LETTERS

The lead isotopic age of the Earth can be explained bycore formation aloneBernard J. Wood1,2 & Alex N. Halliday1

The meaning of the age of the Earth defined by lead isotopes haslong been unclear. Recently it has been proposed1 that the age of theEarth deduced from lead isotopes reflects volatile loss to space atthe time of the Moon-forming giant impact rather than partition-ing into metallic liquids during protracted core formation. Here weshow that lead partitioning into liquid iron depends strongly oncarbon content and that, given a content of 0.2% carbon2,3, experi-mental and isotopic data both provide evidence of strong partition-ing of lead into the core throughout the Earth’s accretion. Earlierconclusions that lead is weakly partitioned into iron arose from theuse of carbon-saturated (about 5% C) iron alloys. The lead isotopicage of the Earth is therefore consistent with partitioning into thecore and with no significant late losses of moderately volatile ele-ments to space during the giant impact.

Both 235/238U–207/206Pb (ref. 4) and 129I–129Xe (ref. 5) chronologyyield apparent ages of loss of the daughter isotopes from the bulksilicate Earth (BSE) about 100 million years (Myr) after the origin ofthe Solar System. Although Xe removal is generally assumed to be dueto atmospheric loss associated with impacts such as the one thoughtto have formed the Moon5,6, the causes of Pb loss are more conten-tious, with opinion divided between Pb being atmophile and lostwith Xe1 or siderophile4,7 and added to the core. Whether Pb wasdominantly atmophile or dominantly siderophile, core formationappears to have affected the abundances of elements of similarvolatility to Pb in the BSE (Fig. 1). This shows that elements suchas S, Ge and Bi, which are siderophile, are up to two orders of mag-nitude more depleted than lithophile B and F.

The short-lived 182Hf–182W chronometer provides a constraint onthe age of the Moon of over 45 Myr, and probably over 60 Myr, after thestart of the Solar System8. A model Rb–Sr age of 70–110 Myr (ref. 9) forthe Moon is in excellent agreement with the age of the early lunarcrust10 and the 235/238U–207/206Pb age of Pb loss from the Moon11. Italso agrees with the two stage 235/238U–207/206Pb and 129I–129Xe ages ofthe BSE. There is now, therefore, broad consistency between the timingof the giant impact, volatile depletion of the Moon, an event thatremoved highly volatile noble gases from the Earth and the loss ofPb from the BSE9 either to the core or to space.

Recent experimental data1,12 (Fig. 2a) purport to show that Pb isinsufficiently siderophile to explain either the observed 238U/204Pb(m) or Pb isotopic composition of the BSE in terms of core formation.An overall partition coefficient DPb( 5 ½Pb�metal=½Pb�silicate) of about25 would be needed to explain the change of m from ,0.7 to ,10 bycore formation alone4. In Fig. 2 the metal–silicate partitioning of Pb isshown relative to that for Fe in terms of an exchange reaction:

PbO(silicate) 1 Fe(metal) 5 Pb(metal) 1 FeO(silicate) (1)

Treating the lead partitioning data in terms of reaction (1) is equival-ent to normalizing the data to constant oxygen fugacity. This enables

results from experiments at different oxygen fugacities to be plottedon the same diagram, provided (as shown in the SupplementaryInformation) that Pb is present as Pb21 in the silicate. To apply equi-librium (1) we calculate an exchange coefficient KD in terms of theweight partition coefficients of Pb and Fe, DPb 5 [Pb]metal/[Pb]silicate

and DFe5 [Fe]metal/[Fe]silicate:

KD 5 DPb/DFe

Given a DFe for the Earth of 13.6 (ref. 13), a logKD of about 0.25 isrequired to explain the U/Pb ratio of the BSE by core formationalone. Because experimentally determined partitioning appears tobe about two orders of magnitude lower, Lagos et al.1 assert that analternative mechanism of Pb loss from the BSE is required. The mostlikely one would be that Pb was atmophile and lost with the noblegases at around the time of the Moon-forming impact.

The experiments that were used to deduce that Pb is not sidero-phile were, however, performed in graphite capsules1 so that the Femetal contained about 5 weight per cent carbon (wt% C). The experi-ments of Malavergne et al.12 are difficult to interpret because manysamples did not undergo complete melting of the silicate. One experi-ment was performed in an MgO capsule, the remainder in C capsulesand the metal also contained large amounts of Si (up to 27 wt%).

1Department of Earth Sciences, Parks Road, Oxford OX1 3PR, UK. 2Department of Earth and Planetary Sciences, Macquarie University, New South Wales 2109, Australia.

Half mass condensation temperature (K)6008001,0001,200

Ab

und

ance

nor

mal

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to

CI c

hond

rites

and

Mg

0.001

0.01

0.1

1

10

ZnFK

Li

Te Se SAu

In

Ag

204Pb

Sn

BiCd

I

Br TlGe

RbB

Cs

Na

Mn Ga

Cu Sb

ClAs

Figure 1 | Volatile element abundances in the BSE. Abundances in Allendecarbonaceous chondrite (white symbols) and in the BSE31 are plotted againsttemperature of 50% condensation (T50) from a gas of solar composition28.The abundances are normalized to CI carbonaceous chondrite and Mg. Redsymbols refer to highly siderophile elements in the BSE , blue symbols referto lithophile and black symbols refer to weakly and moderately siderophileelements. We note that siderophile S, Bi and Ge are one to two orders ofmagnitude more depleted in the BSE than lithophile B and F. Pb is about oneorder of magnitude more depleted than F.

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767Macmillan Publishers Limited. All rights reserved©2010

Carbon and Si dissolved in liquid Fe have long been known toexclude Pb (refs 14, 15) and this leads us to the conclusion thatFig. 2a probably underestimates Pb partitioning into the metal. Weperformed further experiments to test this hypothesis.

Experiments were performed in the piston–cylinder apparatus (at1.5–2 GPa) and in the multi-anvil apparatus at 24 GPa. Startingmaterials were intimate mixtures of silicates and metals, as describedpreviously16. Both graphite and polycrystalline MgO capsules wereused at low pressures, and single-crystal MgO capsules were used inthe multi-anvil apparatus. Products were analysed by electronmicroprobe and by laser inductively coupled plasma mass spectro-metry (ICP-MS)16 (see Methods).

Figure 2b shows a comparison of our values of KD with previous datafrom experiments in carbon capsules. As can be seen, there is goodagreement between our results in carbon capsules and those of Lagos etal.1. Moving to MgO capsules has a dramatic effect on KD, however. KD

is displaced upwards by a factor of 15–20 when C is excluded from thesample. Furthermore, the data from C-free experiments are, as theyshould be, in excellent agreement with the one-atmosphere ther-modynamic data on the endmember exchange reaction (1). Finally,

the observed displacement between C-bearing and C-free experimentsis in excellent agreement with predictions based on the thermo-dynamic properties of the Fe–C–Pb system15. It is clear that Pb is muchmore siderophile than previously believed.

Figure 2b shows that the effect of increasing pressure from 1.5 GPato 24 GPa at constant temperature is extremely small. Adopting thetemperature T effect from the thermodynamic data and fitting forlogKD as a function of pressure P (ref. 17) yields (one standard errorshown in parentheses):

logK D~0:77(0:09){(2022=T )z18(18)(P=T )

This means that, assuming a single stage of core formation with tem-peratures on the silicate liquidus17, a pressure of ,30 GPa, similar tothat observed for other elements17,18, would be consistent with a DPb

(core–mantle partitioning) of 25. The effect of the small amount of Cestimated to be present in the core, about 0.2% (refs 2, 3), would be tolower the partition coefficient of Pb to 0.89 of its value15, whereasexperiments19 on the effect of S on Pb partitioning indicate that the1.7% S estimated to be in the core20 would increase DPb to 1.5 times itsvalue in pure Fe liquid. In neither case would the effect change ourconclusion that lead was siderophile during core formation.

Although the metal–silicate partitioning of Pb is consistent with asingle stage of high-pressure core formation, it has been pointed out21

that the 235/238U–207/206Pb age of the BSE for such a model (45–140 Myr)is inconsistent with that for the 182Hf–182W system (30 Myr) (ref. 22).Consistency between the two systems requires either differences betweenthe degree of equilibration of W and Pb achieved during accretion21 orlate losses of Pb to the core or space after 182Hf was extinct (.45 millionyears). Such further core growth would have occurred at the time of thegiant impact when the final ,10% of the Earth was accreted23.

To test the hypothesis of ‘late’ core growth we modelled thePb-isotopic composition of the BSE by assuming rapid early accre-tion according to the exponential growth model that reproduces the182Hf–182W system22. Material was added to the Earth with a238U/204Pb of 0.7 (ref. 4) and Pb was extracted continuously at equi-librium until the Earth was 90% accreted. The last 10% was added tothe Earth later and 10% of the core was extracted at the same time.

–2.5

–2

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1,600 1,800 2,000 2,200 2,400 2,600 2,800

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MgO capsules 24 GPa

Temperature (K)

Pb partitioning into core

Predicted effectof carbon

ref. 12

ref. 1

a

b

1 atm thermodynamic data

1 atm thermodynamic data

Figure 2 | Experimental data for Pb partitioning into liquid Fe coexisting withliquid silicate. KD is defined as [Pb]metal[Fe]silicate/[Pb]silicate[Fe]metal, wherevalues in brackets are weight concentrations. Error bars refer to 62 standarderrors. a, Data for PbO(silicate)1Fe(metal) 5 Pb(metal)1FeO(silicate) arefrom refs 1 and 12. All experiments were performed in carbon capsules and arecarbon-saturated, apart from the open symbol from ref. 12 (MgO capsule).Metal in the latter contained 9 wt% Si. b, Liquid metal–liquid silicatepartitioning data from this study. Experiments were performed in carbon andMgO capsules. One-atmosphere thermodynamic data for pure liquids32 arecorrected for the activity coefficients of Pb in liquid Fe15. The length of arrowindicates the calculated displacement due to carbon saturation of the Fe liquidat 1,860 K (ref. 15).

15.3

15.4

15.5

15.6

15.7

17.6 17.8 18 18.2 18.4 18.6

K89D84L91GG91DZ79AL89A88ZH88

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yr11

0 M

yr14

0 M

yr

Con

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ent w

ith H

f–W

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b/2

04P

b

206Pb/204Pb

Figure 3 | Calculated and estimated lead isotopic composition of the BSE.The line labelled ‘Geochron’ is the zero age line for fractionation of Pb fromU at the origin of the Solar System. The dotted line labelled ‘Consistent withHf–W’ represents the loci of points for which the timing of U–Pbfractionation would agree with Hf–W fractionation using the exponentialgrowth model22. The dashed lines labelled ‘80 Myr’, ‘110 Myr’ and ‘140 Myr’represent the loci of points calculated for the addition of 10% to the coreduring a giant impact at the corresponding time after the origin of the SolarSystem. The first 90% of accretion was assumed to obey the exponentialgrowth model22 under P,T oxygen fugacity conditions constrained bymantle–core partitioning of a number of siderophile elements18. Symbols arepublished estimates of the BSE composition16.

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Figure 3 shows the calculated Pb-isotopic composition of the BSE forthe addition of the final 10% of the core at different times corres-ponding to the estimated times of the giant impact. We note that theassumed 238U/204Pb or m of the total Earth is 0.7 (ref. 4) throughoutaccretion, but some have argued for a higher value. The effect of this isto increase the discrepancy between the apparent 235/238U–207/206Pband 182Hf–182W ages of the BSE21. The early stage of accretion (,90%)of the Earth was assumed to occur at pressures, temperatures andoxidation states corresponding to those derived from mantle–corepartitioning of Ni, Co, V, Cr, Nb, W and Si (ref. 18). The last 10%of partitioning (at the time of the giant impact) was assumed to takeplace on the silicate liquidus at a pressure of 42 GPa, constrained bythe current Ni and Fe contents of the BSE18. This last 10% of additionshifts the Pb-isotopic composition of the BSE into the region of theindependent estimates and is also consistent with the current Ni, Co,V, Cr, Nb, W and Si contents of the BSE18. A last phase of coreformation 80–140 million years after the origin of the Solar Systemcan clearly explain the observed Pb-isotopic composition of the BSE.

A final argument in favour of Pb depletion by core formation andnot volatile loss is given by the concordance of ages of the giantimpact and Moon formation discussed earlier. The 87Rb–87Sr iso-topic age9 is based on O, Si and W evidence8,24,25 that the lunaraccretion disk isotopically equilibrated with the BSE26. Silicate reser-voirs are expected to produce a range of W isotopic compositionsdepending on the history of accretion and core formation as well asthe particular partitioning of W. For this reason it is not surprisingthat silicate meteorites display considerable diversity in composi-tion27. It is therefore exceedingly unlikely that the BSE and theMoon would have identical W isotopic composition unless theyequilibrated. The only alternative is that the mixture of silicate andmetal from the Moon-forming impactor Theia, together with silicatein the proto-Earth, was identical to the mixture from which theMoon formed, which is inconsistent with dynamical simulations23.Given that tungsten is refractory28 except under very oxidizing con-ditions, Sr isotopes are likely also to have equilibrated with the BSE.On this basis the Sr isotopic composition of the newly formed Moonwas that of the BSE at the time of the giant impact.

This means that we can use the well-defined Rb/Sr of the BSE29 toestimate an age for the Moon, which is 90 6 20 Myr after the start of

the Solar System. This is based on the best data available for the initialSr isotopic composition of the Solar System, namely the data forEfremovka calcium aluminium inclusions (CAIs) and angrites,which are entirely self-consistent. Rubidium is moderately volatilewith a half-mass condensation temperature of 800 K, very similar tothat of Pb (727 K)28. Rb is lighter (its Z 5 37 versus Z 5 82 for Pb) andpossesses an extremely low first ionization potential (4.177 eV or403 kJ mol21 versus 7.417 eV for Pb), so Rb should be lost withXe (ref. 30) to space even more easily than Pb. Losses of Rb at thetime of the giant impact cannot, however, be accommodated withoutdisplacing the apparent time of the giant impact to earlier ages, asshown in Fig. 4. Shifting the Moon’s formation to significantly earlierages, however, leads to discrepancies with W isotopic data and allother indicators of lunar age. For example, Rb loss equivalent to thatrequired to change the U/Pb ratio of the BSE to its current valuewould require the Moon to have formed within the first few millionyears of the Solar System, which is inconsistent with W isotope andother data. Therefore, there is no way of accommodating major lossesof moderately volatile elements from the Earth during the giantimpact, leaving losses to the core as the only way of changing the238U/204Pb of the BSE.

METHODS SUMMARY

Experiments were performed at Macquarie and Oxford Universities in the piston-

cylinder apparatus (at 1.5–2 GPa) and at the Bayerisches Geoinstitut (Universitat

Bayreuth) in the multi-anvil apparatus (24 GPa). Starting materials were intimate

mixtures of silicates and metals with ,0.3% Pb added as PbO. Between 0.1% and

0.3% Cu (and in some cases Ni) as oxide was added to the experiments to provide

internal standards (in addition to Fe) for laser ICP-MS analysis of the quenched

metal phase. At low pressures (1.5–2 GPa) capsules consisted either of graphite or

polycrystalline MgO. To minimize the tendency for the MgO capsule to collapse

into the melt at very high pressure and temperature, the multi-anvil experiments

used capsules made of single-crystal MgO. These maintain their physical integrity

under all experimental conditions used in this study. Metals segregate into a single

ball in most experiments, whereas silicates quench to glass (graphite capsules) or a

mixture of quench crystals and glass (MgO capsules). Products were analysed by

electron microprobe for major elements (and Ni, Cu and Pb in the metal) and by

laser ICP-MS for trace elements using the procedures described in ref. 16

(Methods).

Full Methods and any associated references are available in the online version ofthe paper at www.nature.com/nature.

Received 29 July 2009; accepted 26 March 2010.

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11. Wasserburg, G. J., Papanastassiou, D. A., Tera, F. & Huneke, J. C. Outline of a lunarchronology. Phil. Trans. R. Soc. Lond., A 285, 7–22 (1977).

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120

100

80

60

40

20

00 20 40 60 80 100

Mod

el a

ge o

f Moo

n(M

yr a

fter

sta

rt o

f Sol

ar S

yste

m)

Loss of Rb during giant impact (%)

Based on Efremovka CAIs

Based on SAH99555 angrite

Figure 4 | Rb–Sr model age of the Moon plotted versus the percentage lossof Rb from the Earth at the time of the giant impact. Sr-isotopicequilibration between Earth and Moon was assumed at time of impact. Tosatisfy the Sr-isotopic composition of the Moon using the current Rb/Sr ofthe Earth (0.03), the timing of the Moon-forming impact needs to becomeearlier as the extent of Rb loss from the Earth becomes greater. The two linesare calculated with different values for the initial isotopic composition of theSolar System33. Because a lunar age of 70–110 million years is consistent withW isotopic data for the Moon8 and the age of the earliest crust10, the loss ofRb and similarly volatile elements such as Pb from the Earth during the giantimpact must be small.

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13. Allegre, C. J., Poirier, J.-P., Humler, E. & Hofmann, A. W. The chemicalcomposition of the Earth. Earth Planet. Sci. Lett. 134, 515–526 (1995).

14. The Japan Society for the Promotion of Science and The 19th Committee onSteelmaking. Steelmaking Data Sourcebook Part II, 273–312 (Gordon and Breach,1988).

15. Malhotra, S., Chang, L. & Schlesinger, M. E. Liquid solution thermodynamics in thealumina-saturated Fe-C-Pb system. ISIJ Int. 38, 1–8 (1998).

16. Wood, B. J., Nielsen, S. G., Rehkamper, M. & Halliday, A. N. The effects of coreformation on the Pb- and Tl-isotopic composition of the silicate Earth. Earth Planet.Sci. Lett. 269, 326–336 (2008).

17. Wade, J. & Wood, B. J. Core formation and the oxidation state of the Earth. EarthPlanet. Sci. Lett. 236, 78–95 (2005).

18. Wood, B. J., Wade, J. & Kilburn, M. R. Core formation and the oxidation state of theEarth: additional constraints from Nb, V and Cr partitioning. Geochim. Cosmochim.Acta 72, 1415–1426 (2008).

19. Chabot, N. L., Saslow, S. A., McDonough, W. F. & Jones, J. H. An investigation ofthe behavior of Cu and Cr during iron meteorite crystallisation. Meteorit. Planet.Sci. 44, 505–520 (2009).

20. Dreibus, G. & Palme, H. Cosmochemical constraints on the sulfur content in theEarth’s core. Geochim. Cosmochim. Acta 60, 1125–1130 (1996).

21. Halliday, A. N. Mixing, volatile loss and compositional change during impact-driven accretion of the Earth. Nature 427, 505–509 (2004).

22. Yin, Q. Z. et al. A short timescale for terrestrial planet formation from Hf–Wchronometry of meteorites. Nature 418, 949–952 (2002).

23. Canup, R. M. & Asphaug, E. Origin of the Moon in a giant impact near the end ofthe Earth’s formation. Nature 412, 708–712 (2001).

24. Wiechert, U. et al. Oxygen isotopes and the Moon-forming giant impact. Science294, 345–348 (2001).

25. Georg, R. B., Halliday, A. N., Schauble, E. & Reynolds, B. C. Silicon in the Earth’score. Nature 447, 1102–1106 (2007).

26. Pahlevan, K. & Stevenson, D. J. Equilibration in the aftermath of the lunar-forminggiant impact. Earth Planet. Sci. Lett. 262, 438–449 (2007).

27. Kleine, T. et al. Hf-W chronometry and the accretion and early evolution ofasteroids and terrestrial planets. Geochim. Cosmochim. Acta (in the press).

28. Lodders, K. Solar system abundances and condensation temperatures of theelements. Astrophys. J. 591, 1220–1247 (2003).

29. McDonough, W. F. et al. Potassium, rubidium and cesium in the Earth and Moonand the evolution of the mantle of the Earth. Geochim. Cosmochim. Acta 53,1001–1012 (1992).

30. Zahnle, K. et al. Emergence of a habitable planet. Space Sci. Rev. 129, 35–78 (2007).31. Palme, H. & O’Neill, H. S. C. in The Mantle and Core (ed. Carlson, R. W.) Vol. 2, 1–38

(Elsevier, Amsterdam, 2003).32. Barin, I., Sauert, F., Schultze-Rhonhof, E. & Sheng, W. S. Thermochemical Data of

Pure Substances Parts I and II (CH Verlagsgesellschaft, 1989).33. Nyquist, L. E., Shih, C. Y., Wiesmann, H. & Mikouchi, T. Fossil 26Al and 53Mn in

d’Orbigny and Sahara 99555 and the timescale for angrite magmatism. LunarPlanet. Sci. Conf. XXXIV, 1388 (2004).

Supplementary Information is linked to the online version of the paper atwww.nature.com/nature.

Acknowledgements The assistance of N. Pearson (Macquarie), N. Charnley(Oxford) and J. Day (Cambridge) with microprobe and laser ICP-MS analysis isacknowledged with thanks. B.J.W. acknowledges the support of the AustralianResearch Council through Federation Fellowship FF 0456999 and the NERC (UK)through grant NE/F018266/1. A.N.H. acknowledges support from STFC.Experiments at the Bayerisches Geoinstitut were performed under the EU‘Research Infrastructures: Transnational Access’ Programme (contract number505320; RITA—High Pressure).

Author Contributions B.J.W. performed all the experiments and all the electronmicroprobe and laser ICP-MS analyses, and the Pb-isotopic modelling of Fig. 3.A.N.H. performed the Sr-isotope modelling depicted in Fig. 4. Both authorscontributed to the writing of the manuscript.

Author Information Reprints and permissions information is available atwww.nature.com/reprints. The authors declare no competing financial interests.Readers are welcome to comment on the online version of this article atwww.nature.com/nature. Correspondence and requests for materials should beaddressed to B.J.W. (berniew@earth.ox.ac.uk).

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METHODSStarting materials for the high-pressure experiments consisted of intimate mix-

tures of silicates and iron. The silicate part of the mix was made up of varying

proportions of pre-synthesized CaAl2Si2O8, CaMgSi2O6 and Mg2SiO4, to which

was added Fe0.947O and (in two cases) NaAlSi3O8. Pb and Cu were added as

oxides in concentrations of ,0.3%, together with 0.1–0.3% of up to eight more

trace elements from the following list: V, Cr, Ni, Co, Zn, Ga, Zr, Nb, Mo, Ag, Cd,

In, Sb, Tl and Bi.

For piston-cylinder experiments at Macquarie and Oxford Universities the

starting mixtures were placed in unsealed capsules made either of graphite or of

polycrystalline MgO and inserted in a 12.7-mm-diameter pressure cell made of

concentric cylinders of (from the inside) graphite, SiO2 glass and BaCO3. The

interior of the graphite heater contained the capsule, which was supported by

rods of machinable MgO above and below. The W3Re97/W25Re75 thermocouple

was inserted into the cell and pushed into contact with an alumina disk resting on

top of the capsule. Experiments were performed at 1.5–2 GPa and 1,803–2,223 Kfor periods of between 3 and 120 minutes.

Experiments in the multi-anvil apparatus at the Bayerisches Geoinstitut were

performed in single-crystal MgO capsules using the same starting materials as

those employed in the piston-cylinder apparatus. The assembly used a 10-mm-

edge-length MgO octahedron with a LaCrO3 furnace surrounded by a sleeve of

zirconia. A W3Re97/W25Re75 thermocouple in contact with the capsule moni-

tored temperature. Tungsten carbide cubes with 4-mm-edge-length octahedral

truncations were used to generate pressures of 24 GPa. Because of the extreme

conditions (up to 2,673 K) experiment durations were limited to 2 minutes.

Experimental charges were sectioned and polished using 0.3-mm alumina

powder. The products invariably showed good separation of metal from silicate.

Although the silicate liquids from the experiments performed in graphite cap-

sules quenched to homogeneous glasses, experiments performed in MgO cap-

sules generated non-quenchable liquids. The silicate products in the latter cases

were mixtures of glass and quench crystals. This heterogeneity resulted in a much

wider scatter of major element compositions than was found in graphite capsule

experiments (Supplementary Table 2).

The metallic parts of the products were homogeneous with respect to iron and

moderately siderophile elements such as Ni and Cu. The element of principalinterest, Pb, however came out of metal solution on the quench and formed small

(micrometre-sized) droplets, homogeneously distributed in the metal phase and

observable in some back-scattered electron images. In cases where Tl and/or Bi

was present, these elements appeared to exsolve into the same droplets as the Pb.

Because of the presence of these droplets we used a defocused (10mm) beam for

electron microprobe analysis and employed a second technique capable of aver-

aging the composition of a large volume of material, laser-ablation ICP-MS. The

electron microprobe analysis, which generally had a much higher standard devi-

ation (Supplementary Table 2), was then used as a check on the laser-ICP-MS

data.

Products were analysed using the Cameca SX100 microprobe in the

Geochemical Analysis Unit at Macquarie University (with 15 kV accelerating

voltage) and the JEOL 8800 microprobe (with 20 kV accelerating voltage) at

Oxford University. A range of silicate, oxide and metal standards was used

and count times and beam currents were adjusted for the concentration range

of interest. Typically, major elements were determined using a 20-nA beam

current and a 10–30 s peak count time, while minor elements were analysed

using 40–80 nA beam current and 60 s count time on the peak. Background

counts were collected for half the time of the peak measurements. Results are

given in Supplementary Table 2.

Trace element analyses were obtained using a laser ablation ICP-MS system.

The system principally used was a Merchantek LUV266 laser microprobe con-

nected to an Agilent 7,500 s ICP-MS. Laser operating conditions included a repe-

tition rate of 5 Hz and an output power of ,0.1 mJ per pulse with a spot size

(depending on sample size) of 60–120mm. The ablation was carried out in a

mixture of He (0.25 to 0.3 litres per min) 1 Ar (1.1 to 1.15 litres per min). A

gas background was measured for approximately 60 s before firing the laser and

typical ablation times were 70–100 s. Raw counts were collected on the ICP-MS in

peak-hopping mode (dwell time 30–50 ms) and displayed in time-resolved for-

mat. This allowed each ablation to be monitored to identify heterogeneities such

as small metal inclusions in the silicate and compositional variations with depth.

Further details of this methodology are given in ref. 34. Several of the products

were analysed in the Department of Earth Sciences, Cambridge University, using a

New Wave Research UP213 Nd:YAG laser ablation instrument interfaced to a

Perkin Elmer Elan DRCII ICP-MS. Procedures were similar to those described

above and repeat analysis of four samples demonstrated excellent agreement

between the results from the two instruments.

Data reduction in all cases used the GLITTER (http://www.glitter-gemoc.

com) software package35. Quantification of trace element concentrations by laser

ablation ICP-MS requires an external calibration standard and the concentration

of an internal standard that has been obtained by an independent method. In this

study NIST 610 glass was used for the external standardization and calibrations

were checked using USGS glass BCR-2 before unknowns were measured. Iron

and silicon concentrations (determined by electron microprobe) were used as

the internal standards for metal and silicate respectively. Because the external

and internal standards for metal have vastly different Fe concentration levels

(450 p.p.m. and 90–100% respectively) it was appropriate to use an independent

check of the method. This was provided by electron microprobe analysis of Cu

(and in some cases Ni) in the metal. These two elements were found to be

homogeneously distributed in the quenched metals and are hence appropriate

secondary standards. In almost all cases agreement (Supplementary Table 2)

between microprobe and laser ablation ICP-MS analyses was well within un-

certainty. A final check on the quality of the analysis for Pb in the metal by laser

ablation is given by the doping ratio Pb/Cu in the starting material. This was

0.8–0.87 in all experiments. For those experiments in which Pb behaves in a

siderophile manner we know that, because Cu is siderophile, the Pb/Cu ratio of

the metal must be close to the doping ratio. As can be seen in Supplementary

Table 2, this requirement is fulfilled in the relevant cases.

34. Norman, M. D., Pearson, N. J., Sharma, A. & Griffin, W. L. Quantitative analysis oftrace elements in geological materials by laser ablation ICPMS: instrumentaloperating conditions and calibration values of NIST glasses. Geostand. Newsl. 20,247–261 (1996).

35. Van Achterbergh, E., Ryan, C. G., Jackson, S. E. & Griffin, W. L. in Laser AblationICPMS in the Earth Sciences (ed. Sylvester, P.) Vol. 29, 239–243 (MineralogicalAssociation of Canada, 2001).

doi:10.1038/nature09072

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