Mineralium Deposita, v. 37, p. 240-3. Sangster, D. E, 1972. Precambrian volcanogenic massive sulfide deposits in Canada — a review: Geological Society of Canada, v. 72, p. 22-43. Sato, T, 1973. A chloride complex model for Kuroko mineralisation: Geochemical Journal, v. 7, p. 245-70. Sato, T, 1974. Distribution and geologic setting of the Kuroko deposits: Kozan Chishitsu Special Issue, v. 6, p. 1—9. Sawkins, F. J., and Kowalik, J., 1981. The source of ore metals at Buchans - magmatic versus leaching models, in Swanson, E. A., Strong, D. E, and Thurlow, J. G., eds., The Buchans ore bodies — fifty years of geology and mining: Geological Association of Canada Paper 22, p. 255-67. Schardt, C, Cooke, D. R., Gemmell, J. B., and Large, R. R., 2001. Geochemical modelling of the zones footwall alteration pipe, Hellyer volcanic-hosted massive sulfide deposit, western Tasmania, Australia: Economic Geology, v. 96, p. 1037-54. Schardt, C, Yang, J., and Large, R., in press. Numerical heat and fluid-flow modelling of the Panorama volcanic-hosted massive sulfide district, Western Australia: Economic Geology. Schiffman, P., and Smith, B. M., 1988. Petrology and oxygen isotope geochemistry of a fossil seawater hydrothermal system within the Solea Graben, northern Troodos Ophiolite, Cyprus: Journal of Geophysical Research, v. 93, p. 4612-24. Schiffman, P., Elders, W A., Williams, A. E., McDowell, S. D., and Bird, D. K., 1984. Active metasomatism in the Cerro Prieto geothermal system, Baja California, Mexico — a telescoped low-pressure, low-temperature metamorphic fades series: Geology, v. 12, p. 12-15. Schiffman, P., Smith, B. M., Varga, R. J., and Moores, E. M., 1987. Geometry, conditions and timing of off-axis hydrothermal metamorphism and ore-deposition in the Solea graben (Cyprus): Nature, v. 325, p. 423-5. Schmidt, J., M., 1988. Mineral and whole-rock compositions of seawater-dominated hydrothermal alteration at the Arctic volcanogenic massive sulfide prospect, Alaska: Economic Geology, v. 83, p. 822-42. Schmincke, H-U., and von Rad, U., 1976. Neogene evolution of Canary Island volcanism inferred from ash layers and volcaniclastic sandstones of DSDP Site 397 (Leg 47A): Proceedings of the Ocean Drilling Program, Initial Reports, v. 47, part 1, p. 703-26. Schmincke, H-U., 1967. Fused tuff and peperites in south-central Washington: Geological Society of America Bulletin, v. 78, p. 319-30. Schwartz, G. M., 1959. Hydrothermal alteration: Economic Geology, v. 54, p. 161-83.

2 6 6 | REFERENCES Schweitzer, J. K., and Hatton, C. J., 1995. Chemical alteration within the volcanic roof rocks of the Bushveld Complex: Economic Geology, v. 90, p. 2218-31. Scott, S. D., 1997. Submarine hydrothermal systems and deposits, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits, New York, USA, John Wiley and Sons, p. 797876. Scott, R. B., and Hajash, A., Jr., 1976. Initial submarine alteration of basaltic pillow lavas — a microprobe study: American Journal of Science, v. 276, p. 480-501. Secor, D. T., Jr., 1965. Role of fluid pressure in jointing: American Journal of Science, v. 263, p. 633-46. Sederholm, J. J., 1929. The use of the term 'deuteric': Economic Geology, v. 24, p. 869-71. Seki, Y., 1972. Lower grade stability limit of epidote in the light of natural occurrences: Journal of the Geological Society of Japan, v. 78, p. 405-513. Seki, Y, Oki, Y, Matsuda, T., Mikami, K., and Okumura, K., 1969. Metamorphism in the Tanzawa mountains, central Japan: Journal of the Japanese Association of Mineralogists, Petrologists and Economic Geologists, v. 61, p. 1-24. Selley, D., 1997. Structure and sedimentology of the Dundas Group, western Tasmania: Unpublished PhD thesis, University of Tasmania, 188 p. Seyfried, W. E., Jr., and Bischoff, J. L., 1977. Hydrothermal transport of heavy metals by seawater: the role of seawater-basalt ratio: Earth and Planetary Science Letters, v. 34, p. 71-7. Seyfried, W. E., Jr., and Bischoff, J. L., 1979. Low temperature basalt alteration by seawater - an experimental study at 70°C and 150°C: Geochimica et Cosmochimica Acta, v. 43, p. 193747. Seyfried, W. E., and Mottl, M. J., 1982. Hydrothermal alteration of basalt by seawater under seawater-dominated conditions: Geochimica et Cosmochimica Acta, v. 46, p. 985-1002. Seyfried, W. E., Jr., Mottl, M. J., and Bischoff, J. L, 1978. Seawaterbasalt ratio effects on the chemistry and mineralogy of spilites from the ocean floor: Nature, v. 275, p. 211—13. Seyfried, W. E., Jr., Berndt, M. E., and Seewald, J. S., 1988. Hydrothermal alteration processes at mid-ocean ridges — constraints from diabase alteration experiments, hot spring fluids and composition of the oceanic crust: Canadian Mineralogist, v. 26, p. 787-804. Sharpe, R., 1991. Geology and geochemistry of the barite and siliceous caps, Hellyer Mine, Tasmania: Unpublished Honours thesis, University of Tasmania, 114 p. Sharpe, R., and Gemmell, J. B., 2001. Alteration characteristics of the Archaean Golden Grove Formation at the Gossan Hill deposit, Western Australia — induration as a focusing mechanism for mineralising hydrothermal fluids: Economic Geology, v. 96, p. 1239-62. Sheppard, R. A., and Gude, A. J., 3 rd , 1968. Distribution and genesis of authigenic silicate minerals in tuffs of Pleistocene Lake Tecopa, Inyo County, California, United States Geological Survey Professional Paper 597, 38 p. Sheppard, R. A., and Gude, A. J., 3 rd , 1973. Zeolites and associated authigenic silicate minerals in tuffaceous rocks of the Big Sandy Formation, Mohave County, Arizona, United States Geological Survey Professional Paper P0830, 36 p. Sheppard, R. A., Gude, A. J., 3 rd , and Fitzpatrick, J. J., 1988. Distribution, characterisation, and genesis of mordenite in Miocene silicic tuffs at Yucca Mountain, Nye County, Nevada: United States Geological Survey Bulletin, B1777, p. 1-22. Sherwood, G. J., 1988. Rock magnetic studies of Miocene volcanics in eastern Otago and Banks Peninsula, New Zealand — comparison between Curie temperature and low temperature susceptibility behaviour: New Zealand Journal of Geology

and Geophysics, v. 31, p. 225-35. Shikazono, N., 1999. Rare earth element geochemistry of Kuroko ores and hydrothermally altered rocks — implication for evolution of submarine hydrothermal system at back arc basin: Resource Geology, v. 20, p. 23—30. Shikazono, N., Hoshino, M., Utada, M., Nakata, and Ueda, A., 1998. Hydrothermal carbonates in altered wall rocks at the Uwamuki Kuroko deposits, Japan: Mineralium Deposita, v. 33, p. 346-58. Shiraki, R., and Iiyama, J. T, 1990. Na-K ion exchange reaction between rhyolite glass and (Na, K)C1 aqueous solution under hydrothermal conditions: Geochimica et Cosmochimica Acta, v. 54, p. 923-31. Shiraki, R., Sakai, H., Endoh, M., and Kishima, N., 1987. Experimental studies on rhyolite- and andesite-seawater interactions at 300°C and 1000 bars: Geochemical Journal, v. 21, p. 139-48. Shirozu, H., 1974. Clay minerals in altered wall rocks of the Kurokotype deposits: Kozan Chishitsu = Mining Geology, no. 6, p. 303-10. Sibson, R. H., 1977. Fault rocks and fault mechanisms: Journal Geological Society, v. 133, p. 191—213. Sillitoe, R. H., Hannington, M. D., and Thompson, J. F. H., 1996. High-sulfidation deposits in the volcanogenic massive sulfide environment: Economic Geology, v. 91, p. 204-12. Simmons, S. E, and Browne, P. R. L., 2000. Hydrothermal minerals and precious metals in the Broadlands-Ohaaki geothermal system - implications for understanding low-sulfidation epithermal environments: Economic Geology, v. 95, p. 9 7 1 99. Simmons, S. F., and Christenson, B. W, 1994. Origins of calcite in a boiling geothermal system: American Journal of Science, v. 294, p. 361-400. Simpson, K. A., 2001. Volcanic facies architecture of the Seventy Mile Range Group, Queensland: Unpublished PhD thesis, University of Tasmania, 217 p. Simpson, K. A., andMcPhie, J., 1998. Characteristics of subaqueous basaltic andesite fire fountain deposits - an example from the Mount Windsor Volcanics, northern Queensland, Australia [abs.], in Beams, S. D., ed., Economic geology of northeast Queensland, the 1998 perspective: Townsville, Australia, July 6-10, 1998, Geological Society of Australia Abstracts, v. 49, p. 409. Sippel, R. E, 1968. Sandstone petrology evidence from luminescence petrography: Journal of Sedimentary Petrology, v. 38, p. 530-54. Sise, J. R., and Jack, D. J., 1984. Exploration case history of the Hellyer Deposit, in Baillie, P. W, and Collins, P. L. E, eds., Mineral exploration and tectonic processes in Tasmania: Burnie, Australia, November 1984, Geological Society of Australia, Tasmanian Division, Abstract volume and excursion guide, p. 48—9. Sivell, W. J., 1984. Low-grade metamorphism of the Brook Street Volcanics, D'Urville Island, New Zealand: New Zealand Journal of Geology and Geophysics, v. 27, p. 167—90. Skirrow, R. G., and Franklin, J. M., 1994. Silicification and metal leaching in semi-conformable alteration beneath the Chisel Lake massive sulfide deposit, Snow Lake, Manitoba: Economic Geology, v. 89, p. 31-50. Smith, R. L., I960. Ash flows: Geological Society of America Bulletin, v. 71, p. 795-842. Smith, R. E., 1968. Redistribution of major elements in the alteration of some basic lavas during burial metamorphism: Journal of Petrology, v. 9, p. 191-219. Smith, R. E., 1969. Zones of progressive regional burial metamorphism in part of the Tasman Geosyncline, eastern Australia: Journal of Petrology, v. 10, p. 144-63.

REFERENCES | 2 6 7 Smith, R. E., 1974. The production of spilitic lithologies by burial metamorphism of flood basalts from the Canadian Keweenawan, Lake Superior, in Amstutz, G. C, ed., Spilites and spilitic rocks: New York, USA, Springer-Verlag, p. 403— 16. Smith, R. E., 1977. Petrography and geochemistry of epidote alteration patches in gabbro dykes at Matagami, Quebec - discussion: Canadian Journal of Earth Sciences, v. 14, p. 505-7. Smith, R. B., 1991. Diagenesis and cementation of lower Miocene pyroclastic sequences in the Sulu Sea, Sites 768, 769 and 771: Proceedings of the Ocean Drilling Program, Scientific Results, v. 124, p. 181-99. Smith, R. E., and Smith, S. E., 1976. Comments of the use ofTi, Zr, Y, Sr, K, P and Nb in classification of basaltic magmas: Earth and Planetary Science Letters, v. 32, p. 114—20. Smith, R. E., Perdrix, J. L., and Parks, T. C, 1982. Burial metamorphism in the Hamersley Basin, Western Australia: Journal of Petrology, v. 23, p. 75-102. Solomon, M., 1964. The spilite-keratophyre association of west Tasmania and the ore deposits of Mt Lyell, Rosebery and Hercules: Unpublished PhD thesis, University of Tasmania, 419 p. Solomon, M., 1976. 'Volcanic' massive sulfide deposits and their host rocks — a review and an explanation, in Wolf, K. A., ed., Handbook of stratabound and stratiform ore deposits - Part II, regional studies and specific deposits: Amsterdam, Netherlands, Elsevier, 320 p. Solomon, M., 1981. An introduction to the geology and metallic ore deposits of Tasmania: Economic Geology, v. 76, p. 194— 208. Solomon, M., 1999. 'Mineralisation' use and mis-use: SEG Newsletter, v. 37, p. 48. Solomon, M., and Gaspar, O. C, 2001. Textures of the Hellyer volcanic-hosted massive sulfide deposit, Tasmania — the ageing of a sulphur sediment on the seafloor: Economic Geology, v. 96, p. 1513-34. Solomon, M., and Groves, D. I., 1994. The geology and origin of Australia's mineral deposits: Oxford, UK, Oxford University Press, Monographs in Geology and Geophysics, v. 24, 951 p. Solomon, M., and Khin Zaw, 1997. Formation on the seafloor of the Hellyer volcanogenic massive sulfide deposit: Economic Geology, v. 92, p. 686-95. Solomon, M., and Khin Zaw, 1999. Formation on the seafloor of the Hellyer volcanogenic massive sulfide deposit: Economic Geology, v. 92, p. 686-95. Solomon, M., and Quesada, C, 2003. Zn-Pb-Cu massive sulfide deposits — brine-pool types occur in collisional orogens, black smoker types occur in backarc and/or arc basins: Geology, v. 31, p. 1029-32. Solomon, M., Rafter, T. A., and Jensen, M. L., 1969. Isotope studies on the Rosebery, Mount Farrell and Mount Lyell ores, Tasmania: Mineralium Deposita, v. 4, p. 172—99. Solomon, M., Walshe, J. L., and Palomero, G. F., 1980. Formation of massive sulfide deposits at Rio Tinto, Spain: Institution of Mining and Metallurgy, Transactions, Section B: Applied Earth Science, v. 89, p. 16-24. Solomon, M, Eastoe, C. J., Walshe, J. L, and Green, G. R, 1988. Mineral deposits and sulfur isotope abundances in the Mount Read Volcanics between Que River and Mount Darwin, Tasmania: Economic Geology, v. 83, p. 1307—28. Solomon, M., Tornos, E, and Gaspar, O. C, 2002. Explanation for many of the unusual features of the massive sulfide deposits of the Iberian pyrite belt: Geology, v. 30, p. 87-90. Solomon, M., Gemmell, J. B., and Khin Zaw, 2004. Nature and origin of the fluids responsible for forming the Hellyer Zn-

Pb-Cu, volcanic-hosted massive sulfide deposit, Tasmania, using fluid inclusions, and stable and radiogenic isotopes: Ore Geology Reviews, v. 25, p. 89-124. Sparks, R. S. J., and Wright, J. V, 1979. Welded air-fall tuffs: Geological Society of America Special Paper, no. 180, p. 155-66. Sparks, R. S. J., Stasiuk, M. V., Gardeweg, M., and Swanson, D. A., 1993. Welded breccias in andesite lavas: Journal of the Geological Society of London, v. 150, p. 897-902. Spooner, E. T. C, 1977. Hydrodynamic model for the origin of the ophiolitic cupriferous pyrite ore deposits of Cyprus: London, UK, January 21—22, 1976, Volcanic processes in ore genesis, Proceedings, Institute of Mining and Metallurgy. Spooner, E. T. C, and Fyfe, W S., 1973. Subseafloor metamorphism, heat and mass transfer: Contributions to Mineralogy and Petrology, v. 42, p. 287-304. Spooner, E. T. C, Beckinsale, R. D., Fyfe, W S., and Smewing, J. D., 1974. 8O18 enriched ophiolitic metabasic rocks from E. Liguria (Italy), Pindos (Greece), and Troodos (Cyprus): Contributions to Mineralogy and Petrology, v. 47, p. 41—62. Spry, A., 1976. Metamorphic textures: Exeter, UK, Pergamon Press, 350 p. Stakes, D. S., and Taylor, H. P., Jr., 1992. The northern Samail Ophiolite - an oxygen isotope, microprobe, and field study: Journal of Geophysical Research, v. 97, p. 7043—80. Stanley, C. R., and Madeisky, H. E., 1996. Lithogeochemical exploration for metasomatic zones associated -with hydrothermal mineral deposits using Pearce element ratio analysis: Vancouver, Canada, Mineral Deposit Research Unit, University of British Columbia short course notes, January 1996, 98 p. Stanton, R. L., 1955. Lower Palaeozoic mineralisation near Bathurst, New South Wales: Economic Geology, v. 50, p. 681-714. Stanton, R. L., 1959. Mineralogical features and possible mode of emplacement of the Brunswick Mining and Smelting ore bodies, Gloucester County, New Brunswick: Canadian Mining and Metallurgical Bulletin, v. 570, no. 62, p. 631-42. Stanton, R. L., 1984. The direct derivation of cordierite from a clay-chlorite precursor — evidence from the Geco Mine, Manitouwadge, Ontario: Economic Geology, v. 79, p. 1245— 64. Stanton, R. L., 1985. Stratiform ores and geological processes: Journal and Proceedings of the Royal Society of New South Wales, v. 118, p. 77-100. Stanton, R. L., 1990. Magmatic evolution and the ore type-lava type affiliations of volcanic exhalative ores: Australasian Institute of Mining and Metallurgy Monograph, v. 15, p. 101-7. Starkey, R. J., Jr., and Frost, B. R., 1990. Low-grade metamorphism of the Karmutsen Volcanics, Vancouver Island, British Columbia: Journal of Petrology, v. 31, p. 167-95. Steiner, A., 1953. Hydrothermal rock alteration at Wairakei, New Zealand: Economic Geology, v. 48, p. 1—13. Stolz, A. J., 1995. Geochemistry of the Mount Windsor Volcanics - implications for the tectonic setting of CambroOrdovician volcanic-hosted massive sulfide mineralisation in northeastern Australia: Economic Geology, v. 90, p. 1080-97. Stolz, A. J., Davidson, G. J., and Blake, M., 1996. Barren alteration systems - example, Gydgie Central, Mt Windsor Volcanic Belt, North Queensland: AMIRA/ARC Project P439, Centre for Ore Deposit and Exploration Studies, University of Tasmania, Unpublished Report 3, October 1996, p. 1— 85. Strauss, G. K., Roger, G., Lecolle, M., and Lopera, E., 1981. Geochemical and geologic study of the volcano-sedimentary sulfide ore body of La Zarza, Province of Huelva, Spain: Economic Geology, v. 76, p. 1975-2000. Surdam, R. C, 1968. Origin of native copper and hematite in the

2 6 8 | REFERENCES Karmutsen Group, Vancouver Island, British Columbia: Economic Geology, v. 63, p. 961—6. Surdam, R. C, 1973. Low-grade metamorphism of tuffaceous rocks in the Karmutsen Group, Vancouver Island, British Columbia: Geological Society of America Bulletin, v. 84, p. 1911-22. Surdam, R. C, and Boles, J. R., 1979. Diagenesis of volcanic sandstones: Society of Economic Palaeontologists and Mineralogists Special Publication, v. 26, p. 127-39. Surdam, R. C, and Sheppard, R. A., 1978. Zeolites in saline, alkalinelake deposits, in Sand, L. B., and Mumpton, E A., eds., Natural zeolites — occurrence, properties, use: Oxford, UK, Peregamon Press, p. 145—74. Syddell, M., 2004. Down to earth: Earthmatters (CSIRO Exploration and Mining), v. 5, p. 8-9. Tanimura, S., Date, J.,Takahashi,T., and Ohmoto, H., 1983. Geologic setting of the Kuroko deposits, Japan - Part II, stratigraphy and structure of the Hokuroku district: Economic Geology Monographs, v. 5, p. 24-38. Taranik, J. V., 2001. New applications of remote sensing to mineral exploration [abs.]: Geological Society of America Abstracts with Programs, v. 33, p. 347. Taube, A., 1986. The Mount Morgan gold-copper mine and environment, Queensland — a volcanogenic massive sulfide deposit associated with penecontemporaneous faulting: Economic Geology, v. 81, p. 1322-40. Taylor, J. C. M., 1950. Pore space reductions in sandstones: Bulletin of American Association Petrology Geology, v. 34, p. 701— 6. Taylor, H. P., Jr., 1974. The application of oxygen and hydrogen isotope studies to problems of hydrothermal alteration and ore deposition: Economic Geology, v. 69, p. 843-83 Taylor, H. P., Jr., 1979. Oxygen and hydrogen isotope relationships in hydrothermal mineral deposits, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits, 3 rd edition: New York, USA, John Wiley and Sons, p. 236-77. Taylor, R. G., 1992. Ore textures — recognition and interpretation - Volume 1, infill textures: Townsville, Australia, Economic Geology Research Unit, Geology Department, James Cook University, 24 p. Taylor, R. G., 1996. Ore textures — recognition and interpretation - Volume 3, overprinting textures: Townsville, Australia, Economic Geology Research Unit, Geology Department, James Cook University, 52 p. Taylor, H. P., Jr., and Forester, R. W., 1979. An oxygen and hydrogen isotope study of the Skaergaard intrusion and its country rocks - a description of a 55-m.y. old fossil hydrothermal system: Journal of Petrology, v. 20, p. 355-419. Taylor, M. W., and Surdam, R. C, 1981. Zeolite reactions in the tuffaceous sediments at Teels Marsh, Nevada: Clays and Clay Minerals, v. 29, p. 341-52. Taylor, B., Fujioka, K., Janecek, T. R., Aitchison, J., Cisowski, S., Colella, A., Cooper, P. A., Dadey, K. A., Egeberg, P. K., Firth, J. V, Gill, J. B., Herman, Y., Hiscott, R. N., Isiminger-Kelso, M., Kaiho, K., Klaus, A., Koyama, M., Lapierre, H., Lovell, M. A., Marsaglia, K., Nishimura, A., Pezard, P. A., Rodolfo, K. S., Taylor, R. N., Tazaki, K., and Torssander, P., 1990. Sites 788/789: Proceedings of the Ocean Drilling Program, Initial Reports, v. 126, p. 97-124. Tazaki, K., and Fyfe, W. S., 1992. Diagenetic and hydrothermal mineral alteration observed in Izu-Bonin deep-sea sediments, Leg 126: Proceedings of the Ocean Drilling Program, Scientific Results, v. 126, p. 101-12. Thompson, A. B., 1971. Analcite-albite equilibria at low temperatures: American Journal of Science, v. 271, p. 79— 92. Thompson, G., 1973. A geochemical study of the low-temperature

interaction of seawater and oceanic igneous rocks: EOSTransactions, American Geophysical Union, v. 54, p. 1015— 19. Thompson, A. J. B., and Thompson, J. F. EL, 1996. Atlas of alteration - a field and petrographic guide to hydrothermal alteration minerals: Vancouver, Canada, Geological Association of Canada, 119 p. Thompson, A. J. B., Hauff, P. L., and Robitaille, A. J., 1999. Alteration mapping in exploration - application of shortwave infrared (SWIR) spectroscopy: SEG Newsletter, v. 39, p. 1 and 16-27. Titley, S. R., 1982a. Advances in geology of porphyry copper deposits, southwestern North America: Tucson, USA, University of Arizona Press, 560 p. Titley, S. R., 1982b. The style and progress of mineralisation and alteration in porphyry copper systems; American southwest, in Titley, S. R., ed., Advances in geology of porphyry copper deposits, southwestern North America: Tucson, USA, University of Arizona Press, p. 93-116. Titley, S. R., 1994. Evolutionary habits of hydrothermal and supergene alteration in intrusion-centred ore systems, southwestern North America, in Lentz, D. R., ed., Alteration and alteration processes associated with ore-forming systems, 2: Vancouver, Geological Association of Canada short course notes, May 13-15, 1994, p. 237-260. Titley, S. R., Fleming, A. W., and Neale, T. I., 1978. Tectonic evolution of the porphyry copper system at Yandera, Papua New Guinea: Economic Geology, v. 73, p. 810-28. Tiwary, A., and Deb, M., 1997. Geochemistry of hydrothermal alteration at the Deri massive sulfide deposit, Sirohi District, Rajasthan, NW India: Journal of Geochemical Exploration, v. 59, p. 99-121. Tornos, E, Clavijo, E. G., and Spiro, B., 1998. The Filon Norte ore body (Tharsis, Iberian pyrite belt) — a proximal lowtemperature shale-hosted massive sulfide in a thin-skinned tectonic belt: Mineralium Deposita, v. 33, p. 150-69. Tornos, E, Casquet, C, and Relvas, J. M. R. S., in press. Transpressional tectonics, lower crust decoupling and intrusion of deep mafic sills - a model for the unusual metallogenesis of SW Iberia: Ore Geology Reviews. Torres, M. E., Marsaglia, K. M., Martin, J. B., and Murray, R. W., 1995. Sediment diagenesis in western Pacific basins: AGU Geophysical Monograph, v. 88, p. 241-58. Tsolis-Katagas, P., and Katagas, C, 1989. Zeolites in pre-caldera pyroclastic rocks of the Santorini volcano, Aegean Sea, Greece: Clays and Clay Minerals, v. 37, p. 497-510. Tucker, M. E., 1987. Sedimentary petrology - an introduction: Oxford, UK, Blackwell Scientific Publications, 252 p. Turner, F. J., 1980. Metamorphic petrology —mineralogical, field and tectonic aspects, 2 nd edition: USA, Hemisphere Publishing Corporation, 524 p. Upton, B. G. J., and Wadsworth, W. J., 1970. Early volcanic rocks of Reunion and their tectonic significance: Bulletin Volcanologique, v. 33, p. 1246-68. Urabe, T, 1987. The effect of pressure on the partitioning ratios of lead and zinc between vapour and rhyolite melts: Economic Geology, v. 82, p. 1049-52. Urabe, T, and Sato, T., 1978. Kuroko deposits of the Kosaka Mine, northeast Honshu, Japan - products of submarine hot springs on Miocene seafloor: Economic Geology, v. 73, p. 161-79. Urabe, T, Scott, S. D., and Hattori, K., 1983. A comparison of footwall-rock alteration and geothermal systems beneath some Japanese and Canadian volcanogenic massive sulfide deposits: Economic Geology Monographs, v. 5, p. 345—64. Utada, M., 1970. Occurrence and distribution of authigenic zeolites in the Neogene pyroclastic rocks in Japan: Scientific Papers

REFERENCES I 2 6 9 of the College of General Education, University of Tokyo, v. 20, p. 191-262. Utada, M., 1973. The types of alteration in the Neogene sediments relating to the intrusion of volcano-plutonic complexes in Japan: Scientific Papers of the College of General Education, University of Tokyo, v. 23, p. 167—216. Utada, M., 1991. Zeolitization in the Neogene formations of Japan: Episodes, v. 14, p. 242-5. Utada, M., Minato, H., Ishikawa, T, and Yoshizaki, Y., 1974. The alteration zones surrounding Kuroko deposits in NishiAizu District, Fukushima Prefecture, with emphasis on the analcime zone as an indicator in exploration for ore deposits: Mining Geology, v. 6, p. 291-302. Utada, M., Aoki, M., Inoue, A., and Kusakabe, H., 1988. A hydrothermal alteration envelope and alteration minerals in the Shinzan mineralised area, Akita Prefecture, northeast Japan: Mining Geology Special Issue, v. 12, p. 67—77. Viereck, L. G., Griffin, B. J., Schmincke, H-U., and Pritchard, R. G., 1982. Volcaniclastic rocks of the Reydarfjordur drill hole, eastern Iceland - Part 2, alteration: Journal of Geophysical Research, v. 87, p. 6459-76. von Damm, K. L., 1995. Controls on the chemistry and temporal variability of seafloor hydrothermal fluids: AGU Geophysical Monograph, v. 91, p. 222-47. Walshe, J. L., 1986. A six-component chlorite solid solution model and the conditions of chlorite formation in hydrothermal and geothermal systems: Economic Geology, v. 81, p. 681— 703. Walshe, J. L., and Solomon, M., 1981. An investigation into the environment of formation of the volcanic-hosted Mt Lyell copper deposits using geology, mineralogy, stable isotopes and a six-component chlorite solid solution model: Economic Geology, v. 76, p. 246-84. Walton, A. W, 1975. Zeolitic diagenesis in Oligocene volcanic sediments, Trans-Pecos Texas: Geological Society of America Bulletin, v. 86, p. 615-24. Waring, C. L., Andrew, A. S., and Ewers, G. R., 1998. Use of O, C, and S stable isotope in regional mineral exploration: AGSO Journal of Australian Geology and Geophysics, v. 17, p. 301-13. Waters, J. C, and Wallace, D. B., 1992. Volcanology and sedimentology of the host succession to the Hellyer and Que River volcanic-hosted massive sulfide deposits, northwestern Tasmania: Economic Geology, v. 87, p. 650-66. Weaver, S. D., Gibson, I. L, Houghton, B. E, and Wilson, C. J. N., 1990. Mobility of rare earth and other elements during crystallization of peralkaline silicic lavas: Journal of Volcanology and Geothermal Research, v. 43, p. 57-70. Webster, S. S., and Skey, E. H., 1979. Geophysical and geochemical case history of the Que River deposit, Tasmania, Australia: Ottawa Geological Survey of Canada Economic Geology Report, no. 31, p. 697-720. Weihed, P., Allen, R. L., and Svenson, S. A., 2000. Metallogeny and tectonic evolution of the c. 1.9 Ga Skellefte marine, volcanic arc, northern Sweden [abs.], in Gemmell, J. B., and Pongratz, J., eds., Volcanic environments and massive sulfide deposits, Hobart, Australia, November 16-19, 2000, CODES Special Publication, v. 3, p. 225-226. Whalen, J. B., McNicoll, V. J., Galley, A. G., and Longstaffe, F. J., 2004. Tectonic and metallogenic importance of an Archaean composite high- and low-Al tonalite suite, Western Superior Province, Canada: Precambrian Research, v. 132, p. 275— 301. Whetten, J. T, and Hawkins, J. W, 1970. Diagenetic origin of greywacke matrix minerals: Sedimentology, v. 15, p. 34761. White, D. E., 1970. Geochemistry applied to the discovery,

evaluation and exploitation of geothermal energy resources: Geothermatics Special issue 2, v. 1, p. 58-80. White, M. J., 1996. Stratigraphy, volcanology and sedimentology of the Cambrian Tyndall Group, Mount Read Volcanics, western Tasmania: Unpublished PhD thesis, University of Tasmania, 196 p. White, J. D. L., 2000. Subaqueous eruption-fed density currents and their deposits: Precambrian Research, v. 101, p. 87-109. White, N. C, andHedenquist,J. W, 1990. Epithermalenvironments and styles of mineralisation — variations and their causes, and guidelines for exploration: Journal of Geochemical Exploration, v. 36, p. 445-74. White, M. J., and McPhie, J., 1996. Stratigraphy and palaeovolcanology of the Cambrian Tyndall Group, Mt Read Volcanics, western Tasmania: Australian Journal of Earth Sciences, v. 43, p. 147-59. White, M. J., and McPhie, J., 1997. A submarine welded ignimbritecrystal-rich sandstone facies association in the Cambrian Tyndall Group, western Tasmania, Australia: Journal of Volcanology and Geothermal Research, v. 76, p. 277-95. White, D. E., and Sigvaldason, G. E., 1962. Epidote in hot-spring systems, and depth of formation of propylitic epidote in epithermal ore deposits: United States Geological Survey Professional Paper 450E, p. 80-4. Whitmore, D. R. E., 1969. Geology of the Coronation copper deposit: Geological Survey of Canada Paper, v. 68-5, p. 37— 53. Whitten, D. G. A., and Brooks, J. R. V, 1972. The Penguin dictionary of geology: London, UK, Penguin Books, 493 p. Williams, N. C, 2000. The Basin Lake high-sulfidation alteration system, western Tasmania: Unpublished B.Sc. Honours thesis, University of Tasmania, 102 p. Williams, N. C, and Davidson, G., 2004. Possible submarine advanced argillic alteration at the Basin Lake prospect, western Tasmania, Australia: Economic Geology, v. 99, p. 987-1002. Williams, D., Stanton, R. L., and Rambaud, E, 1975. The PlanesSan Antonio pyritic deposit of Rio Tinto, Spain — its nature, environment and genesis: Institution of Mining and Metallurgy Transactions, v. 84, p. B73-82. Williams, H., Turner, F. J., and Gilbert, C. M., 1982. Petrography — an introduction to the study of rocks in thin section, 2 nd edition: San Fransico, USA, W H . Freeman and Co., 626 p. Williams, E., McClenaghan, M. P., and Collins, P. L. E, 1989. Mid-Palaeozoic deformation, granitoids and ore deposit, in Burrett, C. F., and Martin, E. L., eds., Geology and mineral resources of Tasmania: Geological Society of Australia Special Publication, v. 15, p. 238-92. Wilshire, H. G., 1959. Deuteric alteration of volcanic rocks: Journal and Proceedings, Royal Society of New South Wales, v. 93, p. 105-20. Wilshire, H. G., and Hobbs, B. E., 1962. Structure, sedimentary inclusions, and hydrothermal alteration of a latite intrusion: Journal of Geology, v. 70, p. 328—41. Wilson, A. J., Cooke, D. R., and Harper, B. L., 2003. The Ridgeway gold-copper deposit - a high-grade alkalic porphyry deposit in the Lachlan Fold Belt, New South Wales, Australia: Economic Geology, v. 98, p. 1637—66. Winkler, H. G. E, 1979. Petrogenesis of metamorphic rocks, 5th ed.: New York, USA, Springer-Verlag, 348 p. WoldeGabriel, G., Keating, G. N., and Valentine, G. A., 1999. Effects of shallow basaltic intrusion into pyroclastic deposits, Grants Ridge, New Mexico, USA: Journal of Volcanology and Geothermal Research, v. 92, p. 389-411. Wyman, W. E, 2001. Cambrian granite-related hydrothermal alteration and Cu-Au mineralisation in the southern Mt Read Volcanics, western Tasmania, Australia: Unpublished

2 7 0 | REFERENCES PhD thesis, University of Tasmania, 343 p. Yamagishi, H., 1987. Studies on the Neogene subaqueous lavas and hyaloclastites in southwest Hokkaido: Geological Survey of Hokkaido Report, v. 59, p. 55-117. Yamagishi, H., and Dimroth, E., 1985. A comparison of Miocene and Archaean rhyolite hyaloclastites — evidence for a hot and fluid rhyolite lava: Journal of Volcanology and Geothermal Research, v. 23, p. 337-55. Yamagishi, H., and Goto, Y, 1992. Cooling joints of subaqueous rhyolite lavas at Kuroiwa, Yakumo, southern Hokkaido, Japan: Bulletin of the Volcanological Society of Japan, v. 37, p. 205-7. Yang, K., 1998. Compositional variations of white micas - a literature review: North Ryde, CSIRO, 23 p. Yang, J., and Large, R. R., 2001. Computational modelling of hydrothermal ore-forming fluid migration in complex earth structures, in Xie, H., Wang, Y., and Jiang, Y, eds., Proceedings of the International Symposium on Computer Applications in the Mineral Industries, Beijing, China, April 25-27, 2001, v. 29, p. 115-20. Yang, K., and Scott, S. D., 1996. Possible contribution of a metalrich magmatic fluid to a seafloor hydrothermal system: Nature, v. 383, p. 420-3. Yang, K., Huntington, J. E, Boardman, J. W., and Mason, P., 1999. Mapping hydrothermal alteration in the Comstock mining

district, Nevada, using simulated satellite-borne hyperspectral data: Australian Journal of Earth Sciences, v. 46, p. 420-3. Yardley, B. W. D., 1989. An introduction to metamorphic petrology: Harlow, UK, Longman Scientific and Technical, 248 p. Ylagan, R. E, Altaner, S. P., and Pozzuoli, A., 1996. Hydrothermal alteration of a rhyolitic hyaloclastite from Ponza Island, Italy: Journal of Volcanology and Geothermal Research, v. 74, p. 215-31. Yoshida, K., and Utada, M., 1968. A study on alteration of Miocene green tuff in the Kuroko-type mineralisation area in Odate basin, Akita Prefecture: Mining Geology, v. 18, p. 333-42. Yu, Z., Robinson, P., and McGoldrick, P. J., 2001. An evaluation of methods for the chemical decomposition of geological materials for trace element determination using ICP-MS: Geostandards Newsletter, v. 25, no. 2-3, p. 199-217. Zhang, X-C, Spiro, B., Halls, C, Stanley, C. J., and Yang, K-Y, 2003. Sediment-hosted disseminated gold deposits in Southwest Guizhou, PRC — their geological setting and origin in relation to mineralogical, fluid inclusion, and stable-isotope characteristics: International Geology Review, v. 45, p. 407-70. Zhou, Z., and Fyfe, W. S., 1989. Palagonitization of basaltic glass from DSDP Site 335, Leg 37 - textures, chemical composition and mechanism of formation: American Mineralogist, v. 74, p. 1045-53.

I 271

INDEX

A page number in bold indicates that the reference is to a figure. A bold t indicates that the reference is to a table.

AI-CCPI Alteration box plot see Alteration box plot albite alteration 39, 42-44, 61, 65, 69, 133-5, 165-7, 191, 214,216 Alteration box plot 31-4, 36, 169-70, 245 alteration distribution 22, 63-4 alteration facies describing and defining 15—36 diagenetic 241—2 distribution 63-4 hydrothermal 242-3 metamorphic 242 variables 22 alteration fluids 93, 170-3 Alteration Index (AI) 30-2, 34, 169, 245 alteration indices 26-30, 34, 73, 169-70, 244^5 alteration intensity describing 27t estimation, integrated approach 33—6 explained 25—36 illustrated 28-29, 36 lithogeochemical indications of 243-4 alteration mineral assemblages see mineral assemblages alteration nomenclature 16, 19-22, 2It, 22t, 23t alteration pipes 63, 164-74, 176, 182-3 alteration plumes 63, 167, 168, 191-3 alteration processes 4—6 alteration rates 6 alteration textures deformation textures 52, 54-6, 55-7 described 37-63, 37t, 38t, 62t dissolution textures 41, 50-1, 52 dynamic recrystallisation textures 52 illustrated 39-40, 62, 103-4, 110, 111-13 infill textures 41,48-9 overprinting and false/pseudo textures 37, 54-63, 58-61, 62t recrystallisation textures 52, 53 replacement textures 37-8, 41, 42-7 static recrystallisation textures 52, 53

alteration timing 69-71, 7It, 172 alteration zonation boundaries 64 contact altered zones 5, 64, 66-7, 67, 139, 149-56, 242 diagenetic 64, 105-8 facies model 3 greenschist facies zones 115, 116, H6t, 131, 142, 144-5, 152 halos 66-9 Hellyer deposit 178,182-3,184-93 Henty deposit 178, 212-3, 214-20 Highway Reward deposit 14, 167, 178, 232, 233-40 Hokuroku Basin 119-27 hydrothermal 5-6, 66-8, 164-9, 243 mapping 243 metamorphic 64—5, 66—7 Mount Read Volcanics 128-38 patterns 64-9,98, 165-8 regional deep semi-conformable 66, 142-8 regional metamorphic 115-17,140 Rosebery deposit 178,195,196-201 scales described 64t Thalanga deposit 14, 178, 222, 223-31 veins and fractures 67, 69 Western Tharsis deposit 202-3,204-11 amphibolite facies 115-17,140 Amulet deposit see Noranda district analytical techniques electron microprobe 19,25,88 field observations 18 ICP-AES (inductively coupled plasma atomic emission spectrometry) 76-7 ICP-MS (inductively coupled plasma mass spectrometry 76—7 isotope geochemistry 92—5 HyMap® 246, 247 lithogeochemical sampling 73-87 mineral chemistry analysis 87-91 NAA (neutron activated analysis) 76 petrography 24-5, 33-4 PIMA 25, 33, 245

2 7 2 | INDEX

SWIR spectroscopy 19, 24, 25, 33, 88, 90, 202-3, 243, 245 X-ray diffraction (XRD) 19, 24, 25, 33, 88 X-ray fluorescence spectrometry (XRF) 76-7 anhydrous minerals 97

B Bathurst mining camp 164-5,170 Boco prospect 248, 250 burial-related alteration 97-138 see also diagenesis and submarine environments

c carbonates diagenetic 105 in exploration 90-1, 243-4 hydrothermal 47, 91, 166, 178, 188, 201, 217, 229-30 cataclastic texture 52 CCPI (chlorite-carbonate-pyrite index) see also Alteration box plot explained 31, 34 exploration, uses for 169,245 cementation 97, 102, 105, 108-10, 132 Central Volcanic Complex see Mount Read Volcanics chlorite 19-22, 89, 138, 156, 165-8, 187, 239 closure and alteration indices 26 constant sum effect 243—4 explained 78-9,243-5 mass change anomalies 81 compaction 97, 109-10, 132 compositional nomenclature 20 contact alteration 5, 149-62 corrosion vugs 41, 50—1, 52 crystallisation primary 4 textures 52 of zeolite assemblages 105, 110, 114, 121

D Darwin Granite see Mount Read Volcanics data sheets contents of 36 Darwin Granite 157-62 Hellyer deposit 184-93 Henty deposit 214—20 Highway-Reward deposit 233-40 Hokuroku Basin 122-7 Mount Read Volcanics 133-8 Rosebery deposit 196-201 Thalanga deposit 223-31 Western Tharsis deposit 204—11 deep, semi-conformable altered zones 142—6 deformation textures 52, 54-6, 55-7 detection limit explained 75 deuteric alteration 148 devitrification explained 4 texture 37, 39, 62 zones 151

diagenesis explained 5 Hokuroku Basin 118-27 isotope geochemistry analysis 93-4 and metamorphism 16, 98, 102, 114, 115 Mount Read Volcanics 128-38 in submarine volcanic successions 97, 102—14 diagenetic minerals carbonates 105 genesis of 108-14 layered silicates 102,105 other diagenetic minerals 105 zeolites 105, 110, 118, 120-7 diagenetic zones Hokuroku Basin 118—27 Mount Read Volcanics 128-38 zonation 64-5, 105-8 discharge zone 141—2 dissolution 41, 50-1, 52, 97, 102, 108-10, 114, 132 dynamic recrystallisation textures 52

electron microprobe see analytical techniques element concentrations 32-3 eutaxitic texture 54, 57 exploration Alteration box plot 31-4,169-70 alteration identification as tool in 241—50 Alteration Index (AI) 30-2 isotope geochemistry in 92-5, 246-50 lithogeochemistry in 73-87 mineral chemistry in 87-91 sulfide mapping 243 use of chlorite in 89 use of white mica in 90—1 vectors and proximity indicators 32-3, 94-5, 243-50

false textures see pseudotextures fiamme 54, 57 fluid-rock interaction 169, 170-2, 171, 173 foliation 52, 54, 70, 711 footwall alteration 163-74, 179, 182-9 fused zones 151

geochemistry see isotope geochemistry, lithogeochemistry geothermal gradient 98 geothermometers 92-3 glass alteration in submarine volcanic successions 15, 97-8 common alteration minerals 19t crystallisation 4 diagenesis 102—14 disequilibrium assemblages 24 hydration 4-5, 98-102 reactive quality, 6 Green Tuff Belt see Hokuroku district

INDEX | 2 7 3

H halos 5-6, 38, 41, 66-9, 149-56, 157-62, 163-74, 17881 hanging wall alteration 163, 164, 167-8, 190-3 Hellyer deposit see also Mount Read province alteration fades and zonation 178, 182-3, 184—93 Alteration Index (AI) 31, 183 explained 11, 181-93 exploration 245 geological setting 181—2 ore genesis 170, 183 white mica 90 Henty deposit see also Mount Read province alteration facies and zonation 178, 212-13, 214-20 explained 12 geological setting 212 hanging wall alteration 167 isotopic data 248 ore genesis 213 Hercules deposit see also Mount Read province alteration halo 178 explained 12, 128 geological setting 194 Highway-Reward deposit see also Mount Windsor Subprovince alteration facies and zonation 14, 167, 178, 232,

233-40 explained 14 geological setting 12-14, 232 hanging wall alteration 167,178 ore genesis 232 submarine facies associations 13 Hokoroku Basin see Hokuroku district Hokuroku district alteration 64-5, 67, 118-27, 122-7 geological setting 118—20 Green Tuff Belt 52, 64, 107, 118, 150, 151 Kuroko deposits see Kuroko deposits oxygen isotopes 94-5, 246 size of VHMS deposits 164 hydration of volcanic glass 4-5, 98-102 hydrothermal alteration boundaries between zones 64 chemical reactions 168 and diagenetic alteration 128 discharge zone 141-2,156 discriminating 16—19 explained 4 halos 66-7, 164-74 intensity measures 32-3 intrusion-related 140-61 see also intrusions metamorphic assemblages 174, 175t plagioclase destruction 31, 167 recharge zone 141 subseafloor systems 140—1 syntectonic 6 tectonic deformation 6 VHMS deposits 5-6, 163-240 zones 5-6, 66-8, 164-78, 243 hydrothermal convection 1, 94, 140-1, 140

hydrothermal fluid 67, 172-3 HyMap® 246, 247

Iberian pyrite belt 90, 91, 142, 164, 165, 166, 166, 174 ICP-AES (inductively coupled plasma atomic emission spectrometry) 76—7 ICP-MS (inductively coupled plasma mass spectrometry 76-7 indices alteration 26-30, 34, 73, 244-5 Alteration box plot see Alteration box plot Alteration Index (AI) 30-2, 34, 169, 245 CCPI (chlorite-carbonate-pyrite index) 31, 34, 169, 245 molar proportion alteration 30 multi-component and normalised 26, 30 simple ratio 26 induration 150, 151 infill textures 41,48-9 intrusions halos 66—7 cryptodomes 2, 6, 66, 128, 139, 143, 153, 212, 232 dykes 2, 6, 66, 141, 139, 149, 152, 153 plutons 139, 143, 153 sills 66, 70, 100, 128, 148-9, 152-3, 150, 152, 174, 182, 194,202,212,221,232 in submarine volcanic successions 2 , 3 , 139 synvolcanic 139—62 isotope geochemistry applications 92-5 carbon 76, 77, 248 exploration 94-5, 246-50 hydrogen 76, 77, 93 oxygen 94-5, 246,248-50 stable isotopes 92-5 sulfur 76, 77-8, 92, 141, 180, 246, 248 water-rock ratios 93

K kaolinite in VHMS altered zones 88, 150, 174-75, 178-80 keratophyre 98 K-lens see Rosebery deposit Kuroko deposits see also Hokuroku district alteration 164, 179 Alteration Index (AI) 31 alteration model 166, 178—9

least altered see also alteration intensity explained 26 alteration indices 32 lithification 97, 108-10 lithogeochemistry analytical methods 73-8 carbonates 76-7 chemostratigraphy 79-81 C-H-N elemental analyser 76 closure 78-9, 81 compatible elements 79—80

2 7 4 I INDEX

europium 87 explained 73-8, 74t and exploration 243-5 hydrous minerals 76-7 ICP-AES (inductively coupled plasma atomic emission spectrometry) 76-7 ICP-MS (inductively coupled plasma mass spectrometry 76-7 immobile elements 79-81, 85, 87 inaccuracies in 77 incompatible elements 79 limit of detection 75 LOI (loss on ignition) 76 mass change 73, 81-7, 85-6, 87, 97, 165, 180-1, 245-6 NAA (neutron activated analysis) 76 precision and accuracy required 75 recalculating to volatile free 77-8 REE (rare earth elements) 73, 79, 81, 87 reporting data 77 sampling methods 73—8 summing elements 77 use of reference materials 75 XRF (X-ray fluorescence spectrometry) 76-7

M mass change see lithogeochemistry massive sulfide 163—5, 167 Mattabi deposit 91, 180 metamorphism burial metamorphism and diagenesis 16, 97, 98, 102, 114, 115-17 contact metamorphism 5, 11, 12, 64, 66-7, 149-54, 242 explained 4, 5, 20, 24 regional metamorphism 5, 64-6, 115, 139, 140-8 of VHMS-related altered zones 174-5 metasomatic alteration 4, 5, 144-6 microanalysis 24 microprobe see analytical techniques mineral assemblages alteration assemblages 23—5, 34, 165 burial effects on 97, 109 common assemblages 21t, 22t, 109, 165 disequilibrium 23-4, 63 equilibrium 23-4 in exploration 243 igneous 16, 19 isotopic studies of 92—5 nomenclature 19-20 mineral chemistry 87-91, 245-6 minerals defined 87-8 Mount Lyell field see also Mount Read province deposits 12,202,243 geological setting 202 halos 202 hanging wall alteration 167 HyMap system 246, 247 mineral zonation 243, 247 ore genesis 203 Western Tharsis deposit see Western Tharsis deposit

Mount Read province alteration 7-12, 163-164 Chester deposit 90 history 9, 11-12 Hellyer deposit see Hellyer deposit Henty deposit see Henty deposit Hercules deposit see Hercules deposit Mount Lyell field see Mount Lyell field Mount Read Volcanics see Mount Read Volcanics oxygen isotopic exploration 249—50 Que River deposit see Que River deposit Rosebery deposit see Rosebery deposit size of VHMS deposits 164 Western Tharsis deposit see Western Tharsis deposit Mount Read Volcanics see also Mount Read province AI and CCPI ranges 32,34 alteration 128-32, 133-8 Central Volcanic Complex 9-10, 69, 128, 130-32,

157-62 chemostratigraphic discrimination and correlation 80 compaction effects at 110 Darwin Granite 154-6, 157-62 geology of 7-12, 128, 129 Kershaw Pumice Formation 128,134—6 Mount Black Formation 128, 133, 137-8 pyritic alteration systems 89 metamorphic assemblages 11 Sterling Valley Volcanics 128 Mount Windsor Subprovince alteration 14, 164, 222, 232 geology 12-14, 221-40 Highway-Reward deposit see Highway-Reward deposit Thalanga deposit see Thalanga deposit

N NAA (neutron activated analysis) 76 see also analytical techniques naming altered rock see alteration nomenclature Noranda district 66, 142-7, 167, 179-81 numerical fluid-flow modelling 172

o overprinting textures 37, 70-1, 70t, 711 and false/pseudo textures 37, 54-63, 58-61, 62t relationships 69-71, 7It

palagonite 99, 99-100 Panorama district 180, 246, 248-50 paragenetic sequence 69 perlite 37,40,54, 100-1, 101 PIMA (portable infrared mineral analyser) 25, 33, 245 see also analytical techniques plagioclase destruction 31,167-9 pseudotextures 16, 37, 54-63, 58-61, 62t, 63 pyrophyllite in VHMS systems 174, 202, 207-8

Que River deposit 11-12, 70, 90, 167 see also Mount Read province

INDEX I 2 7 5

R regional metamorphism see metamorphism relict textures 14, 16, 19, 24, 25, 37-38, 141 replacement textures 37-38, 41, 42-7 Rio Tinto deposit see Iberian pyrite belt Rosebery deposit see also Mount Read province alteration 70, 128, 131t, 168, 173, 178, 195, 196201 Chlorite-carbonate-pyrite index (CCPI) 195 geochemical alteration parameters 168 geological setting 194 explained 11-12, 194-201 hydrothermal carbonates 91 hydrothermal fluid flow 174 K-lens 30, 34, 68, 91, 174, 244 lithogeochemical data in exploration 82, 244 sericite 70-1

Scuddles deposit 180 sericite alteration 20-2, 70-1, 165-78, 185-6, 198-9, 20611, 218,225, 231, 238 see also white mica Short-wavelength infrared spectroscopy (SWIR) see analytical techniques siliceous alteration 165-9, 182, 189, 199, 219-20 smectites 102, 109, 117 Snow Lake District 142-6 sodium content in volcanic rocks 34—6, 243—4 solution seams 51, 52 spilite 98 stable isotopes see isotope geochemistry stockwork see stringer zones stringer zones 163, 165, 172-3, 179 stylolites 51, 52 submarine environments 2, 97—138 submarine facies associations Mount Read Volcanics 10-11, 214-20 Seventy Mile Range Group 12-13 submarine volcanic successions 1-14, 3, 97-138, 241-9 SWIR spectroscopy see analytical techniques synvolcanic intrusions see intrusions

Thalanga deposit see also Mount Windsor Subprovince alteration facies and zonation 14, 178, 222, 223-31 explained 14 geological setting 12—14, 221—2 mass change estimations 85—6 ore genesis 222 oxygen isotopic exploration 250 thermodynamic alteration model 170-2

vectors exploration 243—50 isotopic 246—50 lithogeochemical 243-4 mass change 245 mineral chemistry 79, 89-91, 245-6 sulfide 243-4

vein-halo alteration 38,41,43,69 VHMS deposits alteration patterns 164—78 classification 163—6 common features 163—4 comparisons 178-81 exploration 241-50 footwall alteration 163-74 halos 5-6, 67-9, 68, 163-8, 175, 178-81 hanging wall alteration 163—4, 167—8 Hellyer deposit see Hellyer deposit Henty deposit see Henty deposit Highway-Reward deposit see Highway-Reward deposit kaolinite, presence of 174 major VHMS provinces 164 Mount Lyell see Mount Lyell field pyrite in 243 pyrophyllite, presence of 174 regional alteration zones 5-6, 139-62 Rosebery deposit see Rosebery deposit Thalanga deposit see Thalanga deposit Western Tharsis deposit see Western Tharsis deposit VMS deposits see VHMS deposits volcanic facies alteration processes 2—6 alteration in submarine environments 97—138 associations 2 changes in 108 clastic facies 1,19 coherent facies 1, 6, 16, 19 common clay minerals in 102 crystalline facies 6 volcaniclastic facies 1, 6, 12, 16—19 volcanic-hosted massive sulphide deposits see VHMS deposits vugs 41,50-1,52

W water-rock ratios 93—4 Western Tharsis deposit see also Mount Read province see also Mt Lyell field alteration facies and zonation 202—3, 204—11 case study 202-11 geological setting 202 ore genesis 203 white mica 91 white mica 20, 89-91, 165, 246 Woodlawn deposit 178

X X-ray diffraction 19, 24, 25, 33, 88 X-ray fluorescence spectrometry (XRF) 76—7

zeolites 105-6, 106t, 110, 114-16, 118, 120-1 zones see alteration zonation

About the authors

Dr Cathryn Gifkins is a Research Fellow at the Centre for Ore Deposit Research at the University of Tasmania. Cathryn brings to the publication a strong background in mapping, describing and interpreting altered and deformed volcanic rocks in submarine successions. Her current research focusses on the textural, mineralogical and compositional effects of alteration in glassy volcanic rocks, the link between volcanic centres and mineralising hydrothermal systems, and the facies architecture and stratigraphy of the Mount Read Volcanics. ' Walter Herrmann is a Research Fellow in economic geology at the Centre for Ore Deposit Research. Wally's background in mineral exploration in Australian volcanic successions, principally the Mount Read Volcanics and the Mount Windsor Subprovince is a valuable asset to the book. He has a special interest in understanding hydrothermal alteration as a method for discriminating and discovering VHMS and porphyry deposits. Professor Ross Large is Director of the Centre for Ore Deposit Research and has a long and celebrated academic and exploration career. Ross has a comprehensive knowledge of VHMS deposits, and has actively promoted and developed the application of geochemical techniques to mineral exploration. This innovative approach has recently produced the Alteration box plot, an alternative way to relate alteration intensity, mineralogy and geochemistry.