A Mercury-like component of early Earth yields uranium in the core and high mantle (142)Nd.

Recent (142)Nd isotope data indicate that the silicate Earth (its crust plus the mantle) has a samarium to neodymium elemental ratio (Sm/Nd) that is greater than that of the supposed chondritic building blocks of the planet. This elevated Sm/Nd has been ascribed either to a 'hidden' reservoir in the Earth or to loss of an early-formed terrestrial crust by impact ablation. Since removal of crust by ablation would also remove the heat-producing elements--potassium, uranium and thorium--such removal would make it extremely difficult to balance terrestrial heat production with the observed heat flow. In the 'hidden' reservoir alternative, a complementary low-Sm/Nd layer is usually considered to reside unobserved in the silicate lower mantle. We have previously shown, however, that the core is a likely reservoir for some lithophile elements such as niobium. We therefore address the question of whether core formation could have fractionated Nd from Sm and also acted as a sink for heat-producing elements. We show here that addition of a reduced Mercury-like body (or, alternatively, an enstatite-chondrite-like body) rich in sulfur to the early Earth would generate a superchondritic Sm/Nd in the mantle and an (142)Nd/(144)Nd anomaly of approximately +14 parts per million relative to chondrite. In addition, the sulfur-rich core would partition uranium strongly and thorium slightly, supplying a substantial part of the 'missing' heat source for the geodynamo.

Recent 142Nd isotope data indicate that the silicate Earth has an Sm/Nd ratio greater than the supposed chondritic building blocks of the planet. This elevated Sm/Nd has been ascribed either to a “hidden” reservoir in the Earth[1,2] or to loss of an early-formed terrestrial crust by impact ablation[3]. Since removal of crust by ablation would also remove the heat producing elements, K, U and Th, the difficulty of balancing terrestrial heat production with heat flow would be severe[3]. The alternative, a “hidden” reservoir is generally assigned to the silicate lower mantle. We have previously shown, however, that the core is a likely reservoir for some lithophile elements such as Nb[4]. We therefore addressed the question of whether core formation could have fractionated Nd from Sm and also acted as a sink for heat producing elements. We show here that addition of a reduced Mercury (or enstatite-chondrite)-like body rich in sulfur to the early Earth would generate a superchondritic Sm/Nd in the mantle and 142Nd/144Nd anomaly of ~14ppm relative to chondrite. Additionally, the S-rich core would partition U strongly and Th slightly, providing a substantial part of the “missing” heat source for the geodynamo.

Terrestrial rocks were recently found to have higher ratios of radiogenic 142Nd to nonradiogenic 144Nd than the chondritic meteorites generally supposed to be representative of the material from which Earth accreted[1,2]. 142Nd was produced during the early history of the solar system from decay of the extinct radionuclide 146Sm (t1/2=68 Ma[5]) and the presence of a positive 142Nd anomaly of ~20 ppm or ~9 ppm[6] in silicate Earth would require an Sm/Nd ratio higher than chondritic[1,2]. This high Sm/Nd ratio was established early in Earth history while 146Sm was “alive”. A plausible mechanism for generating high Sm/Nd in Earth’s mantle is partial melting and melt extraction to form a crust. Because Nd is less compatible in mantle silicates than Sm[7] partial melts have relatively low Sm/Nd and the solid residue high Sm/Nd. A low Sm/Nd crust could be completely removed from the mantle system by subduction to an inaccessible region of the deep mantle[1] or removed from Earth by impact ablation[3]. The problem with the former hypothesis is the lack of evidence for a hidden silicate reservoir, while the latter suffers from the requirement that much of Earth’s heat production, in the form of radioactive U, Th and K would be removed together with the low Sm/Nd crust. Assuming chondritic abundances of U and Th and a K/U ratio of ~12000 for silicate Earth[8] the heat production in the Earth is only about 0.6 times current heat loss[9]. Reducing the heat sources further by ablation loss would make it even more difficult to reconcile production with heat loss. An additional question in the context of heat production is the energy source for the Earth’s magnetic field[10]. Arising from convection in the core, Earth has had a magnetic field for at least 3.5 Ga. The crystallisation of the inner core is an important source of energy for the geodynamo[11] but most attempts to construct histories of core cooling indicate that the inner core cannot be much older than 1-1.5Ga[10,11] unless a source of radioactive heating is present. Numerous studies have focussed on 40K as a potential core heat source since K, in common with all moderately volatile elements[8] is depleted in silicate Earth relative to chondritic abundance. Furthermore, high pressure experiments[12,13] indicate that K enters sulfide under oxidising conditions and S is believed to be a major component of the core’s complement of ~10% of elements of low atomic number[14]. It appears, however that the maximum possible K content of the core is insufficient to generate more than a small fraction of the 2-5 TW required to generate reasonable core thermal histories[11,13]. The alternative explanation, that U and/or Th provide the energy for core convection has some support from early experiments on sulfide-silicate partitioning[15] but more recent results indicate very little partitioning of U into S-bearing metals even under extreme conditions[16].

We approached the problem of U, Th, Nd and Sm in Earth’s mantle and core from the standpoint of recent work on partitioning between sulfide melts and silicate melts[17]. Kiseeva and Wood (2013) found that the sulfide-silicate partition coefficient for any element i, is dependent on the FeO content of the silicate melt such that for FeS-rich sulfides: where A is a constant and n is a constant dependent on the valency of element i. Therefore, under strongly reducing conditions, where the FeO content of the silicate melt is very low (<1% for example) one would expect Di values to be much higher than under the conditions of MORB crystallisation where the FeO content of the melt is about 8-10%. This hypothesis is consistent with the data of Murrell and Burnett[15] who observed strong partitioning of U into sulfide liquid at low fO2. Given terrestrial accretion models calling for prolonged periods of growth under reduced conditions[18,19], the demonstration that Mercury is a highly-reduced sulfur-rich planet with a liquid core[20,21] and the association of rare earth elements, U and Th with sulphides in enstatite chondrite meteorites[22], we investigated partitioning of U, Th, Sm, Nd and several other lithophile elements into liquid iron sulfide under reducing conditions.

Experiments were performed at 1.5 GPa and temperatures between 1400 and 1650°C using starting materials which were approximately 50:50 mixtures of silicate and FeS doped with a range of lithophile trace elements including U, Th, La, Nd, Sm, Eu,Yb, Ce, and Zr (see Methods). The silicate was a basalt-like composition in the system CaO-MgO-Al2O3-SiO2 with variable FeO. Analysis was by electron microprobe and Laser Ablation ICP-MS (Methods). Table 1 presents a summary of sulfide-silicate partitioning results (see Extended Data Section for complete analyses).

Table1

Summary of sulfide/silicate partition coefficients

Run	P(GPa)	T(°C)	LogFeO(sil)wt(%)	DSm	DNd/DSm	DU/DSm	DTh/DU	DEu/DSm	DLa/DSm	DYb/DSm
421	1.5	1400	0.50	0.005	1.42	2.47	0.036	5.85	1.38	0.16
σ				0.001	0.29	0.37	0.008	0.93	0.23	0.07
428	1.5	1400	0.08	0.013	1.35	1.56	0.038	5.50	1.35	0.16
σ				0.001	0.19	0.17	0.005	1.06	0.20	0.02
427	1.5	1400	−0.25	0.062	1.30	1.81	0.028	2.36	1.23	0.13
σ				0.006	0.19	0.24	0.004	0.40	0.22	0.02
426	1.5	1400	−0.30	2.247	1.25	6.81	0.048	0.14	1.03	0.16
σ				0.333	0.30	1.43	0.010	0.05	0.15	0.09
429	1.5	1400	1.21	0.005	1.04	3.68	0.200	2.04	1.10	0.67
σ				0.0001	0.12	0.89	0.041	0.42	0.47	0.19
461	1.5	1650	0.30	0.023	1.22	1.92	0.058	4.09	1.18	0.21
σ				0.003	0.20	0.32	0.009	0.58	0.18	0.03
462	1.5	1650	−0.21	0.154	1.12	3.58	0.046	1.13	0.92	0.27
σ				0.011	0.13	0.43	0.007	0.14	0.10	0.03
477	1.5	1650	−0.29	0.629	1.10	9.26	0.044	0.37	0.83	0.39
σ				0.038	0.11	0.73	0.043	0.38	0.80	0.39
464	1.5	1650	−0.32	0.751	1.13	9.41	0.035	0.21	0.55	0.28
σ				0.073	0.16	1.18	0.020	0.16	0.44	0.16
1414	1.5	1500	−0.21	0.048	1.31	1.84	0.031	5.95	1.41	0.13
σ				0.004	0.14	0.34	0.009	0.80	0.16	0.02
1415	1.5	1500	−0.39	0.454	1.18	6.99	0.040	0.88	1.04	0.21
σ				0.028	0.13	0.66	0.005	0.11	0.14	0.05
1416	1.5	1500	0.88	0.006	1.39	2.70	0.067	4.92	1.52	0.23
σ				0.0005	0.18	0.25	0.007	0.56	0.19	0.02
1417	1.5	1500	1.06	0.007	1.20	3.11	0.150	2.71	1.19	0.50
σ				0.001	0.24	0.52	0.042	0.48	0.21	0.30

Partition coefficients in weight ratio. σ = calculated from error propagation. n = number of measurements.

Fig 1a shows data from a series of experiments performed at 1400°C. As can be seen, the partition coefficients of U, Nd and Sm are strong functions of the FeO content of the silicate melt, increasing dramatically, as predicted, as the FeO content decreases below 1wt%. The negative slope of logDi as f(log[FeO]sil) reverses at high FeOsil, however, because the sulfide dissolves progressively more oxygen as the FeO content of silicate increases and these lithophile elements follow oxygen into the sulfide. We found similar behaviour in two more series of experiments at higher temperature (Table 1). Other lithophile elements, notably Ti, Nb and Ta (BJW,unpublished data) behave similarly. Importantly we find DU>DNd>DSm for partitioning into sulfide in all experiments. At very low FeO contents all Di become >1 (Fig 1a). Furthermore (Fig 1b) DNd is always significantly greater than DSm with DNd/DSm approaching 1.5 in some cases.

Fig 1

Sulfide-silicate partitioning data

(a) Partition coefficients for U, Nd and Sm at 1.5 GPa/1400°C plotted versus the log of the FeO content of the silicate melt in wt%. (b) The ratio of DNd/DSm plotted versus logFeO. Error bars in both cases are ±2 std errors and, if absent, are smaller than symbol size.

The implications of Fig 1 are that segregation of sulfide (or S-rich metal) from reduced FeO-poor silicate will lead to enrichment of the metallic phase in U and in Nd relative to Sm when compared to the silicate. Addition of such material to the core and mantle respectively of a growing planet would provide a core heat source and a mantle with superchondritic Sm/Nd, Yb/Sm and Yb/La. Although potentially detectable in terms of a mantle 142Nd anomaly, the fractionation of heavy from light REE (Table 1) in the primitive mantle of the body (bulk silicate earth in this case) would have little effect on its overall REE pattern (Extended Figure 3). Similarly, there would be no significant Eu anomaly despite the fact that Eu is probably in the 2+ oxidation state (unlike the other 3+ lanthanides) under these conditions (Extended Figure 3). If such a body represented Earth early in its history then the mantle would have a positive 142Nd anomaly relative to chondrite (as observed) and much of the energy deficit identified for core convection[10] would be supplied by U (and Th). We find that DTh/DU is about 0.1, indicating that U would be accompanied by Th in the S-rich core. Addition of more oxidised material later in accretion would lead to the higher current FeO content of the mantle (8.1%)[8], but could not erase the superchondritic Sm/Nd ratio of the mantle and U content of the core unless there were complete core-mantle re-equilibration.

Extended Fig 3

Rare earth element fractionation at 3.2% and 8.1% sulfur in the core

This figure shows the calculated rare earth element pattern in Bulk Silicate Earth (BSE) for the two extreme cases of Fig 1a (3.2% S) and Extended Fig 2d (8.1% S). Black diamonds represent REE concentrations relative to chondritic abundances and normalized to Yb =1, at 3.2% sulfur in the core (20% reduced mass impactor and 0.15 mass fraction sulfide). White diamonds illustrate the REE fractionation at elevated sulfur content (8.1% sulfur in the core, 35% reduced mass impactor and 0.22 mass fraction sulfide). We assumed DSm(sulfide/silicate) = 1 and Di/DSm ratios for other elements from experiment 464 Both scenarios result in very small depletions of light REE relative to heavy REE in BSE. The trend is broadly consistent with that seen in the depleted MORB-mantle composition (blue diamonds) but much smaller. The effect on the REE pattern of BSE would, as can be seen, be undetectable. Blue diamonds illustrate the measured ratio of depleted MORB mantle (Salters and Stracke, 2004) to bulk silicate earth (Palme and O’Neill, 2003), assuming chondritic abundances of refractory lithophile elements in the latter. Error bars are from propagated error calculation and correspond to 1 standard deviation.

Figure 2 illustrates the impact of adding a highly reduced body rich in sulfide to the growing Earth. The Th/U ratio of silicate Earth would be higher than chondritic (3.8-3.9[8,23]) which provides an important constraint on how much U can be present in Earth’s core. Based on the Pb-isotopic compositions of Archean galenas[24] and of 3.5 Ga komatiites[25] the Th/U ratio of the Archean mantle has been estimated to be ≥ 4.3. Tatsumoto[26] argued, based on the Pb isotopic compositions of basalts, for an early differentiation of the mantle which resulted in a Th/U of 4.2-4.5 in the mantle source regions. Since that time the Th/U ratio of the mantle has decreased probably due to preferential recycling of the more soluble uranium[27].

Fig 2

Core content of U (ppb) and mantle 142Nd anomaly (ppm)

(a) Calculated effect of adding to Earth 20% ME of reduced body containing 0.15 mass fraction sulfide. The sulfide is added to the core and the silicate to the mantle. Sulfide-silicate DU/DSm fixed at 2, DNd/DSm at 1.4 and DTh/DU at 0.1 (Table 1). (b) Same as (a) except mass of reduced body is 45% of Earth mass. (c) and (d) as for (a) and (b) except reduced body contains 0.22 mass fraction sulfide. The reduced body and remainder of Earth each contain 14ppb U and 53.5 ppb Th consistent with chondritic abundances. Sulfide extraction was assumed to take place at the origin of the solar system.

Figure 2 shows 4 models of U content of the core and 142Nd anomaly of the mantle (relative to bulk Earth) based on our partitioning data. We choose a reduced body of 0.15 mass fraction sulfide corresponding to the S content of primitive CI chondrites[8] and use values of DSm (Smsulf/Smsil) close to the observed maximum of 0.8-2.2 noting that DSm values increase with decreasing temperature and that segregation of sulfide from a crystal-melt mush instead of melt alone would increase them further because of the incompatibility of Sm, Nd, U in crystals. As can be seen (Fig 2a), adding 20% of such a body to Earth would lead to 4-5 ppb of U in the core, Th/U of silicate Earth of 4.17 and a 142Nd anomaly in the mantle relative to bulk Earth of ~7ppm. Increasing the reduced body mass to 45% (Fig 2b) leads to about 8 ppb U in the core, Th/U of silicate Earth of 4.5 and mantle 142Nd anomaly of 13.9 ppm relative to bulk Earth. We have performed a sensitivity analysis (Extended Figure 2) and find that, if Th/U of silicate Earth is ≤4.5, the maximum U content of the core is 8 ppb with a Th content of ~8 ppb. These figures increase to ~10 ppb if Th/U of silicate Earth is ≤4.7. The 142Nd anomaly is 13.9 ppm in the former case and ~17 ppm in the latter. The estimated U and Th contents of the core would lead to 2-2.4 TW, sufficient to power the geodynamo[11] even without the potential 0.4-0.8TW from 40K decay[13]. We can reduce the size of the reduced body by increasing its sulphur content (Fig 2c,d) and increase DU/DSm but the overall effects on the core and mantle 142Nd remain close to those summarised above if Th/U of BSE is constrained to be ≤4.5 or ≤4.7. Note that the scenarios shown in Fig 2 refer to a terrestrial core containing between 3.2 and 8.1 wt% S. The concentration of cosmochemically abundant volatile S in the core is unknown, but recent suggestions range from a cosmochemical estimate of 1.7 wt%[14] to ~6 wt%[28] from liquid metal density measurements and 14.7 wt%[29] from high temperature-high pressure equation of state measurements. The range shown in Fig 2 is, therefore appropriate for the current state of knowledge.

Extended Fig 2

Core content of U (ppb) and mantle 142Nd anomaly (ppm) at DU/DSm = 3

Figure illustrates the effect on the Nd and U content using the same parameters (DNd/DSm at 1.4 and DTh/DU at 0.1) as in figure 2 but with a higher DU/DSm ratio. (a), (b) show the calculated effect of adding to Earth 20% ME or 45% ME of reduced body containing 0.15 mass fraction sulphide. (c) and (d) illustrate the same scenario except the reduced body contains 0.22 mass fraction sulfide.

We conclude that a period of growth of the accreting Earth under reduced, S-rich conditions would generate a significant (~14ppm) 142Nd anomaly in silicate Earth, in agreement with observations and add sufficient U and Th to the core to generate 2-2.4 TW required to drive the geodynamo.

Methods

Experimental Methods

Starting materials for high pressure experiments consisted of mixtures of ~50wt% (Fe,Ni)S and ~50% of a synthetic silicate approximating the 1.5 GPa eutectic composition in the Anorthite-Diopside-Forsterite system[30]. The sulfide component was analytical grade FeS doped with 1-3% NiS. Trace elements were added as a mix consisting of Zr, La, Ce, Nd, Sm, Eu, Yb, Th and U as oxides. After adding the trace element mix such that each element was present at 1000-2000 ppm, the silicate and sulfide starting materials were mixed in 50:50 proportions and ground under acetone for 20 minutes before being dried at 110°C prior to the experiment. Starting compositions were loaded into 3 mm O.D. and 1mm I.D. graphite capsules.

Experiments were conducted in a ½-inch diameter piston-cylinder apparatus using external cylinders either of BaCO3-Silica glass (at 1500 and 1650 °C) or CaF2 (at 1400°C) and an 8 mm O.D. graphite furnace with 1 mm wall. The unsealed capsule was separated from the graphite furnace by an interior MgO sleeve, with a 0.5 mm thick alumina disk on top to prevent puncture by the thermocouple. Temperatures were controlled and monitored using a tungsten-rhenium thermocouple (W5%Re/W26%Re), and temperature maintained within ± 1°C. Experimental conditions were 1400°C, 1500°C and 1650°C at 1.5 GPa and with experiment durations between 1 and 4.5 hrs. These times are sufficient to approach equilibrium in small graphite capsules[17]. Experiments were quenched by turning off the power supply. After quench, the capsule was extracted from the furnace, mounted in Acrylic and polished for further analyses with electron microprobe and Laser Ablation ICPMS. All experimental charges contained sulfide blebs embedded in a silicate glass matrix.

Microanalysis

Samples were analyzed on the JEOL 8600 electron microprobe in the Archaeology Department at the University of Oxford. WDS analyses of major element compositions of silicate glasses and sulfides were performed at 15 kV with a beam current of 20 nA and a 10μm defocused beam (Extended Data Table 1,2). At least 20 analyzes were taken of the silicate and sulfide in each experiment. Count times for major elements (Si, Al, Ca, Mg, Fe in silicate, Fe in sulfide) were 30 seconds on the peak and 15 seconds background. Minor elements (S, Ni, O) were analyzed for 60 seconds peak and 30 seconds background. We have previously noted Ni loss from similar experiments[17] and the principal reason for adding Ni was to provide an additional check on Laser Ablation ICP-MS analyses of the trace elements of interest (see below). A range of natural and synthetic standards was used for calibration. Standards for silicate were wollastonite (Si, Ca), jadeite (Al), periclase (Mg) and Hematite (Fe). Standards for sulfides were Ni metal (Ni), galena (S) and Hematite (Fe, O). Oxygen in the sulfides was determined using the Kα peak and LDE crystal.

Extended Table 1

Major element composition of silicate glass

P = 1.5	GPa, T = 1400°C
Run	421		428		427		426		429
n	25	σ	20	σ	23	σ	19	σ	31	σ
SiO2	56.19	0.28	57.69	0.25	59.63	0.80	57.93	0.01	48.14	0.26
Al2O3	20.50	0.14	20.40	0.11	20.21	0.11	19.60	0.09	18.17	0.16
CaO	11.48	0.08	13.02	0.09	10.97	0.07	10.48	0.16	9.39	0.08
MgO	6.66	0.06	6.53	0.06	6.50	0.11	6.22	0.08	5.61	0.05
FeO	3.15	0.07	1.21	0.04	0.56	0.03	0.50	0.23	16.36	0.13
S	0.11	0.01	0.18	0.01	0.34	0.02	6.11	0.24	0.24	0.02
Σ trace	0.81	0.02	0.90	0.02	0.78	0.04	0.30	0.03	0.65	0.05
Total	98.90		99.93		98.99		101.14		98.56

Values in wt%; σ = calculated from error propagation; n = number of measurements; Σ trace= Sum of trace elements measured with LA-ICPMS

Extended Table 2

Major element composition of sulfides

P = 1.5	GPa, T = 1400°C
Run	421		428		427		426		429
n	21	σ	27	σ	23	σ	25	σ	23	σ
O	0.92	0.78	0.36	0.05	0.21	0.04	0.20	0.15	3.06	0.21
S	36.85	1.26	36.00	0.39	36.86	0.33	35.63	1.46	33.24	0.25
Fe	60.97	0.37	62.15	0.52	62.34	0.37	62.51	1.08	62.52	0.18
Ni	0.23	0.43	0.35	0.06	0.29	0.05	0.36	0.10	0.32	0.02
Sm	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.
U	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.	n.m.
Σ trace	0.005	0.0005	0.008	0.001	0.040	0.004	0.556	0.139	0.004	0.002
Total	98.98		98.77		99.74		99.26		99.14

Values in wt%; σ = calculated from error propagation; n = numbers of measurements; n.m. = not measured

We determined Uranium and Samarium contents of 3 product sulfides as a further check on the LA-ICP-MS analyses. In this case we measured the Mα peak for U and Lα peak for Sm using standards of UO2 and SmPO4 respectively and a PET crystal. Operating conditions were 15 kV, 40 nA, and a 10μm beam. Count time for Uranium was 120 seconds on peak and 60 seconds background. Samarium was analyzed for 150 seconds on the peak and 75 seconds background.

Trace elements in silicates and sulfides were measured by Laser Ablation ICPMS employing a NexION 300 quadropole mass spectrometer coupled to a New Wave Research UP213 Nd:YAG laser at the University of Oxford. A laser repetition rate of 10Hz and 25-50 μm spot size were used for silicate glasses and sulphides (Extended Data Table 3,4) with an energy density of ~12 J.cm−2. Operating in time-resolved mode, we employed 20 seconds of background acquisition, followed by ablation for 60 seconds. Between analyses we employed a 60-90 second “wash-out” time. The following masses were counted: 24Mg, 27Al, 29Si, 57Fe, 60Ni, 43Ca, 90Zr, 139La, 140Ce, 142Nd, 152Sm, 153Eu, 174Yb, 232Th, 239U. Our external standard was NIST610 glass and we typically collected 3 spectra of this at the beginning and end of each sequence of 10-15 unknowns. BCR-2G standard was used as a secondary standard to check the accuracy of the calibration. Ablation yields were corrected by referencing to the known concentrations of Si and Ca (silicate glass) and Fe (sulfides), which had been determined by microprobe. Data reduction was performed off-line using the Glitter 4.4.3 software package, which enabled us to identify occasional sulfide inclusions in the silicate analyses. Since the Fe content of the NIST610 standard is only 460 ppm, the background is high and the matrices are very different, cross-checks on the sulfide analyses were required. Therefore Ni was measured with the electron microprobe and LA-ICPMS. In agreement with Kiseeva and Wood[17], we observed no systematic offset between electron microprobe and LA-ICP-MS analyses for Ni (Extended data Tables 2,4). Additionally, as discussed above, the U and Sm contents of the sulfides were measured by electron microprobe in three experimental charges (1415, 464, 477). 20-43 electron probe analyses were performed on each sample. The highest U and Sm concentrations were measured in experiment 1415 with Laser ICPMS (U=2958 ppm, Sm=719 ppm). Comparative measurements with electron probe yielded values of U=3280±490 ppm and Sm=707±110 ppm, (uncertainty=2 standard errors) and therefore show excellent agreement. Two samples with lower U and Sm concentrations were also analysed. ICPMS measurements for 464 yielded U=952ppm and Sm=327ppm, while sample 477 gave U=927 ppm and Sm=300ppm. Electron microprobe concentrations of U=1164±224ppm and Sm=319±87 ppm (464) and U=991±69ppm and Sm=277±38 ppm (477) are also in excellent agreement with the LA-ICP-MS measurements. We conclude that our LA-ICP-MS results have no detectable systematic offset due to matrix effects or calibration errors.

Extended Table 3

Trace element concentration in silicates

P = 1.5	GPa, T = 1400°C
Run	421		428		427		426		429
n	5	σ	5	σ	5	σ	5	σ	5	σ
Zr90	1226	29	1153	43	1171	49	206	55	956	72
La139	641	21	574	8	607	31	247	25	511	39
Ce140	4.0	0.1	4	0.1	3.7	0.4	2	0	3.4	0.4
Nd142	494	15	450	6	474	27	167	25	397	30
Sm152	796	25	728	9	775	39	304	30	645	50
Eu153	709	13	670	11	636	37	452	24	578	47
Yb174	1002	31	903	6	1009	42	590	25	799	57
Th232	1030	21	967	14	984	51	465	37	813	57
U238	961	21	929	22	893	43	117	4	782	52
Ni60	n.m.	n.m.	n.m.	n.m.	34	12	108	3	7	1

values in ppm; n = number of measurements, n.m. = not measured.

Extended Table 4

Trace element concentration in sulfides

P = 1.5	GPa, T = 1400°C
Run	421		428		427		426		429
n	5	σ	5	σ	5	σ	5	σ	5	σ
Zr90	0.5	0.0	1.2	0.2	6	1	1228	58	2.2	1.0
La139	4.1	0.4	10	1	47	7	573	12	2.8	1.7
Ce140	0.1	0.1	0.1	0.03	0.4	0.1	4	0	0.0	0.0
Nd142	3.2	0.5	8	1	38	3	472	16	2.1	1.2
Sm152	3.6	0.5	9	1	48	4	683	13	3.2	2.1
Eu153	19	2	47	8	94	11	140	1	5.9	2.5
Yb174	0.7	0.3	1.8	0.1	8	1	218	12	2.7	2.9
Th232	0.4	0.1	0.7	0.1	3.1	0.3	339	6	3.0	2.8
U238	11	1	19	1	101	6	1792	94	14	3
Ni60	2837	64	3596	138	3320	88	3764	159	3153	51

values in ppm; n = number of measurements.

Partition coefficients for U, Nd and Sm with changing log FeO content in silicate melt (wt%)

(a) represents results for D values of experiments performed 1.5 GPa and 1500°C. (b) D value results at 1.5GPa and 1650°C .