Late accretion history of the terrestrial planets inferred from platinum stable isotopes

However, if the core formed under these conditions, the Pt depletion in the pre-late-veneer mantle would be significantly reduced (e.g., Pt ~ 0.1 ng g^-1 in pre-late-veneer mantle as compared with ca. 7 ng g^-1 in the primitive upper mantle; Becker et al., 2006), and a combination of high-pressure and -temperature core formation with a chondritic late-veneer can explain both the Pt concentrations and Pt stable isotope data.
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Figure 3 [...] The vertical line represents the Pt concentration of primitive upper mantle (PUM; 7.6 ppb; Becker et al., 2006), and the dashed horizontal line marks the intercept of the mantle concentration with PUM, indicating the amount of late-veneer required to reproduce the Pt abundance of PUM.
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Birck, J.L. (2004) An Overview of Isotopic Anomalies in Extraterrestrial Materials and Their Nucleosynthetic Heritage. Reviews in Mineralogy and Geochemistry 55, 25–64.

In contrast to enstatite and ordinary chondrites, carbonaceous chondrites contain significant amounts of refractory inclusions, which are known to preserve nucleosynthetic anomalies for a number of elements (Birck, 2004).
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Borg, L.E., Connelly, J.N., Boyet, M., Carlson, R.W. (2011) Chronological evidence that the Moon is either young or did not have a global magma ocean. Nature 477, 70–72.

Accepting an age of ~4.4 Ga for the timing of the Moon-forming impact (Borg et al., 2011), our data suggest that the delivery, admixing, and homogenisation of chondritic, perhaps volatile-rich, material to the early Earth occurred very shortly after magma ocean crystallisation.
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Nucleosynthetic variability has been documented in relatively low mass siderophile elements, such as Mo and Ru, with significant nucleosynthetic isotope heterogeneity existing between the various classes of chondrites (e.g., Burkhardt et al., 2011; Fischer-Gödde et al., 2015).
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We developed the techniques for the investigation of natural stable isotope fractionation of the HSE platinum (Creech et al., 2013, 2014), which is a novel tool to investigate the late accretion history of the terrestrial planets.
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Creech, J.B., Baker, J.A., Handler, M.R., Bizzarro, M. (2014) Platinum stable isotope analysis of geological standard reference materials by double-spike MC-ICPMS. Chemical Geology 363, 293–300.

This requires efficient mixing of the Hadean mantle, which is most easily understood in the framework of modern-style, mobile-lid tectonics rather than a stagnant-lid regime (Debaille et al., 2013).
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Fischer-Gödde, M., Burkhardt, C., Kruijer, T.S., Kleine, T. (2015) Ru isotope heterogeneity in the solar protoplanetary disk. Geochimica et Cosmochimica Acta 168, 151–171.

Gomes, R., Levison, H.F., Tsiganis, K., Morbidelli, A. (2005) Origin of the cataclysmic Late Heavy Bombardment period of the terrestrial planets. Nature 435, 466–469.

Kimberlite-hosted peridotite xenoliths from this area are considered to represent sub-cratonic lithospheric mantle up to 3.6 Ga in age (Griffin et al., 2004), and all five samples have Late Archean Os model ages (Table S-1).
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Stable isotopic fractionations relating to metal–silicate differentiation have been reported in several stable isotope systems, e.g., Si (Young et al., 2015), Mo (Hin et al., 2013; Burkhardt et al., 2014), Zn (Mahan et al., 2017).
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Theory predicts that the heavy isotopes of an element will tend to be concentrated in the most oxidised component of a system (Schauble, 2004), which is supported by experimental data for Mo (Hin et al., 2013), and therefore heavy Pt stable isotopes might be expected in oxidised silicate mantles of differentiated bodies.
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We interpret this to reflect Pt stable isotopic fractionation during core formation, whereby the heavy isotopes of Pt are preferentially retained in the more oxidised silicate part of the body while the light isotopes are concentrated in the metallic core, which is consistent with qualitative predictions based on stable isotope theory (Schauble, 2004) and Mo experimental data (Hin et al., 2013) as described above.
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Kruijer, T.S., Kleine, T., Fischer-Gödde, M., Sprung, P. (2015) Lunar tungsten isotopic evidence for the late veneer. Nature 520, 534–537.

Based on this, the Pt abundance in Isua samples can be interpreted as reflecting admixing of up to ~50 % of the full complement of late-veneer (Fig. 3), which is consistent with recent ¹⁸²W data indicating that the lunar mantle has a marginally greater ¹⁸²W enrichment relative to the early Archean terrestrial mantle (Kruijer et al., 2015; Touboul et al., 2015).
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Lorand, J.-P., Luguet, A., Alard, O. (2008) Platinum-Group Elements: A New Set of Key Tracers for the Earth’s Interior. Elements 4, 247–252.

Addition of this material can explain the elevated mantle HSE abundances and their broadly chondritic relative proportions (Lorand et al., 2008; Walker, 2009), and provide a mechanism for the delivery of volatiles to Earth (Owen and Bar-Nun, 1995).
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Variable HSE concentrations in kimberlite-hosted xenoliths have been interpreted to represent the sluggish equilibration of Archean mantle with the putative late-veneer (Maier et al., 2012), in keeping with the Pt stable isotopic variability in these samples.
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Alternatively, due to changes in partitioning behaviour of HSE under different physical conditions, the elevated HSE abundances in Earth’s mantle may be the result of core formation at high-temperatures and -pressures (Mann et al., 2012).
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Experimentally determined HSE partition coefficients at high-pressures and -temperatures cannot explain mantle abundances by equilibrium core formation alone (Mann et al., 2012).
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We also find that metabasalts and ultramafic schists from the >3.85 Ga (Nutman et al., 1997) Isua Supracrustal Belt also have heavy δ¹⁹⁸Pt (Fig. 2).
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O’Neill, C., Jellinek, A.M., Lenardic, A. (2007) Conditions for the onset of plate tectonics on terrestrial planets and moons. Earth and Planetary Science Letters 261, 20–32.

More speculatively, the early delivery of volatile-rich material through impact may have promoted rapid formation of the terrestrial hydrosphere and, hence, assisted hydration of the crust that is required for the inception of plate tectonic processes (O’Neill et al., 2007).
View in article

Owen, T., Bar-Nun, A. (1995) Comets, impacts, and atmospheres. Icarus 116, 215–226.

Rankenburg, K., Humayun, M., Brandon, A.D., Herrin, J.S. (2008) Highly siderophile elements in ureilites. Geochimica et Cosmochimica Acta 72, 4642–4659.

Primitive achondrites sample bodies that have undergone varying degrees of metal-silicate segregation, as reflected by HSE abundances that span a range from ~50-1500 ng g^-1 (Warren et al., 2006; Rankenburg et al., 2008) which has been interpreted to reflect the early stages of core formation (Warren et al., 2006).
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Moreover, these samples also preserve ¹⁴²Nd excesses of up to ~10 ppm from the decay of short-lived ¹⁴⁶Sm (t_1/2 ~ 68–103 Ma), indicative of the early differentiation of their source reservoir (Rizo et al., 2013).
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These observations require that progressive admixing of veneering material to Earth’s mantle was initiated prior to the formation of the Isua mantle source at ≳4.3 Ga based on the ¹⁴⁶Sm-¹⁴²Nd decay system (Rizo et al., 2013).
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However, the timing and scale of veneering of the early Earth are poorly constrained and the origin of these Archean enrichments is controversial (e.g., Rizo et al., 2016).
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Schauble, E.A. (2004) Applying Stable Isotope Fractionation Theory to New Systems. Reviews in Mineralogy and Geochemistry 55, 65–111.

Theory predicts that the heavy isotopes of an element will tend to be concentrated in the most oxidised component of a system (Schauble, 2004), which is supported by experimental data for Mo (Hin et al., 2013), and therefore heavy Pt stable isotopes might be expected in oxidised silicate mantles of differentiated bodies.
View in article
We interpret this to reflect Pt stable isotopic fractionation during core formation, whereby the heavy isotopes of Pt are preferentially retained in the more oxidised silicate part of the body while the light isotopes are concentrated in the metallic core, which is consistent with qualitative predictions based on stable isotope theory (Schauble, 2004) and Mo experimental data (Hin et al., 2013) as described above.
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The late-veneer hypothesis is apparently supported by the existence of small enrichments in ¹⁸²W in the early Archean terrestrial rock record (Willbold et al., 2011, 2015; Touboul et al., 2014).
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This is consistent with ¹⁸²W data from Isua and other ancient rocks (Willbold et al., 2011, 2015; Touboul et al., 2014), which have also been interpreted to reflect long-term preservation of pre-late-veneer mantle that had escaped complete mixing with late-veneer.
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Touboul, M., Puchtel, I.S., Walker, R.J. (2015) Tungsten isotopic evidence for disproportional late accretion to the Earth and Moon. Nature 520, doi: 10.1038/nature14355.

Walker, R.J. (2009) Highly siderophile elements in the Earth, Moon and Mars: Update and implications for planetary accretion and differentiation. Chemie der Erde - Geochemistry 69, 101–125.

The late-veneer refers to the sum of material added to Earth’s mantle after the final episode of core formation, which is thought to have comprised a contribution of ~0.5 wt. % of Earth’s mass (~2 ×10²² kg) from chondritic material (Walker, 2009).
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Addition of this material can explain the elevated mantle HSE abundances and their broadly chondritic relative proportions (Lorand et al., 2008; Walker, 2009), and provide a mechanism for the delivery of volatiles to Earth (Owen and Bar-Nun, 1995).
View in article

Willbold, M., Elliott, T., Moorbath, S. (2011) The tungsten isotopic composition of the Earth’s mantle before the terminal bombardment. Nature 477, 195–198.

Willbold, M., Mojzsis, S.J., Chen, H.-W., Elliott, T. (2015) Tungsten isotope composition of the Acasta Gneiss Complex. Earth and Planetary Science Letters 419, 168–177.

Yokoyama, T., Walker, R.J. (2016) Nucleosynthetic Isotope Variations of Siderophile and Chalcophile Elements in the Solar System. Reviews in Mineralogy and Geochemistry 81, 107–160.

However, so far no isotopic anomalies have been documented in heavy siderophile elements with masses more comparable to Pt (e.g., Os, Te, Cd; Yokoyama and Walker, 2016).
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Supplementary Information

Abstract | Letter | Acknowledgements | References | Supplementary Information

Materials and Methods

The techniques utilised in this study for Pt are described in detail in earlier work (Creech et al., 2013, 2014; Creech and Paul, 2015). All sample preparation, digestions and chemical separations were carried out in ultra-clean laboratory conditions at either Victoria University of Wellington, New Zealand, the Centre for Star and Planet Formation, Natural History Museum of Denmark, or at the Institut de Physique du Globe de Paris, France.

Sample information

1. Terrestrial samples. Terrestrial sample information is summarised in Table S-1. Sample masses for peridotite samples were typically 15 g. However, depending on the amount of sample available, some samples were as small as ca. 7 g or as much as ca. 45 g (processed as multiple replicates). Four peridotite mantle xenolith samples from northern Queensland, Australia were analysed in this study. BC1 and BC7 are from Batchelor’s Crater, which is a Plio–Pleistocene alkali basalt vent in the Chudleigh volcanic province. MQ1 and MQ14 are from Mount Quincan in the <3 Ma Atherton Tablelands volcanic province. Major and trace element and Sm–Nd isotopic data for a suite of peridotite xenoliths from this locality, including both of these samples, were presented by Handler et al. (2005). The Mount Quincan xenolith suite represent a fertile mantle source with a ¹⁴⁷Sm–¹⁴³Nd age of ca. 275 Ma (Handler et al., 2005).

Two ophiolite/transitional lherzolite samples emplaced during the Mesozoic rifting of the Italian Alps were measured in this study. LI 138 is from the Internal Liguride Ophiolite (Luguet et al., 2004). L 213 from the Sezia Lanzo HP zone is a spinel/plagioclase lherzolite from the northern part of the Lanzo lherzolite massif that represents fragment of a continental lithosphere that accreted during the early Palaeozoic/late Proterozoic (Bodinier et al., 1991). Mtf37 is an off-craton, basalt-hosted highly fertile lherzolite xenolith from Montferrier, southern Massif Central, France (Alard et al., 2011). TUR 7 is a highly fertile continental orogenic lherzolite from Turon de Técouère, Pyrenees, France (Fabriès et al., 1998). All four samples are LREE-depleted lherzolites showing little evidence of secondary cryptic/modal metasomatism and variable serpentinisation (close to 0 % in TUR 7 and Mtf37, 20 % for LI 138 and close to 100 % for LI 138).

Five cratonic xenoliths from southern Africa were analysed in this study. All five samples have late Archean Os model ages (Table S-1). LMA2 and LTP20 are xenoliths from the Matsoku and Thaba Putsoa kimberlites, respectively, in Lesotho. KBD4, BG and BG2 xenoliths come from diamondiferous kimberlite pipes of the Kimberley area, South Africa. KBD4 and BG2 are described as harzburgitic, whereas BG is lherzolitic (Lorand and Grégoire, 2006). The cratonic xenoliths are characterised by lower Pt concentrations (0.4–2.7 ng g^-1; Table S-1) than other mantle samples analysed in this study.

Eight rocks from the Isua Supracrustal Belt were analysed. A lithologically diverse package of early Archean supracrustal rocks occurs in a ~5 km wide arcuate belt at Isua, Southwest Greenland. In the southwest part of the belt, pillow lava structures are recognised within low deformation zone windows within a sequence of amphibolite rocks. Compositionally, the pillows are high-MgO (12–15 wt. %) tholeiitic basalts and are associated with ultramafic schists of komatiitic composition. A felsic unit which cross-cuts this pillow lava-bearing sequence gives U-Pb SHRIMP ages clustering around ~3.85 Ga (Baadsgaard et al., 1984; Michard-Vitrac et al., 1977; Nutman et al., 1997) suggesting that these amphibolites are amongst the oldest of the Isua supracrustals. Ultramafic schists occur as scattered outcrops within the general vicinity of the pillow basalts and have compositions typical of Archean komatiites (ca. 25–27 wt. % MgO), and therefore most likely represent very high degree melts. We analysed six samples from the metabasalts, as well as two samples from the ultramafic schists (Table S-1). High-precision Nd isotope data have been previously reported for five of the Isua samples analysed here (Rizo et al., 2013), and major and trace element data for these samples are also provided in that work.

2. Meteorite samples. Meteorite sample information is summarised in Table S-2. A total of 40 meteorite samples were analysed in this study. The 14 chondrites samples comprise four ordinary chondrites, two enstatite chondrites and eight carbonaceous chondrites. The 15 primitive achondrites analysed comprise two acapulcoites, one winonaite and 12 ureilites, of which 11 are Antarctic meteorites provided by NASA-JSC. The six iron meteorites analysed comprise two IAB, two IIIAB, one IVA and one IVB.

For most meteorite samples, typical sample sizes were ca. 0.5–1.0 g (Table S-2). For one of the carbonaceous chondrites (Allende), a large fragment (~15 g) was crushed and homogenised with an agate mortar and pestle, from which five replicate digestions of ca. 1 g were each processed to assess the reproducibility of the technique (Table S-2). Previously published cosmic ray exposure (CRE) ages for the iron meteorites are also given in Table S-2, although no correlation between Pt isotopes and CRE ages was observed. Further details about the meteorites analysed can be obtained from the Meteoritical Bulletin Database (http://www.lpi.usra.edu/meteor/).

Sample preparation and digestion (Pt)

Digestions of chondrite, primitive achondrite and terrestrial samples were carried out using a NiS fire assay method as described by Creech et al. (2014). All sample powdering was carried out using agate mortar and pestle or agate mills. Powders were weighed directly into porcelain crucibles for NiS digestion. Prior to powdering, fragments of chondrite and achondrite meteorites were taken from sliced specimens, weighed and then cleaned by rinsing in ultra-pure water, methanol and again in ultra-pure water. The meteorite fragments, typically weighing 0.5–1.0 g, were then placed in a Savillex teflon beaker in ca. 4 mL of 2 M HCl. The beakers were heated to 120 ˚C for ca. 10 min and then ultrasonicated for ca. 3 min. After this gentle acid washing, the samples were rinsed twice with water and then dried in the beaker on a hotplate before crushing and transferring to porcelain crucibles for NiS fire assay digestion. Samples were double-spiked prior to NiS digestion by adding the ¹⁹⁶Pt–¹⁹⁸Pt double-spike (DS) directly to the mixture in the crucible, and allowed to dry before being fused in a furnace. Samples were optimally spiked in approximate proportions of 55 % Pt from DS and 45 % natural Pt. Where sample Pt concentrations were unknown they were estimated based on published results for similar samples. The preparation and fusion of NiS charges and dissolution of NiS beads were performed as described by Creech et al. (2014). Once fully digested, the samples were converted to chloride form by evaporating with concentrated HCl, and finally brought into solution in 10 mL of 0.5 M HCl for anion-exchange chemistry.

Digestion of iron meteorite samples was carried out directly by acid digestion. Small fragments of Fe-Ni metal weighing ca. 0.5–1.0 g were cut from larger iron meteorite specimens using a diamond cutting disc attached to a small handheld drill. Surface contamination and oxidation were removed by abrasion with fine-grained SiC sandpaper, and fragments were cleaned with ultra-pure water and methanol. The metal was then acid washed by adding ca. 4 mL of 1.5 M HCl at room temperature and placing it in an ultrasonic bath for a few minutes, and then by placing it in ca. 4 mL of 0.8 M HNO₃ on a hotplate at 120 ˚C for up to 30 min, before rinsing with ultra-pure water and drying. The Fe–Ni metal was digested in concentrated HCl or aqua regia on a hotplate at 120 ˚C. Iron meteorite samples were double-spiked prior to dissolution of the iron meteorite fragments or in an aliquot of the digested sample. Double-spike was added to have proportions of approximately 55 % of Pt from DS and 45 % natural Pt. Once fully digested and spiked, the samples were evaporated to dryness, converted to chloride form by evaporating with concentrated HCl, and finally brought into solution in 10 mL of 0.5 M HCl for anion-exchange chemistry. The sample and DS were fully equilibrated after this procedure.

Chemical separation of platinum

Purification of Pt from samples was performed using the methods described by Creech et al. (2014), based on separation utilising AG1-X8 (100–200 mesh) anion exchange resin. The major cationic species in rock and meteorite samples are not retained on this resin, and the platinum group elements were eluted using sequentially stronger HCl and HNO₃ acids. Pt is collected in a final step using concentrated HNO₃, with a Pt yield of ≥90 % and purity of ≥95 % (Creech et al., 2014). This level of Pt separation has been shown to be adequate for accurate determination of Pt stable isotope compositions by double-spike MC-ICPMS (Creech et al., 2013). All Pt separation work was conducted in class 10 laminar flow hoods situated in a class 100 clean laboratory at either Victoria University of Wellington, New Zealand, the Centre for Star and Planet Formation, Denmark, or the Institut de Physique du Globe de Paris, France. In order to destroy any potential organic interferences, Pt cuts were treated by refluxing in aqua regia overnight on a hotplate, evaporated to dryness, and then repeatedly (n = 4 or 5) refluxed and then evaporated in HNO₃–H₂O₂ on a hotplate at 150 ˚C until no residue from the resin could be observed. Samples were then reconverted to chloride form by evaporating concentrated HCl, prior to being dissolved in 0.5 M HCl for isotopic measurement by MC-ICPMS.

Mass spectrometry (Pt)

Pt stable isotope measurements were carried out using the techniques described by Creech et al. (2013), using a ¹⁹⁶Pt–¹⁹⁸Pt DS to correct for instrumental mass fractionation, using ¹⁹⁴Pt, ¹⁹⁵Pt, ¹⁹⁶Pt and ¹⁹⁸Pt in the DS inversion. Isotopic measurements were made using a Nu Plasma MC-ICPMS at Victoria University of Wellington, New Zealand, or a Thermo Neptune Plus MC-ICPMS at the Centre for Star and Planet Formation, Denmark. Analytical conditions for both instruments were identical to those described by Creech et al. (2013). Isotope results are reported in delta notation relative to the IRMM-010 Pt isotope standard as the per mil difference in ¹⁹⁸Pt/¹⁹⁴Pt ratios (δ¹⁹⁸Pt). Sample–spike mixtures were monitored in analyses and data points were rejected if sample–spike ratios were found to be outside the ideal range. All data reduction was conducted offline using the double-spike data reduction tool IsoSpike (Creech and Paul, 2015), which is an add-on for the Iolite data reduction package (www.iolite.org.au). IsoSpike is open-source and freely available from www.isospike.org. Pt concentrations of samples were also calculated offline using isotope dilution calculations as described by Creech et al. (2014).

Reproducibility of Pt isotope measurements

Individual analyses have typical internal errors (2 se) on δ¹⁹⁸Pt of ca. 0.020 ‰ or 0.015 ‰ for measurements made in NZ and DK, respectively. In order to monitor the long-term reproducibility of Pt stable isotope measurements, a solution standard (J. T. Baker Pt ICP solution standard) with a published Pt stable isotope composition (δ¹⁹⁸Pt = +0.09 ‰; Creech et al., 2013) was employed as a secondary standard, with double-spiked analyses of the standard made in each measurement session. Repeat analyses of the Baker Pt standard have been made in three different labs, on four different MC-ICPMS instruments and over ~5 years, and yield a reproducibility of ≤0.04 ‰ (2 sd) on δ¹⁹⁸Pt. Specifically during the analytical sessions represented in the data presented here, analyses of the Baker Pt standard yielded δ¹⁹⁸Pt_Baker = 0.09 ± 0.04 ‰ (2 sd; n = 91). Replicate samples from a homogenised powder of the carbonaceous chondrite Allende gave consistent results with a mean δ¹⁹⁸Pt = –0.12 ± 0.06 ‰ (2 sd; Table S-2). These five replicates provide the best estimate of the external reproducibility of our method for NiS processed samples, and thus error bars on all NiS processed samples in figures presented here are either the uncertainty in the weighted mean of the sample or the 2 sd of the Allende replicates (0.06 ‰), whichever is larger. The reproducibility at lower (ppb) Pt concentrations was estimated by pooling data for the Isua suite of samples. While these are separate samples, they are formed from large degrees of partial melting from the same mantle source and therefore are expected to have a common Pt isotopic composition. Treating these eight samples as a single group (excluding one replicate of one sample (460258) as an outlier), these indicate a reproducibility of 0.07 ‰ on δ¹⁹⁸Pt. The homogeneous Pt composition defined by these samples indicates that stable Pt isotopes are robust to secondary events, as the Sm–Nd systematics in these samples (Rizo et al., 2013) suggest that they have been variably affected by post-crystallisation metamorphism.

Reproducibility for iron meteorites processed by acid digestion is estimated based on three replicate digestions of Toluca (processed both in NZ and DK), which give an external reproducibility on δ¹⁹⁸Pt of ±0.05 ‰. As above, error bars in figures represent either the uncertainty in the mean of the sample or the 2 sd of the Toluca replicates (0.05 ‰), whichever is larger.

Platinum blanks

The contribution of Pt blanks was found to be consistent with those reported by Creech et al. (2014). Anion-exchange column Pt blanks were typically ca. ≤0.1 ng. The sizes of the Pt blanks introduced during sample processing were assessed by purifying Pt from multiple blank NiS digestions, and were typically ca. 0.5–2 ng, with an average of 0.8 ng per typical 15 g sample digestion (n = 16; Table S-3). For most samples, the effect of this amount of blank is negligible, but for exceptionally small samples or samples with extremely low Pt content (e.g., <1 ng g^-1), the contribution from blank must be considered. Therefore, efforts were made to determine the effect of the blank on Pt isotope data and to apply a correction.

The approximate Pt isotopic composition of the blank was determined in two ways: (1) adding Pt double-spike to blank ‘samples’ prior to NiS digestions, and (2) adding blank to the IRMM-010 Pt isotope standard. In the former method, a number of blank NiS digestions were made and Pt double-spike was added in the proportions suggested by the average Pt blank given above. NiS beads from up to four such digestions were combined and processed in a single batch through Pt chemistry, and Pt isotopes were subsequently measured by MC-ICPMS. Given the extremely small amounts of Pt processed from these blanks, uncertainties on isotopic measurements are large (Table S-4). However, these data suggest that the composition of the Pt blank, or at least the effect of blank on the DS corrected ratios, is up to δ¹⁹⁸Pt ≈ +1.58 ‰ (Table S-4).

In the latter method, Pt blank was accumulated by combining Pt cuts from several blank NiS digestions. The blank Pt was then added to IRMM-010 Pt isotope standard in varying proportions of up to 30 % blank, and measured by MC-ICPMS (Table S-5). A regression through these data suggests a blank composition or effect of around δ¹⁹⁸Pt ≈ +1.14 ‰.

Note that both of these methods assume that, if the blank is truly fractionated, the isotopic composition of the blank is mass-dependently fractionated from a terrestrial composition.

Assuming the isotopic composition and amount of blank based on the above, a correction can be applied to the data to remove the effect of the blank. This is demonstrated using the NiS-blank-doped data described above (Table S-5; Fig. S-1).

The efficacy of this correction was also tested by taking a peridotite sample (BC-1) for which isotope data is available (Table S-1), and processing multiple replicate NiS digestions with smaller amounts (~1 g per digestion) of sample, such that the blank-to-sample ratio is significantly increased (i.e. up to 20 % Pt from blank). NiS beads from nine such digestions were combined into three batches for Pt chemistry and isotopic analysis (Table S-6). We found that by using the above determined blank composition (δ¹⁹⁸Ptblank = +1.58 ‰) and varying the amount of blank per digestion within the typical range of observed blank amounts (Table S-3), all of the BC-1 batch samples can be corrected to values close to those previously determined for that sample (δ¹⁹⁸Pt = –0.19 ± 0.15 ‰; Table S-6; Fig. S-1).

**Figure S-1** Demonstration of correction for blank Pt on δ¹⁹⁸Pt data for **(a)** mixtures of blank Pt from NiS digested blanks and IRMM-010 Pt isotope standard (δ¹⁹⁸Pt = 0 ‰) and **(b)** the peridotite BC1 (δ¹⁹⁸Pt = -0.09 ± 0.15 ‰), where three samples have been prepared by combining digestions of relatively small amounts of material, thus creating a high blank-to-sample ratio. In each case, uncorrected data are shown in blue and the same data with a blank correction applied are shown in green.

Based on these results, blank corrections were calculated for the entire dataset in Tables S-1 and S-2 (with the exception of the iron meteorites, which were not digested by NiS fire assay), using the average amount of blank (0.8 ng) per typical 15 g sample digestion and the most extreme value for the blank composition (δ¹⁹⁸Pt_blank = +1.58 ‰), as determined above. The size of these additional uncertainties are given in Tables S-1 and S-2 as the ‘Magnitude of blank correction (‰)’. Where the proportion of Pt from blank is ≲4 %, i.e. for almost all samples, the correction is well within the uncertainty of the analysis. This is readily apparent in Figure 2, where extended error bars relating to the uncertainty arising from blank are very small for all but a few samples. Furthermore, in many cases, the magnitude of this correction may be overestimated due to our conservative choices of the composition and amount of blank, as described above.

Note that for samples where less than 15 g of sample were processed, the amount of sample was made up to 15 g with pure silica, and thus the amount of blank per digestion is not affected.

Model details and parameters

Modelling of terrestrial Pt isotopes as presented in Figure 3 was carried out using simple end-member mixing calculations between a hypothetical mantle composition after core formation (pre-late-veneer mantle) and a chondritic late-veneer. The Pt concentration and isotopic composition of the late-veneer are based on chondrite data presented in this work. The Pt concentration of the pre-late-veneer mantle is estimated based on published metal–silicate partition coefficients (i.e. D_Pt^met∕sil), which range from ≥10⁶ at low-pressure and -temperature to ~10⁴ at high-pressure and -temperature (Mann et al., 2012). These parameters are summarised in Table S-7.

Using these parameters, we then calculated the composition of various mixtures between these two end-members with different iterations using different hypothetical δ¹⁹⁸Pt for the pre-late-veneer mantle (Fig. 3). The only constraint on the δ¹⁹⁸Pt of pre-late-veneer mantle is to try and find some value where the mixture between pre-late-veneer mantle and late-veneer goes through the data for Isua samples as well as the post-Archean mantle composition, which assumes these samples are related by different degrees of mixing between those two end-members. This modelling shows that using the most Pt depleted pre-late-veneer mantle corresponding to core formation at low-pressure and -temperature, the Pt isotopic signature of core formation is immediately overwhelmed with a chondritic signature after the addition of minimal amounts of late-veneer, and thus the compositions of the Isua samples cannot be explained by this relationship. Conversely, assuming high-pressure and temperature core formation, the Pt concentrations and isotopic compositions in Isua samples can be explained if the pre-late-veneer mantle was fractionated to δ¹⁹⁸Pt ≳ 4 ‰. Further experimental or theoretical work is required to determine if Pt isotope fractionation of this magnitude is likely under these conditions. We note that the Kaapvaal samples were not used to constrain the model due to uncertainty in the origin of their Pt depletions (i.e. PGE data suggest that these samples have undergone melt extraction, which may have variably depleted Pt to some extent), although the data are also consistent with this model. This could also imply that the Isua and Kaapvaal mantle sources had undergone different degrees of equilibration with late-veneer, which could relate to the relative ages of those mantle sources or spatial heterogeneity in the mixing of late-veneer into the mantle. Subsequent admixing of the remaining late-veneer inventory raised mantle Pt abundances to that observed for the PUM and shifted the post-Archean mantle Pt isotope composition to the chondritic value.

HSE and ¹⁸⁷Os analytical methods

Digestions were carried out on 1 g of whole rock powder spiked with a mixed HSE spike (¹⁸⁵Re, ¹⁰⁶Pd, ¹⁹⁵Pt, ¹⁹¹Ir, ¹⁹⁰Os) and with 7.5 mL inverse aqua regia (2:1 HNO₃:HCl) in an Anton Paar HP-Asher at 220 ˚C and >100 bars for 13.5 hours. After digestion, osmium was extracted from the aqua regia solution into CCl₄ and back-extracted into HBr (Cohen and Waters, 1996) before a final purification step consisting in the micro-distillation of the Os-H₂SO₄-Cr⁶⁺ solution into HBr (Pearson et al., 1998). After dry down, the remaining inverse aqua regia fraction was further digested at 100–120 ˚C using a HF/HNO₃ mixture. After equilibration in 12 N HCl, the HSE were pre-concentrated following an anion exchange chromatography procedure adapted from Pearson and Woodland (2000). Iridium, Pt, Pd and Re concentrations were analysed using the Thermo Finnigan Element 2 ICP-MS at the Steinmann Institute of the University of Bonn. Osmium concentrations and isotopic compositions (i.e. ¹⁸⁷Os/¹⁸⁸Os ratios) were measured using a Thermo-Finnigan Triton N-TIMS at the Department of Earth Sciences, University of Durham (UK), as described by Dale et al. (2008). Aliquots of the DROsS standard solution were run 4 times and are in good agreement with the values obtained by Dale et al. (2008) on 10–100 ng aliquots run in the Faraday mode on the same thermal ionisation mass spectrometer. Total procedural blanks are 0.72 pg Os, 0.54 pg Ir, 10.8 pg Pt, 44.1 pg Pd and 7.2 pg Re.

Supplementary Text

Platinum isotope variability amongst chondrites

Greater variability was observed amongst the eight carbonaceous chondrites analysed than in the other chondrite groups, which is thought to reflect real compositional differences that are more apparent in that group due to the greater proportions of matrix and refractory inclusions (amoeboid olivine aggregates and Ca-Al-rich inclusions). Pt usually resides inside CAI inclusions as refractory metal nuggets of nebular condensation origin as well as low-temperature Pt-Fe alloys and Fe-Ni-sulphides related to asteroidal processes (c.f., Harries et al., 2012; Hewins et al., 2014), and these may contain Pt stable isotope or nucleosynthetic variability that is expressed in samples on this smaller scale. However, the five replicates processed from a single large homogenised sample of Allende are identical to the mean of carbonaceous chondrites as well as that of all chondrite groups, suggesting that the chondrite reservoir that had uniform Pt stable isotope composition. The mean of all of the analysed chondrites gives the chondritic composition of δ¹⁹⁸Pt_chondrite = –0.15 ± 0.09 ‰ (2 sd). However, if the additional variability introduced from the smaller samples of carbonaceous chondrites is excluded, the chondritic composition can be taken as δ¹⁹⁸Pt_chondrite = –0.15 ± 0.03 ‰ (2 sd).

Platinum isotope variability amongst iron meteorites

Despite the relatively high concentration of Pt (typically ~5–15 µg g^-1) in iron meteorites, some iron meteorite samples were found to have poorer reproducibility than chondrites (Fig. S-2; Table S-2). Given the relatively large exposure ages of iron meteorites as compared to chondrites and achondrites, an obvious potential explanation for this variability between iron meteorite samples is mass independent effects arising from thermal neutron capture. The iron meteorites analysed in this study have a range of cosmic ray exposure (CRE) ages (Table S-2). Therefore, it is important to consider mass independent isotopic effects from thermal neutron capture, particularly given that the DS-corrected data in this study did not include unspiked measurements. Two recent studies have suggested using Pt isotope ratios as a dosimeter for correcting cosmogenic effects in Hf–W isotopic data (Kruijer et al., 2013; Wittig et al., 2013). These studies show large anomalies from production of ¹⁹²Pt by neutron capture on ¹⁹¹Ir. This result partially motivated the choice of the ¹⁹⁶Pt–¹⁹⁸Pt DS used in this study, as described by Creech et al. (2013). Of the five stable Pt isotopes, ¹⁹⁵Pt has the largest thermal neutron capture cross section of ca. 28 barns (Mughabghab, 2003; Table S-8), and given the isotopes used in the ¹⁹⁶Pt–¹⁹⁸Pt DS inversion (¹⁹⁴Pt, ¹⁹⁵Pt, ¹⁹⁶Pt and ¹⁹⁸Pt), neutron capture on ¹⁹⁵Pt has the greatest potential to compromise our results.

The effect of thermal neutron capture on Pt isotope data was calculated by taking the natural relative abundances of Pt and Ir (and using a typical Pt/Ir ratio for iron meteorites), applying the mass changes from neutron captures to each isotope in proportion to its neutron capture cross section, and scaling these effects such that the anomalies produced roughly matched those published by Wittig et al. (2013) and Kruijer et al. (2013). We then estimated the sensitivity of DS corrected Pt isotope data to these effects by taking measured Pt isotope data and modifying the raw ratios by the amounts implied by these calculations. These calculations suggest that for values roughly corresponding to anomalies for Tlacotepec (the largest anomaly) in Wittig et al. (2013) and Kruijer et al. (2013), the maximum effect on DS corrected δ¹⁹⁸Pt is ~ –0.13 ‰. While this cannot be directly related to CRE ages, the effect of CRE is evidently to shift DS corrected δ¹⁹⁸Pt towards more negative compositions. As such, the apparently poor reproducibility between iron meteorites and replicate digestions of the same meteorite may in fact reflect real isotopic variations. Given that our iron meteorite data are all either chondritic or more negative than chondritic (Fig. S-2), variable disturbance from cosmic ray exposure is the best explanation for the variability amongst these samples, especially given that the magnitude of this variability is consistent with the calculated influence described above. Cosmic ray exposure induced anomalies in W and Os isotopes have been previously shown to vary on a cm scale in iron meteorites (Leya and Masarik, 2013; Qin et al., 2015). However, we emphasise that most Pt stable isotope data for iron meteorites are within error of the chondritic composition, and thus the variability discussed here has no impact on our interpretations.

**Figure S-2** Pt stable isotope data for individual replicates processed from iron meteorites. The dashed horizontal line and grey field represent the mean and 2 sd of chondrite data. The arrow indicates the expected direction of influence of cosmic ray exposure on DS-corrected Pt stable isotope data.

Platinum isotopic and abundance variations

For some sample groups, such as the Kaapvaal peridotite xenoliths, we observe isotopic variability well beyond the external reproducibility of our method. This variability is attributed to partial re-equilibration with material containing a late-veneer component. If correct, it is expected that samples with the heaviest Pt stable isotope composition should have the lowest Pt abundances. However, this correlated variability is not apparent in the Kaapvaal or Martian datasets. Considering the sizes of the samples involved, we attribute the lack of covariability to nugget effects, resulting in over- or under-estimation of Pt concentrations.

Supplementary Tables

Table S-1 Terrestrial Pt stable isotope and concentration results from double-spike MC-ICPMS, PGE concentrations from ID-ICPMS, and Re–Os isotope data determined by N-TIMS.

Sample	Rock type	Locality	Total sample processed (g)	total Pt (ng)	Pt conc. (ng g^–1)	(n)	δ¹⁹⁸Pt (‰)	±2σ	Blank proportion (%)	Magnitude of blank correction (‰)
Subcontinental Lithospheric Mantle Xenoliths
BC1	Lherzolite	Batchelor's Crater. Queensland. Australia	15.0	103	6.8	2	-0.25	0.04	0.8	-0.01
replicate			17.7	117	6.6	1	-0.14	0.03	0.8	-0.01
average							-0.19	0.15
BC7	Lherzolite	Batchelor's Crater. Queensland. Australia	15.6	78	5.2	1	-0.05	0.03	1.1	-0.02
MQ1	Lherzolite	Mount Quincan. Queensland. Australia	7.2	48	6.6	1	-0.07	0.04	1.7	-0.03
MQ14	Lherzolite	Mount Quincan. Queensland. Australia	15.0	103	6.8	2	-0.17	0.03	0.8	-0.01
replicate			10.7	65	6.1	2	-0.07	0.02	1.2	-0.02
average							-0.12	0.14

Ophiolites/transitional lherzolites
L 213	Lherzolite	Sezia Lanzo HP zone. Alps. Italy	17.6	111	6.3	1	-0.13	0.03	0.8	-0.01
LI 138	Lherzolite	Internal Liguride Ophiolite. Italy	15.1	108	7.3	3	-0.13	0.06	0.7	-0.01

Off craton basalt-hosted xenoliths
Mtf37	Lherzolite	Montferrier. Massif Central. France	15.1	103	6.8	3	-0.03	0.02	0.8	-0.01

Continental orogenic peridotites
TUR 7	Lherzolite	Turon de Técouère. Pyrénées. France	16.5	119	7.2	1	-0.10	0.04	0.7	-0.01

Cratonic xenoliths (Kaapvaal craton)
BG	Lherzolite	Kimberley. South Africa	15.0	21	1.4	2	-0.20	0.05	3.7	-0.07
BG2	Harzburgite	Kimberley. South Africa	13.1	56	4.2	2	0.16	0.03	1.4	-0.02
KBD4	Harzburgite	Kimberley. South Africa	15.3	27	1.7	2	0.27	0.06	3.0	-0.04
replicate			15.1	26	1.7	2	0.26	0.05	3.1	-0.04
replicate			15.6	30	1.9	2	0.46	0.04	2.8	-0.03
average							0.33	0.22
LMA2	Harzburgite	Matsoku kimberlite. Lesotho	11.9	29	2.4	2	-0.08	0.04	2.8	-0.05
LTP20	Lherzolite	Thaba Putsoa kimberlite. Lesotho	7.4	4	0.6	1	0.08	0.12	18.2	-0.33

Isua
460203	Metabasalt	Isua Supracrustal Belt. Greenland	20.25	67	3.3	3	0.13	0.09	1.6	-0.02
460204	Ultramafic schist	Isua Supracrustal Belt. Greenland	22.55	140	6.2	2	0.12	0.02	0.9	-0.01
460217	Metabasalt	Isua Supracrustal Belt. Greenland	20.58	80	3.9	1	0.16	0.16	1.4	-0.02
460218	Metabasalt	Isua Supracrustal Belt. Greenland	25.07	110	4.4	2	0.11	0.02	1.2	-0.02
duplicate			20.69	81	3.9	1	0.07	0.07	1.4	-0.02
average							0.09	0.05
460219	Metabasalt	Isua Supracrustal Belt. Greenland	25.22	112	4.4	2	0.13	0.07	1.2	-0.02
460258	Metabasalt	Isua Supracrustal Belt. Greenland	25.16	115	4.6	2	-0.08	0.06	1.2	-0.02
duplicate			20.53	72	3.5	2	0.04	0.02	1.5	-0.02
average							-0.02	0.17
460275	Metabasalt	Isua Supracrustal Belt. Greenland	20.40	82	4.0	1	0.11	0.05	1.3	-0.02
460276	Ultramafic schist	Isua Supracrustal Belt. Greenland	20.42	77	3.8	2	0.09	0.08	1.4	-0.02

Sample	Os (ng g^–1)	Ir (ng g^–1)	Ru (ng g^–1)	Pt (ng g^–1)	Pd (ng g^–1)	Re (ng g^–1)	Re/Os	¹⁸⁷Os/¹⁸⁸Os	¹⁸⁷Re/¹⁸⁸Os	T_MA (Ma)	T_RD (Ma)
Subcontinental Lithospheric Mantle Xenoliths
BC1

BC7
MQ1^a	1.46					0.214			0.707	113	113
MQ14^a	2.11					0.192			0.439	—	265

Ophiolites/transitional lherzolites
L 213^b	3.25	2.89	5.66	5.78	4.47	0.261			0.387
LI 138^c	3.04	3.06	5.45	7.15	5.95

Off craton basalt-hosted xenoliths
Mtf37^d	5.24	4.6	7.5	8.3	6.6	0.156			0.144

Continental orogenic peridotites
TUR 7^b	4.03	3.59	7.22	7.2	6.54	0.341			0.408

Cratonic xenoliths (Kaapvaal craton)
BG	2.37	2.26	3.26	1.58	0.49	0.217	0.091	0.10661	0.43882		3091
BG2	4.59	4.45	7.01	2.72	0.06	0.023	0.005	0.10865	0.02430	2986	2823
KBD4	2.56	2.16	2.84	1.83	0.06	0.060	0.023	0.10678	0.11276	4107	3069
replicate
replicate
average
LMA2	4.19	3.38	5.00	2.23	0.66	0.042	0.010	0.10724	0.04863	3376	3008
LTP20	4.42	3.15	5.79	0.44	0.08	0.170	0.039	0.10699	0.18496	5195	3041

Pt stable isotope compositions given are weighted means of the number of analyses (n) of each digestion, with the uncertainties reflecting the errors on the weighted means. Where multiple replicates were processed, the mean and 2 sd (n ≤ 2) or weighted mean (n ≥ 3) is also shown along with the values for each replicate. Details of analytical blanks and blank corrections are given in the Supplementary Information. All Re–Os ages for Kaapvaal samples calculated to PUM with a ¹⁸⁷Os/¹⁸⁸Os of 0.1296 and a ¹⁸⁷Re/¹⁸⁸Os of 0.435 after Becker et al. (2006). Model ages (TMA) for Kaapvaal samples are overestimated due to the presence of metasomatic sulphides that affect the Re/Os ratio. PGE and Re–Os data for MQ1, MQ14, L 213, LI 138, Mtf37 and TUR 7 are from the following references: ^aHandler et al. (2005); ^bBecker et al. (2006); ^cLuguet et al. (2004); ^dAlard et al. (2011).

Table S-2 Platinum stable isotope and concentration results for meteorite samples measured by double-spike MC-ICPMS.

Sample	Group	Total sample processed (g)	total Pt (µg)	Pt conc. (µg g^–1)	Ref. Pt conc. (µg g^–1)	δ¹⁹⁸Pt (‰)	±2σ	(n)	Blank proportion (%)	Magnitude of blank correction (‰)	CRE age (Myr)	Structural type‡	Kamacite bandwidth‡ (mm)
Ordinary chondrites
Bovedy	L3	0.52	0.52	1.00		-0.15	0.03	3	0.2	0.00
Begaa	LL3	0.48	0.25	0.52		-0.15	0.03	3	0.3	0.00
Talbachat n'aït Isfoul	LL3	0.96	0.22	0.23		-0.14	0.02	4	0.4	-0.01
replicate		0.49	0.19	0.39		-0.19	0.04	3	0.4	-0.01
average						-0.17	0.07	7
SAH 97172	L5	1.36	1.00	0.79		-0.16	0.04	4	0.1	0.00

Enstatite chondrites
SAH 97096	EH3	0.47	0.63	1.33		-0.15	0.03	4	0.1	0.00
replicate		0.47	0.60	1.27		-0.17	0.01	1	0.1	0.00
average						-0.16	0.03
SAH 97159	EH3	1.24	1.47	1.12		-0.14	0.04		0.1	0.00

Carbonaceous chondrites
Allende	CV3	1.00	1.26	1.26	1.36^a	-0.10	0.04	3	0.1	0.00
replicate		1.14	1.45	1.27		-0.10	0.07	3	0.1	0.00
replicate		1.04	1.24	1.20		-0.13	0.01	5	0.1	0.00
replicate		1.02	1.30	1.29		-0.12	0.04	5	0.1	0.00
replicate		1.05	1.22	1.16		-0.17	0.02	2	0.1	0.00
average				1.24		-0.12	0.06
Gujba	CB3	0.48	2.13	4.42		-0.08	0.04	3	0.0	0.00
NWA 1232	CO3	0.54	0.52	0.98		-0.08	0.03	3	0.2	0.00
NWA 763	CO3	0.48	0.49	1.02		-0.17	0.04	3	0.2	0.00
NWA 1559	CK3	0.66	0.32	0.49		-0.17	0.04	4	0.3	0.00
NWA 1563	CK5	0.50	0.35	0.68		-0.08	0.03	4	0.2	0.00
NWA 723	CV3	0.44	0.51	1.16		-0.24	0.03	3	0.2	0.00
SaU 290	CH3	0.64	0.50	0.78		-0.16	0.03	3	0.2	0.00

Primitive achondrites
Dho 125	Acapulcoite	0.86	1.24	1.45		-0.09	0.02	4	0.1	0.00
replicate		0.86	1.25	1.45		-0.09	0.02	3	0.1	0.00
average						-0.09	0.01
NWA 2871	Acapulcoite	0.88	1.99	2.27		-0.04	0.03	3	0.0	0.00
Tierra Blanca	Winonaite	0.31	0.25	0.82		0.01	0.03	1	0.3	0.00
ALH 84136	Ureilite	1.20	0.31	0.26	0.28^b	-0.05	0.06	3	0.3	0.00
ALHA77257	Ureilite	1.54	0.23	0.15	0.19^b	-0.01	0.03	3	0.3	-0.01
ALHA78019	Ureilite	0.27	0.25	0.90	1.54^b	-0.21	0.04	3	0.3	-0.01
ALHA81101	Ureilite	1.23	0.08	0.06	0.06^b	0.17	0.13	3	1.0	-0.02
EET 87517	Ureilite	1.09	0.54	0.49	0.59^b	-0.05	0.06	3	0.1	0.00
EET 96042	Ureilite	1.47	0.69	0.47	0.57^b	-0.04	0.06	3	0.1	0.00
GRA 95205	Ureilite	1.00	0.54	0.54	0.83^b	-0.15	0.09	3	0.1	0.00
GRA 98032	Ureilite	1.07	0.26	0.25	0.41^b	-0.13	0.11	3	0.3	0.00
GRO 95575	Ureilite	1.45	0.54	0.37	0.31^b	0.10	0.14	3	0.1	0.00
NWA 2234	Ureilite	0.69	0.18	0.26		0.00	0.04	4	0.4	-0.01
PCA 82506	Ureilite	1.78	0.47	0.26	0.47^b	-0.19	0.02	3	0.2	0.00
META 78008	Ureilite	0.74	0.27	0.36	1.04^b	-0.11	0.07	1	0.3	0.00

Iron meteorites
Canyon Diablo	IAB			6.09	8.0^c	-0.24	0.05	1			540^d	Og	2
replicate				6.73		-0.10	0.03	8
average						-0.17	0.20
Toluca	IAB			7.30	5.47^e	-0.18	0.03	2			600^f	Og	1.4
replicate				5.39		-0.17	0.05	4
replicate				2.96		-0.13	0.02	1
average						-0.16	0.05
Henbury	IIIAB			15.24	18.31^e	-0.31	0.03	2			700^f	Om	0.9
replicate				16.55		-0.24	0.03	7
replicate				15.44		-0.18	0.04	8
average						-0.24	0.13
Savik (Cape York)	IIIAB				11.85^e	-0.27	0.02	6			82^g	Og	1.2
replicate						-0.28	0.05	3
average						-0.28	0.01
Gibeon	IVA			3.90	6.91^e	-0.22	0.02	7			400^f§	Of	0.3
replicate				3.90		-0.09	0.02	2
average						-0.16	0.18
Chinga	IVB/ung			7.61	9.56^e	-0.19	0.01	7			845^f§	D	–
replicate				9.39		-0.09	0.03	7
average						-0.14	0.14

Pt stable isotope compositions are weighted means of the number of analyses (n) of each digestion, with the uncertainties reflecting the errors on the weighted means. Where multiple replicates were processed, the mean and 2 sd is also shown along with the values for each replicate. Details of analytical blanks and blank corrections are given in the Supplementary Information. No blank correction was applied to iron meteorite data due to different digestion methods (Supplementary Information). Greater uncertainty in iron meteorite data may reflect cosmogenic effects in these samples, as discussed in the Supplementary Information. ^§CRE age given is weighted average for group. ^aFischer-Gödde et al. (2010); ^bRankenburg et al. (2008); ^cCrocket (1972); ^dMichlovich et al. (1994); ^ePetaev and Jacobsen (2004); ^fScherstén et al. (2006); ^gMathew and Marti (2009).

Table S-3 Determination of Pt blanks from replicate NiS blank digestions.

#	digestions	amount of 'sample' (g)	total blank (ng)	blank per 15g digestion (ng)
1	1	25.61	1.57	0.92
2	2	40.37	4.72	1.75
3	1	15.01	0.22	0.22
4	1	15.10	0.23	0.23
5	1	15.00	1.20	1.20
6	1	15.01	1.04	1.04
7	1	15.00	0.54	0.54
8	1	15.00	0.49	0.49
9	1	15.14	0.46	0.46
10	1	15.07	0.76	0.76
11	4	60.37	4.49	1.12
12	1	20.10	1.31	0.98

			mean	0.81
			median	0.84
			max	1.75

Table S-4 Isotopic composition of blanks from replicate, double-spiked, NiS blank digestions.

	Total 'sample' digested (g)	digestions	δ¹⁹⁸Pt (‰)	±	total blank (ng)	blank per 15g digestion (ng)
DS blank 1	60.37	4	0.83	0.25	4.48	1.11
DS blank 2	20.1	1	1.58	0.39	1.31	0.98

Table S-5 Pt isotope data and NiS blank correction of mixtures of IRMM-010 and NiS blank Pt.

	blank Pt proportion (%)	δ¹⁹⁸Pt (‰)	±	δ¹⁹⁸Pt_BC (‰)
IRMM-010	0	0.00	0.03	0.00
IRMM-010–blank mixture #1	10	0.12	0.07	0.02
IRMM-010–blank mixture #2	30	0.33	0.07	0.00

Table S-6 Pt isotope data and NiS blank correction of low-Pt replicate digestions of peridotite BC1, compared with normal digestions of the same sample.

	Total sample processed (g)	digestions	δ¹⁹⁸Pt (‰)	±	total Pt (ng)	Pt conc. (ng g^-1)	blank per digestion (ng)	total blank (ng)	blank Pt proportion (%)	δ¹⁹⁸Pt_BC (‰)
BC1 low-Pt #1	2.8	3	0.17	0.07	19	6.6	1.5	4.6	20	–0.18
BC1 low-Pt #2	3.0	3	0.18	0.07	20	6.6	1.7	5.0	20	–0.17
BC1 low-Pt #3	3.2	3	0.00	0.07	18	5.6	0.7	2.0	10	–0.17

BC1	15.0	1	–0.25	0.04	103	6.8	1.0	1.0	1	–0.27
BC1	17.7	1	–0.14	0.03	117	6.6	1.0	1.0	1	–0.16

Table S-7 Model parameters used in preparing Figure 3.

Reservoir		D_Pt^met/sil	Pt conc. (ng g^-1)	δ¹⁹⁸Pt (‰)
Pre-late-veneer mantle	– Low P-T	≥106	<1.44 x 10-5
	– High P-T	104	0.144

Post-veneer-mantle			7.63	–0.10 ± 0.10
Late-veneer (chondrite)			982	–0.19 ± 0.14

Table S-8 Thermal neutron capture cross sections for isotopes in the mass range of platinum. ^aIsotopic abundances from Berglund and Wieser (2011). ^bThermal neutron capture cross sections from Mughabghab (2003).

Isotope	¹⁹⁰Pt	¹⁹¹Ir	¹⁹²Pt	¹⁹³Ir	¹⁹⁴Pt	¹⁹⁵Pt	¹⁹⁶Pt	¹⁹⁷Au	¹⁹⁸Pt
Relative abundance^a (%)	0.01	37.27	0.78	62.73	32.97	33.83	25.24	100	7.16
Cross-section^b (barns)	147	954	10	111	~1	28	~0.6	99	3.66

Supplementary Information References

Alard, O., Lorand, J.-P., Reisberg, L., Bodinier, J.-L., Dautria, J.-M., O’Reilly, S.Y. (2011) Volatile-rich Metasomatism in Montferrier Xenoliths (Southern France): Implications for the Abundances of Chalcophile and Highly Siderophile Elements in the Subcontinental Mantle. Journal of Petrology 52, 2009–2045.

Baadsgaard, H., Nutman, A.P., Bridgwater, D., Rosing, M., McGregor, V.R., Allaart, J.H. (1984) The zircon geochronology of the Akilia association and Isua supracrustal belt, West Greenland. Earth and Planetary Science Letters 68, 221–228.

Berglund, M., Wieser, M.E. (2011) Isotopic compositions of the elements 2009 (IUPAC Technical Report). Pure and Applied Chemistry 83, 397–410.

Bodinier, J.-L., Menzies, M.A., Thirlwall, M.F. (1991) Continental to Oceanic Mantle Transition—REE and Sr-Nd Isotopic Geochemistry of the Lanzo Lherzolite Massif. Journal of Petrology Special Volume, 191–210.

Cohen, A.S., Waters, F.G. (1996) Separation of osmium from geological materials by solvent extraction for analysis by thermal ionisation mass spectrometry. Analytica Chimica Acta 332, 269–275.

Creech, J.B., Baker, J.A., Handler, M.R., Bizzarro, M. (2014) Platinum stable isotope analysis of geological standard reference materials by double-spike MC-ICPMS. Chemical Geology 363, 293–300.

Creech, J.B., Paul, B. (2015) IsoSpike: Improved Double-Spike Inversion Software. Geostandards and Geoanalytical Research 39, 7–15.

Crocket, J.H. (1972) Some aspects of the geochemistry of Ru, Os, Ir and Pt in iron meteorites. Geochimica et Cosmochimica Acta 36, 517–535.

Dale, C.W., Luguet, A., Macpherson, C.G., Pearson, D.G., Hickey-Vargas, R. (2008) Extreme platinum-group element fractionation and variable Os isotope compositions in Philippine Sea Plate basalts: Tracing mantle source heterogeneity. Chemical Geology 248, 213–238.

Fabriès, J., Lorand, J.-P., Bodinier, J.-L. (1998) Petrogenetic evolution of orogenic lherzolite massifs in the central and western Pyrenees. Tectonophysics 292, 145–167.

Fischer-Gödde, M., Becker, H., Wombacher, F. (2010) Rhodium, gold and other highly siderophile element abundances in chondritic meteorites. Geochimica et Cosmochimica Acta 74, 356–379.

Handler, M.R., Bennett, V.C., Carlson, R.W. (2005) Nd, Sr and Os isotope systematics in young, fertile spinel peridotite xenoliths from northern Queensland, Australia: A unique view of depleted MORB mantle? Geochimica et Cosmochimica Acta 69, 5747–5763.

Harries, D., Berg, T., Langenhorst, F., Palme, H. (2012) Structural clues to the origin of refractory metal alloys as condensates of the solar nebula. Meteoritics & Planetary Science 47, 2148–2159.

Hewins, R.H. et al. (2014) The Paris meteorite, the least altered CM chondrite so far. Geochimica et Cosmochimica Acta 124, 190–222.

Kruijer, T.S., Fischer-Gödde, M., Kleine, T., Sprung, P., Leya, I., Wieler, R. (2013) Neutron capture on Pt isotopes in iron meteorites and the Hf–W chronology of core formation in planetesimals. Earth and Planetary Science Letters 361, 162–172.

Leya, I., Masarik, J. (2013) Thermal neutron capture effects in radioactive and stable nuclide systems. Meteoritics & Planetary Science 48, 665–685.

Lorand, J.-P., Grégoire, M. (2006) Petrogenesis of base metal sulphide assemblages of some peridotites from the Kaapvaal craton (South Africa). Contributions to Mineralogy and Petrology 151, 521–538.

Luguet, A., Lorand, J.-P., Alard, O., Cottin, J.-Y. (2004) A multi-technique study of platinum group element systematic in some Ligurian ophiolitic peridotites, Italy. Chemical Geology 208, 175–194.

Mathew, K.J., Marti, K. (2009) Galactic cosmic ray-produced 129Xe and 131Xe excesses in troilites of the Cape York iron meteorite. Meteoritics & Planetary Science 44, 107–114.

Michard-Vitrac, A., Lancelot, J., Allègre, C.J., Moorbath, S. (1977) U-Pb ages on single zircons from the Early Precambrian rocks of West Greenland and the Minnesota River Valley. Earth and Planetary Science Letters 35, 449–453.

Michlovich, E.S., Vogt, S., Masarik, J., Reedy, R.C., Elmore, D., Lipschutz, M.E. (1994) Aluminum 26, 10Be, and 36Cl depth profiles in the Canyon Diablo iron meteorite. Journal of Geophysical Research: Planets 99, 23187–23194.

Mughabghab, S.F. (2003) Thermal neutron capture cross sections resonance integrals and g-factors. Report, INDC (NDS)-440, International Atomic Energy Agency.

Pearson, D.G., Shirey, S.B., Harris, J.W., Carlson, R.W. (1998) Sulphide inclusions in diamonds from the Koffiefontein kimberlite, S Africa: constraints on diamond ages and mantle Re–Os systematics. Earth and Planetary Science Letters 160, 311–326.

Pearson, D.G., Woodland, S.J. (2000) Solvent extraction/anion exchange separation and determination of PGEs (Os, Ir, Pt, Pd, Ru) and Re-Os isotopes in geological samples by isotope dilution ICP-MS. Chemical Geology 165, 87–107.

Petaev, M.I., Jacobsen, S.B. (2004) Differentiation of metal‐rich meteoritic parent bodies: I. Measurements of PGEs, Re, Mo, W, and Au in meteoritic Fe‐Ni metal. Meteoritics & Planetary Science 39, 1685–1697.

Qin, L., Dauphas, N., Horan, M.F., Leya, I., Carlson, R.W. (2015) Correlated cosmogenic W and Os isotopic variations in Carbo and implications for Hf–W chronology. Geochimica et Cosmochimica Acta 153, 91–104.

Rankenburg, K., Humayun, M., Brandon, A.D., Herrin, J.S. (2008) Highly siderophile elements in ureilites. Geochimica et Cosmochimica Acta 72, 4642–4659.

Scherstén, A., Elliott, T., Hawkesworth, C., Russell, S., Masarik, J. (2006) Hf-W evidence for rapid differentiation of iron meteorite parent bodies. Earth and Planetary Science Letters 241, 530–542.

Wittig, N., Humayun, M., Brandon, A.D., Huang, S., Leya, I. (2013) Coupled W–Os–Pt isotope systematics in IVB iron meteorites: In situ neutron dosimetry for W isotope chronology. Earth and Planetary Science Letters 361, 152–161.

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Figures and Tables

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Supplementary Figures and Tables

Table S-1 Terrestrial Pt stable isotope and concentration results from double-spike MC-ICPMS, PGE concentrations from ID-ICPMS, and Re–Os isotope data determined by N-TIMS.

Sample	Rock type	Locality	Total sample processed (g)	total Pt (ng)	Pt conc. (ng g^–1)	(n)	δ¹⁹⁸Pt (‰)	±2σ	Blank proportion (%)	Magnitude of blank correction (‰)
Subcontinental Lithospheric Mantle Xenoliths
BC1	Lherzolite	Batchelor's Crater. Queensland. Australia	15.0	103	6.8	2	-0.25	0.04	0.8	-0.01
replicate			17.7	117	6.6	1	-0.14	0.03	0.8	-0.01
average							-0.19	0.15
BC7	Lherzolite	Batchelor's Crater. Queensland. Australia	15.6	78	5.2	1	-0.05	0.03	1.1	-0.02
MQ1	Lherzolite	Mount Quincan. Queensland. Australia	7.2	48	6.6	1	-0.07	0.04	1.7	-0.03
MQ14	Lherzolite	Mount Quincan. Queensland. Australia	15.0	103	6.8	2	-0.17	0.03	0.8	-0.01
replicate			10.7	65	6.1	2	-0.07	0.02	1.2	-0.02
average							-0.12	0.14

Ophiolites/transitional lherzolites
L 213	Lherzolite	Sezia Lanzo HP zone. Alps. Italy	17.6	111	6.3	1	-0.13	0.03	0.8	-0.01
LI 138	Lherzolite	Internal Liguride Ophiolite. Italy	15.1	108	7.3	3	-0.13	0.06	0.7	-0.01

Off craton basalt-hosted xenoliths
Mtf37	Lherzolite	Montferrier. Massif Central. France	15.1	103	6.8	3	-0.03	0.02	0.8	-0.01

Continental orogenic peridotites
TUR 7	Lherzolite	Turon de Técouère. Pyrénées. France	16.5	119	7.2	1	-0.10	0.04	0.7	-0.01

Cratonic xenoliths (Kaapvaal craton)
BG	Lherzolite	Kimberley. South Africa	15.0	21	1.4	2	-0.20	0.05	3.7	-0.07
BG2	Harzburgite	Kimberley. South Africa	13.1	56	4.2	2	0.16	0.03	1.4	-0.02
KBD4	Harzburgite	Kimberley. South Africa	15.3	27	1.7	2	0.27	0.06	3.0	-0.04
replicate			15.1	26	1.7	2	0.26	0.05	3.1	-0.04
replicate			15.6	30	1.9	2	0.46	0.04	2.8	-0.03
average							0.33	0.22
LMA2	Harzburgite	Matsoku kimberlite. Lesotho	11.9	29	2.4	2	-0.08	0.04	2.8	-0.05
LTP20	Lherzolite	Thaba Putsoa kimberlite. Lesotho	7.4	4	0.6	1	0.08	0.12	18.2	-0.33

Isua
460203	Metabasalt	Isua Supracrustal Belt. Greenland	20.25	67	3.3	3	0.13	0.09	1.6	-0.02
460204	Ultramafic schist	Isua Supracrustal Belt. Greenland	22.55	140	6.2	2	0.12	0.02	0.9	-0.01
460217	Metabasalt	Isua Supracrustal Belt. Greenland	20.58	80	3.9	1	0.16	0.16	1.4	-0.02
460218	Metabasalt	Isua Supracrustal Belt. Greenland	25.07	110	4.4	2	0.11	0.02	1.2	-0.02
duplicate			20.69	81	3.9	1	0.07	0.07	1.4	-0.02
average							0.09	0.05
460219	Metabasalt	Isua Supracrustal Belt. Greenland	25.22	112	4.4	2	0.13	0.07	1.2	-0.02
460258	Metabasalt	Isua Supracrustal Belt. Greenland	25.16	115	4.6	2	-0.08	0.06	1.2	-0.02
duplicate			20.53	72	3.5	2	0.04	0.02	1.5	-0.02
average							-0.02	0.17
460275	Metabasalt	Isua Supracrustal Belt. Greenland	20.40	82	4.0	1	0.11	0.05	1.3	-0.02
460276	Ultramafic schist	Isua Supracrustal Belt. Greenland	20.42	77	3.8	2	0.09	0.08	1.4	-0.02

Sample	Os (ng g^–1)	Ir (ng g^–1)	Ru (ng g^–1)	Pt (ng g^–1)	Pd (ng g^–1)	Re (ng g^–1)	Re/Os	¹⁸⁷Os/¹⁸⁸Os	¹⁸⁷Re/¹⁸⁸Os	T_MA (Ma)	T_RD (Ma)
Subcontinental Lithospheric Mantle Xenoliths
BC1

BC7
MQ1^a	1.46					0.214			0.707	113	113
MQ14^a	2.11					0.192			0.439	—	265

Ophiolites/transitional lherzolites
L 213^b	3.25	2.89	5.66	5.78	4.47	0.261			0.387
LI 138^c	3.04	3.06	5.45	7.15	5.95

Off craton basalt-hosted xenoliths
Mtf37^d	5.24	4.6	7.5	8.3	6.6	0.156			0.144

Continental orogenic peridotites
TUR 7^b	4.03	3.59	7.22	7.2	6.54	0.341			0.408

Cratonic xenoliths (Kaapvaal craton)
BG	2.37	2.26	3.26	1.58	0.49	0.217	0.091	0.10661	0.43882		3091
BG2	4.59	4.45	7.01	2.72	0.06	0.023	0.005	0.10865	0.02430	2986	2823
KBD4	2.56	2.16	2.84	1.83	0.06	0.060	0.023	0.10678	0.11276	4107	3069
replicate
replicate
average
LMA2	4.19	3.38	5.00	2.23	0.66	0.042	0.010	0.10724	0.04863	3376	3008
LTP20	4.42	3.15	5.79	0.44	0.08	0.170	0.039	0.10699	0.18496	5195	3041

Table S-2 Platinum stable isotope and concentration results for meteorite samples measured by double-spike MC-ICPMS.

<> Sample

Group	Total sample processed (g)	total Pt (µg)	Pt conc. (µg g^–1)	Ref. Pt conc. (µg g^–1)	δ¹⁹⁸Pt (‰)	±2σ	(n)	Blank proportion (%)	Magnitude of blank correction (‰)	CRE age (Myr)	Structural type‡	Kamacite bandwidth‡ (mm)
Ordinary chondrites
Bovedy	L3	0.52	0.52	1.00		-0.15	0.03	3	0.2	0.00
Begaa	LL3	0.48	0.25	0.52		-0.15	0.03	3	0.3	0.00
Talbachat n'aït Isfoul	LL3	0.96	0.22	0.23		-0.14	0.02	4	0.4	-0.01
replicate		0.49	0.19	0.39		-0.19	0.04	3	0.4	-0.01
average						-0.17	0.07	7
SAH 97172	L5	1.36	1.00	0.79		-0.16	0.04	4	0.1	0.00

Enstatite chondrites
SAH 97096	EH3	0.47	0.63	1.33		-0.15	0.03	4	0.1	0.00
replicate		0.47	0.60	1.27		-0.17	0.01	1	0.1	0.00
average						-0.16	0.03
SAH 97159	EH3	1.24	1.47	1.12		-0.14	0.04		0.1	0.00

Carbonaceous chondrites
Allende	CV3	1.00	1.26	1.26	1.36^a	-0.10	0.04	3	0.1	0.00
replicate		1.14	1.45	1.27		-0.10	0.07	3	0.1	0.00
replicate		1.04	1.24	1.20		-0.13	0.01	5	0.1	0.00
replicate		1.02	1.30	1.29		-0.12	0.04	5	0.1	0.00
replicate		1.05	1.22	1.16		-0.17	0.02	2	0.1	0.00
average				1.24		-0.12	0.06
Gujba	CB3	0.48	2.13	4.42		-0.08	0.04	3	0.0	0.00
NWA 1232	CO3	0.54	0.52	0.98		-0.08	0.03	3	0.2	0.00
NWA 763	CO3	0.48	0.49	1.02		-0.17	0.04	3	0.2	0.00
NWA 1559	CK3	0.66	0.32	0.49		-0.17	0.04	4	0.3	0.00
NWA 1563	CK5	0.50	0.35	0.68		-0.08	0.03	4	0.2	0.00
NWA 723	CV3	0.44	0.51	1.16		-0.24	0.03	3	0.2	0.00
SaU 290	CH3	0.64	0.50	0.78		-0.16	0.03	3	0.2	0.00

Primitive achondrites
Dho 125	Acapulcoite	0.86	1.24	1.45		-0.09	0.02	4	0.1	0.00
replicate		0.86	1.25	1.45		-0.09	0.02	3	0.1	0.00
average						-0.09	0.01
NWA 2871	Acapulcoite	0.88	1.99	2.27		-0.04	0.03	3	0.0	0.00
Tierra Blanca	Winonaite	0.31	0.25	0.82		0.01	0.03	1	0.3	0.00
ALH 84136	Ureilite	1.20	0.31	0.26	0.28^b	-0.05	0.06	3	0.3	0.00
ALHA77257	Ureilite	1.54	0.23	0.15	0.19^b	-0.01	0.03	3	0.3	-0.01
ALHA78019	Ureilite	0.27	0.25	0.90	1.54^b	-0.21	0.04	3	0.3	-0.01
ALHA81101	Ureilite	1.23	0.08	0.06	0.06^b	0.17	0.13	3	1.0	-0.02
EET 87517	Ureilite	1.09	0.54	0.49	0.59^b	-0.05	0.06	3	0.1	0.00
EET 96042	Ureilite	1.47	0.69	0.47	0.57^b	-0.04	0.06	3	0.1	0.00
GRA 95205	Ureilite	1.00	0.54	0.54	0.83^b	-0.15	0.09	3	0.1	0.00
GRA 98032	Ureilite	1.07	0.26	0.25	0.41^b	-0.13	0.11	3	0.3	0.00
GRO 95575	Ureilite	1.45	0.54	0.37	0.31^b	0.10	0.14	3	0.1	0.00
NWA 2234	Ureilite	0.69	0.18	0.26		0.00	0.04	4	0.4	-0.01
PCA 82506	Ureilite	1.78	0.47	0.26	0.47^b	-0.19	0.02	3	0.2	0.00
META 78008	Ureilite	0.74	0.27	0.36	1.04^b	-0.11	0.07	1	0.3	0.00

Iron meteorites
Canyon Diablo	IAB			6.09	8.0^c	-0.24	0.05	1			540^d	Og	2
replicate				6.73		-0.10	0.03	8
average						-0.17	0.20
Toluca	IAB			7.30	5.47^e	-0.18	0.03	2			600^f	Og	1.4
replicate				5.39		-0.17	0.05	4
replicate				2.96		-0.13	0.02	1
average						-0.16	0.05
Henbury	IIIAB			15.24	18.31^e	-0.31	0.03	2			700^f	Om	0.9
replicate				16.55		-0.24	0.03	7
replicate				15.44		-0.18	0.04	8
average						-0.24	0.13
Savik (Cape York)	IIIAB				11.85^e	-0.27	0.02	6			82^g	Og	1.2
replicate						-0.28	0.05	3
average						-0.28	0.01
Gibeon	IVA			3.90	6.91^e	-0.22	0.02	7			400^f§	Of	0.3
replicate				3.90		-0.09	0.02	2
average						-0.16	0.18
Chinga	IVB/ung			7.61	9.56^e	-0.19	0.01	7			845^f§	D	–
replicate				9.39		-0.09	0.03	7
average						-0.14	0.14

Table S-3 Determination of Pt blanks from replicate NiS blank digestions.

#	digestions	amount of 'sample' (g)	total blank (ng)	blank per 15g digestion (ng)
1	1	25.61	1.57	0.92
2	2	40.37	4.72	1.75
3	1	15.01	0.22	0.22
4	1	15.10	0.23	0.23
5	1	15.00	1.20	1.20
6	1	15.01	1.04	1.04
7	1	15.00	0.54	0.54
8	1	15.00	0.49	0.49
9	1	15.14	0.46	0.46
10	1	15.07	0.76	0.76
11	4	60.37	4.49	1.12
12	1	20.10	1.31	0.98

			mean	0.81
			median	0.84
			max	1.75

Table S-4 Isotopic composition of blanks from replicate, double-spiked, NiS blank digestions.

	Total 'sample' digested (g)	digestions	δ¹⁹⁸Pt (‰)	±	total blank (ng)	blank per 15g digestion (ng)
DS blank 1	60.37	4	0.83	0.25	4.48	1.11
DS blank 2	20.1	1	1.58	0.39	1.31	0.98

Table S-5 Pt isotope data and NiS blank correction of mixtures of IRMM-010 and NiS blank Pt.

	blank Pt proportion (%)	δ¹⁹⁸Pt (‰)	±	δ¹⁹⁸Pt_BC (‰)
IRMM-010	0	0.00	0.03	0.00
IRMM-010–blank mixture #1	10	0.12	0.07	0.02
IRMM-010–blank mixture #2	30	0.33	0.07	0.00

Table S-6 Pt isotope data and NiS blank correction of low-Pt replicate digestions of peridotite BC1, compared with normal digestions of the same sample.

	Total sample processed (g)	digestions	δ¹⁹⁸Pt (‰)	±	total Pt (ng)	Pt conc. (ng g^-1)	blank per digestion (ng)	total blank (ng)	blank Pt proportion (%)	δ¹⁹⁸Pt_BC (‰)
BC1 low-Pt #1	2.8	3	0.17	0.07	19	6.6	1.5	4.6	20	–0.18
BC1 low-Pt #2	3.0	3	0.18	0.07	20	6.6	1.7	5.0	20	–0.17
BC1 low-Pt #3	3.2	3	0.00	0.07	18	5.6	0.7	2.0	10	–0.17

BC1	15.0	1	–0.25	0.04	103	6.8	1.0	1.0	1	–0.27
BC1	17.7	1	–0.14	0.03	117	6.6	1.0	1.0	1	–0.16

Table S-7 Model parameters used in preparing Figure 3.

Reservoir		D_Pt^met/sil	Pt conc. (ng g^-1)	δ¹⁹⁸Pt (‰)
Pre-late-veneer mantle	– Low P-T	≥106	<1.44 x 10-5
	– High P-T	104	0.144

Post-veneer-mantle			7.63	–0.10 ± 0.10
Late-veneer (chondrite)			982	–0.19 ± 0.14

Isotope	¹⁹⁰Pt	¹⁹¹Ir	¹⁹²Pt	¹⁹³Ir	¹⁹⁴Pt	¹⁹⁵Pt	¹⁹⁶Pt	¹⁹⁷Au	¹⁹⁸Pt
Relative abundance^a (%)	0.01	37.27	0.78	62.73	32.97	33.83	25.24	100	7.16
Cross-section^b (barns)	147	954	10	111	~1	28	~0.6	99	3.66