31. Rapid Accretion and Differentiation of Iron Meteorite Parent Bodies

Home , Gibeon (meteorite), IIE iron meteorite, Parent body, Tlacotepec meteorite

Earth and Planetary Science Letters 273 (2008) 94–104

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Earth and Planetary Science Letters

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Rapid accretion and differentiation of iron meteorite parent bodies inferred from 182Hf–182W chronometry and thermal modeling

Liping Qin a,b,⁎, Nicolas Dauphas a,b, Meenakshi Wadhwa b,1, Jozef Masarik c, Philip E. Janney b,1 a Origins Laboratory, Department of the Geophysical Sciences and Enrico Fermi Institute, The University of Chicago, 5734 South Ellis Avenue, Chicago, IL 60637, USA b Department of Geology, The Field Museum, 1400 South Lake Shore Drive, Chicago, IL 60605, USA c Komensky University, Department of Nuclear Physics, Mlynska dolina F/1, SK-842 15 Bratislava, Slovakia

ARTICLE INFO ABSTRACT

Article history: New high-precision W isotope measurements are presented for 33 iron meteorites from 8 magmatic groups Received 26 February 2008 (IC, IIAB, IID, IIIAB, IIIE, IIIF, I VA and IVB), 2 non-magmatic groups (IAB–IIICD and IIE), and one ungrouped Received in revised form 3 June 2008 iron (Deep Springs). All magmatic irons have ε182W values that are, within errors, equal to, or less radiogenic Accepted 11 June 2008 than, the Solar System initial of −3.47±0.20. A method was developed to correct the measured ε182W values Available online 25 June 2008 of magmatic iron meteorites for the presence of cosmogenic effects produced during space exposure to Editor: T. Spohn galactic cosmic rays. The corrected data provide new constraints on the timing of metal-silicate differentiation in iron meteorite parent bodies, which must have taken place within a few million years Keywords: (b2 to 6 My) of condensation of calcium–aluminum-rich inclusions (CAIs). Metal-silicate differentiation ages Hf-W model age (from 182Hf–182W systematics) were combined with parent body sizes (from metallographic cooling rates) metal-silicate differentiation into a model of planetesimal heating by 26Al-decay, to constrain the accretion timescale of iron meteorite accretion timescales parent bodies. Accretion of iron meteorite parent bodies most likely occurred within 1.5 My of the formation of CAIs. The fast accretion times of iron meteorite parent bodies are consistent with dynamical models indicating that these objects may have originated in the terrestrial planet-forming region, where the accretion rates were high. Our W isotopic data for non-magmatic IAB–IIICD and IIE irons provide new constraints for their formation mechanisms. In particular, they support formation of IAB–IIICD iron meteorites by melting during a single collision event dated at 4–7 My after formation of the Solar System. © 2008 Elsevier B.V. All rights reserved.

1. Introduction silicate differentiation in the early Solar System (Jacobsen and Harper, 1996; Lee and Halliday, 1996; Horan et al., 1998; Yin et al., 187 187 Several extant and extinct radiochronometers ( Re– Os, 2002; Schoenberg et al., 2002; Kleine et al., 2002; Quitté and Birck, 107 107 53 53 Pd– Ag, Mn– Cr) have been used to constrain the time of 2004; Lee, 2005; Kleine et al., 2005; Scherstén et al., 2006; formation of magmatic iron meteorites (Shen et al., 1996; Smoliar Markowski et al., 2006a,b) because Hf is lithophile while W is et al., 1996; Chen and Wasserburg, 1996; Carlson and Hauri, 2001; moderately siderophile. If bulk planetesimals had chondritic Sugiura and Hoshino, 2003). These systems indicate that accretion, compositions and metal-silicate differentiation occurred after differentiation, and crystallization of iron meteorite parent bodies complete decay of 182Hf, metallic cores should have chondritic occurred early, within the first several tens of My of the formation 182W compositions. If differentiation occurred while 182Hf was still of the Solar System. However, they all have limitations to their alive, a deficit in 182W in cores relative to chondrites should be application as fine scale chronometers for early Solar System produced. Thus, by studying the isotopic abundance of 182Win − events. Hafnium-182 is an extinct radionuclide which β decays magmatic iron meteorites, which are thought to represent 182 into Ta with a half-life of 8.9 My (Vockenhuber et al., 2004). The fragments of planetesimal cores, we can infer the timing of 182 latter is very unstable (t1/2 = 115 d) and rapidly decays to W. The metal-silicate differentiation relative to condensation of refractory 182 182 Hf– W system is useful to date the relative timing of metal- CAIs in the protosolar nebula. Horan et al. (1998) performed the first extensive study of W isotopes in iron meteorites. They showed that magmatic iron ⁎ Corresponding author. Present address: Department of Terrestrial Magnetism, meteorites have ε182W values (relative deviation of the ratio of 182W Carnegie Institution of Washington, 5241 Broad Branch Road, NW, Washington, DC to a stable W isotope from a terrestrial standard in parts per 104) 20015, USA. between −5.1 and −3.1. Recent studies have shown that some iron E-mail address: [email protected] (L. Qin). 182 1 ε − Present address: Center for Meteorite Studies, School of Earth and Space meteorites, such as Tlacotepec, have very negative W values ( 4.4 Exploration, Arizona State University, Box 871404, Tempe, AZ 85287, USA. to −4.0) (Quitté and Birck, 2004; Lee, 2005; Scherstén et al., 2006;

Markowski et al., 2006a,b), lower than the initial solar value of −3.47± exponential law. This law describes well the fractionation occurring in 0.20 defined by CAI and chondrite isochrons (Kleine et al., 2005). a MC-ICPMS (Maréchal et al., 1999). The NIST 3163 W reference Taken at face value, the implication would be that metal-silicate material was used as the W isotope standard. Both the samples and differentiation in these iron meteorite parent bodies predated the standards were run at ion intensities of 3–8×10− 11 A for 184W, formation of CAIs, which goes against the current paradigm that CAIs obtained with 30–80 ppb solutions. The sample and standard are the oldest condensed material in the Solar System. solutions were introduced into the ICP source in the form of aerosols Exposure to galactic cosmic radiation (GCR) can modify the W using an Aridus desolvating nebulizer. The W concentration of the isotopic compositions of iron meteorites (Masarik, 1997; Leya et al., sample was always matched with that of the bracketing standard 2003) through capture of secondary thermal neutrons. This can lower within 3% because we noticed that a mismatch in the W concentration ε182W values and yield false negative ages relative to CAIs (Masarik, can affect the accuracy of the isotope measurements (Qin et al., 2007). 1997; Leya et al., 2003). Markowski et al. (2006b) measured the W The sample measurements were interspersed between those of the isotopic compositions along depth profiles in Grant (IIIB) and Carbo standard, and the internally normalized W isotope composition of the (IID) iron meteorites. They demonstrated that the anomalously low sample was corrected using the mean of the two bracketing standard ε182W values in these meteorites were produced by interactions with analyses. All isotope ratios are presented using the ε notation, always GCR. Thus in those samples, ε182W cannot be interpreted in terms of using 183W as the denominator isotope. A total of 13–20 (n) repeats 182Hf decay alone. It is conceivable that the very negative ε182W values were obtained for each sample and were used to calculate averages measured in Tlacotepec have also been produced by exposure to GCR, and 95% confidence intervals as follows, given the long exposure age (945±55 Ma) of this sample (Voshage and sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 182 –182 n n Feldmann, 1979). A major challenge in establishing the Hf W ɛ ¼ 1 ∑ ɛ F 1 ∑ ðÞɛ −ɛ 2 t0:p95ffiffiffi;n−1 ð1Þ k − k system as a reliable chronometer is to find a way to accurately correct n k¼1 n 1 k¼1 n for the cosmogenic effect. Knowledge of exposure ages alone is not sufficient because the effect of irradiation by GCR is modulated by the where t0.95,n − 1 is student's t-value corresponding to a two-sided 95% depth of burial in the pre-atmospheric object (Masarik, 1997). confidence interval for n−1 degrees of freedom. We studied Hf–W systematics in iron meteorites using an The absolute 182W/183Wand184W/183W ratios (after normalization to improved method, allowing variations of less than 0.1 ε-unit on 186W/183W=1.98594) of the NIST 3163 W reference material averaged 182W to be resolved (Foley et al., 2005; Qin et al., 2007; Qin et al., over a period of two years are 1.85174 (±43, 2σ) and 2.14123 (±63, 2σ), 2008). Increasing the precision of W isotopic measurements in iron respectively. These values agree within error with those reported meteorites can improve the time resolution of core segregation previously by negative thermal ionization mass spectrometry (1.85128± processes, and provide a potential means of correcting cosmogenic 35 and 2.14078±13) and MC-ICPMS (1.85163±8 and 2.14076±7) and nucleosynthetic effects. Several aspects of our approach distin- (Völkening et al., 1991; Lee and Halliday, 1995). guish this work from previous studies: 3. Results (i) This study focused on a subset of samples with well characterized 41K–40K exposure ages (Voshage and Feldmann, 3.1. ε182W results 1979; Voshage et al., 1983; Voshage, 1984). (ii) A precision of ∼±0.1 or better on ε182W was achieved, which is Iron meteorites from both magmatic (IC, IIAB, IID, IIIAB, IIIE, IIIF, I VA comparable to or better than the most recent studies (Scherstén and IVB) and non-magmatic (IAB–IIICD and IIE) groups were studied et al., 2006; Markowski et al., 2006a), but represents a ∼2–5 (Table 1). All samples show deficits in ε182Wof−4.2 to −2.3 relative to the fold improvement compared to earlier efforts (Jacobsen and terrestrial W reference NIST 3163, a proxy for the composition of the bulk Harper, 1996; Lee and Halliday, 1996; Horan et al., 1998; Yin silicate Earth (Fig. 1). These values are also significantly lower than those et al., 2002; Schoenberg et al., 2002; Kleine et al., 2002; Quitté measured in bulk chondrites (−1.9±0.2) (Yin et al., 2002; Schoenberg et al., and Birck, 2004; Lee, 2005). 2002; Kleine et al., 2004). This range of variation is consistent with (iii) A new method is presented for estimating ε182W of magmatic previous work, and demonstrates that 182Hf was alive at the time of metal- iron meteorites prior to GCR exposure. silicate differentiation in these meteorites (Jacobsen and Harper, 1996; Lee (iv) The time of core–mantle differentiation, as inferred from 182Hf– and Halliday, 1996; Horan et al., 1998; Yin et al., 2002; Schoenberg et al., 182W systematics, was used in a thermal model of planetesimal 2002; Kleine et al., 2002; Quitté and Birck, 2004; Lee, 2005; Kleine et al., differentiation to constrain the time of accretion of iron 2005; Scherstén et al., 2006; Markowski et al., 2006a). Resolvable isotopic meteorite parent bodies. variations are present both within and between iron meteorite groups. Some of the samples measured in this study have also been 2. Methods analyzed in previous work, including Nocoleche (IC) (Scherstén et al., 2006), Arispe (IC) (Lee, 2005; Scherstén et al., 2006; Markowski et al., The methods used for purifying W and analyzing its isotopic 2006a), Bendego (IC) (Scherstén et al., 2006), El Burro (IIA) (Scherstén composition have been presented in detail elsewhere (Foley et al., et al., 2006), Carbo (IID) (Markowski et al., 2006a,b), Henbury (IIIAB) 2005; Qin et al., 2007) and will be briefly reviewed here. Iron (Horan et al., 1998; Scherstén et al., 2006; Markowski et al., 2006a), meteorite samples were first leached in 11 N HCl–1 N HF to remove Grant (IIIB) (Lee, 2005; Markowski et al., 2006b), Nelson County (IIIF) surface-sited terrestrial contamination from various sources. The (Markowski et al., 2006a), Duchesne (IVA) (Scherstén et al., 2006), cleaned samples were then dissolved in aqua regia and evaporated to Tawallah Valley (IVB) (Horan et al., 1998; Markowski et al., 2006a), dryness and redissolved in 11 N HCl. Chemical separation was Cape of Good Hope (IVB) (Horan et al., 1998; Lee, 2005; Scherstén achieved by passing the samples through one cation exchange column et al., 2006; Markowski et al., 2006a), Santa Clara (IVB) (Markowski and several anion-exchange columns (Qin et al., 2007). This protocol et al., 2006a), Hoba (IVB) (Scherstén et al., 2006), Tlacotepec (IVB) can accommodate sample sizes up to 1.7 g so that large quantities of (Horan et al., 1998; Quitté and Birck, 2004; Lee, 2005; Kleine et al., clean W (usually 500–2000 ng) can be retrieved for isotopic analyses. 2005; Scherstén et al., 2006; Markowski et al., 2006a), and Watson The final W solutions were analyzed on a Micromass Isoprobe multi- (IIE) (Snyder et al., 2001; Markowski et al., 2006a). There is good collector inductively coupled plasma mass spectrometer (MC-ICPMS) agreement with the present work except for Tlacotepec (Kleine et al., located at the Field Museum, Chicago. The measurements were 2005) and Nocoleche (Scherstén et al., 2006) that can be explained by normalized to 186W/183W=1.98594 (Völkening et al., 1991) using the differences in the degree of shielding from GCR. 96 L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104

Table 1 Measured W isotopic compositions of iron meteorites

Group Sample Exposure⁎⁎ age Pre-atmospheric mass ε182W ε184W Model age N

(Ma) (103 kg) (My) Magmatic IC Nocoleche 250±70 −3.58±0.05 0.03±0.04 −0.8±1.7 13 Arispe 955±90 ∞ (2π) −3.91±0.06 0.04±0.03 −3.1±1.7 18 Bendego 940±90 ∞ (2π) −3.97±0.06 0.08±0.04 −3.5±1.7 16 IIAB Smithsonian⁎ 90±80 −3.42±0.05 0.02±0.06 0.4±1.7 15 Sierra Gorda⁎ 140 ± 110 −3.47±0.08 −0.02±0.07 0.0±1.8 13 Cedartown #1a⁎ 180 ± 80 −3.36±0.05 0.01±0.04 0.9±1.7 17 Cedartown #1b⁎ 180 ± 80 −3.33±0.18 0.04±0.10 1.2±2.3 5 Cedartown #2⁎ 180 ± 80 −3.58±0.10 0.00±0.11 −0.9±1.8 9 Sample mean −3.39±0.04 0.01±0.03 0.6±1.7 El Burro⁎ 165 ± 115 −3.52±0.11 −0.01±0.09 −0.4±1.9 16 IID Carbo 850±140 5 −4.09±0.08 −0.03±0.04 −4.3±1.7 17 IIIAB Ruff's Mountain 120 −3.53±0.06 0.04±0.04 −0.5±1.7 15 Henbury 240 −3.70±0.07 0.01±0.03 −1.8 ± 1.7 15 Sacramento Mountains 315±55 2 −3.46±0.06 −0.02±0.07 0.1±1.7 14 Grant #1(G 01) 695±65 2 −3.54±0.09 0.01±0.08 −0.5±1.7 15 Grant #2 (G05) 695±65 2 −3.66±0.07 −0.01±0.05 −1.5 ± 1.7 14 Grant #3 (G10) 695±65 2 −3.69±0.13 0.01±0.05 −1.7 ± 1.9 14 Grant #4 (G10) 695±65 2 −3.56±0.12 0.03±0.06 −0.7±1.9 13 IIIE Rhine Villa 325±70 5 −3.58±0.09 0.02±0.04 −0.9±1.8 14 Kokstad 470±90 ∞ (2π) −3.65±0.06 −0.02±0.06 −1.4 ± 1.7 15 IIIF Nelson County 490±55 2 −3.30±0.05 −0.05±0.03 1.5±1.7 17 IVA Gibeon #1⁎ 32 −3.36±0.09 0.04±0.07 1.0±1.8 12 Gibeon #2⁎ 32 −3.39±0.06 0.07±0.05 0.7±1.7 14 Gibeon #3⁎ 32 −3.41±0.15 0.02±0.06 0.5±2.1 15 Sample mean −3.38±0.05 0.05±0.04 Duchesne 220±70 −3.33±0.27 0.07±0.14 1.2±2.9 12 Bishop Canyon 220 −3.56±0.11 0.07±0.03 −0.7±1.8 15 Wood's Mountain 280 −3.36±0.14 −0.03±0.07 0.9±2.1 14 Hill City 475±90 3 −3.48±0.18 −0.05±0.09 −0.1±2.2 12 IVB Tawallah Valley #1⁎ 250±85 −3.53±0.06 −0.10±0.02 −0.2±1.7 18 Tawallah Valley #2⁎ 250±85 −3.53±0.05 −0.07±0.05 −0.1±1.7 16 Sample mean −3.53±0.04 −0.10±0.02 Santa Clara⁎ −3.62±0.06 −0.06±0.03 −0.9±1.7 16 Hoba⁎ 340±110 100 −3.43±0.08 −0.07±0.05 0.7±1.8 16 Cape of Good Hope⁎ 775±70 30 −3.71±0.06 −0.12±0.04 −1.5 ± 1.7 15 Skookum⁎ 945±90 1 −3.55±0.1 −0.11±0.05 −0.3±1.8 16 Tlacotepec #1a⁎ 945±55 3 −4.02±0.07 −0.10±0.06 −3.6±1.7 17 Tlacotepec #1b⁎ 945±55 3 −3.95±0.04 −0.05±0.02 −3.2±1.7 7 Tlacotepec #2⁎ 945±55 3 −4.04±0.09 −0.09±0.04 −3.7±1.7 14 Tlacotepec #3⁎ 945±55 3 −4.14±0.07 −0.06±0.05 −4.3±1.7 17 Tlacotepec #4⁎ 945±55 3 −4.22±0.08 −0.07±0.06 −4.8±1.7 13 Tlacotepec #5⁎ 945±55 3 −4.22±0.07 −0.11±0.03 −4.8±1.7 15 Tlacotepec #6⁎ 945±55 3 −4.25±0.05 −0.09±0.02 −4.9±1.7 16 Tlacotepec #7⁎ 945±55 3 −4.11±0.16 −0.05±0.08 −4.2±1.9 11 Tlacotepec #8a⁎ 945±55 3 −4.20±0.07 −0.09±0.05 −4.6±1.7 17 Tlacotepec #8b⁎ 945±55 3 −4.13±0.06 −0.05±0.04 −4.2±1.7 15 Tlacotepec #9⁎ 945±55 3 −4.01±0.05 −0.15±0.04 −3.5±1.7 25 Sample mean −4.10±0.02 −0.08±0.01 Non-magmatic IAB–IIICD BoHumilitz 140±230 −3.02±0.09 0.06±0.07 4.3±2.0 11 Surprise Springs 130±170 −2.94±0.07 0.04±0.04 5.3±1.9 16 Deport 1140±70 50 −3.82±0.05 0.05±0.04 −2.6±1.7 17 Nantan −2.93±0.07 0.01±0.05 5.4±1.9 15 IIE Watson 8 −2.30±0.08 0.02±0.04 17.5±3.9 15 Arlington 150 −3.01±0.08 0.03±0.04 4.5±1.9 15 Ungrouped Deep Springs #1a⁎ 2275±65 5 −3.80±0.06 −0.19±0.05 −1.9 ± 1.7 16 Deep Springs #1b⁎ 2275±65 5 −3.75±0.06 −0.14±0.03 −1.6 ± 1.7 2 Deep Springs #2⁎ 2275±65 5 −3.78±0.05 −0.15±0.06 −1.8 ± 1.7 13 Deep Springs #3⁎ 2275±65 5 −3.77±0.10 −0.26±0.09 −1.8 ± 1.8 12 Deep Springs #4⁎ 2275±65 5 −3.74±0.11 −0.12±0.05 −1.5 ± 1.8 10 Sample mean −3.77±0.02 −0.15±0.02

182 184 i i 183 i 183 4 ε W and ε W are the relative deviations from NIST 3163 W standard; ε W=((W/ W)sample/( W/ W)std −1)×10 . Pre-atmospheric masses are from Voshage (1984). N represents the number of repeats. Arabic numeral sufﬁxes represent separate dissolutions of different pieces of the same meteorite sample. Small letter sufﬁxes represent different aliquots of the same dissolution. ⁎Results from our previous work (Qin et al., 2007; Qin et al., 2008). ⁎⁎Exposure ages from Signer and Nier (1962); Schultz (1967); Nord and Zahringer (1972); Voshage and Feldmann (1979); Villa et al. (1981); Voshage et al. (1983); Voshage (1984); Olsen et al. (1994); Nishiizumi et al. (1995); Lavielle et al. (1997); Lavielle et al. (1999).

All magmatic irons have ε182W values that are, within error, equal ranging from −3.82±0.05 to −2.30±0.08, with a majority at −3.00. This to or less radiogenic than the Solar System initial value (−3.47±0.20) is also consistent with previous work (Horan et al., 1998; Scherstén derived from a CAI isochron (Kleine et al., 2005). This result is et al., 2006; Markowski et al., 2006). One IAB–IIICD iron meteorite, consistent with previous studies. In contrast to magmatic irons, non- Deport, has a very unradiogenic ε182Wof−3.82±0.05 (Table 1; Fig. 1), magmatic irons are characterized by more radiogenic ε182W values consistent with its long GCR exposure age of 1140 Ma. L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104 97

ungrouped iron Deep Springs show systematic deﬁciencies in ε184W of −0.08±0.01 and −0.15±0.02, respectively. The deﬁciencies in ε184W in IVBs have been replicated in different laboratories using 3 different

Fig. 1. ε182W values measured in iron meteorites. Circles and triangles correspond to magmatic and non-magmatic iron meteorites, respectively. Non-magmatic groups are indicated with asterisks. Solid symbols and open symbols represent iron meteorites with long exposure ages (N600 My) and relatively short exposure ages (b600 My) respectively; the 600 My exposure age cut-off is arbitrary. The error bars represent 95% conﬁdence intervals.

Three pieces of the Grant, a group IIIB iron meteorite, were analyzed (Table 1). This meteorite has an exposure age of 695 Ma (Voshage and Feldmann, 1979). The 3 specimens were sampled from a section of Grant that has been well characterized in terms of its cosmic ray exposure (see Fig. 2A for sample locations in the section and (Hoffman and Nier, 1958; Signer and Nier, 1960) for noble gas data). As can be seen in Fig. 2B, our results are comparable with those of Markowski et al. (2006b). A general positive correlation is found, i.e., the ε182W values increase with increasing distance from the center (or decreasing depth from pre-atmospheric surface), although all the data points are within error of each other. Note that Markowski et al. (2006b) adopted a pre-atmospheric center based on a recent noble gas study (Ammon et al., 2006), which differs signiﬁcantly from that suggested in previous work (Hoffman and Nier, 1958; Signer and Nier, 1960). In Fig. 2B, the distances in (Markowski et al., 2006b) were recalculated using the pre-atmospheric center determined in (Hoff- man and Nier, 1958; Signer and Nier, 1960).

3.2. ε184W results Fig. 2. A) Sample locations of three specimens taken from within a slice of Grant. The slice had been cut into a number of bars (labeled A to T). The distances of the specimens G01, G05 and G10 from the reference line (vertical line) are 325, 200 and 84 mm, 184 The ε W results are shown in Fig. 3. The majority of the samples respectively. The black dot shows the location of the pre-atmospheric center, according show no deviation in ε184W from the terrestrial NIST W standard to (Hoffman and Nier, 1958). B) and C) Plots of measured ε182W values vs. distance to the outside of the analytical uncertainty (∼±0.1 ε unit). In particular, the pre-atmospheric center for Grant and Carbo. Also shown are modeled proﬁles for the two meteorites, with three different libraries: ENDF, JEF and KASKAD (see Kollar et al., two largest magmatic iron groups, IIIAB and IIAB have ε184W values of 2006 for details). The circles and squares represent the values obtained in this study and 0 within error bars. This is consistent with a recent study by in (Markowski et al., 2006b), respectively. The exposure ages for Grant and Carbo are Markowski et al. (2006a). However, all IVB iron meteorites and the 695 and 850 My (Voshage and Feldmann, 1979). 98 L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104

some samples have significantly negative model ages, confirming previous findings. The paradoxical negative Hf–W model ages of some iron meteorites can have two possible causes. One is that the actual initial ε182W value in CAIs is lower than −3.47±0.20, and the other is that ε182W values measured in iron meteorites have been altered by exposure to GCR. Palme et al. (1994) and Humayun et al. (2007) found evidence for W redistribution in Allende CAIs accompanying parent body metamorphism and alteration of the carbonaceous chondrites. Palme et al. (1994) showed that W was depleted in opaque assemblages (OAs) from the Allende Egg 6 CAI. They also found complementary W enrichments in surrounding silicates, suggestive of W redistribution during formation of OAs. Humayun et al. (2007) reported a metal vein cross-cutting a silicate mineral grain in the Allende Golfball CAI with a very high W concentration (N100 ppm) compared to surrounding areas, providing evidence for the transport of oxidized W in the CAI. According to Humayun et al. (2007), this could potentially affect the interpretation of the 182Hf-182W isochron reported for Allende CAIs by Kleine et al. (2005), since W in OAs might have isotopically exchanged with radiogenic W in the silicates during metamorphism. In contrast with the ambiguity in the initial ε182W of CAIs, there are strong lines of evidence that ε182W values in some iron meteorites were modified by exposure to GCR (Markowski et al., 2006b). A model simulation (Masarik, 1997) predicted a maximum shift of −0.5 in ε182W for the Toluca iron meteorite (3.9-m radius and 600 My exposure age). Thus GCR effect can explain some (if not all) of the negative model ages. A remaining question is how to correct for these effects on metal-silicate differentiation ages.

4.1. The effect of galactic cosmic ray irradiation and correction of model ages Fig. 3. ε184W values measured in iron meteorites. The symbols are the same as in Fig. 1. See Qin et al. (2008) for the significance of negative ε184W values in some iron meteorite groups. As a first approach towards quantifying the GCR effect on W isotopic compositions, the ε182W values are plotted against exposure ages for magmatic irons in Fig. 4. The exposure ages are from (Signer instruments (Isoprobe at the Field Museum in Chicago, Nu-plasma at and Nier, 1962; Schultz, 1967; Nord and Zahringer, 1972; Voshage and ETH in Zurich, and Neptune at the Thermo Factory in Bremen) and in Feldmann, 1979; Villa et al., 1981; Voshage et al., 1983; Voshage, 1984; the case of IVBs, may reflect incomplete mixing of products of s- and r- Olsen et al., 1994; Nishiizumi et al., 1995; Lavielle et al., 1997; Lavielle process nucleosynthesis in the solar nebula (Qin et al., 2008). Qin et al. et al., 1999), and are mostly 41K–40K ages. The 41K–40K ages are usually (2008) estimated the correction that needs to be applied to ε182W more precise than the exposure ages based on other chronometers but model ages due to this possible nucleosynthetic effect. The conclusion may still suffer from a systematic bias. Nevertheless, it can be seen in is that the age correction is small (b0.5 My; see Qin et al. 2008 for Fig. 4 that ε182W values tend to decrease with increasing exposure details). Note that several other samples may also show non-zero ages. The samples with exposure ages of less than 400 My form a ε184W values (e.g. Bishop Canyon and Deport), but systematic tests and replicates on these samples were not done to assess the significance of these results.

4. Discussion

As has already been pointed out in a number of studies, a major difficulty with 182Hf-182W chronology of iron meteorites is that some specimens yield model ages that pre-date CAIs. Assuming a simple two-stage evolution model, in which only one Hf/W fractionation event associated with core formation occurs, the ε182W values of iron meteorites can be used to calculate metal-silicate differentiation ages, 182 182 1 ɛ WA−ɛ WC Δ − ¼ ; ð2Þ TA B ln 182 182 λ ɛ WB−ɛ WC where λ is the decay constant of 182Hf (0.078±0.002 My− 1) 182 (Vockenhuber et al., 2004), ε WC is the present day value of carbonaceous chondrites (−1.9±0.1 ε-unit) (Kleine et al., 2004), and 182 ε WA is the initial Solar System value, and in practice, this term is usually set to be the initial value inferred from refractory inclusions Fig. 4. Measured ε182W values of magmatic iron meteorites vs. their exposure ages. The (−3.47±0.20 ε-units) (Kleine et al., 2005). Model ages calculated using three lines show the model results (maximum cosmogenic effect, i.e., optimum radius Eq. (2) are shown in Table 1 and Fig. 1. Most magmatic irons have and depth) computed with the three libraries and assuming a pre-exposure ε182W value model ages relative to CAIs that are zero within uncertainties, and of −3.47. L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104 99 plateau, defining an average ε182W value of −3.47±0.20, identical to the initial value of refractory inclusions (Kleine et al., 2005). Magmatic iron meteorites from the same group are thought to sample the same planetesimal core, so they should record the same metal-silicate differentiation event and hence should have the same initial ε182W (prior to GCR exposure). The fact that ε182W is variable within a single magmatic group or sometimes within a single sample is most likely caused by different exposures to GCR. One does not expect a simple correlation between ε182W values and exposure ages within each group because the production rates of cosmogenic nuclides depend on many factors besides exposure age, such as flux of primary and secondary high-energy cosmic ray particles, the chemical composition of the iron meteorite, and the sample location within the meteorite. Thus, it is impossible to correct for the cosmogenic effect in individual samples using the exposure age information alone. Fig. 5. Schematic diagrams illustrating the method applied for correcting the Markowski et al. (2006b) proposed a method for correcting the cosmogenic effect on ε182W in magmatic iron meteorites. A) Cosmogenic effects on GCR effect using the concentration of another cosmogenic nuclide, 3He the W isotopic compositions of four iron meteorites (# 1–4) of the same magmatic 182 (Markowski et al., 2006b). They demonstrated a correlation of ε182W group. The original pre-exposure ε W value of the group is assumed to be −3.47 (a priori unknown). Exposure to GCR can only drive the ε182W values to the low side. value with cosmogenic 3He concentration for both Grant (IIIB) and The lengths of the solid arrows represent the actual shifts due to the cosmogenic effect. Carbo (IID) meteorites. For samples not too far from the pre- The solid circles represent the measured W isotopic compositions. B) The ε182W values 182 3 atmospheric surface, a positive correlation between ε W and He of all four samples were corrected for the maximum cosmogenic effect (open arrows). concentration is expected. They were thus able to extract the pre- The open symbols show the corrected W isotopic compositions. The minimum of the fi exposure ε182W values based on their model results of ε182W variation corrected values (# 2) de nes the upper limit and the maximum of the uncorrected fi ε182 3 3 values (# 1) de nes the lower bound of the pre-exposure W value of this magmatic with He. We did not attempt to correct cosmogenic effects using He iron meteorite group. concentrations for Grant and other samples since the method has several limitations. One is that meteorites are susceptible to 3He loss on heating (Schultz, 1967). Also, different types of nuclear reactions 3 are involved in the production of cosmogenic He and W isotopes. of the effect. The depth- and size-dependent production rates of Helium-3 is primarily produced by spallation processes through cosmogenic W were calculated by integrating the spectra of primary interactions of high-energy GCR particles with target elements such as and secondary particles with excitation functions of relevant nuclear 3 O, Mg, Si, Ca, Fe and Al. As a result, the maximum He production rate reactions. The spectra of primary and secondary particles were occurs at the pre-atmospheric surface of the meteorite. In contrast, the calculated using the LAHET code (Masarik and Reedy, 1994). The production/destruction of W isotopes occurs through capture of differential particle spectra depend on the solar modulation para- secondary thermal neutrons that have relatively low energy. The meter M, which is a measure of solar activity and therefore varies with production of thermal neutrons reaches a maximum at a certain space and time (Leya et al., 2000). The calculated spectra are for depth, as do the neutron capture reaction rates. As a result of these M=550 MeV (Leya et al., 2000). A GCR proton (N10 MeV) flux of 3 ε182 − 2− 1 factors, the correlation between He and W is not very robust. 2.86 cm s was adopted (Kollar et al., 2006). Three main libraries Here, a new method was developed and applied to estimate the (ENDF, JEF, and KASKAD) give cross sections for neutron induced 182 pre-exposure ε W value for each magmatic iron meteorite group, reactions relevant for the production of W isotopes and were used for which relies on the assumption that samples from the same group comparison purposes. The computed effects are not sensitive to Fe and 182 (same planetesimal core) have the same pre-exposure ε W value. Ni concentrations. Fig. 6 shows a plot of maximum cosmogenic shift in The method proceeds as follows (Fig. 5): ε182W against the pre-atmospheric radius, assuming a constant • For each group, the highest measured ε182W value is taken to exposure age of 657 My. For a given pre-atmospheric radius, the ε182 represent a lower bound on the initial ε182W. The reason is that GCR maximum possible effect on W is linearly correlated with can only shift ε182W towards lower values and the highest value exposure age, thus the effect at any exposure age can be calculated represents a minimum for the pre-exposure ε182W in the group. from Fig. 6. The magnitude of the modeled effect with the ENDF • The ε182W values of all samples belonging to a group were corrected library is about 1.52 and 1.89 times higher than those with the JEF and for maximum possible cosmogenic effects. Inevitably, all ε182W KASKAD libraries, respectively. The modeled results obtained with values must have been over corrected because we apply the the JEF library were chosen for correction because of the following maximum conceivable correction (i.e., for the optimum depth to reasons: yield the largest correction at a given exposure age and pre- (i) We computed depth profiles for Carbo (group IID; pre- atmospheric radius). Thus the least radiogenic value among the atmospheric radius 0.53 m; exposure age 850 My) and Grant corrected values must be an upper limit to the pre-exposure ε182W (group IIIB; 40 cm; 695 My) using the ENDF, JEF, and KASKAD in the group. libraries and compared these with measured profiles (Markowski Measurement uncertainties were taken into account when et al., 2006b, Table 2)inFig. 2B and C; ENDF predicts a steeper estimating the values of the lower and upper limits. The maximum slope than what is observed, while JEF and KASKAD give good fits possible cosmogenic effects in individual meteorites were computed to the measurements. by combining our model simulations of neutron capture rates on W (ii) The modeled GCR effects with the ENDF library are unrealis- isotopes, estimated pre-atmospheric radii (when available) (Voshage, tically large, e.g., the maximum effect is up to 1.5 ε for a 650 My 1984), and exposure ages of individual meteorites (see the footnotes of exposure. If we assume that the original ε182W values of iron Table 1 for references). According to Masarik (1997), the maximum − neutron capture rate on W isotopes, and hence the maximum deficit meteorites are around 3.45, the maximum cosmogenic shift in ε182W, occurs at a certain depth beneath the pre-atmospheric would have led to a present day value of ∼−5.0. As can be seen in surface. The pre-atmospheric size of the meteorite may affect the Fig. 4, such a negative value has not been measured in any iron depth where the maximum effect takes place and also the magnitude meteorite with high-precision methods. 100 L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104

Fig. 6. Model results of maximum cosmogenic effects on ε182W for meteorites with various pre-atmospheric radii and constant exposure age of 657 My.

(iii) Although JEF and KASKAD give equally good fits to the measured profiles, the former library was preferred because the uncertainties on pre-exposure ε182W based on that library are more conservative. For samples with unknown pre-atmospheric size, the largest possible correction was applied (corresponding to a pre-atmospheric radius of 0.7 m). Precise W isotope measurements reported in recent work (Scherstén et al., 2006; Markowski et al., 2006a,b) were included in our calculations and the ranges of pre-exposure ε182W values thus calculated for each magmatic iron meteorite group are shown in Table 2 and Fig. 7; these are −3.63 to −3.35, −3.43 to −3.23, −3.95 to −3.24, −3.40 to −3.28, −3.41 to −3.33, −3.35 to −3.00, −3.43 to −3.28 and −3.5 to −3.46 for IC, IIAB, IID, IIIAB, IIIE, IIIF, I VA and IVB irons, respectively. For IID irons, only two samples have been studied so far: Carbo and Hraschina. A value of −3.14± 0.12 was reported for Hraschina (Markowski et al., 2006a), thus indicating a very radiogenic lower bound. However, the depth profile of Carbo suggests a pre- exposure ε182Wof∼−3.65 (Fig. 2C). To be conservative, the least negative value of Carbo analyses was used as the lower bound (which could still have been affected significantly by GCR effect). The fact that Fig. 7. A) Metal-silicate differentiation model ages and B) accretion ages for magmatic irons relative to the time of CAI formation. The metal-silicate differentiation model ages are calculated using Eq. (2) and the estimated ranges for ε182W values corrected for exposure to GCR. In 7B, the dark gray bars represent the preferred age estimates and the light grey bars represent the conservative age estimates.

Table 2 ε182W values in magmatic iron meteorites corrected for cosmogenic effects

182 Group ε W ΔTM–S ΔT′M–S R ΔTaccretion ΔT′accretion no single calculated upper bound is lower than the uncorrected lower (My) (My) (km) (My) (My) boundary suggests that this method works as intended. The estimated ε182 IC −3.63 to −3.35 −1.2 to 1.0 −3.0 to 2.6 52 0–0.7 0–1.6 ranges of pre-exposure W values for all groups overlap with the 182 IIAB −3.43 to −3.23 0.3 to 2.1 −1.4 to 3.6 94 0.1–1.4 0–1.8 initial CAI value. No single group shows pre-exposure ε W that is IID −3.95 to −3.24 −3.4 to 2.0 −5.2 to 3.5 170 0–1.3 0–1.8 significantly more radiogenic than the CAI initial, except possibly the IIIAB −3.40 to −3.28 0.6–1.6 −1.2 to 3.1 41–50 0.4–1.1 0–1.7 IIIF irons. IIIE −3.41 to −3.33 0.5–1.2 −1.2–2.7 120 0.3–0.8 0–1.6 IIIF −3.35 to −3.00 1.1–4.6 −0.7 to 6.1 35 0.7–1.9 0–2.0 IVA −3.43 to −3.28 0.3–1.6 −1.4 to 3.2 7–18 0.15–1.1 0–1.7 4.2. Time scales of metal-silicate differentiation IVB −3.50 to −3.46 −0.2 to 0 −2.0 to 1.6 3–21 0 0–1.1 Metal-silicate differentiation model ages of magmatic groups are ΔTM–S represents the model age of metal-silicate differentiation calculated with Eq. (2) 182 − and adopting an initial ε W CAI value of −3.47. ΔT′M–S is a more conservative metal- calculated relative to the CAI initial of 3.47 (Table 2 and Fig. 7). silicate differentiation age, considering an uncertainty of ±0.2 ε in the initial value of Tight time constraints were obtained for the IC, IIAB, IIIAB, IIIE, IVA CAIs. R stands for the conventional value of the radius of iron meteorite parent body and IVB iron meteorite groups, indicating that core formation on calculated from metallographic cooling rates, based on models without regolith (Haack their parent bodies very likely occurred within 1 to 2 My of the et al., 1990). ΔTaccretion represents the accretion age of individual iron meteorite parent formation of CAIs. Not a single group shows a resolvable negative bodies derived from corresponding curves in Fig. 8 and using ΔTM–S. ΔT′accretion is a more conservative accretion time estimate (see text). model age relative to CAIs. The IIIF group is the only one that stands L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104 101 out with metal-silicate differentiation possibly occurring as late as modeling purposes. It is assumed here that the planetesimals have the 5 My after CAI formation. composition of H-type ordinary chondrites, which was used in a previous thermal model for the HED parent body (Ghosh and 4.3. Modeling the accretion time scales of iron meteorite parent bodies McSween, 1998). Using an Al concentration of 1.13 wt.% (Wasson, 1988), the calculated initial heat production for a planetesimal formed As stated in the above section, all magmatic iron meteorite parent at t=0, is 1.9×10− 7 W/kg. The model parameters are summarized in bodies underwent rapid metal-silicate differentiation. In the case of Table S1. most magmatic groups, this event was predated by the condensation It is assumed that core formation is instantaneous and corre- of refractory inclusions in the early Solar System by at most 5 My. The sponds to the time when the center of the parent body reaches the Hf–W model ages can be combined with thermal modeling to eutectic temperature for the Fe–FeS system. This eutectic tempera- constrain the accretion timescales of their parent bodies. ture at 1 bar is 1261 K (Fei et al., 1997). Electrical conductivity For a spherical planetesimal, the temperature T at any radial measurements show that Fe–S alloy will segregate at temperatures distance R and at any time t after its formation is governed by the heat high enough to melt the alloy (∼1261 K), but below the silicate conduction equation: solidus (Yoshino et al., 2003). The segregation velocity is rapid in comparison with the timescale of core formation (Yoshino et al., AT 1 A AT ArðÞ; t – ¼ R2κ þ ð3Þ 2003). The trigger of Fe S melting is set at the center, because for all A 2 A A t R R R C planetesimals with radii N50 km, Fe–FeS melting at various depths where C is the specific heat capacity, A(r,t) is the heat production per takes place at about the same time, except for an outermost shell of a “ ” κ κ ¼ K few kilometers in thickness. The proportion of this cold shell to the unit mass, and is the thermal diffusivity ( Cρ, where K is the thermal conductivity and ρ is the density). One effective and attested whole planetesimal radius increases with decreasing planetesimal size. heat source for the differentiation of early formed planetesimals is the decay of short-lived radionuclides, such as 26Al (Lee et al., 1976) and Planetesimals accreted at different times have different initial heat 60 production rates At . The time that it takes to initiate Fe–FeS melting Fe (Shukolyukov and Lugmair, 1993; Tachibana and Huss, 2003; 0,Al 26 (onset of core formation) can be calculated analytically using Eq. (5) Mostefaoui et al., 2005; Tachibana et al., 2006). With an initial Al/ 27 − 5 (Fig. 8). For planetesimals accreted at the same time, the larger the Al ratio of 5×10 (Macpherson et al., 1995; Russell et al., 1996; Huss radius is, the sooner the core forms. For planetesimals of b∼5.2 km et al., 2001), a half-life of 0.73 My and a decay energy of 3 MeV/atom, 26 radius, melting of Fe–FeS will never occur, unless a regolith (with very Al could have been an important heat source in the first few million low thermal diffusivity) is present to reduce heat loss (Haack et al., years of the Solar System. Iron-60 (t1/2 =1.49 My) is another possible 1990). A general feature of the melting-accretion relation in Fig. 8 is heat source but its initial abundance is under great debate. Given 60 56 that for each radius, there is a “critical accretion time”, after which the the inferred low and uncertain initial Fe/ Fe ratio (on the order of − 7 planetesimals will not undergo metal-silicate differentiation. The 5×10 , Tachibana and Huss, 2003; Mostefaoui et al., 2005; Tachibana critical accretion time is sensitive to the radius for smaller planeste- et al., 2006; Cook et al., 2006), it is considered separately. Thus, the simals (Rb20 km). Planetesimals formed later than 2.4 My after CAIs term A in Eq. (3) can be represented mathematically as

−λ ðÞ¼; Al t; ð4Þ Art At0;Al e

− λ λ 26 Alt where Al is the decay constant of Al. At0,AL =A0,Ale 0 is the heat 26 production by Al decay at the time of planetesimal formation, t0, where A0,A1 is the heat production at the time of CAI formation. κ, ρ, and C are all temperature-dependent (Yomogida and Matsui, 1983; Ghosh and McSween, 1999). Since we are only interested in the ﬁrst- order correlation between accretion and differentiation timescales, appropriate constant values have been assigned to these parameters. In that case, Eq. (3) can be solved analytically (Carslaw and Jaeger, 1959; Miyamoto et al., 1981):

κ 1=2 At ;Al −λ R sinðrðÞλAl=κ Þ T ¼ T þ 0 e Alt ð −1Þ ð5Þ 0 κλ ð ðÞλ =κ 1=2Þ Al r sin R Al 3 ∞ n 2R A ; ðÞ−1 nπr 2 2 2 þ t0 Al ∑ sin e−κn π t=R π3 ðÞ2−λ 2=κπ2 r K n¼1 nn R R

The values chosen for κ and ρ in (LaTourrette and Wasserburg, 1998; Hevey and Sanders, 2006) were used here. For C, a value of 800 J/kg− 1K− 1 was used (Ghosh and McSween, 1999). T is the initial Fig. 8. Timing of metal-silicate differentiation vs. accretion age based on thermal 0 modeling of the evolution of asteroids with radii between 5.2 and 100 km. The numbers temperature of the planetesimal, which is equal to the ambient adjacent to the solid lines are the radii in kilometers. Given a differentiation age (from temperature. Eq. (5) applies to a system with a fixed boundary 182Hf–182W systematics) and a parent body size (from metallographic cooling rates), it is condition, where the surface temperature is always equal to the possible to derive the accretion age (Table 2, Eq. (5)). Note that the adiabatic curve for – ambient temperature. The ambient temperature varies with time and incipient melting of Fe FeS (computed assuming no heat loss, Eq. (8)) provides a robust upper limit on the accretion age given the differentiation age. The long dashed curve heliocentric distance. A value of 250 K was adopted, which is the same represents adiabatic melting when the initial 60Fe/56Fe=5×10− 7. For the range of metal- as that used in (Hevey and Sanders, 2006), and is within the range of silicate differentiation ages obtained for most iron meteorite groups, the effect on expected temperatures (100–400 K) in the mid-plane of the solar disk inferred accretion time scales of adding such an amount of 60Fe is very small. The at 2.5 AU and at 1 My after formation of the Sun (Woolum and Cassen, dashed-dotted line is adiabatic melting with an Al content of 1.75%. All other curves are 1999). The chemical compositions of iron meteorite parent bodies for modeled using an Al content of 1.13%. The dotted line represents the extreme case in which the time of differentiation is equal to that of accretion. The forbidden zone lithophile elements like Al are uncertain. Previous studies have used reflects the fact that after accretion, differentiation cannot proceed faster than in the Al concentrations of both carbonaceous and ordinary chondrites for case of adiabatic heating. 102 L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104 will never undergo metal-silicate differentiation, regardless of their any given 182Hf–182W metal-silicate differentiation age and any parent sizes. body radius. A conservative accretion time, denoted as ΔT′accretion in The calculations show that for large planetesimals of ≥50 km Table 2, was calculated taking into account both the thermal model radius, melting in the interior is adiabatic. In that case, the time that it uncertainties and an uncertainty of ±0.2 ε in the initial ε182W value of takes to initiate melting in the Fe–FeS system is determined by the the CAIs (−3.47). The accretion times of all magmatic groups are, for energy balance between radiogenic heat produced by decay of 26Al the most part, within 1 to 2 My after the formation of the Solar System. and change in internal energy, It should be noted that the model did not take into account 60Fe as a heat source. The abundance of 60Fe in the planetesimal will affect the ðÞ ¼ ð6Þ − Atdt CdT time of melting. If the initial 60Fe/56Fe ratio is about 5×10 7, the Integrating both sides of Eq. (6) from accretion to core formation, it adiabatic curve will shift towards the right (Fig. 8). The latest accretion follows, time for meteorite parent bodies of any size that would undergo metal-silicate differentiation is ∼2.4 My (Fig. 8). For the range of −λ ∫tC Alt ¼ ∫Teut ; ð7Þ metal-silicate differentiation ages reported for most iron meteorite t A0;Ale dt T CdT 0 0 groups, the effect on inferred accretion time scales of adding such an 60 where tC is the time of core formation. T0 and Teut are the initial amount of Fe is very small. The Al content in iron meteorite parent temperature and eutectic temperature of the Fe–FeS system. tC can be bodies is also uncertain. The highest value of 1.75 wt.% among all written as: chondrite groups (Wasson, 1988) shifts the adiabatic curve to the right, corresponding to a critical accretion time of ∼2.5 My (Fig. 8). 1 −λ ðÞTeut−T0 ¼ − t0 − λ : ð8Þ Our results show that accretion of iron meteorite parent bodies tC λ ln e C A0;Al occurred very early, within 1–2 My of CAI formation. Although the absolute accretion rate of planetesimals in the Solar System is not very Using metallographic cooling rates, Chabot and Haack (2006) well constrained, studies have shown that the accretion timescale recalculated the radii of iron meteorite parent bodies for several increases with increasing heliocentric distance (Weidenschilling, groups of irons, when taking into account the presence of a regolith 1977; Greenberg et al., 1978; Grimm and McSween, 1993). Iron with a thickness of 0.003 times the radius. The new calculated radii are meteorite parent bodies are usually assumed to have formed in the approximately 2-fold lower than conventional values obtained with- main asteroid belt. However, Bottke et al. (2006) showed that out a regolith (Haack et al., 1990). However, we did not incorporate a differentiated meteorite parent bodies may not be indigenous to the regolith in our thermal modeling and used conventional radii instead. main asteroid belt but may be interlopers from the terrestrial planet The effect of this omission to the model results (Fig. 8) is discussed formation region, where fast accretion rates allowed small planete- below. simals to melt early. The very early accretion ages of most magmatic For each magmatic group, the size of the parent body is known and iron meteorite parent bodies implied by the fast core differentiation thus the corresponding evolution curve can be found in Fig. 8. With times are consistent with this scenario. the metal-silicate differentiation age inferred from 182Hf-182W systematics, constraints can be put on the accretion time, denoted 4.4. Clues to the origin of non-magmatic iron meteorites as ΔTaccretion in Table 2. Our preferred accretion timescales are within ∼1.5 My of the formation of CAIs for most magmatic iron groups. The origins of non-magmatic groups IAB–IIICD and IIE are under Several sources of uncertainties for the accretion time need to be debate and it is not clear whether or not samples from the same group addressed: should possess a single ε182W value. IABs and IIICDs belong to a single (i) The differentiation ages computed with the thermal model group of non-magmatic meteorites (Wasson and Kallemeyn, 2002). – correspond to incipient melting of Fe–FeS at the center of the Three out of four analyses of IAB IIICD iron meteorites studied present ε182 − − planetesimal (onset of core formation) while 182Hf–182W model homogeneous W values of 3.01 to 2.80 (Table 1; Fig. 1). This is ages correspond to large-scale metal-silicate differentiation. consistent with previous studies (Scherstén et al., 2006; Markowski The calculated accretion ages based on Fig. 8 and 182Hf–182W et al., 2006a). Markowski et al. found that some IIICDs have resolvably ∼− ∼− model ages must therefore represent upper limits. lower values ( 3.3) than IABs ( 3.0) and concluded that IIICD irons (ii) In the thermal model, the possible existence of a regolith was formed earlier (contemporaneous with CAIs) than IABs. The only IIIC ε182 − not incorporated. Adding a thin regolith will decrease the time sample from this study, Nantan, yields a W value of 2.93±0.07, required to melt the metal and cause the curves in Fig. 8 to shift which is identical with those of the two IABs. It is clear that its W to the right for small parent body sizes (towards the adiabatic isotopic composition is not affected by the cosmogenic effect given its limit). This will increase the accretion age for a given metal- short exposure age and extremely heavy shielding (Nishiizumi et al., silicate differentiation age and a given parent body radius. 1995). However Nantan is one of the most weathered iron meteorites However, the radius of the parent body derived from metallo- and terrestrial contamination associated with weathering cannot be graphic cooling rates with a thin regolith is smaller than the excluded, even though the sample was intensively leached before ε182 − value used here (computed without a regolith). This will shift dissolution. Deport presents the most unradiogenic W( 3.82) – the accretion time in the opposite direction. For large parent among all IAB IIICDs, which clearly resulted from exposure to GCR bodies, adding a regolith will not affect the model results at all irradiation (given its long exposure age of 1140 Ma). Wasson and – since the interiors will melt adiabatically. The bottom line here Kallemeyn (2002) distinguished six subgroups of IAB IIICD irons. is that the adiabatic curve represents a robust upper limit for They suggested that melting occurred because of impact heating and the accretion times. that the slight compositional difference between the main group (MG) (iii) Based on the diverse cooling rates that have been measured for and low-Au subgroups (SLL, SLM, SLH) resulted from independent IVA iron meteorites (Goldstein and Short, 1967; Rasmussen, impact events that occurred either at different locations of a single 1982; Yang et al., 2007), Yang et al. (2007) inferred a larger chondritic body or different chondritic bodies with similar composi- – parent body size of 300-km diameter. Thus, melting in the IVA tions (Wasson and Kallemeyn, 2002). Most of the IAB IIICD samples parent body would be closer to adiabatic melting. studied for W isotopes are from the main group (MG) and the subgroup SLL. The similarity of W isotope compositions of IAB and In brief, when model uncertainties are considered, the accretion IIICD irons is most easily explained with a single impact event dated at time inferred from the adiabatic curve gives a robust upper limit for 4–7 My after CAI formation causing multiple metal pools in a single L. Qin et al. / Earth and Planetary Science Letters 273 (2008) 94–104 103 chondritic parent body. Alternatively, there may have been multiple thank H. Palme and K. Mezger for their insightful reviews of the impacts occurring within a narrow time interval of 4–7 My after CAI manuscript. This work was supported by the France-Chicago Center, a formation on multiple chondritic bodies with similar chemical Packard Fellowship and NASA grants NNG06GG75G (to N. D.) and compositions. NNG05GG22G (to M. W.). For the other non-magmatic group IIE, Watson has the most radiogenic signature (ε182W=−2.30±0.08) among all iron meteorites Appendix A. Supplementary data measured thus far (also see Snyder et al., 2001). The other sample Arlington has a much lower ε182Wof−3.01±0.08. With an exposure Supplementary data associated with this article can be found, in age of only 150 My, it is unlikely to have been affected by cosmogenic the online version, at doi:10.1016/j.epsl.2008.06.018. effects (Lavielle et al., 1999). IIE iron meteorites also contain abundant silicate inclusions. The heterogeneity in ε182W values of IIEs is References consistent with a formation model that involves melting and metal segregation followed by partial mixing of metal and silicate by a Ammon, K., Masarik, J., Leya, I., 2006. Cosmogenic production rates in iron meteorites. Meteorit. Planet. Sci. 41, A16. secondary thermal event (Bogard et al., 2000). Thus, as has been Bogard, D.D., Garrison, D.H., McCoy, T.J., 2000. 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Conclusions 510 pp. Chabot, N.L., Haack, H., 2006. Evolution of asteroidal cores. In: Lauretta, D.S., McSweenJr. Jr., H.Y. (Eds.), Meteorites and the Early Solar System, vol. II. The University of The present study demonstrates that exposure to galactic cosmic Arizona, Tucson, AZ, pp. 747–771. rays can create strong negative shifts in ε182W values by capture of Chen, J.H., Wasserburg, G.J., 1996. Extinct 107Pd in the early solar system and secondary thermal neutrons. This agrees with most recent studies implications for planetary evolution. In: Basu, A., Hart, S. (Eds.), Earth Processes: Reading the Isotopic Code, Geophysical Monograph, vol. 95. AGU, Washington, DC, (Kleine et al., 2005; Scherstén et al., 2006; Markowski et al., 2006a,b). pp. 1–20. After correction of these effects, the ε182W values in various magmatic Cook, D.L., Wadhwa, M., Janney, P.E., Dauphas, N., Clayton, R.N., Davis, A.M., 2006. High groups (except IID and IIIF) are confined to a narrow range of −3.63 to precision measurements of non-mass-dependent effects in nickel isotopes in meteoritic metal via multicollector ICPMS. Anal. Chem. 78, 8477–8484. − 3.23. This implies that core segregation processes in most magmatic Fei, Y., Bertka, C.M., Finger, L.W., 1997. High-pressure iron-sulfur compound, Fe3S2, and iron meteorite parent bodies occurred b2 My after the formation of the melting relations in the Fe–FeS system. Science 275, 1621–1623. Solar System. This result also agrees with earlier studies (Scherstén Foley, C.N., Wadhwa, M., Borg, L.E., Janney, P.E., Hines, R., Grove, T.L., 2005. The early differentiation history of Mars from 182W–142Nd isotope systematics in the SNC et al., 2006; Markowski et al., 2006a). However, IIIF iron meteorites meteorites. Geochim. Cosmochim. Acta 69, 4557–4571. tend to have more radiogenic ε182W compositions, which may imply a Ghosh, A., McSween, H.Y., 1998. 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