Abundance and distribution of platinum-group elements in orogenic lherzolites; a case study in a Fontete Rouge lherzolite (French Pyrénées)

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Chemical Geology 248 (2008) 174 – 194 www.elsevier.com/locate/chemgeo

Abundance and distribution of platinum-group elements in orogenic lherzolites; a case study in a Fontete Rouge lherzolite (French Pyrénées) Jean-Pierre Lorand a,⁎, Ambre Luguet b , Olivier Alard c , Antoine Bezos d , Thomas Meisel e a

c

Laboratoire « Minéralogie, Pétrologie » CNRS UMR 7160 and Museum National d'Histoire Naturelle, 61 rue Buffon, Paris 75005, France b Department of Earth Sciences, University of Durham, South Road, Durham DH1 3LE, UK Laboratoire de Tectonophysique, CNRS UMR 5568 and ISTEEM, Université de Montpellier II, Pl. E. Bataillon, 34095 Montpellier, France d Department of Earth and Planetary Sciences, Harvard University, 20 Oxford Street, Cambridge 02138 MA, USA e Department of General and Analytical chemistry, University of Leoben, Franz-Josef-Str. 18, A-8700 Leoben, Austria Received 15 January 2007; received in revised form 25 June 2007; accepted 29 June 2007

Abstract Orogenic lherzolites exhibit a specific (although reproducible) platinum-group element signature characterized by slight positive deviations of Pd/Ir, Rh/Ir and Ru/Ir ratios from the canonical chondritic model. Such a signature was alternatively considered to be a compositional feature of the primitive upper mantle or resulting from sulfide melt addition by refertilisation reactions in the continental lithosphere. To shed additional light on this conundrum, the distribution of PGE in FON B 93, an unserpentinized, fertile (3.24% Al2O3) orogenic lherzolite (French Pyrenees) used as in-house standard by our group, has been studied to different scales, from the bulk rock to trace minerals. In addition to new ICPMS analyses after separation into NiS and Te coprecipitation, its whole-rock PGE concentrations were redetermined by three different laboratories using very precise ID-ICPMS methods after digestion of the sample in Carius tube at T = 250 °C and 320 °C or in high-pressure asher (HPA-S). The four methods produce reproducible Ru, Rh and Pd contents (7.1 ± 0.18 ppb; 1.43 ± 0.05 ppb; 7.1 ± 0.30 ppb) whereas the non-ID NiS-fire assay method underestimates the Os and Ir concentrations by c.a. 10 and 15% compared to ID-ICP-MS analyses (4.40 ± 0.07 and 4.00 ± 0.17 ppb, respectively). Platinum was the most difficult to analyse. If performed on powder aliquots smaller than 3 g, the IDICP-MS analyses generate strong nugget effects while the non-ID NiS-fire assay method yields statistically lower (but highly reproducible) Pt concentrations (6.92 ± 0.26 ppb). These features reflect the strong partitioning of Pt into trace phases that SEM and laser ablation-ICP-MS analyses on thin sections identify as both high-temperature Pt–Ir–Os alloys and Pt–(Pd) tellurides of likely subsolidus origin; ten to fifteen grains ranging in maximum dimension from a few micrometres to a few hundreds of nanometers, were identified by standard polished thin section. LA-ICP-MS data on base metal sulfides (25 grains analysed), coupled with the whole-rock S concentration (277 ± 10 ppm) and modal composition of the BMS (90% pentlandite + accessory pyrite and secondary pyrrhotite + 10% chalcopyrite) allowed the contribution of the BMS phase to the PGE budget of FON B 93 to be estimated. Except Pt that exhibits a 95% deficit in the BMS phase, the PGE concentrations measured by ID-ICP-MS can be balanced by BMS while Al-spinel is a negligible contributor, accounting for less than 0.5% of the Ru budget. The occurrence of Pt in trace phases may bias the whole-rock Pt concentrations because 1) mechanical collection of Pt-rich trace phases remains problematic with the NiS button, and 2) Pt–Ir–Os alloys may be prone to digestion problems in conventional Carius tube procedures. Since they could be stable at ⁎ Corresponding author. E-mail address: [email protected] (J.-P. Lorand). 0009-2541/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2007.06.030

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mantle depth, such refractory alloys that contain Os and Ir may also have enhanced the heavy PGEs/light PGEs fractionation. Our observations likely pertain to orogenic lherzolites as a whole because BMS assemblages in these mantle rocks record evolution at low sulfur fugacity, which prevents Pt–Ir–Os alloy from entering the Mss in the mantle; moreover, at subsolidus temperature, pentlandite, the major BMS, cannot accommodate 0 valence state Pt. © 2007 Elsevier B.V. All rights reserved. Keywords: Upper mantle; Orogenic lherzolites; Platinum-group elements; Platinum-group minerals; Base metal sulfides

1. Introduction Platinum-group elements (PGE — Os, Ir, Ru, Rh, Pt, Pd) have been extensively analyzed in mantle rocks over the last 10 years, thus challenging the widely used lithophile trace elements as geochemical tracers of the terrestrial mantle (e.g. Handler and Bennett, 1999; Lorand et al., 1999; Lorand and Alard, 2001; Büchl et al., 2002; Lee, 2002; Luguet et al., 2003; Lorand et al., 2004). Like rare-earth elements (REE), PGE display contrasting geochemical properties with respect to igneous petrogenetic processes. Os, Ir and Ru, the iridium-group PGEs-(IPGE) are refractory, strongly compatible elements whereas Rh, Pt, and Pd, the palladium-group PGEs, (PPGE) are less compatible and more volatile (Barnes et al., 1985). Unlike REE however, PGE occur at very low concentration levels (ng/g) in mantle rocks, due to their extraction from the Primitive Earth into the metallic core. As highly siderophile (“iron-loving”) elements, PGE offer great potentiality for discussing global-scale processes such as core–mantle segregation processes and core–mantle exchanges as well as the late accretion history of the Earth which likely involved addition of PGE-rich undifferentiated materials (Morgan, 1986; Becker et al., 2006 and reference therein). Orogenic lherzolites exhibit a specific platinumgroup element signature characterized by slight positive deviations of Pd/Ir, Rh/Ir and Ru/Ir ratios from the canonical chondritic model. Such a signature is far from being understood. Highly reproducible from one occurrence to another, the light PGE (i.e. Ru, Rh, Pd) enrichment of orogenic lherzolites is usually interpreted as a compositional feature of the primitive upper mantle (e.g. Becker et al., 2006 and references therein). By contrast, other authors (e.g. Rehkämper et al., 1999) consider it to be a young feature, reflecting deposition of Fe–Ni–Cu sulfides (i.e. base metal sulfides: BMS) by magma percolating into the continental lithosphere. PGE are strongly chalcophile elements in the presentday metal-undersaturated upper mantle. Thus they partition strongly into accessory (b 0.1 vol.%) BMS,

being mostly excluded from major mantle minerals, except perhaps from Cr-spinel (e.g. Pattou et al., 1996; Burton et al., 1999; Alard et al., 2000). Whether BMS alone control the PGE inventory of orogenic lherzolites is unclear since no detailed mass balance calculation of the contribution of the different PGE mineral carriers has yet been provided. To shed additional light on this conundrum, the distribution of PGE in FON B 93, a fertile orogenic lherzolite from the French Pyrenees used as in-house standard in our lab (Gros et al., 2002), has been studied to different scales. In addition of being crucial for understanding the PGE systematic of orogenic lherzolites, such a study may help to evaluate the reliability of whole-rock PGE analyses. UB-N and GP-13, the two reference matrix-matched rock standards for peridotite analyses (i.e. samples showing concentration ranges and matrix mineralogy similar to the sample to be analysed; Meisel and Moser, 2004) are orogenic lherzolites. Meisel et al. (2003) suggested that part of the IPGE inventory of UB-N resides in Al-spinel and/or Os–Ir–Ru micro-alloys, two minerals which are acidresistant and hard to dissolve, unless perhaps by highpressure high-temperature attack in aqua regia. Micrometric to nanometric Pt-rich discrete minerals were occasionally reported to occur inside the BMS phase of orogenic lherzolites (Alard et al., 2000; Luguet et al., 2001; 2004) supporting previous petrographic observations (Garuti et al., 1984; Lorand et al., 1999). This specific speciation as ultra-trace elements residing in accessory minerals makes whole-rock PGE analyses prone to strong nugget effects from unevenly distributed PGE mineral carriers. While the nugget effect may be considerably reduced compared to analytical problems by analyzing large (usually N 2–3 g) sample powder aliquots (Meisel and Moser, 2004) the efficiency of sample digestion procedures with respect to the recovery of PGEs critically depends on PGE mineral carriers. We report 14 new bulk-rock analyses obtained from four different laboratories working with the two separation procedures currently used: the NiS-fire

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assay method followed by tellurium co-precipitation and the dissolution in aqua regia using either sealed glass tubes (Carius tubes: CT) or high-pressure dissolution device (HPA-S) followed by anion or cation exchange resin separation. Whole-rock analyses have been implemented by a petrographic study of polished thin section and in-situ analyses of BMS by electron microprobe and laser ablation (LAM) ICP-MS to calculate the mass fraction of PGE that is accounted for by BMS. Discrete platinum-group minerals (PGM) were searched for by scanning electron microscope (SEM). Compared to UB-N and GP-13, FON B 93 is virtually devoid of serpentinization minerals; thus, it is expected to have preserved BMS-PGM assemblages reflecting more closely mantle source conditions (Lorand, 1985, 1989a,b). 2. Main petrographic features of matrix minerals FON B 93 is a fertile mantle lherzolite collected in the Fontête Rouge massif, one of the 40 orogenic lherzolite bodies scattered throughout the French Pyrenées (Fabriès et al., 1991). It shows a mediumgrain size (secondary) protogranular texture displaying millimeter-sized spinel–orthopyroxene–clinopyroxene clusters. The modal composition of FON B 93, estimated by a least-square routine from the bulk-rock analysis and analyses of silicate and spinel is (in wt.%): 49% olivine (Fo89.9), 33% orthopyroxene (En89–Fs10– Wo11), 15% Al-rich clinopyroxene (En4.3–Fs47.4– Wo48.3) and 3% Al-rich spinel (Cr/Cr + Al = 0.092) (Lorand, unpublished). FON B 93 was analysed as FON-2 for major elements and Cu by Bodinier et al. (1988) and as FON-B for S by Lorand (1989a). The major element composition is similar to those of UB-N and GP-13 (3.23 wt.% Al2O3 vs. 3.45 in GP-13 and 2.99 in UB-N; 3.18 wt.% CaO vs. 3.00 in GP-13; Govindaraju, 1982; Pearson et al., 2004). FON B 93 is as fresh as mantle xenoliths sampled in alkali lavas. The lack of serpentinization (L.O.I. bb 1 wt.%) and of any metamorphic recrystallization products in FON B 93 are two major differences compared with UB-N which displays metamorphic reactions (garnet replacement by Al-rich spinel) and extensive hydrothermal alteration (see Meisel et al., 2003). 3. Analytical methods Fe–Ni–Cu sulfides and platinum-group minerals (PGM) were searched on four standard-sized (40 × 25 mm) polished thin sections using reflected light microscopy and investigated using scanning electron

microscope (SEM) (JEOL JSM 840, Pierre and Marie Curie University, Paris VI) operating in the backscattered mode and equipped with a energy dispersive Si(L) detector with a resolution of 129 eV full width at half maximum at the Fe peak. The SEM investigations were used to identify platinum-group minerals qualitatively. Platinum-group element contents of base metal sulfides were determined along with a few other trace elements (Te, Bi, Pb) by laser ablation induced coupled plasma mass spectrometry (LA-ICP-MS). The laser employed was a GEOLAS excimer UV laser operating at 193 nm, coupled with a FISONS VG 353 PlasmaQuad PQplus inductively coupled plasma mass spectrometer (ICP-MS) (University of Montpellier II and National Museum of Natural History (MNHN, Paris)). Ablation was carried out using a helium carrier gas, and ablated products were transferred to the plasma quad in a pure Ar atmosphere (0.85 l/min). Analytical conditions included a 30–50 μm beam diameter, 5 Hz laser frequency and a beam energy about 15 mJ/pulse. Sulfur contents determined by electron microprobe were used as internal standard. Raw data were processed on-line using the GLITTER software package (Van Achterbergh et al., 1999). The PGEs were measured in the peak jumping mode on 99 Ru, 101 Ru, 103 Rh, 105 Pd, 106 Pd and 108 Pd, 190 Os, 192 Os, 193 Ir, 195 Pt, 196 Pt. 63 Cu40 Ar interference on 103 Rh (monoisotopic) was corrected by ablating a PGE-free synthetic Cu metal (PROLABO™) several times during the run and determining the production rate of 63 Cu40 Ar (0.0012 ± 0.0006 63 Cu40 Ar CPS/63 Cu CPS). The accuracy of the correction was checked by correcting 105 Pd for 65 Cu 40 Ar interference and comparing it to 106 Pd. 106 Pd is free of major interference, except for 66 Zn40 Ar but Zn abundance in mantle sulfide is generally low b 0.3 wt.%. PGE-A, a PGE-doped NiS sulfide bead similar in matrix to mantle sulfides (Alard et al., 2000) was used as external standard. Typical detection limit, for the conditions described above, is lower than 40 ppb for all PGE, but Ru showing a 70 ppb detection limit. Two kilograms of rock chips cleaned from weathered surfaces was finely ground in an agate mortar for wholerock S, Se, Te and PGE analyses. Sulfur was re-analysed on four powder batches by iodometric titration of the SO2 produced by combustion of 500 mg powder aliquot at the National Museum of Natural History (see Gros et al., 2005 for further details). Selenium and tellurium were analysed at Geosciences Laboratory (Sudbury, Canada) by hydride Generation and ICP-MS. Accuracy of Se and Te analyses were checked against UB-N, included in the sample batch (Table 1).

J.-P. Lorand et al. / Chemical Geology 248 (2008) 174–194 Table 1 Bulk-rock analyses of FON B 93, compared with GP-13 and UB-N

SiO2 Al2O3 Fe2O3 (T) MnO MgO CaO Na2O K2O TiO2 L.O.I. [Mg] S (ppm) Cu (ppm) Se (ppb) Te (ppb)

FON B 93

GP-13

UB-N

46.26 3.23 8.32 0.12 37.90 3.18 0.27 0.002 0.12 0.04 0.90 277 ± 10⁎ 31 87.3⁎; 86.4⁎ 9.2⁎; 9.7⁎

44.91 3.45 8.19 0.15 39.79 3.01 0.3

44.80 3.30 8.34 0.136 40.01 1.36 0.11 0.02 0.12 12.06 0.895 200 ± 66 (138⁎, 141⁎) 28 ± 8.49 112 ± 3 (123⁎) 8.4 (8.9⁎)

0.14 0.90

Data for FON B 93 are from Bodinier et al. (1988) and Lorand (1989a). Data for GP-13 are from Pearson et al. (2004) and from Govindaraju (1982) and Terashima and Imai (2000) for UB-N. ⁎: This study. [Mg] = Mg2+ / [Mg2+ + Fe2+].

Whole-rock platinum-group element analyses were performed following five different sample digestion and separation procedures. 1) PGEs were separated at the MNHN from 4 powder aliquots of 15 g each by a NiS-fire assay — Te coprecipitation procedure as described by Gros et al. (2002). To improve the recovery of PGE, the procedure was nevertheless modified by adding a second NiS-fire assay step: after separation of the NiS button, the borosilicate glass of the first fusion charge was crushed and melted again along with Ni, S and smelting reagents in similar proportions as in the first fusion step to produce a second NiS button weighing c.a. 7.3–7.5 g. Both buttons were then processed as described by Gros et al. (2002), except for the Te coprecipitation stage which involves the addition of 28 ml (instead of 7 ml) Te, to improve the efficiency of the Te coprecipitation stage. The ICPMS measurements were carried out in the peakjumping mode, and each element was counted three times for 40 s after 150 s rinsing and 45 s uptake. Ions counts were converted into concentrations using 5 and 10 ppb external calibration standards. 2) Three powder aliquots of 3 g each were digested in standard Carius tubes (CT), using 9 ml of reverse aqua regia (2/3 HNO3 and 1/3 HCl in vol.) at the MNHN. The digestion was performed in a homemade oven containing five horizontal inox jackets (one CT/jacket) surrounded by heating wire. This design allowed more precise temperature monitoring

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(i.e. 250 °C) inside each inox jacket, and thus, a better attack. The resulting mixture (acids + silicate gel) after the digestion step was transferred into 60 ml Teflon PFA beakers and desilicified using 2 ml 15% HCl and 10 ml 40% HF at 80–90 °C. The residue was then converted into chloride form by drying down twice with 5 ml of 6 M HCl prior to separation of the PGE into three fractions (Ru, Pd and Ir–Pt) by an anion-exchange separation method. PGE concentrations were determined by isotope dilution measuring 99 Ru/101Ru, 191 Ir/193Ir, 198Pt/195Pt and 105 Pd/108Pd on the Montpellier-MNHN ICP-MS. Direct interferences of 108 Cd on 108 Pd, and of 198 Hg on 198 Pt were corrected on line during data acquisition. Isobaric interferences of 177 Hf16O on 193 Ir and 178Hf16O on194Pt were monitored for each batch of analysis and corrected during data reduction despite the low oxide production measured (b1.3–1%). Procedural blanks range between b1 pg/g for Ru and 160 pg/g for Pt for 3 g of sample (Table 3). More detailed explanations of the analytical method are provided in Bézos et al. (2005). 3) Two powder aliquots of 1 g each were processed at the Department of Terrestrial Magnetism (DTM), Carnegie Institution of Washington, USA, with a modified digestion technique using 9 ml reverse aqua regia and quartz glass Carius tubes at 270 °C. After loading the samples and spikes in the Quartz Carius tube and the subsequent sealing of the tubes, the Carius tubes were inserted into a steel pressure vessel containing ≈20 g dry ice. The CO2 pressure that builds up inside the pressure vessel upon warming counterbalances the internal pressure in the Carius tube at 270 °C. After the overnight digestion, osmium was separated from reverse aqua regia by solvent extraction into CCl4 and back extraction into HBr, followed by micro distillation. Re, Ru, Ir, Pt and Pd, present in the remaining aqua regia fraction, were separated by anion exchange chromatography. Detailed digestion procedure and accuracy tests of this high-pressure Carius tube digestion method were described by Becker et al. (2006) and Luguet et al. (2007) while Horan et al. (2003) provided more information on the column chemistry. Osmium concentrations were determined by negative thermal ionization mass spectrometry using the DTM N-TIMS. The PGE cuts were run separately on the DTM P54 ICP-MS. Average procedural blanks are b4 pg except 10 pg for Pt and 23 pg for Pd (Luguet et al., 2007). 4) A fourth batch of three powder aliquots were analysed at the University of Leoben, Department of General and Analytical Chemistry (Austria), by isotope

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dilution ICP-MS using a simple on-line chromatographic column for analyte–matrix separation after high-pressure asher (HPA-S) digestion (Meisel et al., 2003; Meisel and Moser, 2004). 2 g of powder was weighted into 50 ml quartz glass vessels. A mixed Ru, Pd, Os, Ir and Pt spike, 2 ml sub-boiled concentrated HCl and 5 ml sub-boiled concentrated HNO3 were added. The glass vessels were then sealed with Teflon tape and a quartz glass lid. The vials (six samples and one blank) were heated (300 °C) under pressure (125 bar) in the high-pressure asher (HPA-S, Anton Paar, Graz, Austria) for 3 h. Upon cooling, Os was sparged into the ICP-MS system (HP4500 Agilent Technologies). The sample was then centrifuged to remove the solid residue and the remaining liquid was dried down gently to near dryness. The residue was dissolved in 2 ml 0.1 mol− 1 l HCl and loaded after filtering onto a 1-m-long, 13-ml-volume bed cation exchange column, filled with Dowex AG50Wx8 resin (200–400 mesh). HCl (0.1 mol− 1 l) was used as the mobile phase. Matrix cations such as Fe3+ remain on the column whereas the eluent transports the anionic PGE chloro-complexes without retention directly to the peristaltic pump of the sample introduction system. Rh concentrations were calculated through comparison of sample and a standard peak area of known concentration, assuming the same behavior of Rh and other PGE. Average total procedural blanks (n = 12) were below 2 pg for Ru, Rh, Pd, Os and Ir. Platinum blanks were somewhat higher, on average 16 pg. 5) A last batch of two powder aliquots, weighing 2 g and 2.5 g respectively, was processed by ID-ICP-MS at the Arthur Holmes Isotopic Geology Laboratory of University of Durham using an Anton Paar HPA-S. Sample digestion was performed overnight at 300 °C and 120 bars in pressure-sealed quartz vessels containing 7.5 ml reverse aqua regia. After sample digestion, osmium was recovered by CCl4 solvent extraction and back-extracted in HBr before a final microdistillation. The aqua regia fraction was dried down and then redissolved in 0.5 M HCl prior to its loading on the BioRad AG1X8 anion resin for PGE separation. Re, Ru, Ir and Pt were recovered in 13.5 M HNO3 and Pd in 9 M HCl. Details of the column chemistry were provided in Horan et al. (2003) and Pearson et al. (2004). In addition to the whole-rock analyses, two Alspinel fractions (106 and 70 mg) were analysed following the same procedure in order to access precisely the contribution of spinel in the whole-rock PGE budget. The Al-spinels were separated from fresh, ethanol-washed fragments from the core of the rock, lightly crushed in an agate mortar and finally hand-

picked under a binocular microscope at magnification 30× to 100× to avoid selecting any spinel-BMS composite grains. Osmium isotopic composition and concentration were determined by negative thermal ionisation mass spectrometry using the ThermoFinnigan Triton following the procedure outlined in Pearson and Nowell (2004). PGE cuts were analysed by ICP-MS using the Thermo-Finnigan Element 2. Average total procedural blanks (n = 5) are b1.3 pg for Os, Ir, Ru and Pt, 5.2 pg for Re and 18.2 pg for Pd. 4. Results 4.1. Mineralogy of base metal sulfides and platinumgroup minerals BMS are accessory intergranular minerals with respect to silicate and spinel, although some droplets are partly embayed into recrystallized olivine. Their modal content average estimated from the four polished thin section by image analysis of digitized microphotographs is 0.08 ± 0.01 wt.%. BMS are evenly distributed throughout the four polished thin sections. About half of the BMS grains share a contact with olivine, while there is no preferential concentration in the vicinity of clinopyroxene and spinel, in contrast to the BMS distribution in abyssal peridotites (e.g. Luguet et al., 2003). BMS grains are ovoïds, rounded to ellipsoidal droplets, 50 × 100 μm on average, typical of metal-rich sulfide melts in the Cu–Fe–Ni–S system (Bockrath et al., 2004). Larger grains are polyedral with convex inward or convex outward grain boundaries (Fig. 1A–D). The BMS assemblage is characteristic of Pyrenean lherzolites and other weakly serpentinized orogenic lherzolites studied so far (Garuti et al., 1984; Lorand, 1985, 1989b). Blocky pentlandite is by far predominant and compositionally homogeneous (Fe/Niat. = 1.15 ±0.15; N = 47 electron probe microanalyses-EPMA). It is locally intergrown with Co-rich pyrite blebs (not more than a few percent of total volume) partly replaced by polycrystalline hexagonal-type pyrrhotite. The second most important BMS is chalcopyrite (Cu/Feat. = 0.96 ± 0.04; N = 7 EPMA). The whole-rock mean pentlandite/ chalcopyrite modal ratio, estimated from digitized photomicrographs of 150 BMS grains is about 90/10. Chalcopyrite is generally associated with pentlandite in polyphase grains which display pentlandite/chalcopyrite modal ratios ranging between 100/0 and 40/60.The two sulfides are separated by sharp to diffuse contacts (Fig. 1). The chalcopyrite usually contains pentlandite blebs reflecting the extensive mutual solubility of Cu and Ni in the high-temperature sulfide precursor of both sulfides (Hz-ISS; Peregoedova and Ohnenstetter, 2002).

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Fig. 1. Photomicrographs of base metal sulfides and platinum-group minerals (A to B: plane polarized reflected light; C to H: BSE images). A: Dropshaped two-phase BMS grain; pentlandite (upper right) is both blocky and fine-grained; chalcopyrite (lower left hand side) encloses pentlandite blebs. B: Polyedral pentlandite–chalcopyrite BMS grain; C and D: BSE images of two pentlandite grains displaying lobate grain boundaries and very low dihedral angles (b60°), illustrating the high wetting capacity of sulfide melts in the mantle; E: Merenskyite-type PGM on the wall of a chalcopyrite grain; F: Enlargement of Fig. 1D showing a Pt–Ir–(Os) needle inside pentlandite; F and G: Inclusions of Pt–Te–Bi phases in pentlandite and chalcopyrite.

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Chalcopyrite may also occur as a few monophase grains. In such grains, chalcopyrite stringers protrude inside mutual contacts between matrix silicates. The four polished thin sections contain on average 12 to 14 platinum-group minerals (PGM) which were identified during SEM investigation in the BSE mode. PGM grains range in size between b0.1 μm and 7 μm × 0.5 μm, most of them showing maximum dimensions of 1–2 μm × 0.3–0.5 μm (Fig. 1). Because their small size precludes quantitative analysis by electron microprobe, PGM were identified by EDS spectra, combined with laser ablation count vs. time diagrams (see below). PGM grains are monophase and dominated by Pt–Ir–(Os) alloys and Pt tellurides containing subordinate amounts of Bi and Pd. Both were found to occur in massive pentlandite or to coexist in two-phase Pn–Cp grains. However 80% of Pt–Ir–(Os) alloys occurs inside pure pentlandite while a greater proportion of Pt– Te–Bi phases (N50%) was observed in chalcopyrite-rich polyphase grains. In one telluride grain, Pd is the major element and Pt occurs in subordinated amount. Pt–Ir– (Os) alloys tend to be acicular whereas Pt tellurides are platelets, sometimes highly contorted (Fig. 1E–G). The Pd–(Pt) telluride is a sub-equant particle pasted on the margin of a chalcopyrite grain (Fig. 1H).

LA-ICP-MS is a powerful tool for detecting PGM microparticles located inside BMS and not intercepted during the confection of thin sections (Alard et al., 2000; Luguet et al., 2004). Seven LA analyses (over a total of 35) revealed PGE concentration spikes in the count vs. time diagrams depicting the variation of LAM-ICP-MS spectra (Fig. 2). A single BMS grain can contain at least two PGM microparticles. Fig. 2A shows two spikes corresponding to Pt–Ir (Os) alloys in one pentlandite– chalcopyrite composite grain. Such alloys systematically contain Ir and Os in large amounts. As shown by Fig. 2B, Pt–Ir(Os) alloys may coexist with Pt–Te–(Bi) minerals in the same BMS grain. 4.2. Bulk-rock S and PGE concentrations FON B 93 displays S, Se and Te concentrations (and S/Se ratios as well) close to currently accepted primitive mantle estimates (McDonough and Sun, 1995). The four new S analyses give an average S content (277 ± 10 ppm) c.a. 20% higher than that reported by Lorand (1989a). This difference likely results from inhomogeneities between the old and new powder batches, as well as the smaller grain size of the new powder batch, which allowed a better reaction of vanadium pentoxide with the

Fig. 2. CI chondrite-normalized whole-rock PGE abundances. Normalizing values after McDonough and Sun (1995). See text for acronyms.

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BMS during combustion. The bulk-rock Cu/S ratio thus defined (0.11 ± 0.005; Table 1) is therefore consistent with the modal proportion of chalcopyrite in the BMS phase (10 ± 2%) as estimated from image analysis on thin section. The new whole-rock PGE data are compiled in Table 2 and normalized to chondrite in Fig. 3. FON B 93 displays the characteristic chondrite-normalized PGE patterns of orogenic lherzolites. Apart from two replicates biased by Pt nugget effect (see below), Ru, Rh and Pd (i.e. light PGEs) define slight positive anomalies with respect to Os, Ir and Pt (heavy PGE) which preserve chondritic ratios. There are no straightforward relationships between light PGE concentrations and the different digestion techniques Table 2 Bulk-rock PGE contents of the lherzolite FON B 93 (in ppb) Os

Ir

Ru

Rh

Pt

Pd

Method 1: NiS and Te coprecipitation (Gros et al., 2002) 4.05 3.48 6.75 1.36 6.90 6.32 4.04 3.51 6.94 1.42 7.26 6.86 3.96 3.42 7.01 1.38 6.62 6.81 3.82 3.43 7.24 1.42 6.95 6.97 4.00 3.46 7.00 1.40 6.92 6.74 (0.11) (0.04) (0.20) (0.03) (0.26) (0.29) Method 2: CT-ID-ICP-MS at 250 °C (Bézos et al., 4.1 7.27 8.26 3.87 7.24 7.85 4.03 7.15 8.43 4.00 7.22 8.18 (0.12) (0.06) (0.3)

Au 1.80 1.81 1.76 1.67 1.76 (0.06)

2005) 6.25 6.74 6.14 6.38 (0.32)

Method 3: CT-ID-ICP-MS at 270 °C (Becker et al., 2006, this study) 4.33 3.47 7.89 6.92 6.28 4.31 3.70 8.19 14.7 6.42 Method 4: HPA-S-ID-ICP-MS at 300 °C (Meisel and Moser, 2004) 4.37 3.93 7.35 1.43 8.06 6.97 4.51 4.10 7.04 1.50 15.7 7.32 4.44 3.88 7.10 1.48 8.5 7.56 4.44 3.97 7.16 1.47 10.8 7.28 (0.09) (0.12) (0.14) (0.03) (4.32) (0.29) Method 5: HPA-S-ID-ICP-MS at 300 °C (Luguet et al., unpublished data) 3.73 4.21 8.41 8.57 7.33 4.43 4.19 8.10 8.97 7.32 Best « preferred » average 4.40 4.00 7.11 (0.07) (0.17) (0.18)

1.43 (0.05)

8.38 (0.37)

7.1 (0.3)

Weighed mean are in bold and 1σ standard deviations are between brackets. The average Ru, Rh and Pd concentrations determined by all of the tested procedures overlap with 1σ confidence interval. All values (but Ru from method 2 and method 5) were pooled. Methods 3 and 4 Pt data were obviously biased by nugget effects. Method 1 Os, Ir and Pt are systematically too low.

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of the present study. Ruthenium concentrations obtained after NiS fusion (method 1) and low-T CT (method 2) overlap within 1 sigma level with those generated by HPAS digestion (method 4) at Leoben (7.00 ± 0.20, 7.22± 0.06 and 7.16 ± 0.14 ppb respectively). By contrast, methods 3 (high-pressure CT at DTM) and 5 (HPA-S at Durham) yielded about 15% higher (and quite consistent) Ru contents (7.9–8.4 ppb). Those high Ru concentrations do not result from a nugget effect since, unlike heavy PGEs (see below), no Ru rich mineral has been detected in the course of the SEM investigation, nor is the Al-spinel Ru rich. In addition, method 3 and method 5 produced quite reproducible Ru concentrations for variable amount of powder digested (1 g for method 3 and ≥2 g for method 5). Isobaric interference on 101Ru from nickel argides could be a good alternative explanation to the Ru enhancement in methods 3 and 5; however it is not clear to us why methods 3 and 5 analyses would be more prone to such interferences on Ru data than the 3 other analytical methods. The two data set available for Rh (method 1 and method 4) agree at the 1 sigma level (1.40 ± 0.04 vs. 1.47 ± 0.03 ppb). For Pd, Carius tube techniques (methods 2 and 3) yield the lowest concentrations, which however agree at 1 sigma level with the results of method 1 (NiS-fire assay, Te coprecipitation) (6.38± 0.32 ppb vs. 6.74± 0.29 ppb). Method 4 and method 5 (HPA-S digestion) yield 7% higher values (7.28± 0.29 ppb), yet within the 1 sigma of method 1 mean concentration. As reported in previous papers (Pattou et al., 1996; Gros et al., 2002), FON B 93 displays quite reproducible light PGE contents, even for a 2 g powder sample. Methods 1, 2 and 4, for which 3 replicates or more have been performed yield relative standard deviation (R.S.D.) of 1.1–3.0% for Ru, 2% for Rh and 4–5% for Pd. In contrast to light PGEs, heavy PGE concentrations (Os, Ir and Pt) are indisputably affected by analytical bias (Table 2 and Fig. 3). Both high-P CT (method 3, DTM, Washington) and HPA-S digestion techniques (methods 4 and 5, Leoben and Durham university) yielded, on average, 10% higher Os concentrations, compared with the NiS-fire assay procedures (method 1) (4.4 ± 0.1 vs. 4.00 ± 0.11 ppb). Method 1 Os concentrations are systematically too low. The same is also true for Ir, which unlike Os, was analysed in the five digestion techniques and analytical procedures. Method 1 mean Ir concentrations are about 15 ± 2% lower than those obtained after digestion by low-T CT (method 2) or HPA-S (3.46 ± 0.04 ppb vs. 4.00 ± 0.12, 3.97 ± 0.12 and 4.2 ppb, respectively). High-T CT (method 3) yielded more dispersed values for Ir. Note that FON B 93 is highly homogeneous with respect to Os and Ir concentrations. The relative standard deviations (R.S.D)

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Fig. 3. Variation of LAM-ICP-MS spectra (counts vs. time) collected during two analyses of pentlandite grains. (A) displays two spikes of Pt, Ir and Os concentrations corresponding to Pt–Ir–(Os) microphases. (B) corresponds to analysis of a pentlandite–chalcopyrite mix containing a Pt–Ir–(Os) alloy and a Pt–Te–Bi phase.

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obtained for these elements are very low and almost constant down to sample size of 2 g (2.75–2.0 for Os and 1.15–3% for Ir, respectively). In contrast to Os and Ir, Pt concentrations are poorly reproducible for sample size b3 g. Method 3 and method 4 which operated from 1 g and 2 g powder aliquots, yielded two erratic Pt analyses (Table 2). The corresponding CI-normalized PGE patterns display strongly positive Pt anomalies,

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methods 1 and 2 yield much more reproducible Pt concentrations (R.S.D = 3.74 vs. 3.67%). In method 1, Pt was systematically detected (0.19–0.57 ppb) in the four NiS buttons resulting from the second NiS-fire assays, while the five other PGEs were close to or below detection limits (b 0.1 to 0.05 ppb). The amount of Pt thus detected was combined with the value of the first fusion stage to give the results listed in Table 1.

Fig. 4. CI chondrite-normalized PGE abundances in BMS; A: pentlandite; B: pentlandite + chalcopyrite mixes; C: chalcopyrite.

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Fig. 4 (continued ).

However, these results are 10 to 20% lower than those obtained from the other digestion techniques (both CT and HPA-S). 4.3. Analyses of spinel separates and LA-ICP-MS data The two Al-spinel separates analysed with method 5 are PGE-poor. Os and Ir concentrations range between 0.009–0.227 and 0.095–0.261 ppb, respectively; Pt and Pd concentrations are slightly enriched and less variable (0.29–0.63 and 0.22 to 0.46 ppb, respectively). Only one Ru concentration (1.33 ppb) was considered to be reliable, because of potential isobaric interference with Ni–Ar polyatomic ions on the other spinel Ru analyses. FON B 93 Al-spinel is much poorer in PGE than usually thought for Cr-spinel in mantle rocks (Righter et al., 2004 and ref. therein). LAM-ICP-MS analyses are presented in Fig. 4 and were compiled as weighed means in Table 3. The common occurrence of micrometric PGM as well as the small grain size of the BMS compared with the diameter of laser pits (50–80 μm) was the two most important factors limiting LAM-ICPMS analyses in FON B 93. No concentration was calculated from the analyses displaying Pt concentration spikes in their laser ablation spectra. The analyses showing high Ni (N 25 wt.%) and low Cu (b 1.5 wt.%) were considered to be pure pentlandite analyses while those displaying the opposite (i.e. high Cu N20%) and low Ni (b 2 wt.%) were ascribed to

chalcopyrite. Only one analysis of “pure” chalcopyrite fits these criteria; it has been obtained in the largest monophase grain available in the four polished thin sections studied. Eleven analyses of pure pentlandite have been obtained. No replicate analysis was possible on either pentlandite or chalcopyrite. All the other analyses (22) are mixtures between pentlandite and chalcopyrite and were treated as pentlandite–chalcopyrite mixtures. As previously reported for mantle-derived BMS (Alard et al., 2000; Lorand and Alard, 2001; Luguet et al., 2001, 2004), PGE were found in FON B 93 BMS at ppm to tens of ppm concentration levels. The 11 pure pentlandite analyses are light PGE (Ru, Rh and Pd) as well as Os and Ir-enriched relative to Pt (Fig. 4A). Their CI-normalized PGE patterns are characterized by broadly chondritic IPGE (Ir-group PGE: Os, Ir, Ru) relative abundances varying by a factor of 6 (9–57 × CI chondrites), a slight Rh- and Pd-depletion relative to IPGEs (Rh/IrN = 0.44 ± 0.23; Pd/IrN = 0.82 ± 0.35) and a strong negative Pt anomaly with respect to both IPGE and Pd (Pd/PtN = 53.13 ± 60); pure pentlandites display CI chondrite-normalized patterns closely similar to type 1-pentlandite patterns defined in oceanic peridotites (Luguet et al., 2001, 2004) apart from one order of magnitude higher Pd contents. Pentlandite is Se-rich (139 ± 57 ppm). The analyses that encompassed minor py and its breakdown reaction product (secondary pyrrhotite) yielded distinctly lower Se contents (46– 65 ppm, vs. 139 ± 57 ppm Se). There is no Pt vs Te

J.-P. Lorand et al. / Chemical Geology 248 (2008) 174–194 Table 3 Average (± 1σ) in-situ LAM-ICP-MS analyses of base metal sulfides in FON B 93 (ppm)

Cu Os Ir Ru Rh Pt Pd Te Bi Se

Pn

Pn–Cp

Cp

N = 11

N = 13

N=8

9200 ± 7200 14.0 ± 6 12.34 ± 5.5 15.75 ± 6.8 1.16 ± 0.4 0.55 ± 0.44 9 ± 2.7 3.08 ± 1.3 0.22 ± 0.17 139 ± 57

13.86 ± 5.96 8.79 ± 4.2 14.42 ± 4.87 1.73 ± 0.4 0.55 ± 0.49 12.87 ± 5.47 3.20 ± 2.0 0.79 ± 0.30 139 ± 58

0.82 ± 0.37 0.28 ± 0.09 3.55 ± 1.4 2.59 ± 1.6 0.38 ± 0.1 5.45 ± 1.5 7.45 ± 3.7 – 78 ± 12

N: number of LAM-ICP-MS analyses taken into account for calculation of the average.

positive correlation nor Bi vs. Te, suggesting that the weighted mean contents of Te (3.08 ± 1.3 ppm) and Bi (0.22 ± 0.17 ppm) are not biased by discrete PGM. The pure chalcopyrite analyses are characterized by high contents of Ru (3.55 ± 1.4 ppm), Rh (2.59 ± 1.6 ppm) and Pd (5.45 ± 1.5 ppm) relative to Os (0.82 ± 0.37 ppm), Ir (0.28 ± 0.09 pm) and Pt (0.38 ± 0.1 ppm). Its CInormalized PGE pattern displays a light PGE enrichment relative to heavy PGEs, thus resembling chalcopyrite patterns previously reported by Luguet et al. (2001, 2004). The effects of varying chalcopyrite/pentlandite modal ratios on LAM-ICP-MS results can be seen in Fig. 4C and Table 3. The pentlandite–chalcopyrite mixtures are slightly depleted in IPGE, especially Ir (8.72 ppb) and enriched in Rh and Pd (1.73 ± 1 ppb; 12.87 ± 5.4 ppb) compared to pure pentlandite analyses, yet the two corresponding weighed means overlap at 1 sigma level. Many CI chondrite-normalized PGE patterns show the typical enhancement of the light PGE relative to the heavy PGEs (Rh/IrN = 1.21 ± 0.8; Pd/IrN = 1.34 ± 0.67), coupled with a deep negative Pt anomaly (Pd/PtN = 60 ± 45), that was reported in Type-2 pentlandite from oceanic peridotites (Luguet et al., 2001). Chalcopyrite-rich mixes are also depleted in Se relative to pure pentlandites and pentlandite-rich mixes (78 ± 12 ppm). There is no clear trend for Te and Bi, both varying considerably. 5. Discussion

rocks and ultramafic lavas. The efficiency of the NiSfire assay technique as a whole was recently questioned. Comparing PGE analyses of late archean komatiites (Kostomuskha, Baltic shield and Abitibi, Canada), Puchtel et al. (2004) and Puchtel and Humayun (2005) pointed to a deficit ranging from 22 to 25% for Os, Ir and Pt concentrations to 34% for Ru in the analyses that used the NiS-fire assay separation procedure compared with those based on a CT-ID-ICP-MS procedure. For some orogenic lherzolites previously analysed by Lorand et al. (1999), Becker et al. (2006) obtained PGE concentrations in excess of 17 to 42% with method 3 analytical procedure (high-P CT-ID-ICP-MS). These differences were ascribed to the inability to access and dissolve (or inability to dissolve) all PGE carriers in the sample powders by NiS-fire assay, especially spinel, considered as an important PGE carrier. The present study allows a more documented discussion of the possible source of errors in the NiS–Te separation procedure. As far as light PGEs are concerned, the PGE concentrations of FON B 93 are pretty well reproducible regardless of the digestion method. Accordingly, these elements are likely to be hosted in easily digested and homogeneously distributed minerals such as base metal sulfides. It is worth recalling that the NiS beads after the second fire assay step did not contain detectable amounts of Ru, Rh or Pd. BMS are easily decomposed by the NiSfire assay procedure (Hoffmann and MacLean, 1976) and the two Ni sulfides that crystallize in synthetic NiS beads (Ni1 − xS and Ni3 ± xS2; see Gros et al., 2002 for more details) can dissolve thousands of ppm of Ru, Rh and Pd at 900 °C (Makovicky et al., 1986; Peregoedova

Table 4 Mass-balance estimate of the BMS contribution to the PGE and Te budget of FON B 93 (ppm)

Os Ir Ru Rh Pt Pd Te

A consensus is being reached that isotope dilution (ID) coupled with sample digestion in aqua regia at high pressure and high temperature is most suitable for determining low-level PGE concentrations of mantle

Theoretical BMS a

Measured BMS b

5.60 ± 0.06 5.1 ± 0.3 9.07 ± 0.6 1.80 ± 0.07 10.40 ± 0.40 9.00 ± 0.61 11.9 ± 0.5

12.2 ± 5.4 9.4 ± 4.4 13.5 ± 5.3 1.60 ± 0.8 0.5 ± 0.5 10.0 ± 4.0 3.5 ± 1

Recomputed from the preferred “best” whole-rock estimate of Table 2, by assuming that the bulk of the whole-rock PGE budget resides in a BMS phase composed of 90 ± 2% pentlandite and 10 ± 1% chalcopyrite and representing 0.08 ± 0.01% by weight of the rock modal composition. b Recomputed from average compositions of Pn and Cp listed in Table 3, using the BMS modal composition determined by combining image analysis data and bulk-rock Cu/S ratio (90 ± 2% Pn, 10 ± 1% Cp). a

5.1. Mineralogical effects on PGE separation procedures

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and Ohnenstetter, 2002), much more than the maximum amount calculated for the BMS phase of FON B 93 (Table 4). A simple mass balance calculation based on the “best” whole-rock estimate of Table 2 and the sulfide modal abundance (0.08 ± 0.01) indicates that BMS can effectively account for the light PGE content of FON B 93. Thus, two possible sources of errors may explain the analytical bias detected by Becker et al. (2006) and Puchtel and co-workers: small-sized NiS beads may hamper correct collection of PGE during fire assay while modified stoichiometry of the bead can affect the solubility of PGE into the Ni sulfide phases. A second, more important step for non-ID NiS-Te methods is the Te coprecipitation stage after NiS-fire assay. The four-fold increase in the amount of Te in our Method 1 (28 ml vs. 7 ml in the Gros et al., 2002 original procedure) improved the Ru, Rh and Pd contents by 20 ± 1%, 13– 23% and 5–15%, respectively. Heavy PGE concentrations (Os, Ir and Pt) of FON B 93 were indisputably affected by analytical bias from the NiS–Te procedure. By comparison with the high-T CTand the HPA-S digestion procedures, the NiS–Te procedure systematically underestimated the average Os, Ir and Pt contents by 10, 15 and 18%, respectively. Thus, the NiS– Te procedure results were not taken into account when we calculated the « best » estimate for FON B 93 (Table 2). Refractory PGE carriers that are acid-resistant and therefore difficult to dissolve, unless by high-pressure high-temperature attack in aqua regia are Al-spinel and/or Os–Ir–Ru rich minerals, either alloys or sulfides (i.e. minerals from the laurite–erlichmanite series RuS2– OsS2). Multiple digestion tests led Meisel et al. (2003) to suspect both minerals to be responsible for the systematic Os deficit generated by conventional low-pressure CT digestion procedures of UB-N. Up to 37% of the Os budget of UB-N would be residing in such refractory phases. Unlike UB-N, FON B 93 does not display significant differences in the results of low-pressure and high-pressure (e.g. HPA-S) digestion procedures, thus suggesting Os, Ir and Pt to be residing in easily digested PGE carriers. Al-spinel is too poor in PGE and its modal abundance is too low for this mineral to be a significant contributor to the bulk-rock PGE budget of FON B 93. Its contribution would not exceed 0.5%, taking into account Ru, the most concentrated PGE in Al-spinel. The good reproducibility of Os, Ir and Ru analyses even for sample size b 2 g excludes trace Os–Ir–Ru rich PGM heterogeneously distributed within the spinel structure to be the Os, Ir and Ru host minerals. Likewise, there are no coupled Os/ Ir/Ru spikes in the time-integrated LAM signals. Indeed, due to the very high solubility of Os, Ir and Ru in monosulfide solid solutions, Os–Ir–Ru rich PGM are not

expected to coexist with BMS at magmatic temperature (Brenan and Andrews, 2001; Andrews and Brenan, 2002). The differences of potential Os mineral carriers between UB-N and FON B 93 likely reflect the different subsolidus history of the two samples. UB-N is a strongly serpentinized peridotite which experienced extensive metamorphic recrystallization of garnet that produced abundant Al-rich, Cr-poor spinel (100Cr/Cr + Al = 6–9). Its S/Se ratio is twice lower than PM estimates (Table 1), suggesting remobilization of mantle-derived BMS. We may suspect concomitant liberation of Os–Ir and Ru as alloy trapped inside metamorphic Al-spinel. By contrast, both the S/Se ratios, close to PM estimates and the sulfide melt shape of BMS grains argue against significant crustal alteration in FON B 93. Despite evidence for Pt–Ir–Os alloys, the bulk composition recomputed from the modal composition and LAM-ICP-MS analyses of BMS (« Measured BMS » in Table 4), is clearly in excess for Os (200%) and Ir (180%) compared to the hypothetical BMS phase, recomputed from the whole rock PGE budget (« Theoretical BMS » in Table 4). This discrepancy is likely to be due to the large number of pure pentlandite analyses whereas there are comparatively less analyses of pure chalcopyrite. By contrast, the Pt deficit of c.a. 95% in the measured BMS composition is perfectly accounted for by the Pt-rich trace phases detected by SEM and LAM-ICP-MS analyses. Moreover, such Ptrich phases containing more Pt than Os and Ir explain the following two observations: 1) method 1 generates a deficit with respect to the preferred « best » PGE concentration decreasing from c.a. 18% for Pt to 15% for Ir and 10% for Os, and 2) up to 0.5 ppb Pt was detected in the NiS bead of the second fusion step whereas Ir and Os concentrations are below detection limits of the NiS–Te separation procedure. The same Pt-rich phases may be prone to mechanical and/or chemical problems in the NiS–Te separation procedure. The four-fold increase in the amount of Te improved Ir and Pt data by less than c.a. 3–4% and 5– 6%, respectively (vs. 10% for Os) compared to the mean values reported by Gros et al. (2002). This (and the occurrence of Pt inside the NiS bead after two fusion steps) points to Pt, Ir and Os losses prior to the Te coprecipitation stage, i.e. during the NiS-fire assay, either because of incomplete dissolution of Pt-rich microphases and/or inefficient borosilicate slag–NiS bead separation. The latter may be hampered by the very small size of Pt-rich particles with high surface energy. Unlike Os–Ir–Ru rich minerals, Pt-rich minerals are easily dissolved in hot aqua regia (Pascal, 1963). This is

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the reason why low-P Carius tube dissolution runs, HPA-S and HP-CT techniques performed on FON B 93 yielded highly consistent results. The fact that the latter two techniques yielded erratic Pt concentrations for powder aliquots ≤2 g may result from a nugget effect for Pt from Pt-rich minerals. As any random phenomenon, the nugget effect usually produces both under and overestimated concentrations. The other method 3 test effectively yielded underestimated Pt, far below the preferred Pt concentration (6.92 vs. 8.38; Table 2). Conversely, the Pt excess in one of the three method 4 analyses correlates with slightly higher Os and Ir. Alternatively, one could suggest that high Pt contents could reflect either high Pt blanks or even HfO interferences that were not corrected for. However, method 3 analyses were performed using Quartz Carius tubes and generated very low Pt blanks of 10 ± 5 ppt. Moreover, if HfO were responsible of the high Pt contents, anomalously high Ir contents would also be expected as HfO interfere with 193 Ir as well as 195 Pt. The excellent reproducibility of Ir analyses rules out this analytical explanation as origin of the high Pt contents. 5.2. Origin of Pt-rich PGM and their relationships to base metal sulfides 5.2.1. Pt–Pd–Te–Bi phases; exsolution products of Curich sulfide melt differentiates Mantle peridotites were often presumed to be poor in discrete platinum-group minerals (PGM), owing to their low PGE abundances and the systematic occurrence of base metal sulfides that may concentrate thousands of ppm of PGEs. Our study of FON B 93 obviously questions this assumption, The occurrence of two kinds of Pt-rich microphases, i.e. Pt–Ir–(Os) alloys and Pt– Te–Bi (Pd) phases, characterized by very different stability fields but sometimes coexisting inside the same BMS grain, denote a complex behaviour of Pt compared to other PGE. For the sake of the discussion, Pt–Te–Bi– (Pd) phases will be discussed first. Pt–Te–Bi–(Pd) phases are likely moncheite (PtTe2) displaying important substitution toward inziswaite (PtBi2) and merenskyite (PdTe2). Pure moncheite is stable up to 1150 °C while merenskyite has lower thermal stability (725 °C; Kim et al., 1990). Their crystallization temperatures also decrease significantly with increasing substitution of Bi for Te, and of Pd for Pt. Thus, these phases are not expected to be stable in the convecting upper mantle (Harney and Merkle, 1990; Gervilla and Kojonen, 2002). Pt tellurides are common in ophiolitic peridotites but their origin, magmatic or hydrothermal, is debated, owing to the high degree of

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serpentinization of host peridotites (e.g. Ohnenstetter, 1992; Luguet et al., 2004). The virtual absence of serpentinization precludes an hydrothermal origin for FON B 93 Pt–Te–Bi–(Pd) phases. Since they systematically occur within chalcopyrite or chalcopyrite-rich grains, the way these phases may have formed can be examined in the light of the crystallization history of the Cu-rich BMS. Paragenetic succession of BMS has been extensively discussed by Lorand (1989b) for Pyrenean orogenic peridotites as a whole. At 1.5 Gpa, the average pressure of equilibration of Pyrenean peridotites within the lithospheric continental mantle (Fabriès et al., 1991), the bulk-sulfide composition recomputed from modal compositions and average EMP data should be molten at c.a. 1150 ± 50 °C (Bockrath et al., 2004). On cooling down to 1000 °C, this composition is expected to precipitate almost equal proportions of a Ni-rich monosulfide solid solution (22–25 wt.% Ni) with metal/sulfur (M/S) ratio of 0.9, and a Ni-rich sulfide melt more metal-rich than the Mss (Fig. 5). Note that, unlike implicitely assumed in Lorand (1989b), the small amount of Py (b10% by volume) was not integrated in the estimate of the bulk-sulfide composition because both the S/Se ratio measured in the present study (≥ 10,000) and sulfur isotopic compositions (Chaussidon and Lorand, 1990) question its magmatic origin. Whenever it is taken into account, pyrite would increase M/S to 1.06 still projecting the bulk-sulfide composition of FON B 93 inside the two-phase field Mss + Ni-rich sulfide melt at 1000 °C (Fig. 5). In addition, the presence of sulfide melt is supported by the contorted shapes of most sulfide grains and their wetting textures at the contact of olivine. The field of the Ni-enriched sulfide melt protrudes into the Cu–Fe–Ni–S quaternary system so that the sulfide melt is also enriched in Cu until Iss (Intermediate solid solution) crystallizes at 880840 °C (Craig and Kullerud, 1969; Peregoedova and Ohnenstetter, 2002). As soon as Mss starts to crystallize, Pt, Pd, Te as well as Bi, Sb or As, the other semi-metals are rejected along with Cu into the sulfide melt because of their low Dmss/sulfide melt (b 0.3; Li et al., 1996; Peregoedova et al., 2004; Mungall et al., 2005; Ballhaus et al., 2006). Semi-metals form soft ligands that stabilize Pt and Pd in the sulfide melt, thus further decreasing their Dmss/sulfide melt partition coefficients (Ballhaus and Sylvester, 2000; Helmy et al., 2007). The latter authors report a complete miscibility between sulfide melt and telluride melt above 1150 °C in experimental charges doped with c.a. 70,000 ppm Te + Pt + Pd. Both melts become immiscible at lower temperature and the telluride melt is saturated with respect to Pt-telluride

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Fig. 5. Plot of the recomputed bulk composition of BMS (asterisk) in the Fe–Ni–S system at 1000 °C (after Craig and Kullerud, 1969).

at T N 1015 °C while Pd–Ni telluride saturation is shifted down to subsolidus temperatures (b 700 °C). Like mantle-derived BMS as a whole, FON B 93 BMS are much poorer in Pt–Pd–Te (30 ppm at best, considering all these elements as residing in the BMS phase at magmatic temperatures). Thus, saturation in Pt–Te–(Bi) phases is expected to have been achieved at near-solidus temperature (850–890 °C; c.f. Peregoedova and Ohnenstetter, 2002) from the very last (and the most Pt–Pd–Te-enriched) Cu-sulfide melt fractions (c.f. Barnes et al., 2006). Iss, the solidification product of that Cu-rich sulfide melt does not incorporate Pt (nor Pd) (D Iss/sulfide melt b 0.13 at 840 °C; Peregoedova, 1998). The above scenario explains the tendency for Pt–Te–Bi–(Pd) phases to occur mostly in chalcopyrite-rich grains. The occurrence of single-phase cp grains is perfectly consistent with a Cu-rich derivative sulfide liquid whilst local migration of this liquid during the uplift and solid-state emplacement of orogenic lherzolites at deep crustal levels is reflected by grain-by-grain variations of Cp/ Pn modal ratio. The fact that massive pentlandite may enclose Pt–Te–Bi–(Pd) phases can be due to the pentlandite-forming process at 610 °C which involves both Mss and the high temperature form of heazlewoodite (Hz), the solidification product of the Ni-rich sulfide melt in the Fe–Ni–S system (Craig and Kullerud, 1969). Hz and Iss are now assumed to be a continuous solid solution at 839 °C (Hz–Iss; Peregoedova and Ohnenstetter, 2002) yielding to bi-

phase Cp and Pn grains at T b 600 °C by reaction with Mss. One may infer that Pt-rich tellurides were accidentally trapped within massive pentlandite at this stage of subsolidus recrystallization. Conversely, due to their low temperature of stability, the merenskyite-type grain at the margin of a pentlandite grain is more likely an exsolution product. 5.2.2. Pt–Ir–(Os) alloys, subsolidus exsolution products of Mss or exotic PGM collected by sulfide melts? In contrast to Pt–Te–Bi–(Pd) phases, Pt–Ir (Os) alloys are high-temperature phases (Fleet and Stone, 1991), potentially stable from a redox state point of view, in a mantle more reduced than FMQ, as assumed for the asthenospheric mantle (c.f. Ballhaus, 1995; Mungall et al., 2005; C. Ballhaus, pers. communication to J.P. Lorand). Because of change in space group between Os–Ru alloys (closed-packed hexagonal) and Pt–Ir alloys (body-centered cubic), extensive substitution of Os for Pt is not possible at low-temperature, at least over the T range of pentlandite stability (Harris and Cabri, 1991). Makovicky and Karup-Môller (2002) found up to 29% Ir in Pt3Fe at 1100 °C in the Fe–Pt–Ir– S system; flame-like Ir inclusions are common in natural Pt–Fe alloys (Cabri et al., 1996). The lack of Pt–Ir–(Os) alloy within the silicate matrix imposes to consider a subsolidus origin from the BMS first. Their preferential occurrence in massive pentlandite must be interpreted in the light of the pentlanditeforming reaction from the assemblage Mss + Hz–Iss at

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T b 600 °C. As has been discussed above, the Hz–Iss phase is presumed to have been depleted in Pt by hightemperature formation of Pt–Te–Bi–(Pd) phases. Thus, the Pt that formed alloys was necessarily brought up by the Mss, either as solid solution or as discrete phases. For reasons of lower valence state (+ 1 or 0), Pt (unlike Pd) does not enter the octahedral site of pentlandite (nor chalcopyrite). Peregoedova and Ohnenstetter (2002) provided strong experimental evidence that, in metalrich (b50 at.% S) base metal sulfide assemblages precipitating mainly massive pentlandite, Pt systematically forms its own phases. At 550 °C, the bulk-sulfide composition of FON B 93 plots inside a domain of the Fe9S8–Ni9S8–Cu9S8 plane which makes Pt–Fe/Pt3Fe alloys to coexist with Pd-bearing pentlandite (Fig. 6). By contrast, pentlandite can accommodate Pd, Rh and Ru (up to 12.5 wt.% of each element at 600 °C; Makovicky et al., 1986). Pd is assumed to be moderately compatible into Iss as well as in its breakdown products (chalcopyrite, cubanite; Ballhaus and Sylvester, 2000; Lorand and Alard, 2001; Luguet et al., 2004; Barnes et al., 2006). The average Rh content measured in FON B 93 chalcopyrite (2.59 ± 1.6 ppm) is consistent with the concentrations reported for chalcopyrite and cubanite from Norilsk PGE ores by Barnes et al. (2006) who used a collision cell to suppress the well known 63Cu40Ar interference on 103Rh. Due to their very small size (smaller than the beam diameter of electron microprobes) and their occurrence in Fe-rich matrix sulfides, it was not possible to determine Fe in FON B 93 Pt–Ir–Os alloys. However, if Pt–Ir–(Os) alloys had exsolved from pentlandite,

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then, the Pt contents measured in pentlandite would represent the Pt saturation level at temperatures of final subsolidus reequilibration of BMS assemblages (a few tens of degrees; Lorand, 1989b). Compared with magmatic sulfide ores that (like our sample) effectively equilibrated down to low-T and were saturated in Pt-rich phases (c.f. Barnes et al., 2006), FON B 93 pentlandite appears to be Pt-depleted by at least one order of magnitude, thus suggesting pentlandite precipitation from a Pt-depleted Mss + Hz–Iss assemblage. Like Pt– Te–Bi–(Pd) phases in pentlandite, a few Pt–Ir (Os) alloys were accidentally trapped in chalcopyrite during Iss breakdown at 600 °C, because neither chalcopyrite nor the Iss carry Os, Ir and Pt in solid solution. Another interpretation should consider the Pt–Ir– (Os) alloys observed in FON B 93 as high-temperature exsolutions from the Mss. Due to their rod-shaped morphology, these minerals look like Pt–Ir–Fe exsolutions experimentally obtained by Mss desulfidation reactions at 1000 °C (Peregoedova et al., 2004). To be acceptable, this interpretation must explain why Pt–Ir– (Os) alloys and Pt–Te–Bi–(Pd) phases occur in roughly equal proportions over the four polished thin sections investigated. In other words, Pt must have been equally partitioned between the Mss and the Cu– Ni-rich sulfide melt. For simplicity, this partitioning is assumed to obey a simple mass balance equation: Fmss Cmss + (1 − F)meltCmelt = C0 where Fmss is the weight fraction of Mss, Cmss = concentration of Pt in the Mss, Cmelt = concentration of Pt in the metal-rich sulfide melt and C0 = the theoretical concentration of Pt in the BMS, calculated by assuming the bulk of the whole-

Fig. 6. Stability fields of Pt and Pd phases on the background of principal BMS assemblages (solid lines; field boundaries; stippled: one-phase fields) in the pseudoternary system, Cu9S8–Fe9S8–Ni9S8 at 550 °C (Makovicky, 2002). Broken lines with log fS2 delineate stability fields of the platinum minerals. Other symbols as in Fig. 6.

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rock Pt content dissolved in the initial, homogeneous sulfide melt precursor to the Mss (Table 4). The equation may be rearranged to solve for Cmss mss=melt

Cmss ¼ C0 DPt

mss=melt

=FDPt

þ ð1  F Þ

ð1Þ

Assuming F = 0.5 (see above), C0 = 10.4 ± 0.4 ppm (the theoretical Pt content calculated in Table 4) and an mss/melt average DPt of 0.05 (Li et al., 1996; Barnes et al., 2001; Mungall et al., 2005; Ballhaus et al., 2006), then the Mss is expected to dissolve a maximum Pt content of 1 ± 0.5 ppm Pt and the coexisting sulfide melt about 20 ppm, respectively. This theoretical Pt content is consistent with measurements in metal-rich (M/S = 1) Mss (Lorand and Alard, 2001). As has been discussed above, the sulfide melt precipitated Pt–Te–Bi–(Pd) phases; thus, its Pt content must be consistent with both modal ratios between Pt–Te–Bi–(Pd) and Pt–Ir–(Os) phases and the budget of Te. However, SEM observations suggest roughly similar modal abundances for Pt–Te–Bi–(Pd) and Pt–Ir–(Os) phases in FON B 93. Because tellurium is as chalcophile as PGE in the mantle (sulfide melt/silicate melts partition coefficient N104, Yi et al., 2000), it has been recalculated to 100 wt.% BMS, by dividing the whole-rock concentration (9.5 ppm) by the mean BMS modal content (0.08 ± 0.01 wt.%). Subtracting the amount of Te determined for Pn and Cp (3.5 ±1 ppm; Table 4) gives a deficit of 8.3 ± 1 ppm Te in the BMS assemblage. Natural moncheites display Te/Pt weighed ratios ranging between 1.5 to 2.0 (Ohnenstetter, 1992; Gervilla and Kojonen, 2002). Thus the Pt content bonded to Te probably amounted 4–5 ppm, i.e. was too low by c.a. 50% compared to the 10 ppm predicted by Eq. (1) for F = 0.5 mss/melt and DPt of 0.05. To force the mass balance equation to satisfy the Te and Pt budget (and petrographic data as well) requires considerable changes in the F and D parameters in order to increase the Pt content of the Mss and to decrease that of the sulfide melt. Concerning D, it is well known that PGE solubility in mss/pyrrhotite increases as the sulfur/metal ratio decreases, this latter closely reflecting the number of octahedral vacancies in crystal lattice. This relationship was documented in natural samples (Ballhaus and Ulmer, 1995; Lorand and Alard, 2001) and experimentally reproduced by Makovicky et al. (1986), Li et al. (1996) and Barnes et al. (2001). Majzlan et al. (2002) reported 5 wt.% Pt in Fe0.9 S, decreasing down to detection limits of conventional EMP procedures (500 ppm) for FeS. Eq. (1) yields roughly similar amount of Pt in the mss and the sulfide melt if mss/melt DPt = 0.3 and F = 0.9. Such a high D value is the highest yet reported for strongly non-stoichiometric mss (x, the number of octahedral vacancies ≥0.12–0.14;

Barnes et al., 2001). However, the resulting hypothetical sulfide composition would be too S-rich to precipitate 90% pentlandite as observed in FON B 93. To summarize, Pt–Ir–(Os) alloys are seemingly too abundant to be explained by subsolidus decomposition of BMS, whether pentlandite or precursor Mss. Petrographic data and the mass balance constraints on the Pt budget can be reconciled by assuming that Pt–Ir– (Os) alloys and BMS are not co-genetic. Our assumption is motivated by models that interpret orogenic lherzolites as refertilization products of refractory lithospheric harzburgites by pervasive percolation of tholeiitic basaltic melts. Such models first tested in a PGE study of plagioclase lherzolites from the Horoman peridotites (Japan; Rehkämper et al., 1999) has gained further audience for the past three years from petrological studies of Lanzo (Italy; Müntener et al., 2004), Beni Bousera (Northern Morocco) (Pearson et al., 2004) and very recently, the Lherz massif (France), which belongs to the same cluster of Eastern Pyreneean lherzolites as the Fontête Rouge massif that provided FON B 93. At Lherz, Le Roux et al. (2007) provided indisputable microstructural and trace element evidence that lherzolites are younger than the harzburgites they enclose and bear the inprint of melt-rocks reactions from reactive porous melt flow. By contrast, the harzburgites display PGE and S systematics of solid residues from a high (N20%) degree of partial melting (IrN = 0.012; Ru/ IrN = 1.15; Pt/IrN ≤ 1; Pd/IrN = 0.07–0.3; N = CI chondrite-normalized; 0 b S b 23 ppm; Lorand, 1989a; Lorand et al., 1999; Luguet et al., 2007); as shown by Luguet et al. (2007), the whole-rock PGE budget is hosted in discrete PGM of the laurite–erlichmanite series, Pt–Ir–(Os) rich alloys and complex Cu–Pt sulfides of likely subsolidus origin. Transposed to FON B 93 the refertilization model imposes to consider Pt–Ir–(Os) alloys as inherited from the harzburgitic protolith and BMS from the percolating tholeiitic melt, respectively. This model perfectly accounts for 1) the exclusively intergranular location of BMS in the interstitial pores of the rocks, 2) the lack of trapped Mss inside olivine, the only survival to melt-rock reactions, and 3) the shape of molten sulfides and mineral assemblage of BMS blebs. Sulfur-saturated mantle melts are now assumed to transport S as Cu–Nisulfide melt with M/S N 1 (up to 1.16; Ballhaus et al.; 2006), i.e. similar in bulk compositions to the BMS phases of FON B 93 and more generally to BMS assemblage in abyssal and ophiolitic peridotites that display evidence of such percolation reactions (Luguet et al., 2001, 2003, 2004). Such sulfide melts in equilibrium with primitive tholeiites resulting from

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191

melting degrees of c.a. 15 ± 5% are expected to have been fractionated as regard Pd/Ir (55 ± 20) and Pt/Ir too (30 ± 10) (see Alard et al., 2000; Bézos et al., 2005). Thus, they are thought to have brought up the bulk of the Pd in FON B 93 while contributing to a fraction of the Pt and Rh budget. Of course, minerals of the laurite–erlichmanite series were likely the main contributor to the IPGE budget, despite no such mineral was identified by SEM in FON B 93: the solubility of Os, Ir and Ru in base metal sulfide melt is by far too high (10,000 ppm; Brenan and Andrews, 2001; Andrews and Brenan, 2002 vs. 5–8 ppm in « theoretical » BMS, Table 4) for Os–Ir–Ru rich PGM to coexist with BMS in mantle rocks. Such PGM likely dissolved into the sulfide melt, thus producing the Os, Ir and Ru enrichment over Rh and Pt that characterize CI chondrite-normalized PGE patterns of Cu-poor pentlandite. Unlike the IPGEs, platinum is expected to be stable as alloy (and not as sulfides) in the convecting mantle. Moreover, Pt- and Ir-bearing phases have experimentally been shown to coexist with S-poor pyrrhotite (FeS), even when the latter was doped with trace quantities of Pt or Ir (10 ppm, similar to the Pt concentration in BMS calculated from the whole-rock Pt data; Table 4) (Peregoedova et al., 2004). In the Peregoedova et al. experiments Mss, Cu–Pd-rich sulfide melt and Pt–Ir alloys coexist for log fS2 of − 6.8 to − 7.1 at 1000 °C (Fe– FeS buffer). Redox conditions imposed by the metal-rich sulfide melt likely prevent Pt–Ir–(Os) alloys to dissolve into sulfide melts. The broadly chondritic average Pt/Ir ratio of FON B 93 (2.1 ± 0.15 vs. 2.17 in CI chondrites; McDonough and Sun, 1995) and of orogenic lherzolites as a whole (Becker et al., 2006 and reference therein) could be considered as a benchmark of the stability of Pt–Ir–(Os) alloys in refertilization processes.

Ni-rich sulfide melts, provided that the oxygen fugacity, and in corollary, the sulfur fugacity, are below the FMQ–FeS buffer. The occurrence of Pt as Pt–Ir–(Os) alloys is likely a rule, rather than an exception, in orogenic lherzolites because these latter, like FON B 93, suggest a cooling history of the BMS assemblages within the metal-rich part of the Fe–Ni–(Cu) system. Reflected light microscope studies in the Pyrenean Massifs (Lorand, 1989a,b), Lanzo, (Lorand et al., 1993), Beni Bousera and Ronda (Spain) (Lorand, 1985) as well as in Baldissero and Balmuccia (Italy) (Garuti et al., 1984) identified by order of decreasing modal contents 1) massive pentlandite (70–90 vol.%) 2) chalcopyrite in weakly serpentinized samples (N10 vol.% as attested by bulk-rock Cu/S ratios ≥ 0.1), 3) more metal-rich Cu sulfides such as bornite, whereas pyrrhotite and pyrite are subordinate. To date, Pt-rich minerals were reported at Lherz (Lorand et al., 1999; Alard et al., 2000), at Baldissero (Garuti et al., 1984) and Pt–Te at Zabargad (Jedwab, 1992); we may surmise that PGM were probably overlooked, due to the lack of detailed SEM studies. The different sets of whole-rock PGE analyses

HPAS ID-ICP-MS3

2g

5.3. Conclusions: How does the PGM mineralogy of FON B 93 (and its effect on bulk-rock PGE analyses) apply to other orogenic peridotites

HPAS ID-ICP-MS4

1g

CT-ID-ICP-MS5

2g

Except Pt that exhibits a 95% deficit in the BMS phase, the PGE concentrations measured in FON B 93 by ID-ICP-MS can be balanced by BMS while Al-spinel is a negligible contributor, accounting for less than 0.5% of the Ru budget. FON B 93 could be used as a peridotite reference material for method validation (Meisel and Moser, 2004), provided homogeneity for Pt is tested at N 2 g test portion level. Pt–Te–Bi–(Pd) phases are crystallization products of Cu-rich derivative liquids at near-solidus temperatures in the continental lithosphere. Conversely, Pt–Ir–(Os) alloys are high-temperature alloys, potentially stable in the upper mantle in presence of metal-rich Mss and Cu–

UB-N HPAS ID-ICP-MS3

2g

Table 5 Average PGE concentrations for GP-13 and UB-N obtained by various analytical methods Method

P. Al. N

GP 13 CT-ID-ICP-MS1

Os

8 3.87 4.4 9

CT-ID-ICP-MS2

CT-ID-ICP-MS6 HP CT-ID-ICP-MS7 1 g HP CT-ID-ICP-MS8 1 g

7 4.06 0.6 4 3.61 2.75 5 3.62 3.2

Ir

Ru

3.56 6.97 9.2 3.4 3.38 6.92 4.2 11 3.33 6.25 2.8 6.6 3.27 6.84 4.64 1.79 3.32 6.88 4.6 4.4

Pt

Pd

7.00 7.5 6.43 5.5 6.69⁎ 10.0 6.45 7.51 6.80 7

5.64 6.3 5.42 6.2 5.68 4.66 5.88 2.90 5.54 6.5

14 3.71 3.38 6.30 7.42 7.1 6.4 4.6 4.0 7 3.64 3.52 7.234 7.48 11.1 10.0 10.0 4.3 6 3.66 3.24 6.48 8.07 4.0 9.6 4.5 15 4 3.85 3.37 6.92 7.00 3.4 6.2 4.5 3.9

6.11 2.9 6.1 4.0 6.17 4.0 5.85 2.8

Relative standard deviation (1 sigma/mean) are in italic. other ±2 s aver. Pearson et al. (2004) reported the Durham data set with a 2 R.S.D. confidence level.⁎: one value at 8.04. 1: Pearson et al. (2004), 2: DLC, 3: Meisel and Moser (2004), 4: Luguet et al. (unpublished data), 5: Puchtel et al. (2004), 6: Puchtel and Humayun (2005), 7: Luguet et al. (2007), 8: Becker et al. (2006).

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published for the other two matrix-matched orogenic lherzolite standards provide indirect evidence of Pt-rich discrete phases (Table 5). For GP-13, the relative standard deviation indicates more scattered Pt concentrations at subsample size ≤ 2 g with an outlier at 8.04 ppb (Pearson et al., 2004; Meisel and Moser, 2004). UB-N seems to be more homogeneous down to subsample scale of 2 g yet the Luguet et al. (2007) analyses at subsample size of 1 g gave high relative standard deviation (15%) for Pt. Meisel and Moser (2004) reported a greater scatter of the data (61% RSD) for Pt in JP1, a BMS-poor Horoman harzburgite likely reflecting segregation of Pt-rich alloys in such lithologies (Luguet et al., 2007), The potential occurrence of Pt–Ir–(Os) alloy in an upper mantle more reduced than FMQ is of prime concern for models that interpret fertile orogenic lherzolites as pervasively refertilized harzburgites. Such alloys likely play a role in maintaining broadly chondritic ratios between heavy PGE (Os, Ir and Pt) while decoupling them from the light PGE that are exclusively hosted in BMS. Acknowledgements We are grateful to M. Marot for thin sections, Omar Boudouma for his help with SEM analyses (UFR “Sciences de la Terre, University Pierre and Marie Curie), H. Rémy and M. Fialin (Camparis) for electron microprobe data, and Simone Pourtalès and Olivier Brugier who operated the ICP-MS facility in Montpellier. Michel Gros is gratefully acknowledged for his help during PGE analyses performed at the MNHN. Acknowledgements are extended to Mary Horan, Steven Shirey, Richard Carlson and Graham Pearson for their help and fruitful discussions on PGE analytical methods during A. Luguet postdoctoral stays. Financial support was provided by CNRS grants to UMR 7160 (JPL), and an EEC MarieCurie postdoctoral fellowship and a Carnegie Institution of Washington postdoctoral fellowship to AL. We acknowledge Harry Becker and Frank Melcher for their thoughtful reviews and Graham Pearson for editorial suggestions that greatly helped to prepare a more readable final version. References Alard, O., Griffin, W.L., Lorand, J.P., Jackson, S., O'Reilly, S.R., 2000. Non-chondritic distribution of highly siderophile elements in mantle sulfide. Nature 407, 891–894. Andrews, D.R., Brenan, J.M., 2002. The solubility of ruthenium in sulfide liquid: implications for platinum group mineral stability and sulfide melt–silicate melt partitioning. Chem. Geol. 192, 163–181. Ballhaus, C., 1995. Is the upper mantle metal-saturated? Earth Planet. Sci. Lett. 132, 75–86.

Ballhaus, C., Ulmer, P., 1995. Platinum-group elements in the Merensky reef: II. experimental solubilities of platinum and palladium in Fe1 −xS from 950 to 450 °C under controlled fS2 and fH2. Geochim. Cosmochim. Acta 59, 4881–4888. Ballhaus, C., Sylvester, P., 2000. Noble metal enrichment processes in the Merensky Reef, Bushveld Complex. J. Petrol. 41, 545–561. Ballhaus, C., Bockrath, C., Wohlgemuth-Ueberwasser, C., Laurenz, V., Berndt, J., 2006. Fractionation of the noble metals by physical processes. Contrib. Mineral. Petrol. 152, 667–684. Barnes, Sarah Jane, Naldrett, A.J., Gorton, M.P., 1985. The origin of the fractionation of Platinum-group elements in terrestrial magmas. Chem. Geol. 53, 303–323. Barnes, Sarah-Jane, van Achterbergh, E., Makovicky, E., Li, C., 2001. Proton microprobe results for the partitioning of platinum-group elements between monosulfide solid solution and sulphide liquid. S. Afr. J. Geol. 104, 275–286. Barnes, Sarah-Jane, Cox, R.A., Zientek, M.L., 2006. Platinum-group element, gold, silver and base metal distribution in compositionally zoned sulfide droplets from the Medvezky Creek Mine, Noril'sk, Russia. Contrib. Mineral. Petrol. 152, 187–200. Becker, H., Horan, M.F., Walker, R.J., Gao, S., Lorand, J.-P., Rudnick, R.L., 2006. Highly siderophile element compositions of the earth's primitive mantle. Geochim. Cosmochim. Acta 70, 4528–4550. Bézos, A., Lorand, J.-P., Hummler, E., Gros, M., 2005. Platinumgroup element systematics in Mid-Oceanic Ridge Basaltic glasses from the Pacific, Atlantic and Indian Oceans. Geochim. Cosmochim. Acta 69, 2613–2627. Bockrath, C., Ballhaus, C., Holzheid, A., 2004. Fractionation of the Platinum-group elements during mantle melting. Science 305, 1951–1953. Bodinier, J.-L., Dupuy, C., Dostal, J., 1988. Geochemistry of Eastern Pyrenean peridotites. Geochim. Cosmochim. Acta 52, 2893–2907. Brenan, J.M., Andrews, D., 2001. High-temperature stability of laurite and Ru–Os–Ir alloy and their role in PGE fractionation in mafic magmas. Can. Mineral. 39, 341–360. Büchl, A., Brügmann, G.E., Muenker, C., Hofmann, A.W., 2002. Melt percolation monitored by Os-isotopes and PGE abundances: a case study from the mantle section of the Troodos ophiolite. Earth Planet. Sci. Lett. 204, 385–402. Burton, K.W., Schiano, P., Birck, J.L., Allègre, C.J., 1999. Osmium isotope disequilibrium between mantle minerals in a spinellherzolite. Earth Planet. Sci. Lett. 172, 311–322. Cabri, L.J., Harris, D.C., Weiser, T.W., 1996. Mineralogy and distribution of platinum-group mineral (PGM) placer deposits of the world. Explor. Min. Geol. 5, 73–167. Chaussidon, M., Lorand, J.-P., 1990. Sulphur isotope composition of orogenic spinel lherzolite massifs from Ariège (Northeastern Pyrennes, France): an ion microprobe study. Geochim. Cosmochim. Acta 54, 2835–2846. Craig, J.R., Kullerud, G., 1969. Phase relations in the Cu–Fe–Ni–S system and their application to magmatic ore deposits. Econ. Geol. Monogr. 4, 344–358. Fleet, M.E., Stone, W.E., 1991. Partitioning of platinum group elements in the Fe–Ni–S system and their fractionation in nature. Geochim. Cosmochim. Acta 55, 245–253. Fabriès, J., Lorand, J.-P., Bodinier, J.-L., Dupuy, C., 1991. Evolution of the upper mantle beneath the Pyrenees: evidence from orogenic spinel lherzolite massifs. In: Menzies, M.A., Dupuy, C., Nicolas, A.J. (Eds.), J. Petrology, Special Lherzolites Issue, pp. 55–76.

J.-P. Lorand et al. / Chemical Geology 248 (2008) 174–194 Garuti, G., Gorgoni, C., Sighinolfi, G.P., 1984. Sulfide mineralogy and chalcophile and siderophile element abundances in the IvreaVerbano mantle peridotites (Western Italian Alps). Earth Planet Sci. Lett. 70, 69–87. Gervilla, F., Kojonen, K., 2002. The platinum-group minerals in the upper section of the Keivitsansarvi Ni–Cu–PGE deposit, northern Finland. Can. Mineral. 40, 377–394. Govindaraju, K., 1982. Report (1967–1981) on four ANRT rock reference samples: diorite DR-N, serpentinite UB-N, bauxite BXN and disthene DT-N. Geostand. Newsl. 6, 91–94. Gros, M., Lorand, J.-P., Luguet, A., 2002. Analysis of platinum group elements and gold in geological materials using NiS fire assay and Te coprecipitation: the NiS dissolution step revisited. Chem. Geol. 185, 179–190. Gros, M., Lorand, J.-P., Bezos, A., 2005. Determination of total sulfur contents in international rock standard SY-2 and other mafic and ultramafic rocks using an improved scheme of combustion/ iodometric titration. Geoanal. Geostand. Newsl. 29, 123–130. Handler, M., 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. Harney, D.M.W., Merkle, R.K.W., 1990. Pt–Pd minerals from the upper zone of the eastern Bushveld Complex, South Africa. Can. Mineral 28, 619–628. Harris, D.C., Cabri, L.J., 1991. Nomenclature of Platinum-group element alloys: review and revision. Can. Mineral. 29, 231–237. Helmy, H., Ballhaus, C., Berndt, J., Bockrath, C., WohlgemuthUeberwasser, C., 2007. Formation of Pt, Pd and Ni tellurides: experiments in sulfide–telluride systems. Contrib. Mineral. Petrol. 153, 493–524. Hoffmann, E., MacLean, W.H., 1976. Phase relations of michenerite and merenskyite in the Pd-Bi-Te system. Econ. Geol. 71, 1461–1468. Horan, M.F., Walker, R.J., Morgan, J.W., Grossman, J.N., Rubin, A.E., 2003. Highly siderophile elements in chondrites. Chem. Geol. 196, 5–20. Jedwab, J., 1992. Platinum group minerals in ultrabasic rocks and nickeliferous veins from Zabargad Island (Egypt). C.R. Acad. Sci. 314, 157–163. Kim, Won-Sa, Chao, G.Y., Cabri, L.J., 1990. Phase relations in the Pd–Te system. J. Less-Common Met. 162, 61–74. Lee, C.-T.A., 2002. Platinum-group element geochemistry of peridotite xenoliths from the Sierra Nevada and the Basin and Range, California. Geochim. Cosmochim. Acta 66, 3987–4005. Le Roux, V., Bodinier, J.-L., Tommasi, A., Alard, O., Dautria, J.-M., Vauchez, A., Riches, A.J.V., 2007. The Lherz spinel lherzolite: refertilized rather than pristine mantle. Earth Planet. Sci. Lett. 259, 599–612. Li, C., Barnes, Sarah-Janes, Makovicky, E., Rose-Hansen, J., Makovicky, M., 1996. Partitioning of nickel, copper, iridium, rhenium, platinium, and palladium between monosulfide solid solution and sulfide liquid: effects of composition and temperature. Geochim. Cosmochim. Acta 60, 1231–1238. Lorand, J.-P., 1985. The behaviour of the upper mantle sulfide component during the incipient serpentinization of "alpine-type" peridotites as exemplified by the Beni Bousera (Northern Morocco) and Ronda (Southern Spain) ultramafic bodies. Tschermak's Mineral. Petrogr. Mitt. 34, 183–209. Lorand, J.-P., 1989a. Abundance and distribution of Cu–Fe–Ni sulphides, sulphur, copper and platinum group elements in orogenic-type spinel peridotites of Ariege (Northeastern Pyrenees, France). Earth Planet Sci. Lett. 93, 50–64.

193

Lorand, J.-P., 1989b. Mineralogy and chemistry of Cu–Fe–Ni sulfides in mantle-derived spinel peridotite bodies from Ariège (Northeastern Pyrenees, France). Contrib. Mineral. Petrol. 103, 335–345. Lorand, J.-P., Alard, O., 2001. Geochemistry of platinum-group elements in the sub-continental lithospheric mantle; in-situ and whole-rock analyses of some spinel peridotite xenoliths, Massif Central, France. Geochim. Cosmochim. Acta 65, 2789–2806. Lorand, J.-P., Keays, R.R., Bodinier, J.-L., 1993. Copper- and noble metal enrichment across the asthenosphere–lithosphere mantle diapirs: the Lanzo lherzolite massif. J. Petrol. 34, 1111–1140. Lorand, J.P., Pattou, L., Gros, M., 1999. Fractionation of platinum group element in the upper mantle: a detailed study in Pyrenean orogenic lherzolites. J. Petrol. 40, 957–981. Lorand, J.-P., Delpech, G., Grégoire, M., Moine, B., O'Reilly, S.Y., Cottin, J.Y., 2004. Highly siderophile elements and the multi-stage metasomatic history of Kerguelen lithospheric mantle (south Indian ocean). Chem. Geol. 208, 195–215. Luguet, A., Alard, O., Lorand, J.-P., Pearson, N.J., Ryan, C.G., O'Reilly, S.Y., 2001. LAM-ICPMS unravel the highly siderophile element geochemistry of abyssal peridotites. Earth Planet Sci. Lett. 189, 285–294. Luguet, A., Lorand, J.-P., Seyler, M., 2003. A coupled study of sulfide petrology and highly siderophile element geochemistry in abyssal peridotites from the Kane Fracture Zone (MARK area, MidAtlantic ridge). Geochim. Cosmochim. Acta 67, 1553–1570. Luguet, A., Lorand, J.-P., Alard, O., Cottin, J.Y., 2004. A multitechnique study of platinum-group elements systematic in some Ligurian ophiolitic peridotites, Italy. Chem. Geol. 208, 175–194. Luguet, A., Shirey, S., Lorand, J.-P., Horan, M.F., Carlson, R.C., 2007. Residual platinum-group minerals from highly depleted harzburgites of the Lherz massif (France) and their role in HSE fractionation of the mantle. Geochim. Cosmochim. Acta 71, 3082–3097. Makovicky, E., 2002. Ternary and Quaternary phase systems with PGE. In: Cabri, L.J. (Ed.), The Geology, Geochemistry, Mineralogy and Mineral Beneficiation of Platinum-Group Elements. Canadian Institute of Mining, Metallurgy & Petroleum, Montreal, pp. 131–178. Makovicky, E., Karup-Môller, 2002. Phase relations in the metal-rich portions of the phase system Pt–Ir–Fe–S at 1000 °C and 1100 °C. Mineral. Mag. 64, 1047–1056. Makovicky, M., Makovicky, 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, pp. 415–423. Majzlan, J., Makovicky, M., Makovicky, E., Rose Hansen, J., 2002. The system Fe–Pt–S at 1100 °C. Can. Mineral. 40, 509–517. McDonough, W.F., Sun, S.S., 1995. The chemical composition of the Earth. Chem. Geol. 120, 223–253. Meisel, T., Moser, J., 2004. Reference materials for geochemical PGE analysis: new analytical dataa for Ru, Rh, Pd, Os, It, Pt and Re by isotope dilution ICP-MS in 11 geological reference materials. Chem. Geol. 208, 319–338. Meisel, T., Reisberg, L., Moser, J., Carignan, J., Melcher, F., Brügmann, G., 2003. Re–Os systematics of UB-N, a serpentinized peridotite reference material. Chem. Geol. 201, 161–179. Morgan, J.W., 1986. Ultramafic xenoliths; clue to the earth's late accretionary history. J. Geophys. Res. 91, 12375–12387. Mungall, J., Andrews, D.R., Cabri, L.J., Sylvester, P., Rubrett, M., 2005. Partitioning of Cu, Ni, Au, and platinum-group elements between monosulfide solid solution and sulfide under controlled oxygen and sulfur fugacities. Geochim. Cosmochim. Acta 69, 4349–4360.

194

J.-P. Lorand et al. / Chemical Geology 248 (2008) 174–194

Müntener, O., Peeke, T., Desmurs, M., Meier, M., Schaltegger, U., 2004. Refertilization of mantle peridotite in embryonic ocean basins; trace element and Nd isotopic evidence and implications for crust–mantle relationships. Earth Planet Sci. Lett. 221, 293–318. Ohnenstetter, M., 1992. Platinum-group element enrichment in the upper mantle peridotites of the Monte Maggiore ophiolitic massif (Corsica, France): mineralogical evidence for ore–fluid metasomatism. Mineral. Petrol. 46, 85–107. Pascal, 1963. Nouveau traité de chimie minérale, tome XVII, 2nd fascicule, Masson (editor), Paris. 896p. Pattou, L., Lorand, J.P., Gros, M., 1996. Non-chondritic platinum group element ratios in the Earth's mantle. Nature 379, 712–715. Pearson, D.G., Nowell, G.M., 2004. Re/Os and Lu/Hf isotope constraints on the origin and age of pyroxenites from the Beni Bousera Peridotite Massif: implications for mixed peridotite– pyroxenite mantle sources. J. Petrol. 45, 439–455. Pearson, D.G., Irvine, G.J., Ionov, D.A., Boyd, F.R., Dreibus, G.E., 2004. Re–Os isotope systematics and platinum group element fractionation during mantle melt extraction: a study of massif and xenolith peridotite suites. Chem. Geol. 208, 29–59. Peregoedova, A.V., 1998. The experimental study of the Pt–Pd partitioning between monosulfide solid solution and Cu–Nisulfide melt at 900–840 °C. 8th International Platinum Symposium Abstracts. Geol. Soc. South Africa and South African Inst. Min. Metall. Symposium Series, vol. S18, pp. 325–327. Peregoedova, A.V., Ohnenstetter, M., 2002. Collectors of Pt, Pd and Rh in a S-poor Fe–Ni–Cu sulfide system at 760 °C: experimental data and application to ore deposits. Can. Mineral. 40, 527–561. Peregoedova, A.V., Barnes, Sarah-Janes, Baker, D.R., 2004. The formation of Pt–Ir alloys and Cu–Pd-rich sulfide melts by partial

desulfurization of Fe–Ni–Cu sulfides: results of experiments and implications for natural systems. Chem. Geol. 208, 247–264. Puchtel, I., Humayun, M., 2005. Highly siderophile element geochemistry of 187Os-enriched 2.8 Kostomukscha komatiites. Geochim. Cosmochim. Acta 69, 1607–1618. Puchtel, I., Humayun, M., Campbell, A.J., Sproule, R.A., Lesher, C.M., 2004. Platinum-group element geochemistry of komatiites from the Alexo and Pyke Hills areas, Ontario, Canada. Geochim. Cosmochim. Acta 68, 1361–1383. Rehkämper, M., Halliday, A.N., Alt, J., Fitton, J.G., Zipfel, J., Takazawa, E., 1999. Non-chondritic platinum-group element ratios in oceanic mantle lithosphere: petrogenetic signature of melt percolation? Earth Planet Sci. Lett. 172, 65–81. Righter, K., Campbell, A.J., Humayun, M., Hervig, R.L., 2004. Partitioning of Ru, Rh, Pd, Re, Ir, and Au between Cr-bearing spinel, olivine, pyroxene and silicate melts. Geochim. Cosmochim. Acta 68, 867–880. Terashima, S., Imai, N., 2000. Determination of selenium in fifty two geochemical reference materials by hydride generation atomic absorption spectrometry. Geostand. Newsl.: J. Geostand. Geoanal. 24, 83–86. Van Achterbergh, E., Ryan, C.G., Griffin, W.L., 1999. Glitter: on line intensity reduction for the laser ablation inductively coupled plasma mass spectrometry. 9th Goldschmidt Conf., p. 305. Boston M.A. J. Conf. Abstr., Cambridge. Yi, A., Halliday, A., Alt, J.C., Lee, D.C., Rehkämper, M., 2000. Cadmium, indium, tin, tellurium, and sulfur in oceanic basalts; implications for chalcophile element fractionation in the Earth. J. Geophys. Res. 105 (B8), 18,927–18,948.