Formation Of Podiform Chromitite Deposits-implications From Pge

Chemical Geology 208 (2004) 217 – 232 www.elsevier.com/locate/chemgeo

Formation of podiform chromitite deposits: implications from PGE abundances and Os isotopic compositions of chromites from the Troodos complex, Cyprus Anette Bu¨chl a,b,*, Gerhard Bru¨gmann a,c, Valentina G. Batanova a,d a Max-Planck Institut fu¨r Chemie, Postfach 3060, 55020 Mainz, Germany Institut fu¨r Mineralogie, Universita¨t Mainz, Becherweg 21, 55128 Mainz, Germany c Institut fu¨r Mineralogie, Universita¨t Mu¨nster, Corrensstr. 24, 48149 Mu¨nster, Germany d Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences, Kosygin Street 19, Moscow 117975, Russia b

Abstract Podiform chromitite deposits occur in the mantle sequences of many ophiolites that were formed in supra-subduction zone settings (SSZ). We have measured PGE abundances and Os isotopic compositions of three major chromitite deposits (Kannoures, Hadji Pavlou, Kokkinorostos) and associated mantle peridotites from the Troodos Ophiolite Complex in order to investigate the petrogenesis of these rocks, and their genetic relationships and to examine the geochemical behaviour of the PGE. Spinels from the chromitite deposits have flat chondrite-normalized PGE patterns, but have distinct negative Pt anomalies. Thus, Pd, Os, Ru and Ir concentrations are very high compared to the Pt concentrations (Os: 13.7 – 104 ng/g, Ir 11.3 – 19.0 ng/g, Ru 34.3 – 83.6 ng/g, Pt 0.41 – 9.07 ng/g, Pd 11.1 – 76.8 ng/g). With the exception of Pd, this appears to be a general feature of chromitites from ophiolites worldwide. However, Pd concentrations determined in this study are high compared to other studies where whole rock samples were analysed. There is no simple explanation for this difference because mass balance constraints would not allow that this is solely due to Pd-depletion in the interstitial component. Rather, it implies that chromitites display large variations of relative PGE abundances, even on a local scale. Podiform chromitite deposits form as a result of the interaction of fluid-rich, percolating melts with surrounding mantle peridotites. Osmium, Ir, Ru and Cr concentrations decrease systematically from harzburgite to dunite surrounding the deposits. In addition, dunites and chromites have complementary PGE distribution patterns. Thus, the mantle peridotite is the source of these metals in chromitites. This also indicates that these elements behave incompatibly and are mobilized during continuous melt percolation. However, the low Pt concentrations in the chromitites suggest that Pt is not as effectively mobilized during melt percolation. Uniformly high Pt concentrations in harzburgite and dunite (ca. 11 ppb) also imply that most Pt remains in the mantle peridotite. This can be explained if residual Pt-rich phases, such as PtFe alloys, limit the mobility of Pt. PGE and Cr become concentrated when chromite and sulfide liquids precipitate as a result of the mixing of percolating melts in magma pools near the crust – mantle boundary. The 187Os/188Os ratios of the chromite separates (0.1265 – 0.1301) are less variable than those of the associated peridotites (0.1235 – 0.1546). The average isotopic composition of the chromites (187Os/188Os: 0.1284 F 0.0021) is superchondritic compared with the carbonaceous chondrite value (187Os/188Os: 0.1260 F 0.0013 after Geochim. Cosmochim. Acta 66 (2002)

* Corresponding author. Current address: Department Earth Sciences, University of Bristol, Wills Memorial Building, Queens Road, Bristol BS8 1RJ, UK. E-mail address: [email protected] (A. Bu¨chl). 0009-2541/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2004.04.013

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329; Geochim. Cosmochim. Acta 66 (2002) 4187) and similar to the average value measured for podiform chromitites worldwide (0.12809 F 0.00085 after Geochim. Cosmochim. Acta 66 (2002) 329; Geochim. Cosmochim. Acta 66 (2002) 4187). Radiogenic melts/fluids derived from the subducting slab trigger partial melting in the overlying mantle wedge and add significant amounts of radiogenic Os to the peridotites. Mass balance calculations suggest that a melt/rock ratio of approximately 17:1 (melt: 187Os/188Os: 0.163, Os: 0.01 ng/g, mantle peridotite: 187Os/188Os: 0.127, Os 4.2 ng/g) is necessary in order to increase the Os isotopic composition of the chromitite deposits to its observed average value. This value implies a surprisingly low average melt/rock ratio during the formation of chromitite deposits. The percolating melts likely have variable isotopic composition and PGE concentration. However, in the chromitite pods the Os from many melts is pooled and homogenized, which is the reason why the chromitite deposits show such a small variation in their Os isotopic composition. The results of this study suggest that the 187Os/188Os ratio of chromitites is not representative for the DMM, but only reflects an upper limit. D 2004 Elsevier B.V. All rights reserved. Keywords: Podiform chromitite deposits; Os isotopes; Platinum-group elements; Ophiolite; Troodos

1. Introduction The formation of podiform chromitite deposits has been discussed for many decades (e.g. Lago et al., 1982; Paktunc, 1990; Prichard and Lord, 1990; McElduff and Stumpfl, 1990; Zhou et al., 1994). A general consensus has emerged that such deposits form in the mantle section of ophiolites from SSZ environments during melt/rock or melt/melt interaction (Zhou and Robinson, 1997; Ballhaus, 1998; Zhou et al., 1994, 1996, 1998; Melcher et al., 1999). These processes mobilize Cr and platinum-group elements (PGEs), and these elements are subsequently concentrated again during pooling of the percolating melts and fluids (e.g. Irvine, 1977; Matveev and Ballhaus, 2002). So far, no chromitite deposits have been observed in abyssal peridotites. PGEs and chromium behave compatibly during dry partial melting (e.g. Mitchell and Keays, 1981; Dick and Bullen, 1984) and these metals therefore show restricted mobility at mid-ocean ridges. In contrast, mantle fluxing by hydrous fluids and melts is a typical feature of supra-subduction zone (SSZ) environments. In this environment, mantle peridotite can be melted to a higher degree than beneath mid-ocean ridges, because the mantle wedge is fluxed by fluids released from the subducting oceanic lithosphere (e.g. Pearce et al., 1984; Roberts and Neary, 1993). Thus, in order to understand the formation of chromitite deposits, it is important to know the nature of these fluids and the behaviour of Cr and PGE during the interaction among fluids, silicate melts and mantle peridotites.

Podiform chromitites are mainly composed of spinel and olivine with occasional subordinate pyroxene. Sulfide grains are common in most of the chromitite deposits (e.g. McElduff and Stumpfl, 1990). Most podiform chromitites have Os, Ir and Ru concentrations of between 0.1 and 0.01 times chondritic (Page et al., 1982; Talkington and Watkinson, 1986; Leblanc, 1991) and lower chondrite-normalized abundances of Pt and Pd. Very few podiform chromitites are enriched in Pt and Pd relative to the other PGEs. These exceptional cases include the chromitites of Greece (Konstantopoulou and Economou, 1991), the Zambales ophiolite in the Philippines (Bacuta et al., 1990) and the Shetland ophiolite in Scotland (Prichard et al., 1996). In these cases, the high Pt and Pd values are believed to reflect a contribution from magmatic sulfides (e.g. Prichard et al., 1996). Walker et al. (2002a,b) reviewed the Os isotopic composition of chromitite deposits from ophiolites worldwide and observed a well defined average 187 Os/188Os value of 0.12809 F 0.00085 (2r). These authors suggested that this value is representative of the DMM and that the addition of 187Os from the dehydrating oceanic crust has no significant effect on the Os isotopic composition of the chromitites. In this paper, we present PGE abundances, Os isotopic compositions and Cr numbers of the chromitite deposits and of surrounding mantle rocks from the Troodos Ophiolite, Cyprus, in order to investigate the influence of percolating fluids/melts on the behaviour of PGEs and Cr in the Earth’s upper mantle. The study

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also re-evaluates the importance of the imprint of a radiogenic signature from percolating melt on the isotopic composition of mantle peridotites and chromitites in ophiolites.

2. Samples and methods 2.1. Samples The Troodos Ophiolite Complex represents oceanic lithosphere, which formed in a supra-subduction

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zone environment 90 Ma ago (e.g. Robinson and Malpas, 1990; Mukasa and Ludden, 1987). The mantle sequence of the Troodos Ophiolite Complex can be divided into two parts (Batanova and Sobolev, 2000). The eastern part (Unit 1 in Fig. 1) consists mainly of spinel – lherzolite with subordinate cpx-bearing harzburgites and dunites. The western part (Unit 2 in Fig. 1) is composed of harzburgites and dunites and contains three chromitite deposits: Kannoures, Kokkinorotsos and Hadji Pavlou (Fig. 1). The chromitites occur (e.g. at Kannoures) in dunite lenses in the harzburgite and (e.g. at Kokkinorotsos) at the base

Fig. 1. Location of the chromitite deposits Hadji Pavlou, Kannoures, Kokkonorotsos and associated mantle peridotites from the Troodos Ophiolite Complex. Peridotite sample locations are shown with white squares circles (Bu¨chl et al., 2002), chromitite samples by sample numbers.

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2.2. Analytical methods

of a dunite unit as shown in Fig. 1 (Greenbaum, 1977; Prichard and Lord, 1990). Unit 2 has been overprinted by melt percolation (Batanova and Sobolev, 2000; Bu¨chl et al., 2002). Major element compositions, Os isotopes and PGE abundances of spinel were determined in one chromitite sample from Kannoures, in three samples from Kokkinorotsos and in two samples from Hadji Pavlou (Fig. 1). In addition, one sample of a dunite body from Unit 1 containing chromite ‘‘schlieren’’ was analysed (T14). Samples were crushed, and fresh spinels were handpicked and washed with deionised water in an ultrasonic bath before analysis. The data of Bu¨chl et al. (2002) and Bu¨chl et al. (2003, GCA in revision) complement the present data set providing additional data for harzburgites and dunites enclosing the chromitite deposits (Fig. 1).

Major and trace element concentrations were determined on glass and powder pellets with a Phillips PW 1404 X-ray fluorescence spectrometer at the University of Mainz. Electron microprobe (Jeol JXA 8900 RL) analyses of spinel were made at the University of Mainz using the routine standard procedure. For PGE analysis, 50 mg of handpicked and washed spinel separates were completely dissolved in a quartz vessel together with a mixed PGE/Re isotope tracer ( 185 Re, 190 Os, 191 Ir, 101Ru, 198 Pt, 106 Pd), conc. HCl and conc. HNO3 (2:3) for 16 h in a high-pressure asher at 100 bar and 300 jC. The PGE were separated from the spinel matrix using the procedure described by Bru¨gmann et al. (1999). Osmium was separated from the sample solution by

Table 1 PGE abundances, Os isotopic compositions and spinel mineral data from handpicked spinels from the chromitite deposits from the Troodos Ophiolite Complex Sample mine

T6

T4e

98-3

Hadji Pavlou Whole rock

Spinel

187 Os/188Os 2 sigma duplicate Os (ng/g) Os duplicate Ir (ng/g) Ir duplicate Ru (ng/g) Pt (ng/g) Pt duplicate Pd (ng/g) Re (ng/g) Ir (normalized) Os (normalized) Ru (normalized) Pt (normalized) Pd (normalized) Re (normalized) Cr# in spinel Mg# in spinel MgO FeO Cr2O3 Al2O3

0.1301 0.1297 19.38 13.99 11.49 17.92 34.87 0.41 0.03 11.13 1.35 0.03 0.04 0.05 0.00 0.02 0.04 0.69 0.62 14.11 15.65 53.16 16.21

0.1284 0.0002 23.35

T10a

98-16

98-17

Kokkinorotsos 0.1270

5.89

0.1265

18.62

0.1265

288.56

0.1298

5.84

18.99 83.62 9.07 28.61 0.04 0.04 0.12 0.01 0.05 0.64 0.60 13.80 16.35 48.96 18.81

0.76 0.62 12.95 17.36 57.19 11.84

0.75 0.53 11.84 19.07 54.73 12.14

0.75 0.54 12.02 18.72 55.69 12.20

T11

T14

Kannoures

U1 picked

0.1294

0.1296

0.1289 15.52 12.01 11.41 11.11 34.30 1.41 0.08 76.80 0.33 0.02 0.02 0.05 0.00 0.14 0.01 0.76 0.62 13.75 14.92 58.52 12.23

Primitive mantlea

28.11

3.40

24.31

3.20

19.26 6.04

5.00 7.10

48.03

3.90 0.28 0.007 0.007 0.007 0.007 0.007 0.007

0.05 0.06 0.03 0.01 0.09 0.75 0.50 10.71 19.40 56.72 12.83

Abbreviations: U1 = Unit 1; U2 = Unit 2; Cr# = Cr/(Cr + Al); Mg# = Mg/(Mg + Fe), hzb = harzburgite. Spinel mineral data determined by electron microprobe, Os isotopes by N-TIMS and PGE abundances by isotope dilution with MC-ICPMS. a Data for the primitive mantle from McDonough and Sun (1995).

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solvent extraction with bromine and purified by microdistillation (Birck et al., 1997). Afterwards, Ru, Pd, Re, Ir and Pt were sequentially extracted from the solution by using anion exchange columns, applying a technique modified after Rehka¨mper and Halliday (1997). Osmium isotope measurements were carried out by N-TIMS with a Finnigan MAT 262 mass spectrometer at the Max-Planck-Institut in Mainz. Fourteen procedural blanks (determined by isotope dilution) ranged from 67 to 558 fg for Os. The 2r external precision (the reproducibility of the isotopic ratios) of the 187Os/188Os was 0.3% based on repeated measurements of a standard (n = 77) containing 35 – 70 pg Os. Duplicate analyses of two samples were performed, each starting with digestion of a separate aliquant of sample powder. The results of these duplicates agreed within 0.5% for 187Os/188Os and within 28% for Os concentrations. The Pt, Pd, Ir and Ru concentrations were determined by isotope dilution using the Micromass Isoprobe, a second-generation multicollector ICPMS at the University of Mu¨nster. The 2r external precision based on repeated measurements of a standard solution were 0.36% for 198 Pt/ 194 Pt, 0.24% for 106 Pd/ 108 Pd, 0.16% for 101 Ru/99Ru and 0.30% for 191Ir/193Ir. Procedural blanks (spiked and determined by isotope dilution) ranged for Ru (n = 5) from 0.12 to 0.84 ng, for Ir (n = 5)

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from 0.009 to 0.034 ng, for Pt (n = 4) from 0.07 to 0.25 ng and for Pd (n = 4) from 0.26 to 0.79 ng. These blank concentrations are at least 25 times lower for Os, Ru, Ir and Pd and three times lower for Pt, than those observed in the samples; thus blank corrections are generally negligible. Repeated complete digestions of a fresh aliquant of the UBN standard indicated a reproducibility of 10% for the highly siderophile elements (Bru¨gmann et al., in preparation).

3. Results The Cr numbers (Cr/(Cr + Al)) of the studied spinels range from 0.64 to 0.76 (Table 1; Fig. 2) and lie within the range previously measured in chromitite deposits from ophiolites worldwide (e.g. Zhou et al., 1996, 1998; Melcher et al., 1997). Fig. 2 shows that the Mg numbers (Mg/(Mg + Fe)) in spinel vary significantly and are lower than those observed in previous studies. The Mg# is dependent on the relative proportions of spinel and olivine, because Mg – Fe are exchanged between olivine and spinel. A higher modal percentage of olivine leads to higher Mg# in spinel. Hence, variation of the relative proportions of olivine and chromite may explain the variable Mg# in spinel.

Fig. 2. Cr number and Mg number in spinel of the chromitite deposits and the associated mantle peridotites. The Cr numbers of the chromites are at the high end of the range displayed by harzburgites and dunites from the mantle sequence. Data for the mantle peridotites are from Bu¨chl et al. (2004, GCA in press).

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Fig. 3. (A) PGE patterns of separated spinels from the chromitite deposits. The patterns are in the range of worldwide measured chromitite deposits from ophiolites with Pt depleted relative to Os, Ir and Ru. However, Pd has an unusually high concentration compared to most other studies. The data from Prichard and Lord (1990) and McElduff and Stumpfl (1990) are from whole rock chromitite samples. PGE concentrations for the primitive mantle are from McDonough and Sun (1995). (B) PGE patterns of chromites from the chromitite deposits and surrounding harzburgites and dunites (average concentrations revealed). Chromites and dunites show complementary PGE patterns, whereas the PGE patterns of the harzburgites are mantle-like. Data for the harzburgites and dunites are from Bu¨chl et al. (2004, GCA in press).

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Osmium, Ir, Ru and Pd concentrations of chromites from all three chromitite deposits and of the dunite (T14) are very high compared with those of the primitive mantle, but Pt concentrations are low (Table 1; Fig. 3A). The Os concentrations vary from 5.8 to 288.5 ng/g, Ir from 11.2 to 24.3 ng/g, Ru from 19.2 to 83.2 ng/g, Pt from 0.4 to 9.0 ng/g, and Pd from 11.1 to 76.8 ng/g (Table 1). Such large variations in the PGE concentrations are typical for chromitite deposits (e.g. Zhou et al., 1998; Melcher et al., 1999). High concentrations for Os, Ir and Ru, but low concentrations for Pt and Pd have also been observed in whole rock samples of chromitites from the Troodos Ophiolite by Prichard and Lord (1990) and McElduff and Stumpfl (1990) and from other ophiolite chromitite deposits (e.g. Page and Talkington, 1984; Talkington and Watkinson, 1986; Leblanc, 1991; Crocket, 1981, 2002; Agiorgitis and Wolf, 1978; Zhou et al., 1998). The Pd concentrations in the chromites from this study lie in the range of Os, Ir and Ru concentrations observed for deposits worldwide, but are high if compared with results of the Troodos study of Prichard and Lord (1990) and McElduff and Stumpfl (1990).

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The 187Os/188Os ratios of the chromites from three chromitite deposits vary from 0.1265 to 0.1305, with an average of 0.1284 F 0.0021 2r (Fig. 4). This value agrees well with the average of worldwide measured podiform chromitites (187Os/188Os: 0.12809 F0.00085 2r, Walker et al., 2002a,b). It is somewhat higher than estimates for the carbonaceous chondritic reservoir (187Os/188Os: 0.127; Shirey and Walker, 1998; 187 Os/ 188 Os: 0.1260 F 0.0013; Walker et al., 2002a,b), and slightly lower than the value estimated for the primitive upper mantle (187Os/188Os: 0.1296 F 0.0008 2r; Meisel et al., 2001), but overlaps within uncertainty with the ratios proposed for both of these reservoirs. The samples show no correlation between major elements, PGE concentrations and ratios and Os isotopic compositions.

4. Discussion In detail, it is still poorly understood how Cr and PGE are mobilized and transported in the ophiolite mantle and subsequently become concentrated in

Fig. 4. 187Os/188Os ratios and Os concentrations of the spinels from the chromitite deposits and associated mantle rocks. The chromites have a much smaller range in their Os isotopic composition and higher Os concentrations compared with the peridotites. A mixing line between mantle and melt suggest that a melt/rock ratio of 17:1, on average, is necessary to explain the 187Os/188Os ratio of the chromitite deposits. Data for the mantle peridotites are from Bu¨chl et al. (2004, GCA in press).

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chromitite deposits. This is also true for the role of fluids during chromitite formation. 4.1. Origin of chromium in the chromitite deposits In order to form chromitite deposits, large amounts of Cr have to be mobilized in the Earth’s upper mantle. The chromitite deposits of the Troodos Ophiolite as well as most of the chromitite deposits worldwide are enclosed in dunite envelops (Roberts, 1988, Prichard and Lord, 1993; Lago et al., 1982; Melcher et al., 1999) and surrounded by moderately depleted harzburgites (Pearce et al., 1984; Roberts, 1988; Nicolas, 1989; Leblanc and Nicolas, 1992). Gradual lithological changes from dunites to harzburgites to lherzolites have also been described around chromitite pods in a number of ophiolites (Zhou et al., 1996). Such dunites formed by melt percolation and not by partial melting, because the liquidus temperature of the dunite is not approached in the Earth’s upper mantle during partial melting. During melt percolation, dunites and harzburgites can form from lherzolites by clinopyroxene dissolution and incongruent melting of orthopyroxene, and precipitation of olivine from the melt (Kelemen et al., 1997; Suhr, 1999). Chromium behaves compatibly during igneous

fractionation processes as long as orthopyroxene (KDCropx/liq = 4.6 – 29 after Jones and Layne, 1997), clinopyroxene (KDCrcpx/liq = 8.1 – 36 after Jones and Layne, 1997) or spinel (KDCrsp/liq = 77 after Ringwood, 1970) are fractionating phases. However, Cr is incompatible in olivine (KDCrol/liq = 0.58 – 0.657 after Gaetani and Grove, 1997). During partial melting in the presence of orthopyroxene, clinopyroxene and spinel Cr will behave compatibly. Therefore, chromitite deposits cannot form under these conditions. However, during continuous melt percolation in the sub-arc mantle pure Ol-residues form and Cr therefore could be mobilized by the melts. In the mantle peridotites enclosing the chromitite deposits of the Troodos Ophiolite, Cr concentration systematically decreases from harzburgites (average: 2528 ppm) to dunites (average: 1758 ppm) (Fig. 5). Thus, the dunites enclosing the chromitite deposits and the dunite melt channels occurring throughout the mantle section are the source for Cr in the chromitite deposits. 4.2. Origin of PGEs in the chromitite deposits With the exception of Pd, the PGE patterns of the spinels determined in this study are similar to those of whole rock chromitites observed in other studies of

Fig. 5. Cr and Os concentrations from peridotites associated with the chromitite deposits. Cr and Os both behave incompatibly during the melt percolation process. Data are from Bu¨chl et al. (2004, GCA in press).

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the Troodos Ophiolite (McElduff and Stumpfl, 1990; Prichard and Lord, 1990) (Fig. 3A). They also are within the range of PGE patterns observed worldwide in chromitite deposits (Fig. 3A) (e.g. Zhou et al., 1998; Melcher et al., 1999). The depletion of Pt relative to Ir, Os and Ru is also a characteristic feature of chromitites from ophiolite complexes. This suggests that podiform chromitite deposits concentrate the PGE by similar processes and under comparable thermodynamic conditions. Variable relative abundances of Os, Ir and Ru may reflect the presence and heterogeneous distribution of Os – Ir alloys and sulfide phases. For example, the presence of laurite inclusions could provide an explanation for the high Ru concentration in sample 10a (Fig. 3A). This also provides an explanation for the non-chondritic element ratios of the samples (Table 1). Our spinel samples have higher Pd concentrations (0.02 –0.15 times C-1) than those observed in previous studies of the Troodos Ophiolite Complex (0.0015 and 0.01 times C-1; McElduff and Stumpfl, 1990; Prichard and Lord, 1990) (Fig. 3A). All previous studies analyzed whole rock powder, whereas in this study, handpicked and purified spinel samples were analyzed. The enrichment in Os, Ir and Ru relative to Pt and Pd observed in previous studies has been explained by the early removal of Os, Ir and Ru with chromite from the melt, whereas the more incompatible Pt and Pd remain in the silicate melt. Alternatively, it has been suggested that the low Pd and Pt contents require that the source material lost Pd and Pt prior to the formation of the chromite deposits (Melcher et al., 1999; Crocket, 1981; Barnes and Naldrett, 1985; Edwards, 1990; Keays, 1995; Zhou et al., 1998). Our study shows that in fact the enrichment of Pd is similar to that of Os, Ir and Ru and it is not necessary to invoke such processes in order to explain the PGE pattern. However, it is difficult to explain the different Pd abundances in mineral separates and whole rock samples. It appears to be obvious that the matrix between the spinel grains makes the difference. If this is the case, then the matrix should be depleted in Pd relative to Ir, Os and Ru. The matrix between the fresh spinel grains is strongly altered, mainly serpentinised. Thus, it is likely that the most mobile PGE, namely Pd, has been lost during the alteration process. Alternatively, the Pd depletion is a primary feature as suggested by previ-

225

ous studies (Melcher et al., 1999; Crocket, 1981; Barnes and Naldrett, 1985; Edwards, 1990; Keays, 1995; Zhou et al., 1998). However, mass balance constraints indicate that Pd depletion in the matrix cannot be the only reason for the difference. The measured whole rock Pd/Pt ratios of Troodos chromitites are about 10 times lower than those of the chromite separates. This implies that the great majority of the Pd, and presumably the other PGE, was hosted by the matrix before alteration. As the modal chromite content of these rocks is about 85%, this means that the matrix must have been highly enriched in PGE relative to the chromite. For example, if the whole rock chondrite-normalized Pd content before alteration was 0.1, and after alteration was 0.01, and we assume that all of the remaining Pd is in the chromite, then the matrix originally had a chondritenormalized Pd content of f 0.6 compared to a value of 0.0118 in the chromite. If this original Pd content is representative of the chondrite-normalized contents of Os, Ir, and Ru, then the chromitite whole rock concentrations of these elements should be about 10 times higher than those of the chromite separates. Many whole rock samples in Fig. 3a have such high Ir, Os and Ru concentrations. The concentration range of whole rocks from the Troodos complex (Prichard and Lord, 1990; McElduff and Stumpfl, 1990), however, overlaps with that of our chromite separates. Therefore, Pd depletion in the matrix is not the complete explanation for the difference between the whole rock and chromite separate results. Prichard and Lord (1990) suggested that the high PGE concentrations in the chromitite deposits are the result of enhanced partial melting of the peridotites. However, during partial melting, Os, Ru, and Ir behave compatibly even at the high degrees of partial melting necessary to form komatiitic and picritic melts (>30%; Morgan, 1986; Bru¨gmann et al., 1987, 2000; Lorand et al., 1999). Enhanced partial melting, therefore, cannot be responsible for the high PGE concentration of the chromitite deposits. In the mantle peridotites from Troodos, Ir, Os, and Ru concentrations systematically decrease from the harzburgites to the dunites as observed by Bu¨chl et al. (2002) and shown for Os in Fig. 5. In fact, chromites and dunites have complementary PGE patterns (Fig. 3b). This pattern also supports our suggestion that Pd is mobilized in the mantle peridotites as well as Os, Ir

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and Ru. Bu¨chl et al. (2002) suggested that these elements tend to behave incompatibly during melt percolation because first sulfide, and eventually pyroxenes and chromite are dissolved by the percolating melt. This is supported by the observation that the dunites (no sulfides visible) have a lower sulfide content than the harzburgites ( < 0.001% after Batanova and Sobolev, 2001). Prichard and Lord (1990) also noted the virtual absence of visible sulfides in the Troodos ultramafic rocks. The PGE pattern of the spinels show a distinct depletion in Pt compared with the other PGEs and this is a typical feature of chromitites from ophiolite complexes worldwide (Fig. 3A). Interestingly, the associated dunites and harzburgites have very high Pt concentrations varying from 9.11 to 14.95 ng/g. In addition, Pt concentrations in harzburgites (11.8 ng/g) and dunites (10.7 ng/g) are similar on average (Bu¨chl et al., 2002; Fig. 3B). This suggests that Pt was not mobilized with the other PGEs during melt migration through mantle peridotites. The discrete behaviour of Pt relative to the other PGEs has been described by other authors. Handler and Bennett (1999) suggested, based on separated spinel and silicates of Australian peridotite xenoliths, that Pt and Pd occur mostly as discrete Pt- and Pd-rich PGE phases that are the cause for the poor reproducibility of Pt and Pd whole rock analyses and the lack of correlation between these elements and Ir in bulk rock analyses. Laser-ablation analyses of mantle rocks and abyssal peridotites from the Mid-Atlantic and South West Indian ridges showed that the whole rock budget of Os, Ir, Ru, Rh and Pd is balanced by the concentrations measured in low-temperature sulfide assemblages (Alard et al., 2000; Luguet et al., 2001). In contrast, Pt shows a deficit if compared to the measured whole rock concentrations. These authors also observed Pt concentration peaks along laser-ablation profiles in Curich pentlandite. Alard et al. (2000) suggested that Pt occurs as disseminated Pt-rich micronuggets, too small in size to be properly analysed, and that these may be relatively low-temperature exsolution products. However, Luguet et al. (2001) concluded that they could also represent primary igneous minerals. Lorand and Alard (2001) suggested, based on laserablation analyses of Massif Central xenoliths, that if Pt-rich discrete microphases really exist they probably exsolved during subsolidus decomposition of mantle

sulfide. Pendlandite may theoretically accommodate Ru, Rh, and Pd in its octahedral sites while rejecting Pt (cf. Mackovicky et al., 1986; Czamanske et al., 1992; Ballhaus and Ryan, 1995; Ballhaus and Sylvester, 2000). Thus, Pt alloys could be stable in the presence of pendlandite ( < 600 jC; fS2 < 10
7 atm; Vaughan and Craig, 1978). Its fS2 dependency could result in a complex partitioning behaviour of Pt between Mss, Cu-sulfides, and perhaps a small amount of Pt alloys. At 900 jC, the Cu – Ni-rich sulfide liquid, the high-temperature precursor phase of coarse-grained pentlandite and Cu-sulfides, can dissolve up to 15% Pt (Mackovicky et al., 1986), and there is no large difference of Dmss/liq for Pt and Pd (Li et al., 1996). The Eggler and Lorand (1993) sulfide barometer calibrated for the P –T conditions of the lithospheric mantle (950 jC, 1.2 GPa) supports the occurrence of Pt as sulfide rather than as alloys. However, the dunite PGE pattern of this study (Fig. 3B) indicates that Pt has been retained in the absence of sulfide. In detail, however, the formation and solubility of discrete PGE minerals as a function of pressure, temperature, fS2 and fO2 is not well known. The present study shows that the mobilization of Pt has been inhibited during the melt percolation process in the mantle peridotites of the Troodos complex. This suggests that Pt-rich phases are also present in the upper mantle. We suggest that the PGEs (except Pt) and Cr are mobilized in the mantle by percolating melts. The incompatible behavior of Ir, Os, Ru and Pd during melt percolation could be due to the presence of fluids in the supra-subduction zone environment. Johan and Le Bel (1978) and Johan (1986) also suggested that chromitite deposits form during the interaction of magma with reducing fluids. Matveev and Ballhaus (2002) suggested that basaltic melts parental to podiform chromitite deposits need to be water saturated in order to produce podiform chromitite deposits. Their experiments indicate that in conjugate basalt – water systems (at 1150– 1200 jC and 0.5 GPa hydrostatic pressure) silicates, oxides, and metallic phases crystallize together and may be fractionated from each other by purely physical processes. For example, the authors observed that dispersed PGE nuggets are mechanically concentrated along with chromite in the exsolving fluid phase. However, this model does not explain the Pt-depletion observed

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in podiform chromitites. In addition, Pd is not known to form discrete mineral phases in such deposits. Thus, the similar enrichment of Pd, Ir, Os and Ru consistently observed in our study does not support an accidental concentration of different PGE phases. It rather suggests the contemporaneous concentration of all PGE and that the PGE distribution in chromites reflects that of the percolating melts. The only known mantle phase which has high and similar partition coefficients for all PGE is a sulfide liquid. We therefore propose that sulfide liquids play a major role as collectors of PGE during the formation of chromitite deposits. Chromitite deposits form in the mantle if chromite becomes the major liquidus phase of migrating melts. Irvine (1977) proposed that the chromium solubility in silicate melts depends on the silica activity. This can be effectively changed by assimilating silica during contamination with continental crust, which is, however, not a feasible process in the ophiolite mantle. Alternatively, silica activity can be changed during magma mixing. This process has been claimed to produce chromite layers in mafic– ultramafic intrusions (Irvine, 1977). Similarly, experiments made by Ballhaus (1998) showed that chromitite deposits can form during mingling and mixing of silicate magmas. Naldrett and Duke (1980) and Naldrett et al. (1990) suggested that this process may also trigger the formation of an immiscible sulfide liquid that collects the PGE. Thus, we propose that that during the pooling and homogenization of percolating melts near the mantle – crust boundary chromite and sulfides coprecipitate from the hybrid magmas. These phases may be further concentrated in a fluid phase, as proposed by Matveev and Ballhaus (2002), eventually forming a podiform chromitite deposit. Variable amounts of droplets of immiscible sulfide liquids will be scavenged by chromite during its crystallization and this explains the PGE pattern and variable metal abundances observed in our spinels. Post-magmatic recrystallization of chromite would cause desulfurization due to the transfer of Fe2 + from the sulfides into vacancies of the chromite (Naldrett et al., 1989). This would favour the formation of metal alloys as inclusions in chromite and explains the occurrence of discrete mineral phases of Ir, Os, and Ru often observed in podiform chromitites. In contrast, sulfides in the groundmass between chromite grains suffer a

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strong alteration caused by late stage fluids whereby sulfides and Pd are mobilized again.

5. Implications of the Os isotope systematics in chromitite deposits Chromites from ophiolites worldwide have suprachondritic Os isotopic compositions (average 187Os/ 188 Os: 0.12809 F 0.00085 (2r) after Walker et al., 2002a,b) if compared with the carbonaceous chondrite ratio of 0.126. Only the Jormua Ophiolite has subchondritic 187Os/188Os ratios probably due to the involvement of subcontinental lithospheric mantle (SCLM) (Tsuru et al., 2000). Walker et al. (2002a,b) argued that the Os isotopic composition of chromitites represents an integrated value of the depleted mantle. As described above, chromitites may form during the pooling and mixing of percolating melt. These melts most probably have radiogenic Os isotope compositions (e.g. Woodland et al., 2002; Borg et al., 2000). Likewise the boninites (187Os/188Os: 0.163) and pyroxenite veins (cpx-veins: 0.129 –0.130, opx-veins: 0.166 – 0.184) (Bu¨chl et al., 2003, GCA in revision) from the Troodos Ophiolite have a radiogenic initial Os isotopic composition. Even though most of their chromitite samples are from SSZ environments, Walker et al. (2002a,b) believe that radiogenic melts or fluids derived from subducting slabs cannot cause a significant bias towards higher 187Os/188Os. We suggest that the influence of radiogenic melts from the subducting slab on the isotopic composition of the mantle wedge cannot be ignored. Osmium can be mobilized during the dehydration of the oceanic crust, and the influence of slab derived 187Os on the Os isotopic composition of mantle peridotites has been demonstrated by several studies (e.g. Brandon et al., 1996, 1999; Parkinson et al., 1998; McInnes et al., 1999; Borg et al., 2000). For example, Brandon et al. (1996) explain a 16% increase of 187Os in some peridotite xenoliths from areas lying above recent subduction zones by adding slab-derived Os. Mass balance calculations which include the whole mantle wedge (Walker et al., 2002a,b) tend to overestimate the amount of melts/fluids necessary to overprint the Os isotopic composition of the mantle peridotites. Melting in the mantle wedge is caused by lowering the melting temperature of the peridotites due to the

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invasion of melt/fluids. Therefore, partial melting occurs only in those parts of the mantle that are invaded by melt/fluids. In this case, the influence of the radiogenic melts from the subducting slab on the peridotites is not negligible. The proportion of melt necessary to elevate the Os isotopic composition of the mantle wedge peridotites can be calculated from the following mixing equation: amix ¼ ðaendmember1  Cendmember1  ð1
f Þ þ aendmember2  Cendmember2  f Þ =ðCendmember1  ð1
f Þ þ Cendmember2  f Þ

The variables are defined as follows: amix represents the Os isotopic ratio of the mixture of melt and peridotite, here the average 187Os/188Os ratio of the chromitite deposits from the Troodos Ophiolite: 0.1284; aendmember1 is the Os isotopic ratio of the melt, here the initial (90 Ma) of boninites from the Troodos Ophiolite (187Os/188Os: 0.163). Comparison of the trace element composition of the cpx from a websterite vein from the mantle section enclosing the chromitite deposits with that of cpx phenocrysts from the upper pillow lava boninites from the Troodos Ophiolite shows a close match (Bu¨chl et al., 2002). This observation allows us to suggest that a melt with a composition similar to that of the boninites has percolated through this part of the mantle section; aendmember2 is the Os isotopic ratio of the peridotite, here the spinel –lherzolite of the Troodos Ophiolite (most primitive mantle rock in the mantle section) (187Os/188Os: 0.127); Cendmember1 is the concentration of Os in the melt, here the boninites from the Troodos complex with 0.01 ng/g Os; Cendmember2 is the concentration of Os in the peridotite, here the spinel– lherzolite from the Troodos Ophiolite with 4.2 ng/g Os; f: is the mass fraction of endmember2. The melt/rock ratio necessary to elevate the Os isotopic composition of the mantle wedge peridotites (spinel – lherzolite from the Troodos Ophiolite with 187 Os/188Os of 0.127) to the average value of the chromitite deposits of 0.1284 is 17:1, on average. Our calculations imply that the chromitites form during percolation processes with a melt/rock ratio of < 1:1 – 41:1 (Fig. 4). These values were obtained using the simple two-component mixing equation given above;

more realistic melt percolation models will in fact yield slightly lower melt/rock ratios. Kelemen et al. (1995, 1997) suggested an average melt/rock ratio for dunite conduit formation of 8:1 to 20:1. Our calculated mean value lies within this range. Kelemen et al. (1995, 1997) also determined that the integrated melt/ rock ratio for both chromitites and surrounding dunites in the Oman Ophiolite must have been >300. This is based on the fact that the solubility of Cr-spinel is low in silicate melts and chromitites must have scavenged Cr from 300 to 400 times their mass of liquid (Leblanc and Ceuleneer, 1992). While the exact value of this estimate depends on the solubility of Crspinel and thus perhaps on the fluid content of the melt phase, the main point is unlikely to change. That is, it is easy to attain very high melt/rock ratios in the mantle wedge environments where most chromitites form. The melt/rock ratio required to significantly alter the Os isotopic ratio of the peridotite will of course depend on the Os concentrations and isotopic compositions chosen for the endmembers. Thus, assuming lower Os concentration ( < 0.01 ng/g) or/and a less radiogenic Os isotopic composition for the melt component (187Os/188Os < 0.163) would result in a significantly higher melt/rock ratio (e.g. 180 for a melt Os concentration of 0.001 ng/g). Nevertheless, given the very extensive degree of melt percolation proposed on the basis of Cr solubility, the melt/rock ratios required to significantly modify the 187Os/188Os ratios seem quite plausible. In summary, we propose that melts passing through harzburgite and dunite conduits mobilize Os when all mantle sulfides are dissolved. During the formation of the chromitite deposits melts with different Os signatures are pooled, mixed and homogenized on a large scale. This is the reason for the relatively homogeneous Os isotopic composition in the chromitite deposits. The Os isotopic composition of the chromitite deposits is buffered by the Os component of the peridotites. However, in detail, the 187Os/188Os ratios of the spinels in the chromitite deposits show some variation (Fig. 4), depending on the relative amount of Os of melts and peridotites and their Os isotopic compositions. The range in 187Os/188Os ratios is mainly a function of the initial compositions of melts and peridotites and the melt/rock ratio. Therefore, we conclude that podiform chromitites do not represent the Os isotopic composition of the upper convecting

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mantle, but that they define an upper limit. This implies that prior to melt percolation, the mantle source of the ophiolites had a significantly lower 187 Os/188Os than the PUM.

6. Conclusions We present a data set of PGE abundances, Os isotopic compositions, and Cr numbers of chromites from chromitite deposits and associated mantle peridotites from the Troodos Ophiolite Complex. This study shows that the surrounding mantle peridotites are the sources of the PGE and Cr in the chromitite deposits. This is because Cr and PGE— with the exception of Pt—are mobilized in peridotites during the interaction with percolating, probably fluid-rich, melts. At higher stratigraphic levels, the metals are precipitated from the melt and become concentrated when chromite and sulfide liquids precipitate as a result of magma mixing in magma pods now represented by the chromitite deposits. Thus, the PGE distribution observed in a chromitite reflects that of the average percolating melt. With the exception of Pd, the PGE patterns of spinels separated from chromitites of the Troodos complex are similar to those of whole rock analyses of podiform chromitite deposits worldwide. Typical features are the high concentrations of Ir, Os, and Ru and negative Pt anomalies. The stability of Pt-rich phases in the ophiolite mantle inhibits Pt mobilization during melt percolation and eventually causes the negative Pt anomalies in the chromitite deposits. However, Pd concentrations in separated spinel are higher than those of whole rock chromitites. Pddepletion in the matrix cannot entirely explain this difference and this implies significant variations of PGE abundances in chromites even on a local scale. The 187Os/188Os ratios of the spinel from podiform chromitite deposits from the Troodos Ophiolite Complex range from 0.1265 to 0.1305. The average value is similar to that of chromitite deposits worldwide (Walker et al., 2002a,b). The variation of the 187 Os/188Os ratio of the chromitite deposits reflects a mixture of Os derived from the mantle peridotites with Os from the subducting slab. The latter component is transported along with percolating melts and fluids and probably has a radiogenic composition

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typical of many arc basalts. Model calculations imply that a melt/rock ratios of < 1:1 – 41:1 are necessary in order to increase the 187Os/188Os of the spinel– lherzolites from the Troodos Ophiolite to the observed values in the chromitites. These values are similar to or lower than independent estimates of melt/rock ratios in ophiolites. Thus, podiform chromitites do not represent the Os isotopic composition of the upper convecting mantle. However, they do define the upper limit, and this implies that the ophiolite mantle has a significantly lower 187Os/188Os than PUM and has therefore suffered a long-term depletion of Re.

Acknowledgements We thank Klaus Mezger (University of Mu¨nster) for providing access to the Isoprobe. This study also benefited from many discussions with Chris Ballhaus and Alexander Sobolev. G. Bru¨gmann is most grateful to Costas Xenophontos for introducing him into the geology of the Troodos Ophiolite. We would also like to thank the editor Laurie Reisberg and the reviewers Jim Crocket, Monica Handler, and an anonymous reviewer for their critical comments which significantly helped to improve the quality of the manuscript. This work has been supported by the Graduierten Kolleg ‘‘Stoffbestand und Entwicklung von Mantel und Kruste’’ at the Johannes Gutenberg University of Mainz. [RR]

References Agiorgitis, G., Wolf, R., 1978. Aspects of osmium, ruthenium and iridium contents in some Greek chromites. Chem. Geol. 23, 267 – 272. Alard, O., Griffin, W.L., Lorand, J.P., Jackson, S.E., O’Reilly, S.Y., 2000. Non-chondritic distribution of the highly siderophile elements in mantle sulfides. Nature 407, 891 – 894. Bacuta, G.C., Kay, R.W., Gibbs, A.K., Bruce, R.L., 1990. Platinumgroup element abundance and distribution in chromite deposits of the Acoje Block, Zambales ophiolite complex, Philippines. J. Geochem. Explor. 37, 113 – 145. Ballhaus, C., 1998. Origin of podiform chromite deposits by magma mingling. Earth Planet. Sci. Lett. 156, 185 – 193. Ballhaus, C., Ryan, C.G., 1995. Platinum-group elements in the Merensky Reef: I. PGE in solid solution in base metal sulfides and the down-temperature equilibration history of Merensky ores. Contrib. Mineral. Petrol. 122, 241 – 251.

230

A. Bu¨chl et al. / Chemical Geology 208 (2004) 217–232

Ballhaus, C., Sylvester, P., 2000. Noble metal enrichment processes in the Merensky Reef, Bushveld complex. J. Petrol. 41, 545 – 561. Barnes, S.J., Naldrett, A.J., 1985. Geochemistry of the J-M (Howland) Reef of the Stillwater Complex, Minneapolis adit area: I. Sulfide chemistry and sulfide – olivine equilibrium. Econ. Geol. 80, 709 – 729. Batanova, V.G., Sobolev, A.V., 2000. Compositional heterogeneity in subduction-related mantle peridotites, Troodos massif, Cyprus. Geology 28, 55 – 58. Batanova, V.G., Sobolev, A.V., 2001. Sulfide minerals from the upper-mantle section of Troodos massif, Cyprus. J. Conf. Abs. 6. Birck, J.L., Roy-Barmann, M., Capmas, F., 1997. Re – Os isotopic measurements at the femtomole level in natural samples. Geostand. Newsl. 20, 19 – 27. Borg, L.E., Brandon, A.D., Clynne, M.A., Walker, R.J., 2000. Re – Os isotopic systematics of primitive lavas from the Lassen region of the Cascade arc, California. Earth Planet. Sci. Lett. 177, 301 – 317. Brandon, A.D., Creaser, R.A., Shirey, S.B., Carlson, R.W., 1996. Osmium recycling in subduction zones. Science 272, 861 – 864. Brandon, A.D., Becker, H., Carlson, R.W., Shirey, S.B., 1999. Isotopic constraints on time scales and mechanisms of slab material transport in the mantle wedge: evidence from the Simcoe mantle xenoliths, Washington, USA. Chem. Geol. 160, 387 – 408. Bru¨gmann, G.E., Arndt, N.T., Hofmann, A.W., Tobschall, H.J., 1987. Noble metal abundances in komatiic suites from Alexo, Canada, and Gorgona Island, Columbia. Geochim. Cosmochim. Acta 51, 2159 – 2169. Bru¨gmann, G.E., Yang, J.-H., Snow, J.E., Rehka¨mper, M., Mezger, K., 1999. Gemeinsame Bestimmung der Os-Isotopie und der Edelmetall-konzentrationen in Peridotiten, Basalten, Sulfiden und Eisenmeteoriten. Kolloquium des DFG-Schwerpunktprogramms ‘Ocean drilling Programm and Deep Sea Drilling Project’, Tagungsprogramm und Abstracts, 21 – 22. Bru¨gmann, G.E., Hanski, E.J., Naldrett, A.J., Smolkin, V.F., 2000. Sulphide segregation in ferropicrites from the Pechenga Complex, Kola Peninsula, Russia. J. Petrol. 41 (12), 1721 – 1742. Bu¨chl, A., Bru¨gmann, G., Batanova, V.G., Mu¨nker, C., Hofmann, A.W., 2002. Melt percolation monitored by Os isotopes and HSE abundances: a case study from the mantle section of the Troodos Ophiolite. Earth Planet. Sci. Lett. 204, 385 – 402. Bu¨chl, A., Bru¨gmann, G., Batanova, V.G., Hofmann, A.W., 2004. Os mobilization during melt percolation: the evolution of Os isotope heterogeneities in the mantle of the Troodos Ophiolite. Geochim. Cosmochim. Acta (in press). Crocket, J.H., 1981. Geochemistry of the platinum-group elements. In: Cabri, L.J. (Ed.), Platinum-Group Elements: Mineralogy, Geology, Recovery. Canadian Institute of Mining, Metallurgy and Petroleum, vol. 24, pp. 47 – 64. Crocket, J.H., 2002. The geology, geochemistry, mineralogy and mineral benefication of platinum-group elements. In: Cabri, L.J. (Ed.), Canadian Institute of Mining, Metallurgy and Petroleum, vol. 54, pp. 177 – 210. Czamanske, G.K., Kulinov, V.E., Zientek, M.L., Cabri, L.J., Likhachev, A.P., Oscarson, R.L., 1992. A proton microprobe study of

magmatic sulfide ores from the Norils’k Talnakh district, Siberia. Can. Mineral. 30, 249 – 287. Dick, H.J.B., Bullen, T., 1984. Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatially associated lavas. Contrib. Mineral. Petrol. 86, 54 – 76. Edwards, S.J., 1990. Harzburgites and refractory melts in the Lewis Hill massif, Bay of Islands ophiolite complex: the base-metals and precious-metals story. Can. Mineral. 28, 537 – 552. Eggler, D.H., Lorand, J.P., 1993. Mantle sulfide oxybarometry. Geochim. Cosmochim. Acta 57, 2213 – 2222. Gaetani, G.A., Grove, T.L., 1997. Partitioning of moderately siderophile elements among olivine, silicate melt, and sulfide melt: constraints on core formation in the Earth and Mars. Geochim. Cosmochim. Acta 61, 1829 – 1846. Greenbaum, D., 1977. The chromitiferous rocks of the Troodos ophiolite complex, Cyprus. Econ. Geol. 72, 1175 – 1194. Handler, M.R., Bennett, V.C., 1999. Behaviour of platinum-group elements in the subcontinental mantle of eastern Australia during variable metasomatism and melt depletion. Geochim. Cosmochim. Acta 63, 3597 – 3618. Irvine, T.N., 1977. Origin of chromitite in the Muskox intrusion and other stratiform intrusions: a new interpretation. Geology 5, 273 – 277. Johan, Z., 1986. Chromite deposits in the Massif du Sud ophiolite, New Caledonia: genetic consideration. Chromites. UNESCO’s IGCP 197 Project. Theophrastos, Athens, pp. 311 – 339. Johan, Z., Le Bel, L., 1978. Sur la genese des couches et podes de chromite dans les complexes ophiolitiques, Resume des principaux resultats science et technique. Serv. Geol., 96 – 99. Jones, R.H., Layne, G.D., 1997. Minor and trace element partitioning between pyroxene and melt in rapidly cooled chondrules. Am. Mineral. 82, 534 – 545. Keays, R.R., 1995. The role of komatiitic and picritic magmatism and S-saturation in the formation of ore deposits. Lithos 34, 1 – 18. Kelemen, P.B., Shimizu, N., Salters, V.J.M., 1995. Extraction of mid-ocean-ridge basalt from the upwelling mantle by focused flow of melt in dunite channels. Nature 375, 747 – 753. Kelemen, P.B., Hirth, G., Shimizu, N., Spiegelman, M., Dick, H.J.B., 1997. A review of melt migration processes in the adiabatically upwelling mantle beneath oceanic spreading ridges. Philos. Trans. R. Soc. Lond., A 355, 283 – 318. Konstantopoulou, U.G., Economou, M., 1991. Distribution of platinum-group elements and gold within the Vourinos chromite ores, Greece. Econ. Geol. 86, 672 – 1682. Lago, B., Rabinowicz, M., Nicolas, A., 1982. Podiform chromite ore bodies: a genetic model. J. Petrol. 23, 103 – 125. Leblanc, M., 1991. Platinum-group elements and gold in ophiolite complexes: distribution and fractionation from mantle to oceanic floor. In: Peters, T., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer Academic Publishing, Dordrecht, pp. 231 – 260. Leblanc, M., Ceuleneer, G., 1992. Chromite crystallization in a multicellular magma flow: evidence from a chromitite dike in the Oman ophiolite. Lithos 27, 231 – 257. Leblanc, M., Nicolas, A., 1992. Ophiolitic chromites. Int. Geol. Rev. 34, 653 – 686.

A. Bu¨chl et al. / Chemical Geology 208 (2004) 217–232 Li, C., Barnes, S.J., Mackovicky, E., Rose-Hansen, J., Mackovicky, M., 1996. Partitioning of nickel, copper, iridium, rhodium, platinum and palladium between monosulfide solid solution and sulfide liquid; effects of composition and temperature. Geochim. Cosmochim. Acta 60, 1231 – 1238. Lorand, J.P., Alard, O., 2001. Platinum-group element abundances in the upper mantle: new constraints from in situ and whole-rock analyses of Massif Central xenoliths (France). Geochim. Cosmochim. Acta 65, 2789 – 2806. Lorand, J.-P., Gros, M., Pattou, L., 1999. Fractionation of platinumgroup element in the upper mantle: a detailed study in Pyrenean orogenic peridotites. J. Petrol. 40, 951 – 987. Luguet, A., Alard, O., Lorand, J.P., Pearson, N.J., Ryan, C., O’Reilly, S.Y., 2001. Laser-ablation microprobe (LAM)-ICPMS unravels the highly siderophile element geochemistry of the oceanic mantle. Earth Planet. Sci. Lett. 189, 285 – 294. Mackovicky, M., Mackovicky, E., Rose-Hansen, J., 1986. Experimental studies and distribution of platinum-group elements in base metal sulfides in platinum deposits. In: Galagher, M.J., Ixer, R.A., Neary, C.R., Prichard, H.M. (Eds.), Metallogeny of Basic and Ultrabasic Rocks. Institution of Mining and Metallurgy, London, pp. 415 – 423. Matveev, S., Ballhaus, C., 2002. Origin of podiform chromite deposits. Earth Planet. Sci. Lett. 203, 235 – 243. McDonough, W.F., Sun, S.-S., 1995. Composition of the Earth. Chem. Geol. 120, 223 – 253. McElduff, B., Stumpfl, E.F., 1990. Platinum-Group Minerals from the Troodos Ophiolite, Cyprus. Mineral. Petrol. 42, 211 – 232. McInnes, B.I.A., McBride, J.S., Evans, N.J., Lambert, D.D., Andrew, A.S., 1999. Osmium isotope constraints on ore metal recycling in subduction zones. Science 286, 512 – 516. Meisel, T., Walker, R.J., Irving, A.J., Lorand, J.-P., 2001. Osmium isotopic compositions of mantle xenoliths: a global perspective. Geochim. Cosmochim. Acta 65, 1311 – 1323. Melcher, F., Grum, W., Simon, G., v. Thalhammer, T., Stumpfl, E.F., 1997. Petrogenesis of the ophiolitic giant chromite deposits of Kempirsai, Kazakhstan: a study of solid and fluid inclusions in chromite. J. Petrol. 38 (10), 1419 – 1458. Melcher, F., Grum, W., Thalhammer, T.V., Thalhammer, O.A.R., 1999. The giant chromite deposits at Kempirsai, Urals: constraints from trace element (PGE, REE) and isotopic data. Miner. Depos. 34, 250 – 272. Mitchell, R.H., Keays, R.R., 1981. Abundance and distribution of gold, palladium and iridium in some spinel and carnet lherzolites: implications for the nature and origin of precious metalrich intergranular component in the upper mantle. Geochim. Cosmochim. Acta 45, 2425 – 2442. Morgan, J.W., 1986. Ultramafic xenoliths; clue to the earth’s late accretionary history. J. Geophys. Res. 91, 12375 – 12387. Mukasa, S.B., Ludden, J.N., 1987. Uranium – lead ages of plagiogranites from the Troodos Ophiolite, Cyprus, and their tectonic significance. Geology 15, 825 – 828. Naldrett, A.J., Duke, J.M., 1980. Platinum metals in magmatic sulfide ores. Science 208, 1417 – 1424. Naldrett, A.J., Lehmann, J., Auge, T., 1989. Spinel non-stoichiometry and reactions between chromite and closely associated sulphides, with examples from ophiolite complexes. In:

231

Prendergast, M.D., Jones, M.J. (Eds.), Magmatic Sulphides— The Zimbabwe Volume. The Institution of Mining and Metallurgy, England, pp. 221 – 227. Naldrett, A.J., Bru¨gmann, G.E., Wilson, A.H., 1990. Models for the concentration of PGE in layered intrusions. Can. Mineral. 28, 389 – 408. Nicolas, A., 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Kluwer Academic Publishing, Dordrecht. Page, N.J., Talkington, R.W., 1984. Palladium, platinum, rhodium, ruthenium, and iridium in peridotites and chromitites from ophiolite complexes in Newfoundland. Can. Mineral. 22, 137 – 149. Page, N.J., Cassard, D., Haffty, J., 1982. Palladium, platinum, rhodium, ruthenium, and iridium in chromitites from the Massif du Sud and Tiebaghi Massif, New Caledonia. Econ. Geol. 77, 1571 – 1577. Paktunc, A.D., 1990. Origin of podiform chromite deposits by multistage melting, melt segregation and magma mixing in the upper mantle. Ore Geol. Rev. 5, 211 – 222. Parkinson, I.J., Hawkesworth, C.J., Cohen, A.S., 1998. Ancient mantle in a modern arc: osmium isotopes in the Izu – Bonin – Mariana forearc peridotites. Science 281, 2011 – 2013. Pearce, J.A., Lippard, S.J., Roberts, S., 1984. Characteristics and tectonic significance of supra-subduction zone ophiolites. In: Kokelaar, B.P., Howells, M.F. (Eds.), Marginal Basin Geology. Spec. Publ.-Geol. Soc., vol. 16. Oxford, Blackwell, pp. 77 – 94. Prichard, H.M., Lord, R.A., 1990. Platinum and palladium in the Troodos ophiolite complex, Cyprus. Can. Mineral. 28, 607 – 617. Prichard, H.M., Lord, R.A., 1993. An overview of the PGE concentrations in the Shetland ophiolite complex. In: Prichard, H.M., Alabaster, T., Harris, N.B.W., Neary, C.R. (Eds.), Magmatic Processes and Plate Tectonics. Special Publication Geological Society (London), vol. 67, pp. 273 – 294. Prichard, H.M., Lord, R.A., Neary, C.R., 1996. A model to explain the occurrence of platinum- and palladium-rich ophiolite complexes. J. Geol. Soc. 153, 323 – 328. Rehka¨mper, M., Halliday, A.N., 1997. Development and application of new ion-exchange techniques for the separation of the platinum-group and other siderophile elements form geological samples. Talanta 44, 663 – 672. Ringwood, A.E., 1970. Petrogenesis of Apollo 11 basalts and implications for lunar origin. J. Geophys. Res. 75 (6), 453 – 479. Roberts, S., 1988. Ophiolitic chromitite formation: a marginal basin phenomenon. Econ. Geol. 83, 1034 – 1036. Roberts, S., Neary, C.R., 1993. Petrogenesis of ophiolitic chromitite. In: Prichard, H.M., Alabaster, T., Harris, N.B.W., Neary, C.R. (Eds.), Magmatic Processes and Plate Tectonics. Spec. Publ.-Geol. Soc., vol. 76, pp. 257 – 272. London. Robinson, P.T., Malpas, J., 1990. The Troodos ophiolite of Cyprus: new perspectives on its origin and emplacement. In: Malpas, J. (Ed.), Ophiolites-Oceanic Crust Analogues. Proceedings of the Symposium ‘‘Troodos 1987’’, Nicosia. Cyprus Geological Survey, pp. 13 – 26. Shirey, S.B., Walker, R.J., 1998. Re – Os isotopes in cosmochemistry and high-temperature geochemistry. Annu. Rev. Earth Planet. Sci. 26, 423 – 500. Suhr, G., 1999. Melt migration under oceanic ridges: interferences

232

A. Bu¨chl et al. / Chemical Geology 208 (2004) 217–232

from reactive transport modeling of upper mantle hosted dunites. J. Petrol. 40, 575 – 599. Talkington, R.W., Watkinson, D.H., 1986. Whole rock platinumgroup element trends in chromite-rich rocks in ophiolitic and stratiform igneous complexes. In: Gallagher, M.J., Ixer, R.A., Neary, C.R., Prichard, H.M. (Eds.), Metallogeny of Basic and Ultrabasic rocks. Symp. Vol., Inst., Min. Metall., pp. 427 – 440. London, UK. Tsuru, A., Walker, R.J., Kontinen, A., Peltonen, P., Hanski, E., 2000. Re – Os isotopic systematics of the Jormua Ophiolite Complex, NW Finland. Chem. Geol. 164, 123 – 141. Vaughan, D.J., Craig, J.R., 1978. Mineral Chemistry of Metal Sulfides. Cambridge Univ. Press, Cambridge, UK. Walker, R.J., Prichard, H.M., Ishiwatari, A., Pimentel, M., 2002a. The osmium isotopic composition of the convecting upper mantle deduced from ophiolite chromites. Geochim. Cosmochim. Acta 66, 329 – 345. Walker, R.J., Horan, M.F., Morgan, J.W., Becker, H., Grossman, J.N., Rubin, A.E., 2002b. Comparative 187Re – 187Os systematics of chondrites: implications regarding early solar system processes. Geochim. Cosmochim. Acta 66, 4187 – 4201.

Woodland, S.J., Pearson, D.G., Thirlwall, M.F., 2002. A platinum group element and Re – Os isotope investigation of siderophile element recycling in subduction zones: comparison of Grenada, Lesser Antilles arc, and the Izu – Bonin arc. J. Petrol. 43 (1), 171 – 198. Zhou, M.-F., Robinson, P.T., 1997. Origin and tectonic environment of podiform chromite deposits. Econ. Geol. 92, 259 – 262. Zhou, M.-F., Robinson, P.T., Bai, W.-J., 1994. Formation of podiform chromitites by melt/rock interaction in the upper mantle. Miner. Depos. 29, 98 – 101. Zhou, M.-F., Robinson, P.T., Malpas, J., Zijin, L., 1996. Podiform chromitites in the Luobusa Ophiolite (Southern Tibet): implications for melt – rock interaction and chromite segregation in the upper mantle. J. Petrol. 37 (1), 3 – 21. Zhou, M.-F., Sun, M., Keays, R.R., Kerrich, R.W., 1998. Controls of platinum-group elemental distributions of podiform chromitites: a case study of high-Cr and high-Al chromitites from Chinese orogenic belts. Geochim. Cosmochim. Acta 62, 677 – 688.