Tungsten stable isotope compositions of terrestrial ...

CHEMGE-18194; No of Pages 10 Chemical Geology xxx (2016) xxx–xxx

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Tungsten stable isotope compositions of terrestrial samples and meteorites determined by double spike MC-ICPMS Nadine Krabbe ⁎, Thomas S. Kruijer, Thorsten Kleine Institut für Planetologie, University of Münster, Wilhelm-Klemm Straße 10, 48149 Münster, Germany

a r t i c l e

i n f o

Article history: Received 12 October 2016 Received in revised form 13 December 2016 Accepted 17 December 2016 Available online xxxx Keywords: Tungsten Double spike Stable isotope fractionation Meteorites

a b s t r a c t Tungsten stable isotopes hold great potential to examine a variety of physical and chemical processes operating during the accretion and differentiation of asteroids and terrestrial planets, such as core formation and mantle– crust differentiation. To assess the magnitude and origin of W isotope fractionations, we determined the W stable isotopic compositions of six USGS geological reference materials, a NIST steel, and 24 iron meteorites and chondrites. The W isotope data were obtained using multi-collector inductively coupled plasma mass spectrometry (MC-ICPMS) and using a 180W–183W double spike. Chondrites and iron meteorites exhibit a very narrow range in W stable isotope compositions, resulting in a mean δ184/183W = 0.027 ± 0.007‰ (95% conf.) relative to the NIST 3163 W standard. This value represents a good estimate for the W stable isotope composition of bulk planetary bodies from the inner solar system. The δ184/183W of some iron meteorites slightly deviates from this value, most likely due to W isotope fractionations induced during crystallization of the metal cores of iron meteorite parent bodies. The investigated terrestrial silicate rocks exhibit a narrow range in δ184/183W, which for most samples is indistinguishable from the mean value of chondrites and iron meteorites. However, felsic samples tend to be isotopically lighter than mafic samples, indicating that magmatic processes on Earth induced W isotope fractionations. These fractionations are possibly related to the fluid-mobility of W in subduction zones, but more data are needed to test this hypothesis. Given that most terrestrial igneous rocks are isotopically indistinguishable from chondrites and iron meteorites, core formation on Earth does not seem to have induced a measurable isotopic fractionation for W. However, more data are needed to firmly arrive at this conclusion. © 2016 The Authors. Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).

1. Introduction The investigation of non-traditional stable isotopes in terrestrial and extraterrestrial materials has emerged as a new tool for investigating a variety of high-temperature processes, such as core formation, magmatic differentiation and crust formation (e.g., Georg et al., 2007; Savage et al., 2011; Armytage et al., 2011; Shahar et al., 2011; Wang et al., 2012; Hin et al., 2013; Burkhardt et al., 2014; Bezard et al., 2016; Bonnand et al., 2016; Millet et al., 2016). While mass-dependent isotope variations have been investigated for several elements, W so far has received little attention. Instead, W isotope studies primarily focused on applications of the short-lived 182Hf–182W chronometer (t1/2 = 8.9 Myr) to assess the timescales of planetary accretion and differentiation (e.g., Kleine et al., 2009). Nevertheless, W is a promising element to investigate for mass-dependent isotope variations, because it is strongly fractionated among Earth's major geochemical reservoirs. As a refractory siderophile element, W is enriched in the metal phase and thus predominantly occurs in the cores of differentiated planetary bodies. During magmatic ⁎ Corresponding author. E-mail address: [email protected] (N. Krabbe).

processes on Earth, W is one of the most incompatible elements and is therefore strongly enriched in Earth's continental crust (Newsom et al., 1996; Shearer and Righter, 2003; Arevalo and McDonough, 2008). Moreover, W is fluid-mobile in subduction zones and is selectively enriched in the sub-arc mantle by fluids derived from the down-going slab (König et al., 2008; Bali et al., 2012). Thus, crust formation and subduction zone processes could potentially be associated with massdependent W isotope fractionations. Finally, in silicate systems W adopts the 4+ or 6+ oxidation states (e.g., O'Neill et al., 2008; Wade et al., 2013), and so the partitioning of W between metal and silicates during core formation may have induced isotope fractionation, akin to that found for Mo (Hin et al., 2013). Utilizing mass-dependent W isotope fractionations as a geochemical tracer requires techniques for the precise and accurate measurement of W stable isotope compositions. In addition, to assess the usefulness of W stable isotope variations as a geochemical tracer, a reconnaissance study of mass-dependent W isotope variations among terrestrial and extraterrestrial materials is needed. However, until now only a few studies examined the extent of W isotope fractionations among natural samples (Irisawa and Hirata, 2006; Breton and Quitté, 2014; Abraham et al., 2015). Of these studies, Breton and Quitté (2014) investigated the

http://dx.doi.org/10.1016/j.chemgeo.2016.12.024 0009-2541/© 2016 The Authors. Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).

Please cite this article as: Krabbe, N., et al., Tungsten stable isotope compositions of terrestrial samples and meteorites determined by double spike MC-ICPMS, Chem. Geol. (2016), http://dx.doi.org/10.1016/j.chemgeo.2016.12.024

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N. Krabbe et al. / Chemical Geology xxx (2016) xxx–xxx

largest sample set, which also included data for two chondrites and one iron meteorite. In general, these authors found relatively large W isotope variations spanning a total range of ~0.4‰ for both terrestrial samples and meteorites. However, two other studies reported much smaller W isotope variations for terrestrial samples, but also investigated a smaller sample set (Irisawa and Hirata, 2006; Abraham et al., 2015). The reasons for these disparate results remained unclear, and so the extent of mass-dependent W isotope variations among terrestrial rocks and meteorites is not well constrained. To address these issues, we have developed a new analytical technique for the precise and accurate measurement of W stable isotope compositions by double spike MCICPMS, and applied this method to several geological reference materials as well as a comprehensive set of iron meteorites and chondrites. 2. Analytical methods 2.1. Double spike design and calibration Application of the double spike technique requires measuring the abundances of four isotopes from a given element, two of which are enhanced through the addition of the double spike. By measuring the isotopic composition of the double spike–sample mixture, and using the known double spike composition, the true mass-dependent isotope signature of a sample corrected for instrumental mass fractionation can be calculated by algebraic inversion. The precision achievable by this technique is predominantly governed by the choice of the double spike composition, and the proportion into which the double spike and the sample are mixed (e.g., Dodson, 1970; Hofmann, 1971; Galer, 1999; Rudge et al., 2009). Here we used the Double Spike Toolbox (Rudge et al., 2009) to assess the optimal double spike composition for W. This approach shows that using an 180W–183W double spike and 184W and 186W for the double spike inversion provides a low intrinsic uncertainty for the W stable isotope analyses. We purposely excluded 182W as a potential spike or normalizing isotope, because in meteorites the 182W abundance varies due to the decay of now-extinct 182Hf. Thus, for meteorites the independent measurement of an unspiked sample aliquot would have been necessary to determine mass-dependent W isotope variations. In contrast, using the 180W–183W double spike this information can be obtained from a single measurement. A disadvantage of the 180 W–183W double spike is the low natural abundance of 180W (~0.12%) combined with the high abundance of 180Hf (~35%), requiring the efficient chemical separation of Hf from W prior to isotope measurement (see Section 2.3). Two single spikes of 180W and 183W were purchased from the Oak Ridge National Laboratory (ORNL) and were delivered as 10 mg metal powders. These powders were weighed separately into 15 ml Savillex® vials and completely dissolved in a 1:2 mixture of concentrated HFHNO3 on a hotplate at 100 °C for 1 day. The fully dissolved spikes were then diluted using 2 M HNO3–0.5 M HF to solutions containing ~ 10 ppm W. Before mixing the single spikes together, their isotope compositions were determined by MC-ICPMS (for details see Section 2.4). For this, the single spikes were diluted to ~ 100 ppb W and were measured multiple times relative to a certified solution standard (NIST SRM 3163 W). The optimal double spike composition, as given by the Double spike toolbox, was obtained for a 7:3 mixture of the 180W and 183 W spikes, and the lowest uncertainty is obtained if this double spike is mixed with the sample in equal proportions (1:1). The two single 180W and 183W spikes were mixed together in the double spike proportion inferred above, diluted to a 1 l solution containing ~180 ppb W (in 2 M HNO3–0.5 M HF), and then calibrated against the NIST 3163 W solution standard. The isotope composition of the final double spike is given in Table 1. To assess the accuracy of the double spike calibration, and the sensitivity of our method to over- and under-spiking, we analyzed the NIST 3163 standard as well as AGV-2 admixed with varying amounts of double spike. While the NIST 3163-double spike mixtures were directly

Table 1 Isotope compositions of 180W–183W spike and NIST 3163 standard reference material. 180

W double spike 95% conf. NIST W 3163 95% conf.

W/183W

0.18599 0.00001 0.00831 0.00001

182

W/183W

0.83279 0.00003 1.85173 0.00003

184

W/183W

0.59830 0.00002 2.14110 0.00004

186

W/183W

0.35524 0.00002 1.98625 0.00004

measured, the spiked AGV-2 samples were first processed through the full chemical separation procedure (see below). The δ184/183W values (where δ184/183W is the permil deviation from NIST 3163; see below) obtained for all the different mixtures are indistinguishable (Fig. 1), demonstrating that the calibration of the double spike is accurate (note that the W stable isotope composition of AGV-2 is indistinguishable from that of NIST 3163). These data also show that the 180 W–183W double spike provides accurate results over a relatively wide range of spike-to-sample ratios (Fig. 1). 2.2. Samples Six geological reference materials from the United States Geological Survey (USGS) spanning a range of SiO2 contents were investigated for this work, including two basalts (BCR-2 and BHVO-2), an andesite (AGV-2), a granodiorite (GSP-2), a rhyolite (RGM-2), and a shale (SBC-1). We also analyzed an industrially produced high sulfur steel (NIST129C). In addition, we measured the W stable isotope compositions of 24 meteorites, including 4 carbonaceous chondrites (CB, CM, CV), 5 ordinary chondrites (L5, L6, H5), 2 enstatite chondrites (EL6) and 13 iron meteorites from different chemical groups (IC, IIAB, IID, IIIAB, IVA, IVB). The iron meteorite samples were sawn using a diamond saw, polished with abrasives (SiC) and ultrasonically cleaned in ethanol to remove any saw marks and adhering dust. Chondrite samples were cleaned with ethanol in an ultrasonic bath and then crushed to fine grained powders using an agate mortar. The two investigated Allende samples are aliquots of the MS-A and MS-B powders, which were prepared from ~100 and ~40 g slices of Allende in an agate mill. 2.3. Sample preparation and chemical separation of W The samples (~ 0.05–0.5 g) and an appropriate amount of double spike were weighed together into Savillex® vials. Spiked silicate samples were digested according to sample mass in 6–24 ml concentrated

Fig. 1. δ184/183W results for reference materials with variable amounts of admixed double spike. Shown are results for the NIST 3163 W solution standard (blue diamonds) and for AGV-2 (red squares). The shaded area shows the external reproducibility of ±0.03‰ (2 s.d.) obtained for various reference materials. Error bars represent the internal precision (2 s.e.) of each individual measurement obtained from within-run statistics. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)



HF–HNO3 (2:1) at ~180 °C for ~5 days on a hotplate, while the iron meteorite samples, Gujba and NIST129C were digested in 4–16 ml of concentrated HNO3–HCl–HF (2:1:0.5) at 130 °C for at least 24 h. To promote efficient digestion of the carbonaceous chondrite samples, several drops of concentrated HClO4 were added. Following digestion, samples were evaporated to dryness and then converted by adding concentrated HNO3 and HCl several times. The separation of W involves a three-stage anion exchange chromatography following methods from earlier studies (Kleine et al., 2012; Kruijer et al., 2013, 2014a). The first chromatography step serves to separate W from most matrix elements, and uses 4 ml anion exchange resin (BioRad® AG1X8, 200–400 mesh) filled in columns made out of transfer pipettes (~ 0.5 cm in diameter and 30 cm length). Samples weighing ~ 0.15–0.2 g were loaded onto the columns in 25 ml of 0.5 M HCl–0.5 M HF, while larger samples (~ 0.2–0.5 g) were loaded in 75 ml 0.5 M HCl–0.5 M HF. Following a 10 ml 0.5 M HCl–0.5 M HF rinse, W, Ti and other high field strength elements (HFSE) were eluted in 15 ml of 6 M HCl–1 M HF. The second chromatography step separates W from any elements remaining after the first column, especially Ti and other HFSE such as Ta, Nb, Hf, and Zr. For the second-stage chemistry we used pre-cleaned BioRad® columns filled with 1 ml anion resin (BioRad® AG1X8, 200–400 mesh). After loading samples onto the column in 6 ml 0.6 M HF–0.4% H2O2, Hf, Zr, and Ti were eluted using 10 ml 1 M HCl–2% H2O2 and 9 ml 8 M HCl–0.01 M HF. Tungsten was finally recovered in 8.5 ml 6 M HCl–1 M HF. This second separation step was conducted twice to facilitate efficient removal of Hf from the W cuts. This is important because of the isobaric interference of 180Hf on 180W, especially given the high natural abundance of 180Hf (35.1%) and the low abundance of 180W (0.12%). The final W cuts typically had 178Hf/184W b 3 × 10−5, and only a few samples had higher 178Hf/184W (up to ~7 × 10−4). The magnitude and effects of the Hf-interference on 180W is discussed in detail below (Section 2.4). Osmium and Ta potentially also form isobaric interferences on masses 180, 184 and 186, but these two elements are almost quantitatively removed by the chemical separation, resulting in 188 Os/184W b 4 × 10−5 and 181Ta/180W b 4 × 10−4 in the analyzed samples. For these ratios, Os and Ta interferences on W masses are inconsequential for the W stable isotope measurements. After each chromatography step, the W cuts were repeatedly evaporated to dryness in HClO4 to eliminate organics, converted by adding concentrated HNO3 several times, and finally dissolved in ~ 0.5 ml 0.56 M HNO3–0.24 M HF measurement solutions. The W yields of the chemical procedure were typically 60–90%. Total procedural blanks were 100–150 pg W and inconsequential for our results given the large amounts of W analyzed (~100 ng W per sample). 2.4. Isotope measurements and data reduction The W isotope measurements were performed using the ThermoScientific® Neptune Plus MC-ICPMS in the Institut für Planetologie at the University of Münster. The samples were introduced into the mass spectrometer using an ESI® self-aspirating PFA nebulizer connected to a Cetac® Aridus II desolvating system. The W isotope measurements were performed in low resolution mode at total ion beam intensities of approximately 1–2 × 10− 10 A (using 1011 Ω amplifiers), which were obtained for ~100 ppb W solutions at an uptake of ~60 μl/ min and using standard Ni ‘H’ cones. Each measurement consisted of an on-peak baseline measurement for 90 s using the same 0.56 M HNO3–0.24 M HF solution that is used for the dilution of the samples and standards, followed by the W isotope measurement comprised of 100 cycles of 4.2 s integration time each. Ion beams at 180W, 182W, 183 W, 184W and 186W were simultaneously collected in static mode using Faraday cups connected to amplifiers with 1011 Ω feedback resistors. In addition, ion beams at 178Hf, 181Ta and 188Os were monitored to account for isobaric interferences from Hf and Ta (on mass 180) as well as Os (on masses 184 and 186). For the Hf and Ta interference monitors Faraday cups connected to amplifiers with 1012 Ω resistors were used,

3

whereas Os was measured using cups connected to 1011 Ω resistors (Table 2). The data reduction for all measurements was performed off-line using the Double Spike Toolbox from Rudge et al. (2009). Using the measured ion beams (178Hf, 180W, 183W, 184W, 186W) and the W double spike composition as input, this software iteratively solves for three unknowns using MATLAB's non-linear equation solving routine: 1) the natural fractionation factor α, 2) the instrumental mass fractionation factor β (assuming that the mass fractionation in the mass spectrometer follows the exponential law), and 3) the molar proportion of double spike p within the double spike–sample mixture. Using this output, the W concentration in a sample was calculated from the molar proportion of spike, and the mass-dependent W isotope fractionation was obtained from the natural fractionation factor α. Consistent with previous studies, we observed that the α values measured for the reference standard (NIST 3163 W) showed small variations between different analytical sessions (Siebert et al., 2001; Schoenberg et al., 2008; Hin et al., 2013; Burkhardt et al., 2014). Such variations are attributed to instrumental drift between analytical sessions that does not accurately follow the exponential law for mass fractionation, and which results in slight variations in absolute measured isotope ratios between analytical sessions. To correct for such non-exponential drifts, the measured α of the samples were normalized to the α obtained for the reference standard (NIST 3163) within the same analytical session (Hin et al., 2013): δ184=183 W ¼ −1000 α Sample −α NIST3163 ln ðm184 =m183 Þ

ð1Þ

where m184 and m183 are the atomic weights of 184W and 183W. Here δ184/183W is defined as the permil deviation of the 184W/183W ratio from the composition of NIST 3163: " 184 δ184=183 W ¼

W=183 W sample

#

ð184 W=183 WÞNIST3163

−1 1000

ð2Þ

To facilitate a reliable Hf-interference correction, we implemented a script into the Double Spike Toolbox that runs the double spike deconvolution repeatedly—each iteration step employs the β factors obtained from the previous step—until the output α and β values converge to constant values. To assess the accuracy of the Hf interference correction, we performed tests using Hf-doped double spike–standard (NIST 3163 W) mixtures having variable 178Hf/184W ratios of between ~9 × 10−4 and ~2 × 10−2 (Fig. 2a). The results of this test demonstrate that Hf interferences can be accurately corrected for solutions having 178 Hf/184W b 1 × 10−3 (Fig. 2b), corresponding to an interference correction of ~4 δ-units. Note that for all the W stable isotope data of natural samples reported here, the Hf amounts in the analyzed sample solutions were significantly smaller (Table 3; Fig. 2b), resulting in typical interference corrections of ~ 0.1 δ-units for most samples. Only for some samples interference corrections as high as 3.4 δ-units were necessary, but even these corrections are smaller than those that can be Table 2 Cup configuration, amplifiers, natural isotope abundances and interferences. Faraday cup Amplifier (Ω) Monitored mass Interfering isotopes Natural abundances (%) W Hf Ta Os

L4 10

L3 12

178

Hf

27.8

10

L2 11

180

W

180

Hf, 180Ta

0.12 35.08 0.01

10

L1 12

181

Ta

10

C 11

182

W

H1 11

10

183

W

10

11

H2 10

11

184

W

186

W

184

Os

186

Os

H3 1011 188 Os

26.50 14.31 30.64 28.43 99.99 0.02

1.59

13.24


4


a

Table 3 Tungsten stable isotope data for terrestrial samples. Sample

ID

Na

W (ppb)

GSP-2

65 66 69 101 102

4 4 3 3 3 17

347.0 362.2 347.5 385.9 351.2 358.7

61 62

4 4 8

1598 1601 1600

48 89 90 91

3 3 3 3 12

483.9 473.5 479.3 484.3 480.3

35 47

2 3 5

221.7 220.0 220.9

70 71 92

4 4 3 11

490.1 493.2 498.6 494.0

73 74 75 76

3 3 3 3 12

1460 1454 1458 1450 1455

63 64 53

3 3 2 8

722.7 716.2 743.0 727.3

Mean (±2 s.d.) (±95% conf.) SBC-1 Mean (±2 s.d.) (±95% conf.) BCR-2

b

Mean (±2 s.d.) (±95% conf.) BHVO-2 Mean (±2 s.d.) (±95% conf.) AGV-2

Mean (±2 s.d.) (±95% conf.) RGM-2

Mean (±2 s.d.) (±95% conf.) NIST 129C

Fig. 2. δ184/183W results for Hf-doped double spike–standard (NIST 3163) mixtures uncorrected for Hf interference (open symbols, white) and corrected for Hf-interference on 180W (closed symbols, blue). Also shown are δ184/183W results of the three samples with the highest 178Hf/184W remaining in the analyzed W cut (closed symbols, green). Uncertainties on individual data points represent the 2 s.e. of individual sample runs and are smaller than the symbol size in most cases. The shaded area denotes the external reproducibility of ± 0.03‰ (2 s.d.) obtained for most geological reference materials using our method. Fig. 2b is a magnification of Fig. 2a and shows that the Hf interference correction is accurate for 178Hf/184W b 0.001. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

accommodated. Thus, our double spike method can accurately account for typical Hf amounts remaining after the chemical separation. Note that for all samples the analyzed W cuts were essentially free of Ta and Os, and so no interference corrections were necessary for these elements.

3. Results The accuracy and precision of the W isotope measurements were evaluated by analyses of the spiked NIST 3163 standard solution that was processed through the full chemical separation procedure, and by repeated measurements of geological reference materials. The analyses of the processed NIST 3163 yielded a mean δ184/183W of + 0.002 ± 0.020‰ (2 s.d., n = 12), demonstrating that the chemical separation does not induce any significant isotope fractionation that would not be corrected by the double spike. Repeated measurements of different digestions of several USGS geological reference materials (AGV-2, BCR2, BHVO-2, GSP-2, RGM-2, SBC-1) provide an external reproducibility (2 s.d.) between ±0.012 and ±0.040‰ on δ184/183W. This difference in reproducibility probably reflects sample heterogeneities, as is also apparent from scatter in measured W concentrations (see below). For most samples, however, a reproducibility of about ± 0.030‰ (2 s.d.)

Mean (±2 s.d.) (±95% conf.)

δ184/183W ±2σb 0.008±0.027 0.036±0.020 0.011±0.030 0.021±0.030 0.046±0.030 0.024±0.040 ±0.010 0.031±0.009 0.030±0.013 0.030±0.012 ±0.005 0.042±0.030 0.040±0.030 0.046±0.030 0.027±0.030 0.039±0.023 ±0.007 0.037±0.030 0.040±0.030 0.039±0.039 ±0.024 −0.001±0.021 −0.020±0.033 0.011±0.030 −0.005±0.039 ±0.013 −0.009±0.030 0.001±0.030 −0.001±0.030 0.000±0.030 −0.002±0.031 ±0.010 0.178±0.030 0.185±0.030 0.195±0.030 0.185±0.029 ±0.012

178

Hf/184Wc

pd

0.00045 0.00034 0.00050 0.00005 0.00006

0.515 0.504 0.514 0.481 0.505

0.00002 0.00002

0.505 0.505

0.00016 0.00006 0.00004 0.00031

0.489 0.517 0.514 0.512

0.00038 0.00021

0.527 0.540

0.00044 0.00025 0.00007

0.507 0.508 0.509

0.00027 0.00049 0.00051 0.00049

0.583 0.584 0.583 0.585

0.00003 0.00001 0.00002

0.495 0.499 0.489

a

N = number of measurements. For solution replicates with n b 4 the external reproducibility of 0.030‰ obtained for the geological reference materials is given; for n ≥ 4 the 95% conf. interval from replicate measurements is given. c Ratio in the analyzed W cut. d Molar proportion of spike. b

was obtained. This or a better reproducibility was also obtained for most meteorites that were analyzed multiple times (Table 4); only for Allende and two iron meteorites (Arispe, Henbury) more scatter is observed, which again most likely is attributable to sample heterogeneities. Taken together, a reproducibility of about ± 0.030‰ (2 s.d.) provides a good estimate of the uncertainty for a single analysis of a rock sample. Note that most samples investigated in the present study were analyzed multiple times, so that for these samples δ184/183W is known more precisely. In general, all samples investigated in the present study, including terrestrial igneous rocks as well as chondrites and iron meteorites, have very similar W stable isotope compositions (Figs. 3, 4). The largest fractionation is observed for the high sulfur steel NIST 129C with δ184/183W = 0.185 ± 0.012‰ (95% conf., n = 8) (Table 3; Fig. 3); its heavy W isotopic composition was most likely induced during the production of this steel. The geological reference materials exhibit a narrow range in δ184/183W from −0.005 ± 0.013 to +0.039 ± 0.007‰ (Figs. 3, 4; Table 3), where felsic rocks (AGV-2, RGM-2) tend to be isotopically lighter compared to mafic rocks (BCR-2, BHVO-2). The chondrites analyzed in this study, including carbonaceous, ordinary and enstatite chondrites, exhibit indistinguishable δ184/183W from ~+0.02 to ~+0.05‰, with a mean δ184/183W = +0.031 ± 0.021‰ (95% conf., n = 11) (Table 4), very similar to the values observed for the USGS reference materials (Fig. 4). The iron meteorites also exhibit a narrow range in δ184/183W from ~0.00 to ~+0.07‰ (Table 4), similar to the values obtained for the chondrites, but with slightly



5

Table 4 Tungsten stable isotope data for meteorites. Sample

Classification

ID

Na

W (ppb)

Chondrites Gujba Jbilet Winselwan Allende MS-A

CB CM2 CV3

52 133 43 103 105 113 123

Mean (±2 s.d.) (±95% conf.) Allende MS-B

1 1 2 1 1 2 3 9

586.0 130.5 167.5 166.8 167.6 182.8 177.5 172.5

CV3

126 127

Mean (±2 s.d.) ALHA 81021 LAP 1030 Forest City

EL6 EL6 H5

130 129 110 111

Mean (±2 s.d.) (±95% conf.) Gao-Guenie

1 2 3 1 1 2 2 4

164.5 171.5 168.0 130.7 139.2 189.0 166.5 177.8

H5

45 106 107

Mean (±2 s.d.) (±95% conf.) Ergheo Kunashak

2 2 2 6

182.8 168.8 164.2 171.9

L5 L6

135 108 109

L6

131

1 2 1 3 1

122.2 114.0 101.9 107.9 109.2

Iron meteorites Arispe

IC

80 135

Mean (±2 s.d.) (±95% conf.) Bendego Ainsworth Negrillos

4 4 8

2115 2088 2101

IC IIAB IIAB

81 117 84 85

Mean (±2 s.d.) (±95% conf.) Sikhote Alin

3 2 4 3 7

1398 606.8 3638 3640 3639

IIAB

51 87

Mean (±2 s.d.) (±95% conf.) Carbo Rodeo Boxhole Henbury

3 2 5

733.3 742.5 737.9

IID IID IIIAB IIIAB

121 83 27 86 28 188 189

Mean (±2 s.d.) (±95% conf.) Gibeon

2 5 2 2 2 2 4 10

2833 2395 1254 1252 1260 1258 1211 1245

IVA

77 78

Mean (±2 s.d.) (±95% conf.) Muonionalusta

4 4 8

549.5 549.6 549.5

IVA

26 79

1 3 4

402.7 417.1 409.9

IVB IVB

118 122

3 4

2940 2944

Mean (±2 s.d.) Leedey Mean of chondrites (±2 s.d.) (±95% conf., n = 11)

Mean (±2 s.d.) (±95% conf.) Hoba Tawallah Valley Mean iron meteorites (±2 s.d.) (±95% conf., n = 12) Mean chondrites + iron meteorites (±2 s.d.) (±95% conf., n = 23)

δ184/183W±2σb 0.035±0.030 0.028±0.030 0.024±0.030 0.063±0.030 0.077±0.030 0.032±0.030 0.018±0.030 0.036±0.045 ±0.017 0.028±0.030 0.041±0.030 0.034±0.030 0.054±0.030 0.025±0.030 0.015±0.030 0.016±0.030 0.016±0.014 ±0.011 0.014±0.030 0.025±0.030 0.038±0.030 0.026±0.027 ±0.014 0.038±0.030 0.032±0.030 0.044±0.030 0.038±0.030 0.022±0.030 0.031±0.021 ±0.007 −0.016±0.017 0.041±0.005 0.012±0.063 ±0.026 0.042±0.030 0.065±0.030 0.004±0.007 0.000±0.030 0.003±0.007 ±0.003 0.042±0.030 0.024±0.030 0.033±0.023 ±0.015 0.001±0.030 0.025±0.009 0.009±0.030 −0.011±0.030 0.072±0.030 0.004±0.030 0.052±0.003 0.031±0.067 ±0.024 0.010±0.002 0.020±0.009 0.015±0.013 ±0.005 0.017±0.030 0.004±0.030 0.007±0.018 ±0.014 0.057±0.030 0.040±0.006 0.023±0.036 ±0.011 0.027±0.030 ±0.007

178

Hf/184Wc

pd

0.00001 0.00003 0.00006 0.00019 0.00009 0.00004 0.00001

0.511 0.516 0.504 0.507 0.507 0.502 0.509

0.00004 0.00015

0.502 0.490

0.00003 0.00009 0.00003 0.00002

0.506 0.490 0.518 0.549

0.00013 0.00002 0.00003

0.525 0.527 0.533

0.00002 0.00003 0.00007

0.603 0.618 0.644

0.00005

0.630

0.00008 0.00001

0.537 0.514

0.00008 0.00002 0.00005 0.00006

0.567 0.519 0.513 0.513

0.00003 0.00011

0.536 0.540

0.00003 0.00007 0.00036 0.00016 0.00021 0.00002 0.00000

0.506 0.498 0.488 0.532 0.534 0.505 0.512

0.00010 0.00011

0.529 0.530

0.00067 0.00019

0.483 0.476

0.00003 0.00001

0.531 0.533

a

N = number of measurements. For solution replicates with n b 4 the external reproducibility of 0.030‰ obtained for the geological reference materials is given; for n ≥ 4 the 95% conf. interval from replicate measurements is given. c Ratio in the analyzed W cut. d Molar proportion of spike. b


6


seem to be smaller than ~0.1‰ per amu. This result contrasts with those of a previous study, which found widespread variations among geological reference materials and meteorites, ranging from ca. -0.05 up to ca. +0.36‰ for δ184/183W (Breton and Quitté, 2014). In particular, Breton and Quitté (2014) reported heavy W stable isotope compositions for AGV-2 (δ184/183W = 0.36 ± 0.01‰) and Allende (δ184/183W = 0.32 ± 0.03‰), as well as for Gibeon (δ184/183W = 0.16 ± 0.02‰) and BCR-1 (δ184/183W = 0.14 ± 0.01‰). In contrast, for the same set of samples, we obtained a much smaller range in δ184/183W values, which also are much closer to the composition of the NIST 3163 standard (Fig. 5). Of note, Breton and Quitté (2014) consistently observed heavier isotopic compositions compared to those determined in the present study, indicating that the different results of both studies cannot be due to sample heterogeneities. Moreover, given the very narrow range of δ184/183W observed in the present study, it is highly unlikely that small-scale heterogeneities, if they exist, were preferentially and selectively sampled in a single study. The major difference between the present study and that of Breton and Quitté (2014) is the analytical approach used for determining W

Fig. 3. δ184/183W results for multiple measurements of geological reference materials analyzed in this study. The shaded areas show the 2 s.d. obtained for replicate measurements of each sample and error bars on data points denote the internal precision (2 s.e.) obtained from individual sample measurements. Different digestions of each individual standard are shown with different colors. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

more scatter (Fig. 4). Nevertheless, the mean δ184/183W = 0.023 ± 0.036‰ (95% conf., n = 12) of the iron meteorites (excluding Ainsworth; see Section 4.2) is very similar to that of the chondrites (Table 4). The W concentrations obtained from the double spike measurements typically reproduce to better than ±2% (2 s.d.), with the exception of GSP-2, Muonionalusta and the two Allende powders. The poorer reproducibility for these samples most likely reflects sample heterogeneities. Note that for GSP-2 and Allende we also obtained a poorer reproducibility for the δ184/183W measurements, suggesting that this poorer precision also results from sample heterogeneity. 4. Discussion 4.1. Comparison to previous studies Our data reveal only small W stable isotope fractionations among chondrites, iron meteorites and terrestrial igneous rocks, which in general

Fig. 4. δ184/183W results for geological reference materials, iron meteorites, and chondrites from this study. Error bars denote the external reproducibility of sample measurements (2 s.d. or 95% conf.; see Table 4). Also shown are the 2 s.d. (hashed area) as well as the 95% confidence limits (shaded area) of the mean obtained for the meteorite samples. Note that Ainsworth was excluded from calculating mean W isotope compositions, because this sample probably exhibits a heavy W stable isotope composition resulting from fractional crystallization of the IIAB core (see text and Fig. 7 for details).



7

4.2. Chondrites and iron meteorites and the W stable isotope composition of bulk planets The uniform δ184/183W values observed for the investigated chondrites, which comprise samples of petrologic types 2 to 6, suggests that parent body processes, such as aqueous alteration or thermal metamorphism, did not generate significant δ184/183W variations. Moreover, the investigated sample suite includes chondrites formed under more reducing (e.g., enstatite chondrites) as well as more oxidizing conditions (e.g., CM chondrites), reflecting distinct formation processes and conditions for these chondrites in the solar nebula (Krot et al., 2014). Consequently, the uniform δ184/183W of these samples demonstrates that the distinct conditions during accretion of chondrite parent bodies did not induce resolvable mass-dependent W isotope variations. Note that the uncertainty on the average δ184/183W of the investigated chondrites of ±0.021‰ (2 s.d.) is similar to the external reproducibility of the W stable isotope measurements of ±0.030‰ (2 s.d.). Thus, any potential W isotope fractionation among different groups of chondrites is likely to be smaller than the external precision of our measurements,

a

Fig. 5. δ184/183W results of this study compared to previous results (Irisawa and Hirata, 2006; Breton and Quitté, 2014; Abraham et al., 2015). The shaded area represents the total range in δ184/183W obtained for different meteorites (diamonds) and geological reference materials (round symbols) in the present study. Error bars on datapoints from previous studies represent 2σ uncertainties obtained from replicate analyses.

stable isotope compositions. While we employed the double spike technique, Breton and Quitté (2014) utilized the conventional standardsample bracketing technique. The advantage of the double spike technique over standard-sample bracketing is that it is less prone to analytical artifacts induced due to different sample matrices and to mass-dependent W isotope fractionations induced during sample preparation. For instance, the recovery of W during ion exchange chromatography often is not quantitative and we found that especially during re-dissolution of W after dry downs in Savillex beakers, substantial W losses can occur (e.g., Kruijer et al., 2012). Such W losses may be associated with W isotope fractionation and would then lead to a heavier isotopic composition in the analyzed sample. If in this case a standardsample bracketing technique were employed, the measured isotopic composition would be heavier than the true isotopic composition of the sample. In contrast, the double spike technique fully accounts for such isotope fractionations induced during sample preparation, as long as they follow the exponential law. Thus, W isotope fractionation induced during sample preparation is a potential cause for the discrepant results observed between our study and that of Breton and Quitté (2014). Nevertheless, the exact cause of this discrepancy is difficult to assess, but we note that the narrow range in δ184/183W we determined for a wide range and large set of terrestrial rocks and meteorites makes it very unlikely that large W stable isotope variations exist among these samples. This is consistent with results from two other W stable isotope studies, which, although they investigated a much smaller sample set than the present study, also did not find large variations in W stable isotope compositions in terrestrial samples (Irisawa and Hirata, 2006; Abraham et al., 2015) (Fig. 5).

b

Fig. 6. Potential effects of cosmic rays and nucleosynthetic isotope anomalies on δ184/183W of iron meteorites. (a) δ184/183W values of several iron meteorites plotted versus the Δε182WGCR of each sample (i.e., the shift in ε182W as a result of secondary neutron capture). The Δε182WGCR values were calculated using the measured ε182W of these samples and the pre-exposure ε182W of the respective iron meteorite group as reported in Kruijer et al., 2014b (i.e., Δε182WGCR = ε182Wmeas. − ε182Wpre-exposure). Solid line shows the modelled GCR effects on δ184/183W, calculated using the model of Leya and Masarik (2013) [for Os/W = 12, Re/W = 0.9 and a cosmic-ray exposure age of 1200 Ma; see Cook et al. (2014) for full details]. (b) δ184/183W values of several iron meteorite samples from this study plotted versus the nucleosynthetic ε183W signatures of IID and IVB irons (Kruijer et al., 2014b). The solid line shows the calculated effects of nucleosynthetic W isotope anomalies on δ184/183W, based on the relation between ε182W and ε183W as given in Kruijer et al. (2014a). ε182W and ε183W values are the parts-per-104 deviations of internally normalized 182W/184W and 183W/184W ratios from terrestrial standard values.


8


Fig. 7. δ184/183W versus W concentration for IIAB iron meteorites. Late-crystallized IIAB irons have lower W contents and tend to have a slightly heavier W isotope composition than early crystallized samples.

and so resolving such potential variations will require the repeated analyses of multiple digestions of several chondrite samples. As for the chondrites, the investigated iron meteorites are also characterized by rather uniform δ184/183W, albeit with slightly more scatter (Fig. 4). Prior studies have shown that the W isotopic composition of iron meteorites can vary due to neutron capture effects induced during prolonged exposure of the iron meteoroids to galactic cosmic rays (GCR) (e.g., Kruijer et al., 2013; Leya and Masarik, 2013; Wittig et al., 2013). To assess the consequence of GCR-effects on measured δ184/183W, we used the neutron capture model of Leya and Masarik (2013) and simulated the effects of neutron capture on individual W isotopes for six iron meteorites with different irradiation histories (the IIAB irons Sikhote Alin & Negrillos IIAB, the IID irons Rodeo & Carbo IID, and the IVB irons Tawallah Valley & Hoba). The neutron capture–affected W isotope compositions of these irons were then used to generate a virtual double spike mass spectrometry run, and the results were processed through the double spike deconvolution. The results of these calculations indicate that neutron capture would generate a deficit in δ184/183W, but the magnitude of this effect is very small (≤0.01‰ δ184/183W) and, hence, insignificant compared to the analytical precision of the W stable isotope data (Fig. 6a). This is consistent with the observation that Carbo and Rodeo, two IID irons with very different irradiation histories (Kruijer et al., 2013), have indistinguishable δ184/183W values. Moreover, although Ainsworth, one of the most strongly irradiated iron meteorites (Kruijer et al., 2014b), shows a slightly higher δ184/183W than most other iron meteorites examined here, GCR-effects would lead to lower and not higher δ184/183W (Fig. 6a). Thus, GCR-effects cannot be the cause for the slightly elevated δ184/183W observed for Ainsworth compared to other IIAB irons. Taken together, GCR-effects are inconsequential for the determination of accurate δ184/183W values for iron meteorites. Another potential source of scatter in the δ184/183W of iron meteorites are nucleosynthetic W isotope variations. Such nucleosynthetic W isotope anomalies have been observed for IID and IVB iron meteorites (Qin et al., 2008; Kruijer et al., 2013). To assess the effect of nucleosynthetic W isotope anomalies on measured δ184/183W, we simulated a double spike measurement by modifying the W isotope abundances of a terrestrial standard according to the proportions expected from s– and r–process heterogeneity (Fig. 6b). The results of this test show that such nucleosynthetic heterogeneities generate only very small δ184/183W variations on the order of ~ 0.007‰. Thus, nucleosynthetic isotope anomalies also do not affect the δ184/183W measurements, at least not for the magnitude of nucleosynthetic W isotope anomalies so far observed for iron meteorites. A third possibility to explain variations in δ184/183W among iron meteorites is fractional crystallization of metallic magma during core solidification. During this process, W is mildly compatible and therefore preferentially partitions into the solid (e.g., Wasson, 1999; McCoy et al., 2011). Thus, during crystallization, the remaining liquid becomes

increasingly depleted in W, and so late-crystallized irons from a given group have lower W contents than early-crystallized samples (Fig. 7). Fractional crystallization can in principle induce stable isotope fractionation (e.g., Chernonozhkin et al., 2016) and these would be most pronounced in late-crystallized samples. Thus, the slightly heavier W isotopic composition observed for Ainsworth, a late-crystallized IIAB iron, compared to Negrillos, an early-crystallized IIAB, probably is the result of W stable isotope fractionation during crystallization of the IIAB core (Fig. 7). Because Ainsworth is one the latest crystallized IIAB irons (Wasson et al., 2007) and is the iron meteorite with the heaviest W isotope composition investigated in the present study, we excluded this sample from calculating a mean δ184/183W composition for the irons. Excluding Ainsworth, the overall range in δ184/183W observed for the iron meteorites is small compared to our analytical resolution, and so iron meteorites (mean δ184/183W = + 0.023 ± 0.011‰, 95% conf., n = 12, excl. Ainsworth) essentially exhibit the same W stable isotope composition as chondrites (mean δ184/183W = + 0.031 ± 0.007‰, 95% conf., n = 11) (Fig. 4). We note that as for the chondrites, resolving potential δ184/183W variations among the iron meteorites, or a difference between chondrites and iron meteorites, will require more precise W stable isotope data. The indistinguishable W stable isotope compositions observed for iron meteorites and chondrites is not unexpected, given that most of the W in a planetary body is hosted in the metal core. Thus, iron meteorites should have the same δ184/183W as the bulk planet, which in turn should be similar to the W stable isotope composition measured for chondrites, as observed previously for Mo stable isotopes (Burkhardt et al., 2014). Thus, the mean δ184/183W of iron meteorites and chondrites of 0.027 ± 0.007‰ (95% conf., n = 23; excl. Ainsworth) provides a good

a

b

Fig. 8. δ184/183W versus SiO2 and MgO for USGS geological reference materials. SiO2 and MgO contents were obtained from the USGS online database. The bulk inner solar system δ184/183W (shaded area; 95% conf.) is shown for comparison.



estimate for the bulk δ184/183W of planetary bodies in the inner solar system. 4.3. Tungsten isotope fractionations in terrestrial samples The USGS geological reference materials from this study were primarily analyzed to assess the accuracy and reproducibility of the W double spike method, and to provide precise δ184/183W values for easily accessible reference materials that can be used for interlaboratory comparisons in future W isotope studies. As such the USGS reference materials are not the ideal samples for assessing the extent and origin of mass-dependent W isotope variations in terrestrial samples, which would be better done using a suite of well-characterized and co-genetic igneous rocks. Nevertheless, the results for the USGS reference materials exhibit small but resolvable δ184/183W variations on the order of ~ 0.04‰, suggesting that magmatic processes on Earth resulted in mass-dependent W isotope variations. These small variations between different rock types only become apparent because for each sample several digestions have been repeatedly analyzed, resulting in precisely defined mean δ184/183W values for each sample (Fig. 4). For instance, for AGV-2 we obtained δ184/183W = − 0.005 ± 0.013‰ (95% conf.), which is resolved from δ184/183W = 0.039 ± 0.007‰ (95% conf.) obtained for BCR-2. In general, the δ184/183W values of the terrestrial samples show a weak positive correlation with SiO2 and a weak negative correlation with MgO contents (Fig. 8), suggesting that chemically more evolved igneous rocks exhibit a lighter W isotope composition than more mafic samples. One problem when interpreting the significance of this observation is that many of the geological reference materials have been produced in very large batches using hardened steel equipment. As such, these samples are prone to contamination with certain metals, including W. So, the variability observed among the USGS reference materials could in principle reflect contamination during preparation of these materials. However, there is no correlation between δ184/ 183 W and W concentration, indicating that the δ184/183W variations do not simply reflect binary mixing between a homogeneous sample composition and an isotopically distinct contaminant. For instance, samples with similar W concentrations such as BCR-2 and AGV-2 have different δ184/183W values of 0.039 ± 0.007‰ and −0.005 ± 0.013‰ (95% conf.), which is an unlikely outcome if the δ184/183W variations would solely reflect W contamination. Moreover, such contamination would be more pronounced for mafic samples, which have lower W contents than more felsic samples. However, the δ184/183W values of the mafic samples investigated here are indistinguishable from those of chondrites and iron meteorites, whereas the more felsic and W-rich samples deviate to lighter W isotope compositions. Thus, to explain the observed variations by contamination would imply that W-rich samples are more strongly contaminated than W-poor samples, which is the opposite of what would be expected. The observed δ184/183W variations, therefore, are unlikely to reflect contamination of the sample powders, but rather suggest that W stable isotope fractionations occur during the formation of igneous rocks. The nature of these W isotope variations is difficult to assess using the data of the present study. It is unlikely, however, that the W isotope fractionations are related to mineral fractionation during crystallization, because as one of the most incompatible elements W is not incorporated into most rock-forming minerals. One possibility is that the light W isotope composition of more evolved magmas is associated with the fluid mobility of W in subduction zones (König et al., 2008; Bali et al., 2012), and that during the release of W from the down-going slab, the light W isotopes preferentially partitioned into the sub-arc mantle, leading to lower δ184/183W for arc-related magmas. Clearly more work is needed to assess the influence of subduction zone processes, as well as of other igneous processes, on W stable isotope compositions. Nevertheless, our data show that small W isotope fractionations are present in igneous rocks, making W stable isotopes a potential new tool to study magmatic processes on Earth and potentially also on other planets.

9

5. Conclusions The narrow range of W stable isotope compositions of chondrites and iron meteorites implies that bulk planetary bodies in the inner solar system have a uniform δ184/183W = 0.027 ± 0.007‰ (relative to NIST 3163). We cannot exclude that small deviations from this value exist, but identifying such variations will require more high-precision W stable isotope data for meteorites. Terrestrial igneous samples analyzed in the present study have similar W stable isotope compositions compared to chondrites and iron meteorites, but chemically more evolved samples tend to be isotopically lighter than more mafic samples. These data show that magmatic processes on Earth induce small W stable isotope fractionations. Although the nature of these processes remains poorly constrained, one possibility is that the mobilization of W in subduction zones leads to the preferential incorporation of light W isotopes into arc-related magmas. Clearly, more work is needed to test this hypothesis and to assess the effects of igneous processes on the W stable isotope composition of magmas. Nevertheless, the presence of W isotope variations in terrestrial magmatic rocks demonstrates the potential of W stable isotopes as a useful new tracer for investigating igneous processes on Earth. Overall, the W stable isotope composition of terrestrial rocks does not seem to be very different from that of bulk Earth as inferred from chondrites and iron meteorites. This observation suggests that Earth's mantle and core have similar W stable isotope compositions, either because there is no W isotope fractionation between metal and silicate during core formation or because temperatures during core formation were too high to induce a measurable isotope fractionation. More data are needed, however, to firmly establish the absence of small W isotope variations that may potentially have been induced by core formation.

Acknowledgments We are grateful to Rachel Bezard, Gerrit Budde, Christoph Burkhardt and Mario Fischer-Gödde for discussions and support, to Benedikt Ahrens for support in generating the MATLAB® script, and to Ulla Heitmann for help with sample preparation. Some of the meteorite samples were provided by the Field Museum of Natural History (Chicago) and the Smithsonian National Museum of Natural History (Washington D.C.). These sources of the meteorite specimens are gratefully acknowledged. Constructive comments by Matthias Willbold, an anonymous reviewer and editor Klaus Mezger helped to improve the manuscript. This research received funding to Thorsten Kleine from the European Research Council under the European Community's Seventh Framework Program (FP7/2007–2013 Grant Agreement 616564 ‘ISOCORE’). References Abraham, K., Barling, J., Siebert, C., Belshaw, N., Gall, L., Halliday, A.N., 2015. Determination of mass-dependent variations in tungsten stable isotope compositions of geological reference materials by double-spike and MC-ICPMS. J. Anal. At. Spectrom. 30: 2334–2342. http://dx.doi.org/10.1039/C5JA00210A. Arevalo Jr., R., McDonough, W.F., 2008. Tungsten geochemistry and implications for understanding the Earth's interior. Earth Planet. Sci. Lett. 272:656–665. http://dx.doi. org/10.1016/j.epsl.2008.05.031. Armytage, R.M.G., Georg, R.B., Savage, P.S., Williams, H.M., Halliday, A.N., 2011. Silicon isotopes in meteorites and planetary core formation. Geochim. Cosmochim. Acta 75: 3662–3676. http://dx.doi.org/10.1016/j.gca.2011.03.044. Bali, E., Keppler, H., Audetat, A., 2012. The mobility of W and Mo in subduction zone fluids and the Mo–W–Th–U systematics of island arc magmas. Earth Planet. Sci. Lett. 351– 352:195–207. http://dx.doi.org/10.1016/j.epsl.2012.07.032. Bezard, R., Fischer-Gödde, M., Hamelin, C., Kleine, T., 2016. The effects of magmatic processes and crustal recycling on the molybdenum stable isotopic composition of Mid-Ocean Ridge Basalts. Earth Planet. Sci. Lett. 453:171–181. http://dx.doi.org/10. 1016/j.epsl.2016.07.056. Bonnand, P., Williams, H.M., Parkinson, I.J., Wood, B.J., Halliday, A.N., 2016. Stable chromium isotopic composition of meteorites and metal–silicate experiments: implications for fractionation during core formation. Earth Planet. Sci. Lett. 435:14–21. http://dx. doi.org/10.1016/j.epsl.2015.11.026.


10


Breton, T., Quitté, G., 2014. High-precision measurements of tungsten stable isotopes and application to earth sciences. J. Anal. At. Spectrom. 29:2284–2293. http://dx.doi.org/ 10.1039/C4JA00184B. Burkhardt, C., Hin, R.C., Kleine, T., Bourdon, B., 2014. Evidence for Mo isotope fractionation in the solar nebula and during planetary differentiation. Earth Planet. Sci. Lett. 391: 201–211. http://dx.doi.org/10.1016/j.epsl.2014.01.037. Chernonozhkin, S.M., Goderis, S., Costas-Rodríguez, M., Claeys, P., Vanhaecke, F., 2016. Effect of parent body evolution on equilibrium and kinetic isotope fractionation: a combined Ni and Fe isotope study of iron and stony-iron meteorites. Geochim. Cosmochim. Acta 186:168–188. http://dx.doi.org/10.1016/j.gca.2016.04.050. Cook, D.L., Kruijer, T.S., Leya, I., Kleine, T., 2014. Cosmogenic 180W variations in meteorites and re-assessment of a possible 184Os–180W decay system. Geochim. Cosmochim. Acta 140:160–176. http://dx.doi.org/10.1016/j.gca.2014.05.013. Dodson, M.H., 1970. Simplified equations for double-spiked isotopic analyses. Geochim. Cosmochim. Acta 34:1241–1244. http://dx.doi.org/10.1016/0016–7037(70)90060-8. Galer, S.J.G., 1999. Optimal double and triple spiking for high precision lead isotopic measurement. Chem. Geol. 157:255–274. http://dx.doi.org/10.1016/S0009-2541(98)00203-4. Georg, R.B., Halliday, A.N., Schauble, E.A., Reynolds, B.C., 2007. Silicon in the Earth's core. Nature 447:1102–1106. http://dx.doi.org/10.1038/nature05927. Hin, R.C., Burkhardt, C., Schmidt, M.W., Bourdon, B., Kleine, T., 2013. Experimental evidence for Mo isotope fractionation between metal and silicate liquids. Earth Planet. Sci. Lett. 379:38–48. http://dx.doi.org/10.1016/j.epsl.2013.08.003. Hofmann, A., 1971. Fractionation corrections for mixed-isotope spikes of Sr, K, and Pb. Earth Planet. Sci. Lett. 10:397–402. http://dx.doi.org/10.1016/0012-821X(71)90087-2. Irisawa, K., Hirata, T., 2006. Tungsten isotopic analysis on six geochemical reference materials using multiple collector-ICP-mass spectrometry coupled with a rhenium-external correction technique. J. Anal. At. Spectrom. 21:1387–1395. http://dx.doi.org/ 10.1039/b607945h. Kleine, T., Touboul, M., Bourdon, B., Nimmo, F., Mezger, K., Palme, H., Jacobsen, S.B., Yin, Q.Z., Halliday, A.N., 2009. Hf-W chronology of the accretion and early evolution of asteroids and terrestrial planets. Geochim. Cosmochim. Acta 73:5150–5188. http://dx. doi.org/10.1016/j.gca.2008.11.047. Kleine, T., Hans, U., Irving, A.J., Bourdon, B., 2012. Chronology of the angrite parent body and implications for core formation in protoplanets. Geochim. Cosmochim. Acta 84: 186–203. http://dx.doi.org/10.1016/j.gca.2012.01.032. König, S., Münker, C., Schuth, S., Garbe-Schönberg, D., 2008. Mobility of tungsten in subduction zones. Earth Planet. Sci. Lett. 274:82–92. http://dx.doi.org/10.1016/j.epsl. 2008.07.002. Krot, A.N., Keil, K., Scott, E.R.D., Goodrich, C.A., Weisberg, M.K., 2014. Classification of meteorites and their genetic relationships. In: Davis, A.M. (Ed.), Meteorites and Cosmochemical Processes, second ed. Treatise on Geochemistry vol. 1. Elsevier, Oxford, pp. 1–63. Kruijer, T.S., Sprung, P., Kleine, T., Leya, I., Burkhardt, C., Wieler, R., 2012. Hf–W chronometry of core formation in planetesimals inferred from weakly irradiated iron meteorites. Geochim. Cosmochim. Acta 99:287–304. http://dx.doi.org/10.1016/j.gca.2012. 09.015. 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 Planet. Sci. Lett. 361:162–172. http://dx.doi.org/10.1016/j. epsl.2012.10.014. Kruijer, T.S., Touboul, M., Fischer-Gödde, M., Bermingham, K.R., Walker, R.J., Kleine, T., 2014a. Protracted core formation and rapid accretion of protoplanets. Science 344: 1150–1154. http://dx.doi.org/10.1126/science.1251766. Kruijer, T.S., Kleine, T., Fischer-Gödde, M., Burkhardt, C., Wieler, R., 2014b. Nucleosynthetic W isotope anomalies and the Hf–W chronometry of Ca–Al-rich inclusions. Earth Planet. Sci. Lett. 403:317–327. http://dx.doi.org/10.1016/j.epsl.2014.07.003. Leya, I., Masarik, J., 2013. Thermal neutron capture effects in radioactive and stable nuclide systems. Meteorit. Planet. Sci. 48:665–685. http://dx.doi.org/10.1111/maps. 12090.

McCoy, T.J., Walker, R.J., Goldstein, J.I., Yang, J., McDonough, W.F., Rumble, D., Chabot, N.L., Ash, R.D., Corrigan, C.M., Michael, J.R., Kotula, P.G., 2011. Group IVA irons: new constraints on the crystallization and cooling history of an asteroidal core with a complex history. Geochim. Cosmochim. Acta 75:6821–6843. http://dx.doi.org/10.1016/j.gca. 2011.09.006. Millet, M.-A., Dauphas, N., Greber, N.D., Burton, K.W., Dale, C.W., Debret, B., Macpherson, C.G., Nowell, G.M., Williams, H.M., 2016. Titanium stable isotope investigation of magmatic processes on the Earth and Moon. Earth Planet. Sci. Lett. 449:197–205. http://dx.doi.org/10.1016/j.epsl.2016.05.039. Newsom, H.E., Sims, K.W.W., Noll Jr., P.D., Jaeger, W.L., Maehr, S.A., Beserra, T.B., 1996. The depletion of tungsten in the bulk silicate earth: constraints on core formation. Geochim. Cosmochim. Acta 60:1155–1169. http://dx.doi.org/10.1016/00167037(96)00029-4. O'Neill, H.S.C., Berry, A.J., Eggins, S.M., 2008. The solubility and oxidation state of tungsten in silicate melts: implications for the comparative chemistry of W and Mo in planetary differentiation processes. Chem. Geol. 255:346–349. http://dx.doi.org/10.1016/ j.chemgeo.2008.07.005. Qin, L., Dauphas, N., Wadhwa, M., Markowski, A., Gallino, G., Janney, P.E., Bouman, C., 2008. Tungsten nuclear anomalies in planetesimal cores. Astrophys. J. 674: 1234–1241. http://dx.doi.org/10.1086/524882. Rudge, J.F., Reynolds, B.C., Bourdon, B., 2009. The double spike toolbox. Chem. Geol. 265: 420–431. http://dx.doi.org/10.1016/j.chemgeo.2009.05.010. Savage, P.S., Georg, R.B., Williams, H.M., Burton, K.W., Halliday, A.N., 2011. Silicon isotope fractionation during magmatic differentiation. Geochim. Cosmochim. Acta 75: 6124–6139. http://dx.doi.org/10.1016/j.gca.2011.07.043. Schoenberg, R., Zink, S., Staubwasser, M., Blanckenburg, F.V., 2008. The stable Cr isotope inventory of solid Earth reservoirs determined by double spike MC-ICP-MS. 249: pp. 294–306. http://dx.doi.org/10.1016/j.chemgeo.2008.01.009. Shahar, A., Hillgren, V.J., Young, E.D., Fei, Y., Macris, C.A., Deng, L., 2011. High-temperature Si isotope fractionation between iron metal and silicate. Geochim. Cosmochim. Acta 75:7688–7697. http://dx.doi.org/10.1016/j.gca.2011.09.038. Shearer, C.K., Righter, K., 2003. Magmatic fractionation of hafnium and tungsten: constraints on the timing of core formation and differentiation in the Moon and Mars. Geochim. Cosmochim. Acta 67:2497–2507. http://dx.doi.org/10.1029/ 2002GL015523. Siebert, C., Nägler, T.F., Kramers, J.D., 2001. Determination of molybdenum isotope fractionation by double-spike multicollector inductively coupled plasma mass spectrometry. Geochem. Geophys. Geosyst. 2:1032–1047. http://dx.doi.org/10.1029/ 2000GC000124. Wade, J., Wood, B.J., Norris, C.A., 2013. The oxidation state of tungsten in silicate melt at high pressures and temperatures. Chem. Geol. 335:189–193. http://dx.doi.org/10. 1016/j.chemgeo.2012.10.011. Wang, K., Moynier, F., Dauphas, N., Barrat, J.-A., Craddock, P., Korliss, K.S., 2012. Iron isotope fractionation in planetary crusts. Geochim. Cosmochim. Acta 89:31–45. http:// dx.doi.org/10.1016/j.gca.2012.04.050. Wasson, J.T., 1999. Trapped melt in IIIAB irons; solid/liquid elemental partitioning during the fractionation of the IIIAB magma. Geochim. Cosmochim. Acta 63:2875–2889. http://dx.doi.org/10.1016/S0016-7037(99)00283-5. Wasson, J.T., Huber, H., Malvin, D.J., 2007. Formation of IIAB iron meteorites. Geochim. Cosmochim. Acta 71:760–781. http://dx.doi.org/10.1016/j.gca.2006.09.032. 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 Planet. Sci. Lett. 361:152–161. http://dx.doi.org/10.1016/j.epsl. 2012.10.013.
