Highly siderophile element and osmium isotope evidence for postcore formation magmatic and impact processes on the aubrite parent body


  • David van ACKEN,

    Corresponding author
    1. Department of Earth and Atmospheric Sciences, University of Houston, 312 Science & Research 1, Houston, Texas 77204–5007, USA
    2. NASA-Lyndon B. Johnson Space Center, 2101 NASA Parkway, Houston, Texas 77058, USA
    3. Department of Earth and Atmospheric Sciences, University of Alberta, 1-26 ESB, Edmonton, Alberta T6G 2E3, Canada
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  • Alan D. BRANDON,

    1. Department of Earth and Atmospheric Sciences, University of Houston, 312 Science & Research 1, Houston, Texas 77204–5007, USA
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  • Thomas J. LAPEN

    1. Department of Earth and Atmospheric Sciences, University of Houston, 312 Science & Research 1, Houston, Texas 77204–5007, USA
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Corresponding author. E-mail: vanacken@ualberta.ca


Abstract– Aubrites exhibit a wide range of highly siderophile element (HSE—Re, Os, Ir, Ru, Rh, Pt, Pd, Au) concentrations and 187Os/188Os compositions. Their HSE concentrations are one to three orders of magnitude less than chondrites, with the exception of the Shallowater and Mt. Egerton samples. While most aubrites show chondritic HSE abundance ratios, significant enrichments of Pd and Re relative to Os, Ir, and Ru are observed in 12 of 16 samples. Present-day 187Os/188Os ratios range from subchondritic values of 0.1174 to superchondritic values of up to 0.2263. Half of the samples have 187Os/188Os ratios of 0.127 to 0.130, which is in the range of enstatite chondrites. Along with the brecciated nature of aubrites, the HSE and Re-Os isotope systematics support a history of extensive postaccretion processing, including core formation, late addition of chondritic material and/or core material and potential breakup and reassembly. Highly siderophile element signatures for some aubrites are consistent with a mixing of HSE-rich chondritic fragments with a HSE-free aubrite matrix. The enrichments in incompatible HSE such as Pd and Re observed in some aubrites, reminiscent of terrestrial basalts, suggest an extensive magmatic and impact history, which is supported by both the 187Re-187Os isotope system and silicate-hosted isotope systems (Rb-Sr, K-Ar) yielding young formation ages of 1.3–3.9 Ga for a subset of samples. Compared with other differentiated achondrites derived from small planetary bodies, aubrites show a wide range in HSE concentrations and 187Os/188Os, most similar to angrites. While similarities exist between the diverse groups of achondrites formed early in solar system history, the aubrite parent body(ies) clearly underwent a distinct evolution, different from angrites, brachinites, ureilites, howardites, eucrites, and diogenites.


Aubrites, or enstatite achondrites, are mostly brecciated stony meteorites that formed under highly reducing conditions compared with other meteorite groups (IW-6 to IW-8; e.g., Mittlefehldt et al. 1998). They consist mainly of enstatite (ferrosillite content 0.4–1.3%; Watters and Prinz 1979), plagioclase, diopside, olivine, as well as minor amounts of FeNi metal, troilite, and exotic sulfides. The sulfides such as alabandite (MnS), oldhamite (CaS), and daubreelite (FeCr2S4) contain nominally lithophile elements as major components. Based on a variety of lithic clasts with igneous textures, an igneous origin for the aubrite precursor material is generally favored (Watters and Prinz 1979; Biswas et al. 1980; Wolf et al. 1983; Wheelock et al. 1994), as opposed to earlier suggestions of an origin as condensates from the solar nebula (Wasson and Wai 1970; Sears 1980). While the breccia fragments have been interpreted as products of fractional crystallization (Watters and Prinz 1979; Okada et al. 1988), their specific igneous histories are largely unknown. Along with enstatite chondrites, aubrites are among the few meteorites whose oxygen isotopic compositions fall on the terrestrial fractionation line (Clayton et al. 1984), raising interest in enstatite meteorites as being among the potential building blocks of Earth (e.g., Javoy et al. 2010). The details of the genetic linkages between aubrites and enstatite chondrites are still debated. Some studies suggest a formation of aubrites from igneous processing of either EL or EH material with a basaltic partial melt component being removed (Watters and Prinz 1979; Biswas et al. 1980; Brett and Keil 1986; Casanova et al. 1993; Wheelock et al. 1994), while other studies suggest an origin on different, compositionally similar parent bodies (Keil 1969, 1989; Brett and Keil 1986; Keil et al. 1989). Aubrites appear to have formed on one or more parent body(ies) after core formation (Casanova et al. 1993), which is supported by subchondritic highly siderophile element (HSE) concentrations (Wolf et al. 1983) and consistent with removal of metal into a core (e.g., Righter 2003).

While there have been petrologic and lithophile element and isotope investigations of bulk aubrites (Keil 1969; Wolf et al. 1983; Easton 1985; Floss et al. 1990; Lodders et al. 1993), few studies have attempted to characterize their HSE concentrations (Ru, Rh, Pd, Re, Os, Ir, Pt, Au) and distributions (Wolf et al. 1983; Keil et al. 1989; Casanova et al. 1993). The HSE partition preferentially into metal phases or sulfides with Dmetal/silicate typically greater than 106 (e.g., Kimura et al. 1974; O’Neill et al. 1995; Righter 2003; Brenan and McDonough 2009), and thus offer a valuable complementary geochemical tool to lithophile trace elements such as the rare earth elements. The HSE include the long-lived Re-Os radiogenic isotope system (λ187Re = 1.6668 · 10−11 a−1; Selby et al. 2007). The processes of planetary accretion and differentiation have been successfully constrained by the use of HSE and Os isotopes in a variety of terrestrial and extraterrestrial samples, but have not been applied to aubrites (e.g., Morgan 1986; Brandon et al. 2000, 2005a, 2012; Morgan et al. 2001; Walker et al. 2002; Horan et al. 2003; Day et al. 2007, 2012; Rankenburg et al. 2007, 2008; Walker 2009; Dale et al. 2012; Riches et al. 2012).

Chronology of aubrites and the aubrite parent body(ies) is not well established, and different isotope systems yield different results. Ages for Shallowater (I-Xe age: 4.563 ± 0.001 Ga; Gilmour et al. 2006), Khor Temiki (I-Xe age: 4.566 ± 0.002 Ga; Busfield et al. 2008), Bishopville (I-Xe age: 4.54 ± 0.01 Ga; Podosek 1970), and Peña Blanca Spring (Mn-Cr age: 4.563 ± 0.003 Ga; Shukolyukov and Lugmair 2004) confirm formation in the earliest stages of the solar system. Less precise Rb-Sr and Ar-Ar ages confirm formation of most aubrites around 4.5 Ga (Norton County, Rb-Sr: 4.7 ± 0.1 Ga, Ar-Ar: 4.2-4.5 Ga, Bogard et al. 1967, 2010; Norton County, Rb-Sr: 4.48 ± 0.04 Ga, Minster and Allegre 1976; Shallowater, Ar-Ar: 4.53 Ga, McCoy et al. 1995). However, some samples, notably Bishopville, Cumberland Falls, and Norton County, show signs of later disturbance, resulting in younger Rb-Sr and Ar-Ar ages or precluding obtainment of statistically meaningful ages (Compston et al. 1965; Bogard et al. 2010). Because Re and Os are hosted in trace phases that may be more or less affected by processes that can disturb isotope characteristics in silicates, the long-lived 187Re-187Os decay system may add some complementary insight in the formation and differentiation history of aubrites and their parent body(ies).

Existing HSE concentration data for aubrites show variation over several orders of magnitude from 0.001 to 1 times chondritic values (Wolf et al. 1983; Keil et al. 1989). However, these older data sets include some samples that are no longer considered aubrites (Bencubbin, Horse Creek) and do not report concentrations of Pt or Ru. Measurements of 187Os/188Os for a comprehensive data set of aubrites have not been reported to date. Here, we present the first data set for Os isotopic compositions of 16 aubrites, including the anomalous nonbrecciated and metal-rich samples Mount Egerton and Shallowater, along with a comprehensive overview of their HSE concentrations. These data are used to address aubrite formation by a combination of igneous parent body processes and planetary processes, such as core formation, silicate differentiation, parent body breakup and reassembly, and inclusion of chondritic fragments, as well as achondrite material from other differentiated bodies.


Fusion crust and sample chips containing terrestrial alteration evident from Fe-oxides were discarded prior to grinding of samples in an Al2O3 mortar and pestle. For analysis, approximately 20–500 mg of sample powder was digested in quartz glass Carius tubes for a minimum of 60 h at 230 °C with 3 ml 12 N HCl and 6 ml 14 N HNO3. A mixed 99Ru-110Pd-185Re-190Os-191Ir-198Pt spike (UChicago #000601) was added prior to digestion. Total procedural blanks were determined using a diluted 99Ru-105Pd-185Re-190Os-191Ir-198Pt spike that was processed exactly as a sample. After opening the tubes, Os was extracted using a CCl4/HBr liquid extraction procedure following the methods of Cohen and Waters (1996). Osmium was further purified using a H2SO4/H2CrO4 micro-distillation procedure (Birck et al. 1997). After Os extraction, the remaining HCl-HNO3 fraction containing all other HSE was dried down. Samples were converted to chlorides by repeated dissolution in 6 N HCl and drying down before final dissolution in 0.2 N HCl. For HSE separation, a cation column with Eichrom 5 × 8 100–200 mesh resin was used. Subsequent to column separation, cuts containing the HSE were dried down and then dissolved in 3 ml 2% HNO3.

The Os fractions were loaded in HBr on ESPI™ Pt filaments and measured as OsO3 on a ThermoFinnigan Triton at the Johnson Space Center, Houston, TX. Measurements were conducted on either Faraday cups in static mode or SEM in dynamic mode, depending on signal intensity. Repeated measurements of 70–105 ng loads of the UMD Johnson Matthey Os standard were conducted to monitor reproducibility of isotope ratio measurements and cup drift. Over the course of the measurement campaign (April 2009–November 2009, n = 41), the measured 187Os/188Os was 0.1137915 ± 20, with 192Os/188Os = 3.083 (Brandon et al. 2005b) as the reference for the instrumental mass fractionation correction, using an exponential law.

Isotope dilution measurements for HSE concentrations were conducted on the Varian 810 quadrupole ICP-MS at the University of Houston in high sensitivity mode. Analyses involved 50 integrations with 10 scans of each peak per integration on masses 90Zr, 97Mo, 99Ru, 100Ru, 101Ru, 102Ru, 103Rh, 105Pd, 106Pd, 108Pd, 110Pd, 111Cd, 178Hf, 185Re, 187Re, 189Os, 191Ir, 193Ir, 194Pt, 195Pt, 196Pt, 198Pt, and 199Hg. Rinse time between samples was 240 s; sample uptake time was 50 seconds. On-peak zeroes were monitored for the course of the analytic campaign, and were insignificant. An in-house multi-element standard solution was measured every five samples to monitor machine drift and instrumental mass fractionation. Instrumental mass fractionation was corrected with the bracketing standard measurements. Oxide species were monitored during instrument tune-up and CeO/Ce were typically <0.01.

Total procedural blanks were processed with each batch of samples, and totaled 1.4 ± 0.8 pg for Os (1 SD), with 187Os/188Os of 0.17 ± 0.03 (1 SD). One blank had significantly more radiogenic 187Os/188Os of 0.73 with 0.9 pg of Os. Other HSE blanks varied strongly between batches of Carius tubes used, and were 17 ± 10 pg for Ru, 11 ± 10 pg for Ir, 21 ± 19 pg for Pt, 70 ± 40 pg for Pd, and 16 ± 9 pg for Re (uncertainties as 1 SD). Measured sample values were corrected for an average of all blanks obtained during the measurement campaign. This resulted in blank corrections below 15% for most samples (Table S2), with some of the low abundance samples (Peña Blanca Spring, Mayo Belwa, Bishopville) as notable exceptions, with blank corrections for single elements >50%. Reproducibility for HSE concentrations is hard to constrain due to sample heterogeneity and nugget effect. Additionally, for the low concentration aubrites (ALHA78113, Bishopville, Bustee, Khor Temiki, LAP 03719, LAR 04316, Norton County, and Peña Blanca Spring), the uncertainty in blank values for Pt, Pd, and especially Re factors in critically, resulting in uncertainties of up to 50% for Re/Os ratios in these samples.


Results for 16 aubrite samples are presented in Table 1. Because of the expected heterogeneity due to the brecciated nature of aubrites, duplicate analyses of sample masses between approximately 20 and 500 mg were conducted for almost all meteorites. Concentrations range over several orders of magnitude, from 0.31 to 474 ppb for Os, 0.14 to 403 ppb for Ir, 0.38 to 710 ppb for Ru, 0.24 to 865 ppb for Pt, 0.30 to 694 ppb for Pd, and 0.10 to 37.3 ppb for Re. The samples Shallowater and Mt. Egerton represent the upper end of the concentration range, while Peña Blanca Spring and Bustee contain the least HSE. Earlier studies that reported HSE concentrations (Os, Ir, Pd, Re) in aubrites and their components analyzed by INAA (Wolf et al. 1983; Keil et al. 1989) found similar results in terms of sample heterogeneity and range of HSE concentrations over the suite of samples. For Shallowater, Keil et al. (1989) reported Ir concentrations between 280 and 710 ppb, which overlap with the values of 218–325 ppb obtained in this study. Wolf et al. (1983) reported Ir concentrations between 0.09 and 0.72 ppb for Bishopville, compared to 0.26–0.74 ppb in this study, 0.822–1.12 ppb for Aubres (3.89 ppb in this study), 0.0058–0.007 for Norton County (0.45–13.8 ppb in this study) and 0.0207–0.81 for Peña Blanca Spring (0.14–4.2 ppb in this study). This variation highlights the nugget effect that is also seen in several samples from the present study (Bustee, Peña Blanca Spring, Mt. Egerton, Norton County). This is especially evident when duplicate analyses of the same seemingly homogenized powder aliquot show HSE concentrations different by an order of magnitude or more, yet with only moderate variations in element ratios of HSE, as seen, for example, in Peña Blanca Spring (0.31–5.5 ppb Os, Os/Ir = 1.3–2.21), Norton County 0.497–18.9 ppb Os, Os/Ir = 1.11–1.36), and Mt. Egerton (3.0–474 ppb Os, Os/Ir 1.13–1.14; Table 1). On the other hand, HSE concentrations in five samples are more reproducible between different powder aliquots (e.g., Larkman Nunataks 04316 and LaPaz Icefield 03719; Tables 1 and S2). Four samples (ALHA 78113, ALH 84008; Bishopville, LAP 03719) show enrichments in Pd and Re, that could not be reproduced between replicates (Fig. 1; Table 1). Uncertainties on individual measurements are thus controlled by counting statistics (less than 5% uncertainty) and the significant blank contributions of >80% on low concentration samples (see Table S2). The heterogeneity in bulk HSE concentrations can be attributed to heterogeneous distribution of isolated Fe-Ni metal grains on the thin section scale (Wasson and Wai 1970; Graham 1978; Watters and Prinz 1979; Easton 1986; Casanova et al. 1993; van Acken et al. 2012).

Table 1.   Highly siderophile element concentration and Os isotopic composition of aubrites.
SampleInstitutionSpecimenFall/findbrecciatedOs(ppb)Ir(ppb)Ru(ppb)Pt(ppb)Pd(ppb)Re(ppb)Re/Os 187Os/188Os2 s.e. 187Re/188Os 187Os/188Os (4.56Ga)Sample weight (g)
  1. Institutions: MWG/JSC: Meteorite Working group at Lyndon B. Johnson Space Center, NASA, Houston, TX; NHML: Natural History Museum London; ASU: Arizona State University, Tempe, AZ; USNM: National Museum of Natural History, Washington, DC; UNM: University of New Mexico, Albuquerque, NM; WAM: Western Australian Museum, Perth; TCU: Monnig Collection, Ft. Worth, TX

  2. reg. indicates regolith breccia.

  3. Data for Shallowater was previously published in van Acken et al. (2011).

Duplicate         1.971.953.733.922.110.180.0920.127990.000020.4410.09310.52387
ALH 84007MWG/JSC,112Findy33.327.455.051.435.93.950.1190.122450.000020.5720.07730.22438
ALH 84008MWG/JSC,84Findy25.318.833.
ALH 84009MWG/JSC,22Findy5.063.959.39.910.92.230.4410.128240.000032.13−0.03960.23138
Duplicate         1.000.741.
BusteeNHMLBM.32100Fally (reg.)0.3450.340.72 4.10.310.8990.22630.00014.39−0.12030.27505
Duplicate         1.501.423.338.62.420.270.1800.133120.000120.8680.06460.46358
Khor TemikiNHMLBM.1934,781Fally (reg.)3.323.
Duplicate 0.110.0510.12020.00040.2450.10080.25305
LAP 03719MWG/JSC,20Findn8.57.210.615.34.90.640.0750.128070.000080.3640.09930.10538
Duplicate         12.410.410.415.
LAR 04316MWG/JSC,35Findy2.−0.28870.21698
Duplicate         2.753.455.66.410.30.520.1890.170040.000110.9180.09760.29930
Mayo Belwa darkUSNMUSNM 5873Fally2.632.474.055.11.783.681.4020.128850.000086.76−0.40490.18308
Repeat    2.53      0.129520.00018   
Mayo Belwa whiteUSNMUSNM 5873Fally0.7980.460.700.731.051.802.2560.117430.0001110.87−0.74040.16521
Mt. EgertonWAM Falln47440370186569437.30.0790.129870.000010.3800.09990.06241
Duplicate         93.582.316517286.57.50.0800.129960.000060.3880.09940.02355
Duplicate         3.02.687.253.728.50  0.130170.00001  0.18404
Norton CountyUNM9491-95Fally0.4970.451.291.676.90.110.2210.136860.000111.0680.05250.19397
Duplicate         18.913.850.930.
Peña Blanca SpringTCU Fally0.3100.140.390.240.30  0.14240.0003  0.20575
PesyanoeUSNMUSNM 1425Findy (reg.)     0.18607
Duplicate         9.807.86.97.513.50.440.0450.122360.000030.2160.10530.36928
Duplicate         2362183474452716.30.0270.128700.000020.1290.11850.01751
Figure 1.

 CI-chondrite-normalized HSE abundance patterns; a) samples with mostly chondritic relative highly siderophile element (HSE) abundances; b) samples showing deviation from chondritic relative HSE abundances in one or more elements; CI chondrite concentrations from Horan et al. (2003).

Chondrite-normalized HSE patterns are predominantly flat with some noted exceptions (Fig. 1). The samples ALHA78113, ALH 84009, Bishopville, Bustee, Khor Temiki, LAP 04316, Mayo Belwa (both white inclusion and dark matrix), and Norton County show distinct enrichments in Re relative to average chondrites, with Re/IrN ratios of greater than 3 and up to 56 (Table S2; Fig. 2). In contrast, Shallowater is depleted in Re with Re/IrN as low as 0.35. Palladium is enriched over compatible HSE (e.g., Ir) with Pd/IrN ratios that are greater than 3 in Bustee, Norton County, Bishopville, and Pesyanoe, with a maximum value of 12 in Norton County. In contrast, Pd is depleted relative to compatible HSE with Pd/IrN below 0.7 in ALH 84008, LAP 03719, and dark matrix in Mayo Belwa.

Figure 2.

 Highly siderophile element (HSE) versus Ir, all concentrations in ppb. Parts (f–j) represent details of parts a–e, respectively. R2-values for parts (f–j) calculated without high concentration samples controlling the R2-values in (a–e). Closed symbols: this study, open symbols: data from Wolf et al. (1983). High degrees of correlation support mineralogical control of all HSE by one phase, presumably Fe-Ni metal.

Interelement HSE concentration correlations for the aubrite suite are generally well defined (R> 0.8, Fig. 2) and the slopes of the covariations are controlled by the high HSE samples Shallowater and Mt. Egerton. Correlations of Os, Pt, Ru, and Ir concentrations of only the low-HSE portion of the aubrite sample set have R> 0.94 (Fig. 2). Correlations between Pd and Re and Ir are less with R2 approximately 0.76 and 0.4, respectively. Because of heterogeneous HSE distribution within each sample (Wasson and Wai 1970; Graham 1978; Watters and Prinz 1979; Easton 1986; Casanova et al. 1993; van Acken et al. 2012) and the resulting nugget effect (e.g., Mt. Egerton), duplicates of aubrite samples are plotted as individual samples on all plots.

Measured 187Os/188Os ratios range between 0.1174 ± 0.0001 (2σ) for Mayo Belwa (white inclusion) and 0.2263 ± 0.0001 for Bustee (Table 1). Sixteen of 31 samples have 187Os/188Os ratios that are comparable to enstatite chondrites, which have a range from 0.127 to 0.130 (Fig. 3). Fifteen samples show 187Os/188Os ratios outside the range of chondrites (Fig. 3; Table 1). Reproducibility of measured 187Os/188Os ratios of duplicate samples is typically poor for the same reasons outlined above for HSE concentrations. For example, small sample aliquots (0.2 g) of Peña Blanca Spring, LAR 04316, Bustee, Norton County, and Khor Temiki yielded 187Os/188Os of 0.1368 ± 0.0001 to 0.2263 ± 0.0001, while duplicates of larger sample aliquots (0.4–0.5 g) of the same samples yielded less radiogenic 187Os/188Os ratios of 0.1201 ± 0.0001 to 0.1331 ± 0.0001. Smaller but significant differences in duplicate measurements of 187Os/188Os ratios were observed in ALH 84007 (0.1224 ± 0.0001 versus 0.1264 ± 0.0001) and Bishopville (0.1341 ± 0.0001 versus 0.1683 ± 0.0002). Good agreements for 187Os/188Os ratios within duplicates, however, were obtained for LAP 03719 (0.1275 ± 0.0001 to 0.1280 ± 0.0001), Shallowater (0.1279 ± 0.0001 to 0.1287 ± 0.0001), and Mt. Egerton (0.1299 ± 0.0001 to 0.1301 ± 0.0001), despite variable Os concentrations (Table 1).

Figure 3.

 Distribution of measured 187Os/188Os values among aubrites; n: number of samples in the respective compositional interval. Light gray: isotopic range for all chondrites; dark gray: isotopic range for enstatite chondrites (Walker et al. 2002; Horan et al. 2003; Brandon et al. 2005; Fischer-Goedde et al. 2010; van Acken et al. 2011). Values in the bin with 187Os/188Os > 0.140 encompass values of up to 0.226 (see Table 1).


Terrestrial Alteration and/or Contamination

Terrestrial crustal materials and fluids can have relatively high 187Os/188Os ratios of greater than 1, high Re/Os ratios of greater than 8, and low total Os concentrations of around 0.05 ppb (e.g., Esser and Turekian 1993). Meteorites, especially those with low-HSE abundances, can be susceptible to disturbance of primary signatures by terrestrial alteration processes that add or remove HSE (e.g., Huber et al. 2006; Brandon et al. 2012). This is especially true for aubrites, which contain minerals such as oldhamite (CaS) or alabandite (MnS) that are altered easily under the oxidized conditions in the terrestrial atmosphere. As most aubrite samples in this study display some evidence for terrestrial oxidation in the form of brown Fe-Oxide coating, some of the samples with measured 187Os/188Os ratios of >0.14 (Table 1) could reflect influence of terrestrial weathering instead of primary signatures. Mixing models of meteoritic Re and Os in the typical range of aubrites from this study (1, 10, and 100 ppb Os; Re/Os = 11; 187Os/188Os = 0.12) with terrestrial Re and Os (0.05 ppb Os, 0.4 ppb Re, 187Os/188Os = 1) shows that some of the variation in 187Os/188Os, especially the elevated values in samples like Bustee, can potentially be attributed to terrestrial contamination (Fig. 4; Table 1). These relationships also hold true for elevated Re/Os, indicating potential addition of not only radiogenic Os, but also of Re. The most extreme Re-Os signatures as seen in Bustee require >70% of Os and Re to be added, while less radiogenic signatures, such as seen in Bishopville, require around 30–50% (Fig. 4).

Figure 4.

187Os/188Os versus a) Re/Os, b) Os (ppb) c) 1/Os (ppb) d) Pd/Ir. To evaluate potential sample contamination by terrestrial alteration, mixing between terrestrial crustal contaminant (0.05 ppb Os, 0.39 ppb Re, 187Os/188Os = 1) with meteoritic material similar to typical aubrites (solid line: 1 ppb Os, 0.08 ppb Re, 187Os/188Os =0.12; dashed line: 10 ppb Os, 0.8 ppb Re, 187Os/188Os = 0.12) was modeled. Horizontal error bars in panel a) show error on the Re/Os ratio introduced by blank variability for low concentration samples.

However, there are several arguments that suggest that most, or possibly even all, of the observed Os signatures are primary. All aubrites other than the Antarctic samples, Mt. Egerton, Pesyanoe, and Shallowater are observed falls, minimizing the length of exposure time within the terrestrial environment (Foshag 1940; Lonsdale 1947; Beck and LaPaz 1951; Keil 1989). With the exception of ALHA78113, all the finds have high Os concentration (>10 ppb for Os), limiting the impact of terrestrial alteration from material with Os concentration below 0.05 ppb (Esser and Turekian 1993; Huber et al. 2006). Rhenium concentrations are as low as 0.18 ppb in ALHA78113, making Re more susceptible to terrestrial contamination, in addition to the issue of blank contribution to these values. Weathered Antarctic chondrites have no resolvable shift in Os isotopic composition compared with less weathered falls and finds, suggesting a minor impact of terrestrial alteration on the Re-Os isotopic system in meteorites (Huber et al. 2006).

In a recent study of Martian meteorites, Brandon et al. (2012) found significant addition of terrestrial Re and Os (up to 48% Re and 2% Os, with radiogenic 187Os/188Os as high as 0.34) in arid desert finds. We were not able to perform similar leaching experiments as were carried out in Brandon et al. (2012) to evaluate the magnitude of terrestrial contamination on HSE because of the unstable behavior of the exotic sulfides. Day et al. (2012a) suggested open-system behavior of Re during terrestrial alteration based on large variability of Re/Os compared to Ir/Os; resulting in mobilization and redistribution of Re within the brachinites and brachinite-like achondrites they examined (e.g., Northwest Africa 5400), rather than addition of Re or Os from external terrestrial sources. As in the case for these meteorites examined by Day et al. (2012a), the limited range in Ir/Os and 187Os/188Os compared with Re/Os in aubrites indicates mobility of Re rather than Os during terrestrial weathering. While the latter process is conceivable for finds with longer, unknown, residence times on Earth relative to falls, it seems unlikely for falls, some of which have been recovered within days or weeks (Lonsdale 1947; Beck and LaPaz 1951). However, because of their exotic mineral composition, aubrites may be more susceptible to terrestrial alteration than less reduced meteorites.

Additionally, fractionation of Re and Os by terrestrial alteration would require largely different relative mobilities of Re and Os, because both are primarily hosted in NiFe metal in aubrites (van Acken et al. 2012). The exotic sulfides do not significantly contribute to the HSE budget of aubrites (van Acken et al. 2012), and hence can only have limited involvement in terrestrial alteration. Fractionation of Re from Os by weathering would have to occur over a period of months for the observed fall samples, which is likely an unreasonably short time for removing both elements from their metal host or adding them from the surroundings, and fractionating Re from Os in the magnitude observed. Furthermore, terrestrial alteration removes Re more efficiently than Os (e.g., Walker et al. 2002; Fischer-Goedde et al. 2010) and should result in lower Re/Os, in contrast to the Re enrichments observed in aubrites. Osmium appears unaffected, as excellent correlations with Ir, Ru, and Pt are maintained (Fig. 2). To take potential terrestrial contamination and uncertainty introduced by large blank contributions into account, the Re/Os and 187Os/188Os values in low concentration samples (ALHA78113, Bishopville, Bustee, Khor Temiki, LAP 03719, LAR 04316, Norton County, and Peña Blanca Spring) are considered with caution during the following discussion.

Neutron Capture

Another potential cause for the observed enrichment in Re may be enrichment in 187Re by neutron capture of 186W, as outlined for lunar crustal rocks by Day et al. (2010). If this process is operating, the isotopic composition of “common” Re used during the isotope dilution calculation may be wrong and Re concentrations would be overestimated. The potential for such an interaction of aubrites with cosmic neutron radiation is significant. Aubrites have cosmic ray exposure (CRE) ages among the longest for all known meteorites (Eberhardt et al. 1965; Lorenzetti et al. 2003; Herzog et al. 2011), and neutron capture by 149Sm and 157Gd, which have large neutron capture cross-sections, has been shown to occur in aubrites, resulting in anomalous Sm and Gd isotopic composition comparable to those in lunar regolith (Hidaka et al. 2006). However, to develop excess in 187Re sufficient to disturb measured Re signatures from neutron capture of 186W, a W/Re >17,000 is required, as observed in lunar soils (Day et al. 2010). While no bulk W data for aubrites exist, Re concentrations determined independently by neutron activation (Wolf et al. 1983) would require a bulk W concentration around 10 ppm for these W/Re ratios to be achieved. This bulk W concentration for aubrites is unreasonably high, given the W concentrations of 0.3–0.7 ppm and W/Re ratios of less than 200 in aubrite metals, the primary host for both elements (van Acken et al. 2012).

Sample Heterogeneity and HSE fractionation

The excellent correlations between concentrations of HSE (Fig. 2) are consistent with all HSE being hosted in the same phase, presumably FeNi metal (Wolf et al. 1983; Casanova et al. 1993; van Acken et al. 2012). Highly siderophile element concentrations and 187Os/188Os ratios of multiple aliquots of the same sample can yield poor reproducibility (Table 1) perhaps resulting from primary sample heterogeneity, physical mixing during brecciation, and/or nugget effects due to the heterogeneous distribution of metal (Casanova et al. 1993; van Acken et al. 2012). Mechanical mixing of fragments during brecciation is a plausible explanation because most aubrites are breccias, and some have been identified as regolith breccias (Keil 1989). However, unless metal grains in aubrites represent different populations from sources with different Re/Os and 187Os/188Os, this “nugget effect” does not account for the large differences between 0.1331 and 0.2263 in measured 187Os/188Os ratios between some sample duplicates (Table 1). These duplicates may represent mixtures of different breccia fragments in various proportions. However, the mm-to-cm sizes of the fragments from which the analyzed powders were prepared are smaller than the several cm average sizes of breccia fragments such that this is unlikely to explain the large 187Os/188Os differences (e.g., Lonsdale 1947; Neal and Lipschutz 1981). The chondritic interelement HSE ratios and 187Os/188Os and isotopic compositions of two aubrites (ALH 84007-009, LAP 03716) suggest a relatively minor influence of aubrite parent body(ies) processes on HSE in these samples.

A “mode effect,” nonrepresentative sampling of the whole rock resulting in increased variability of chemical measurements, has been suggested for lunar samples (Spicuzza et al. 2007) and brachinites (Day et al. 2012). Given the very coarse-grained nature of aubrites (Foshag 1940; Lonsdale 1947; Beck and LaPaz 1951; Keil 1989, 2010), nonrepresentative sampling is likely to also have contributed to the poor reproducibility in some samples (Peña Blanca Spring, Norton County, Mt. Egerton).

In addition to variations in HSE contents within samples, twelve samples (ALHA78113, ALH 84008, ALH 84009, Bishopville, Bustee, Khor Temiki, LAP 03719, LAR 04316, Mayo Belwa, Mt. Egerton, Norton County, Pesyanoe) display fractionated HSE signatures relative to average chondrites (Table 1; Figs. 1 and 2). Deviations from chondritic HSE/Ir ratios are especially present in the more incompatible HSE such as Pd and Re (Figs. 1b and 2; Table S2). The samples which show the most fractionated, nonchondritic HSE patterns (Fig. 1) with enriched Re and Pd also show indications of disturbed Ar-Ar and Rb-Sr isotope systems, which give young ages of 3.5 to 4.5 Ga (Bogard et al. 2010). Norton County, which has Pd/IrN ratios of 4.02 to 11.98 and Re/IrN of 0.93 to 3.61 shows Ar-Ar ages between 1.3 and 4.5 Ga (Bogard et al. 2010), and a Rb-Sr age of 4.47 Ga (Minster and Allegre 1976). Bishopville has Pd/IrN of 1.18 to 4.41 and Re/IrN of 1.91 to 50.8 (Table S2), with reported Ar-Ar ages of 1.5 to 4.2 Ga (Bogard et al. 2010) and Rb-Sr ages of 3.5 to 3.9 Ga (Compston et al. 1965). Shallowater has an Ar-Ar age of 4.53 ± 0.05 Ga (McCoy et al. 1995) and a I-Xe age of 4.562.8 ± 0.3 Ga (Gilmour et al. 2006). Assuming 4.563 ± 0.001 Ga as the formation age of Shallowater (Gilmour et al. 2006) and using this age as a reference point, the calculated initial 187Os/188Os ratios in aubrites are between –0.7475 and 0.1036. Calculations of initial values in low-HSE concentration samples such as Bustee and Norton County are significantly affected by Re blank uncertainty, which can result in spurious initial 187Os/188Os values different by as much as 0.1 (Fig. 5). Even taking the uncertainties introduced by Re blank contribution into account, these impossible initial values support disturbance in the Re-Os isotope system and other isotope systems (Ar-Ar, Rb-Sr) in samples such as Bustee, Mayo Belwa, Norton County, or Bishopville (Table 1; Fig. 5).

Figure 5.

 Temporal evolution of aubrites, using measured 187Os/188Os as present-day composition and measured Re/Os to calculate 187Os/188Os back to solar system formation. Red lines: enstatite chondrite (Walker et al. 2002), black lines: aubrites with approximately enstatite chondritic development; gray lines: aubrites with disturbance of the Re-Os isotope system, quickly developing unreasonably low 187Os/188Os. Dashed gray lines: minimum and maximum calculated development lines for low-Re samples Norton County and Bustee considering uncertainty caused by blank variability.

The enrichment in Pd and Re by a factor of up to 12, excluding the high, nonreplicable values in Mayo Belwa and Bishopville, seen in several samples (e.g., Norton County, Bustee), is commonly interpreted as a partial melt signature in the absence of metal in the terrestrial mantle (e.g., Morgan 1986). Whether this interpretation can be applied to a smaller, much more reduced parent body is questionable. The relative incompatibility of Re, Pd, and, to a lesser extent, Pt in the terrestrial mantle is governed by their partitioning in sulfides and silicates (Ballhaus et al. 2006; Fonseca et al. 2007; Mallmann and O’Neill 2007), and the relative compatibility in sulfide and silicate melts during partial melting (Ballhaus et al. 2006; Mallmann and O’Neill 2007). In most aubrite sulfides, HSE concentrations are below detection limits for LA-ICP-MS (<1–5 ppb, Casanova et al. 1993; van Acken et al. 2012), rendering them insignificant as a HSE carrier phase.

Enrichments of Pd and Re can potentially be explained by a number of processes. As Pd and Re enrichments do not always occur in the same sample, different processes may be the underlying cause. Palladium enrichments could be explained by less siderophile behavior of Pd during metal segregation (Mann et al. 2012), leaving the silicate residue with elevated Pd/Os. Addition of fractionated material, as seen in some metal grains from chondritic inclusions in the Cumberland Falls aubrite (Pd/Os up to 60, Re/Os up to 0.25) or troilite grains from Shallowater (Pd/Os and Re/Os of up to 3; van Acken et al. 2012), would produce small enrichments in Pd and Re. In addition to minor Re addition by terrestrial processes (Fig. 4a), addition of metal and sulfide with fractionated HSE signature could serve to explain the enrichments, although the amount of material needed to reproduce the bulk aubrites HSE signatures might well exceed the amount of inclusion material available (see discussion about chondrite addition below).

Core Formation and Remixing in the Aubrite Parent Body(ies)

Excluding the anomalous metal-rich samples, Shallowater and Mt. Egerton, which may have an origin on reduced bodies different from the aubrite parent body(ies) (Keil 1989, 2010; Keil et al. 1989), all aubrites have HSE concentrations at least an order of magnitude lower than is typical for chondrites. Metal/silicate partition coefficients of HSE (Dmetal/silicate) are typically greater than 106 (e.g., Kimura et al. 1974; O’Neill et al. 1995; Righter 2003; Brenan and McDonough 2009; Mann et al. 2012); hence, HSE are segregated into metal cores. Because of the reduced oxidation state of aubrites, HSE are exclusively hosted in metal phases (Casanova et al. 1993; van Acken et al. 2012) and the abundance of metal from traces to 3.7 wt% (Watters and Prinz 1979) thus controls their bulk HSE concentrations. If representative of the silicate portions of the aubrite parent body(ies), these HSE concentrations are orders of magnitude too high to result from metal/silicate equilibrium partitioning during core formation, which would predict HSE concentrations of 3 ppt or less in the silicate portion. Another explanation for the high and variable HSE concentrations in aubrites is thus required.

Keil et al. (1989) advocated a complex history of impact-induced breakup and reassembly for the parent body of the anomalous Shallowater aubrite based on the petrology and trace element contents of sulfides and silicates. In this case, elevated HSE concentrations could be the result of remixed core material during reassembly with the silicate portions of the body after initial differentiation and parent body destruction. Alternatively, impact-induced shearing stress may draw core material back into the mantle (Rushmer et al. 2005). Different scenarios of remixing of core material into previously depleted silicate can be envisioned: remixing of bulk core metal, either liquid or solid from a differentiated core, or a mixture of the latter two in proportions different from the bulk core. To evaluate these possibilities, the composition of core forming metal on the aubrite parent body(ies) needs to be constrained. Siderophile element partitioning between silicate, liquid and solid metal during metal segregation and differentiation is controlled by the temperature, pressure, and the minor element content of the metal, most notably S, C, Si, and P (e.g., Chabot and Jones 2003).

For modeling, either EL or EH chondrites were assumed as a proxy for undifferentiated starting materials, forming a perfectly segregated metal core making up 40% of the total mass (metal content in enstatite chondrites; Mason 1966), which leads to HSE concentrations in the silicate portion of the resulting differentiated body between 0.1 and 3 ppt. Measured HSE concentrations in aubrites as representative of the silicate portion of the aubrite parent body(ies) are 2–5 orders of magnitude higher than HSE concentrations in the silicate portion after core formation (Table 1; Fig. 1); and planetary processes in addition to accretion and core formation are required to explain HSE systematics in aubrites.

If a differentiated core is taken into account, the influence of minor element concentrations in metal on HSE partitioning behavior between solid and liquid metal becomes important (e.g., Jones and Drake 1983; Chabot and Drake 2000; Chabot et al. 2003, 2006, 2008, 2010; Corrigan et al. 2009; Hayden et al. 2011). For high S content around 25%, all HSE with the exception of Pd partition into solid metal over S-rich melts (Chabot et al. 2003). For lower S concentrations around 5%, all elements are moderately compatible in solid metal over metal liquid. The low S concentrations in aubrite metals below 0.15% support removal of an S-rich melt phase (van Acken et al. 2012). However, bulk HSE abundances (Table 1; Fig. 1) argue against removal of S-rich melt as a dominant process, as Pd is not or only marginally depleted. The effect of C, P, and Si on HSE solid metal/liquid metal partition coefficients is small (Chabot and Drake 2000; Chabot et al. 2006, 2010; Corrigan et al. 2009), although Ru shows greater affinity for P-rich metal. Effects of minor elements on HSE partitioning between solid and liquid metal are negligible over the range of minor element content (Si, S, C, P) encountered in aubrite metals (Casanova et al. 1993; van Acken et al. 2012). The exceptions are high contents of S (approximately 25%), which lead all HSE except Pd to accumulate in metal solid, and C (approximately 4%), which results in fractionation of Ru from the other HSE (Chabot et al. 2006, 2008).

As with addition of chondritic material, addition of small amounts (0.1–1%) of segregated metal of either a differentiated or undifferentiated core back into an HSE-depleted silicate portion of the aubrite parent body(ies) can match the observed HSE concentrations in aubrites reasonably well (Fig. 6). However, the high Dmetal/silicate and uniform Dliquid metal/solid metal values do not cause significant fractionation of HSE from each other, with the possible exception of Pd in S-rich systems, and remixing of core material with a silicate portion of the aubrite parent body(ies) stripped of its HSE cannot account for fractionation of HSE. While remixing of <0.1% to 3% of core material is consistent with the HSE concentrations observed (Fig. 6) and with the broadly chondritic HSE ratios for the more compatible HSE Os, Ir, Ru, and Pt (Fig. 1a), the enrichments in Pd and Re observed in some samples require additional processes to operate during the formation of aubrites (Fig. 1b).

Figure 6.

 a–h) Highly siderophile element (Pt, Ru, Pd, Re) versus Os. Solid lines denote mixing of core material. Parts (e–h) represent details of parts (a–d). Core composition was calculated from an EL starting composition (Horan et al. 2003; Brandon et al. 2005b; Fischer-Goedde et al. 2010; van Acken et al. 2011), assuming a core comprising 40% of the body mass, corresponding to metal content in EL meteorites. Models were calculated with two sets of Dmetal/silicate parameters: For the first set, metal/silicate partitioning coefficients were set to 106 for all HSE (solid lines with long tick marks); for the second set, the high-pressure parameters from Mann et al. (2012) were used (experiment Z538, dashed line with short tick marks). Core remixing models were calculated for both remixing of bulk core material and mixtures of core material having undergone differentiation into a liquid and a solid portion, which yield no resolvable difference.

Aubrite metals in Cumberland Falls, Mt. Egerton, and Aubres show fractionated HSE patterns, with a marked depletion in refractory HSE (Os, Ir, Ru, Pt), which is interpreted to be a quenched metal liquid signature (van Acken et al. 2012). However, the bulk HSE concentrations in Aubres and Mt. Egerton do not reflect depletion of Os, Ir, Ru, and Pt compared with Pd and Re (Fig. 1), raising the need for a complementary HSE-rich phase, potentially heterogeneously distributed HSE alloys, as seen in the metal phases of chondrites (e.g., Rambaldi et al. 1983; Campbell and Humayun 2003; Horan et al. 2009). The presence of these alloys, so far undetected by in situ studies, would also serve to explain the strong nugget effect of over two orders of magnitude, as observed in samples like Mt. Egerton and Norton County.

Chondritic Inclusions

Thirteen of 16 aubrites samples from this study (excluding LAR 04316, Mt. Egerton, and Shallowater) are breccias. Thus, past research has put a special emphasis on various inclusions and clasts, especially in ALHA78113, Cumberland Falls, and Khor Temiki (Neal and Lipschutz 1981; Verkouteren and Lipschutz 1983; Lipschutz et al. 1988; Newsom et al. 1996; Keil et al. 2011). Some of these clasts have been described as chondritic (Neal and Lipschutz 1981; Verkouteren and Lipschutz 1983; Lipschutz et al. 1988), and have been shown to have chondritic HSE concentrations (Wolf et al. 1983). Some aubrites (ALH 84007–84009, LAP 03719, Aubres, Peña Blanca Spring) show broadly chondritic HSE ratios (Table S2), consistent with a primordial origin of the HSE-bearing phases in these samples. The HSE concentrations in these samples are less than chondritic by one to two orders of magnitude (Table 1; Wolf et al. 1983). Brecciated aubrites could thus be regarded as a mixture between the chondritic, HSE-rich inclusions and an aubritic, almost HSE-free matrix. The resulting HSE patterns would then represent “diluted” chondrite patterns and the absolute HSE abundances would be determined by how much of each of the two components are within each sample fraction measured. For example, to explain the Pt and Os concentrations of the aubrite samples, and assuming chondritic concentrations of HSE for the inclusions (Horan et al. 2003; Brandon et al. 2005; Fischer-Goedde et al. 2010; van Acken et al. 2011) and no HSE contents in the orthopyroxene-dominated matrices, between 0.1 and 7.5% by total weight needs to be made up of chondritic inclusion fragments (Fig. 7). For ALHA78113, in which chondritic inclusions are present (Lipschutz et al. 1988), less than 0.5% of chondritic material is sufficient to produce the observed concentrations of 1–2 ppb Os, 2.1–2.2 ppb Pd, and 2.2–3.9 ppb Pt (Table 1; Fig. 7). The percentage of inclusions present within ALHA78113 is difficult to estimate and has not been determined. Three inclusions were studied by Lipschutz et al. (1988) for a total sample mass of 299g for ALHA78113; inclusions were both larger and more abundant in Cumberland Falls (Neal and Lipschutz 1981; Verkouteren and Lipschutz 1983; Lipschutz et al. 1988), where they make up to 10% of the total sample in some fragments.

Figure 7.

 a, c) Pt (ppb) versus Os (ppb); b, d) Pd (ppb) versus Os (ppb); (c and d) represent shaded area in (a and b), respectively. Line denotes mixing of chondritic material (Horan et al. 2003; Fischer-Goedde et al. 2010), as seen in chondritic inclusions in aubrites (Neal and Lipschutz 1981; Rubin 2010) with aubrite silicate mantle material after formation of a core on the aubrite parent body(ies). Tick marks represent the percentage of chondritic material added. Concentration of aubrite silicate calculated using Dmetal/silicateHSE of 106 (e.g., Righter 2003).

The type(s) of chondritic clasts found in aubrites cannot be resolved here due to the similarity of HSE concentrations for the different chondrite groups (Horan et al. 2003; Brandon et al. 2005; Fischer-Goedde et al. 2010; van Acken et al. 2011). While some clasts have been described as ordinary chondrites with variable degrees of equilibration with the aubrite matrix (Rubin 2010), other clasts have been labeled “forsterite chondrites,” with an affiliation to K chondrites, winonaites, and IAB irons (Neal and Lipschutz 1981). It is also noteworthy that metals, which are the primary HSE carriers in aubrites, have unfractionated HSE signatures outside of the chondritic inclusions, while displaying fractionated signatures within, which has been attributed to solid metal/liquid metal partitioning (van Acken et al. 2012). Addition of these metal grains to some aubrites could produce a slightly fractionated HSE pattern, as seen in bulk samples for these elements. In summary, addition of metal from chondritic inclusions to an aubritic matrix can explain some of the HSE concentrations observed in aubrites, and potentially provide an explanation for the fractionated patterns, in particular, the Pd enrichments over chondrites (Mayo Belwa, Bustee, Bishopville, Norton County). However, the bulk inclusions show chondritic Pd/Os and subchondritic Re/Os (Wolf et al. 1983), so whether the metal grains with fractionated patterns can be considered representative for the chondrite inclusions remains an open question.

Late accretion, the addition of chondritic material to a planetary body after core formation, but prior to crystallization of the crust, is thought of as the process governing HSE concentrations in the silicate portions of terrestrial planets (e.g., Kimura et al. 1974). Recent studies show that smaller, differentiated bodies may have been affected by late accretion of chondritic material, albeit to lesser extent than larger bodies (Dale et al. 2012; Day et al. 2012b). The predicted effect of addition of a late chondritic accretion on the bulk HSE systematics of meteorites originating from the silicate portion of such a planetary body is indistinguishable from addition of chondritic fragments by late impact processes as discussed above.

Comparison with Re-Os Isotopes and HSE in Meteorites from Other Small, Differentiated Bodies

To understand differentiation of small planetary bodies such as the aubrite parent body(ies), the results from aubrites need to be put in context with results from other groups of meteorites. A large body of HSE and Os isotope data exists for chondrites (Walker et al. 2002; Horan et al. 2003; Brandon et al. 2005; Fischer-Goedde et al. 2010; van Acken et al. 2011); ureilites (Rankenburg et al. 2007, 2008); howardites, eucrites, and diogenites (HEDs, Dale et al. 2012; Day et al. 2012b); shergottites (Brandon et al. 2000, 2012); angrites (Dale et al. 2012; Riches et al. 2012); and brachinites (Day et al. 2012a). Although all of these groups are significantly different from aubrites with respect to mineral assemblage, oxidation state, and oxygen isotopic composition, they are all thought to have experienced differentiation during the earliest stages of the solar system (e.g., Mittlefehldt 2003).

Concentrations of HSE in these meteorite groups vary over more than seven orders of magnitude from <1ppt in diogenite Meteorite Hills 00424 (Day et al. 2012b) to several 1000 ppb (brachinite NWA 5400, Day et al. 2012a; Fig 8a). Aubrites have Os concentrations comparable to shergottites, HEDs, and angrites between 0.1 and a few tens of ppb, with the anomalous samples Shallowater and Mt. Egerton showing values of a few 100 ppb, in the range of chondrites, ureilites, and brachinites. Concentrations in angrites, HEDs, and shergottites extend up to three orders of magnitude below those seen in aubrites to below 1 ppt (Fig. 8a). This range of concentrations underscores the differences between aubrites and ureilites, which have been suggested to have undergone loss of metallic liquid and not have formed a core (Rankenburg et al. 2008), and brachinites and brachinite-like achondrites, which are interpreted as residues from partial melting of an undifferentiated volatile-rich body, with the anomalous meteorites Graves Nunatak 06128 and 06129 forming complementary felsic crusts (Day et al. 2009, 2012a), while pointing to similarities in the differentiation histories of aubrites, angrites, HEDs, and shergottites.

Figure 8.

 Comparison of HSE and 187Os/188Os signatures in aubrites with other meteorite groups. a) Pt/Os versus Os (ppb), b–c) Pt/Os versus Pd/Os, panel c) represents shaded area in panel b, d) 187Os/188Os versus Re/Ir. Data sources: aubrites: this study, chondrites: Walker et al. (2002), Horan et al. (2003), Brandon et al. (2005), Fischer-Goedde et al. (2010), van Acken et al. (2011); angrites: Dale et al. (2012), Riches et al. (2012); shergottites: Brandon et al. (2012); brachinites and brachinite-like achondrites: Day et al. (2009, 2012a); HEDs: Dale et al. (2012), Day et al. (2012b). Single samples with extreme values were excluded for scale reasons (e.g., shergottite Los Angeles, 187Os/188Os = 0.74088, Pd/Os = 2919, Pt/Os = 1603, shergottite Zagami, Pd/Os = 155, Pt/Os = 197, Brandon et al. [2012]; diogenite Talampaya, Pt/Os = 42, Dale et al. [2012]).

As absolute concentrations in HSE are dependent on the amount of metal present and hence highly variable in samples with inhomogeneous distribution of metal (e.g., Mt. Egerton, Table 1), interelement HSE ratios are a more suitable tool to study planetary differentiation processes. Because Pt, Pd, and Re behave more incompatibly than Os, Ir, and Ru during partial melting of sulfide-bearing silicate lithologies, ratios such as Pt/Os, Pt/Ir, Re/Ir, and Re/Os (approximated by 187Os/188Os) are useful tracers for magmatic differentiation. Elevated Pt/Os, Pt/Ir, Re/Ir, and 187Os/188Os compared with chondritic reference values (Pt/Os: 1.61–2.20, Pd/Os: 0.90–1.49, Re/Ir: 0.07–0.11; Horan et al. 2003; Brandon et al. 2005; Fischer-Goedde et al. 2010; van Acken et al. 2011) are interpreted as melt signatures, whereas lower values reflect melt residues (e.g., Morgan 1986). Aubrites cover a range in Pt/Os (0.76–5.87), Pd/Os (0.56–14), and Re/Ir (0.03–4.7, see Table S2) comparable to angrites (Pt/Os: 0.79–39, Pd/Os: 0.07–7.9, Re/Ir: 0.03–16.7; Dale et al. 2012; Riches et al. 2012), but narrower than shergottites (Pt/Os: 0.30–22.7, Pd/Os: 0.28–15.8, Re/Ir: 0.02–50, excluding Los Angeles and Zagami; Brandon et al. 2012) and HEDs (Pt/Os: 0.74–42, Pd/Os: 0.07–55, Re/Ir: 0.03–8; Dale et al. 2012; Day et al. 2012b), and wider than both ureilites (Pt/Os: 0.83–1.86, Pd/Os: 0.10–0.98, Re/Ir: 0.03–0.11; Rankenburg et al. 2008) and brachinites (Pt/Os: 0.17–2.56, Pd/Os: 0.06–0.79, Re/Ir: 0.03–0.25; Day et al. 2009, 2012a; Figs. 8b–d). The very low Pd/Os and Pt/Os values below 0.5 as seen in achondrites from all other groups are not seen in aubrites (Fig. 8c). Aubrites do show extreme Re/Ir as high as 4.7, as also seen in single shergottites and HEDs, which are interpreted as products of igneous differentiation and formation of different mantle reservoirs on Mars and 4 Vesta, respectively. Aubrites show a wider range in 187Os/188Os, reflective of long-term development of Re/Os, than chondrites (0.123–0.131; Walker et al. 2002; Horan et al. 2003; Brandon et al. 2005; Fischer-Goedde et al. 2010; van Acken et al. 2011; Fig. 3), ureilites (0.121–0.131; Rankenburg et al. 2007), and brachinites (0.120–0.131; Day et al. 2009, 2012a), and comparable to angrites (0.106–0.212; Dale et al. 2012; Riches et al. 2012), HEDs (0.117–0.206; Dale et al. 2012; Day et al. 2012b), and shergottites (0.116–0.246, with the exception of Los Angeles, Brandon et al. 2012; Fig. 8d).

The similar range of Pt/Os and Pd/Os seen in angrites, which crystallized from basaltic magma under oxidizing conditions (IW + 1 – IW + 2; Mittlefehldt et al. 1998), compared to aubrites, albeit with a few samples with Pd/Os as low as 0.07 (Riches et al. 2012), indicates similar igneous processing operating on both the aubrites and angrite parent bodies, despite the large difference in oxidation state. The lesser range in Pt/Os and Pd/Os in aubrites compared with shergottites and HEDs can be interpreted as a lesser degree of magmatic differentiation occurring within the aubrite parent body(ies) compared with Mars and the HED parent body, possibly reflecting faster cooling due to smaller body size. The onset of planetary differentiation and formation of different mantle reservoirs of Mars has been placed around 4.5 Ga (Brandon et al. 2000; Kleine et al. 2002; Debaille et al. 2007; Dauphas and Pourmand 2011). Younger crystallization ages of shergottites and ALH84001 indicate a prolonged period of igneous activity in Mars from 4.1–0.15 Ga (Lapen et al. 2010; Brandon et al. 2012; and references therein). In contrast, aubrite ages much younger than 4.45 Ga (Compston et al. 1965; Bogard et al. 2010) are considered disturbed and without geological meaning. Superchondritic 187Os/188Os from 0.136 to 0.163 in some mafic shergottites (e.g., Dar al Ghani 476, NWA1068, NWA5789; Brandon et al. 2012) reflects long-term elevated Re/Os of the shergottite mantle sources resulting from early differentiation around 4.5 Ga, while similarly superchondritic values in aubrites (up to 0.226 in Bustee; Fig. 8d) probably cannot be attributed to formation of distinct mantle reservoir because of the smaller parent body size, and hence faster cooling time of the aubrite parent body compared with Mars (Zellner et al. 1977).

In summary, while aubrites have HSE characteristics that are distinct from primitive meteorites such as chondrites, their HSE signatures are by no means unique among evolved achondrites. Similarities in the degree and range of HSE depletion are seen in a significant range of redox states, from extremely reduced (IW-6 to IW-8, aubrites) to oxidized (IW+1 to IW+2, angrites) and thus apply to planetary bodies forming and differentiating at various distances from the Sun (e.g., Wasson and Kallemeyn 1988). The closest comparison of aubrites is to angrites and HEDs, which share the wide range in 187Os/188Os ratios, HSE concentrations and ratios seen in aubrites (Fig. 8) (Dale et al. 2012; Day et al. 2012b; Riches et al. 2012). Despite a number of overlapping HSE characteristics with other achondrite groups, aubrites maintain distinct HSE signatures lacking the extreme depletion in HSE concentration seen in angrites. HEDs, and shergottites, and the low Pt/Os and Pd/Os ratios of brachinites and ureilites, point to a unique history of the aubrite parent body(ies).

Concluding Remarks

The new HSE concentration and 187Os/188Os data presented here further highlight the heterogeneity recognized within the aubrites by earlier studies. While some of the heterogeneity may be ascribed to nugget effects caused by the HSE being hosted in trace metal phases and the brecciated nature of most aubrites, the magnitude of difference in HSE concentrations and multiple mechanisms of fractionation underscore the complexities of the formation of the aubrites. The HSE in aubrites record a parent body history distinct from other achondrites, albeit with some similarities. Angrites and HEDs are the most similar groups in terms of the range of HSE concentrations and 187Os/188Os, but angrites show depletion in Pd and Re instead of the observed enrichment in aubrites. Ureilites and brachinites show much smaller range in 187Os/188Os than aubrites, and both groups also show depletions in Pd (ureilites and brachinites), Pt, and Re (brachinites only), different from aubrites. While the ages obtained for all groups using lithophile element isotope systems fall within a narrow range, the aubrite parent body(ies) clearly evolved differently than the parent bodies of other differentiated achondrites.

Remixing of core material with the silicate mantle portion, either by core reflux (Rushmer et al. 2005) or by mixing of core fragments during breakup and reassembly, can account for the HSE concentrations as well as for the heterogeneity inferred for the aubrite parent body, although not for the observed fractionation. Postaccretion addition of chondritic material, either as late accretion or mixing of fragments of a chondritic impactor with HSE-depleted mantle of the aubrite parent body(ies) during reassembly of the aubrite parent body(ies), can account for some characteristics of HSE signatures, and some metal grains in chondritic inclusions (van Acken et al. 2012) possess sufficient enrichments in Pd and Re to account for the observed fractionations, although the question remains whether these grains can be considered representative for the HSE signature of added chondritic material.

Acknowledgments—  Samples for this study were kindly provided by the Meteorite Working Group, Johnson Space Center, Arizona State University, University of New Mexico, the Monnig Collection, Texas Christian University, Smithsonian National Museum of Natural History, Natural History Museum London, and Western Australian Museum. D. v. A. was supported by an ORAU postdoctoral fellowship. We appreciate discussions with T. McCoy, A. Riches, J. M. D. Day, and M. Humayun. We thank M. Righter, J. T. Shafer, Y. Gao, and J. I. Simon for technical support; J. M. D. Day, M. Horan, and C. Dale for their constructive reviews; and N. Chabot for editorial handling. This work was supported by the NASA Cosmochemistry program through grants NNX10AB37G and NNX12AD06G to A. D. Brandon and NNX09AC06G to T. J. Lapen.

Editorial Handling—  Dr. Nancy Chabot