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The Pb isotopic evolution of the Martian mantle

constrained by initial Pb in Martian meteorites

J. J. Bellucci1, A. A. Nemchin1,2, M. J. Whitehouse1, J. F. Snape1, P. Bland2,3, and G. K. Benedix2,3 1

Department of Geosciences, Swedish Museum of Natural History, Stockholm, Sweden,2Department of Applied Geology,

Curtin University, Perth, Western Australia, Australia,3Department of Earth and Planetary Sciences, Western Australia

Museum, Welshpool, Western Australia, Australia

Abstract

The Pb isotopic compositions of maskelynite and pyroxene grains were measured in ALH84001 and three enriched shergottites (Zagami, Roberts Massif 04262, and Larkman Nunatuk 12011) by secondary ion mass spectrometry. A maskelynite-pyroxene isochron for ALH84001 defines a crystallization age of 4089 ± 73 Ma (2σ). The initial Pb isotopic composition of each meteorite was measured in multiple maskelynite grains. ALH84001 has the least radiogenic initial Pb isotopic composition of any Martian meteorite measured to date (i.e.,206Pb/204Pb = 10.07 ± 0.17, 2σ). Assuming an age of reservoir formation for ALH84001 and the enriched shergottites of 4513 Ma, a two-stage Pb isotopic model has been constructed. This model links ALH84001 and the enriched shergottites by their similarμ value (238U/204Pb) of 4.1–4.6 from 4.51 Ga to 4.1 Ga and 0.17 Ga, respectively. The model employed here is dependent on a chondriticμ value (~1.2) from 4567 to 4513 Ma, which implies that core segregation had little to no effect on theμ value(s) of the Martian mantle. The proposed Pb isotopic model here can be used to calculate ages that are in agreement with Rb-Sr, Lu-Hf, and Sm-Nd ages previously determined in the meteorites and confirm the young (~170 Ma) ages of the enriched shergottites and ancient,>4 Ga, age of ALH84001.

1. Introduction

Radiogenic isotopic systems can be used to constrain the time-integrated bulk composition of the Martian silicate mantle and crust, timing of differentiation, and the behavior of the parent and daughter elements of a given isotopic system. Currently, there are six known types of Martian meteorites: shergottites, nakhlites, chassignites (collectively, SNCs), Alan Hills (ALH) 84001 (an orthopyroxenite), Northwest Africa (NWA) 8195 (an augite basalt), and NWA7533 (a regolith breccia) and its pairs [e.g., Agee et al., 2013, 2014; Humayun et al., 2013; Mittlefehldt, 1994; Nyquist et al., 1995, and references therein]. ALH84001 is a unique orthopyroxenite that has a crystallization age of 4.074 ± 0.099 Ga determined by a Pb-Pb isochron [Bouvier et al., 2009], which is in agreement with the Lu-Hf age of 4.091 ± 0.030 Ga (2σ [Lapen et al., 2010]). The SNCs comprise mafic and ultramafic rocks with a relatively limited range in composition and87Rb-87Sr,

176Lu-176Hf, and147Sm-143Nd ages that span 1.3–0.6 Ga [e.g., Borg et al., 2003, 2005; Bouvier et al., 2005,

2008; Lapen et al., 2008; Misawa et al., 2006; Nyquist et al., 1995; Shih et al., 2009; Shafer et al., 2010, and references therein]. The shergottites can be subdivided into basaltic and lherzolitic types [e.g., Nyquist et al., 1995] and even further divided, on the basis of light rare Earth element concentrations and radio-genic isotope compositions, into a spectrum from depleted to enriched varieties [e.g., Borg and Draper, 2003; Borg et al., 1997, 2003; Symes et al., 2008]. The enriched shergottites generally have a younger age (~170 Ma) than the depleted shergottites (~300–580 Ma) [e.g., Borg et al., 1997, 2003, 2005; Nyquist et al., 1995; Moser et al., 2013; Symes et al., 2008]. Pb isotopic data obtained from solution analysis of shergottite mineral separates, which yield arrays in 207Pb/204Pb versus 206Pb/204Pb diagrams corresponding to apparent>4 Ga ages, are explained either as (1) true crystallization ages resulting from in situ decay of U to Pb [Bouvier et al., 2005, 2008, 2009] or (2) mixing of two end-members with highly contrasted Pb isotopic compositions either terrestrial or Martian [Bellucci et al., 2015; Borg et al., 2005; Gaffney et al., 2007]. The source reservoir for the enriched shergottites and ALH84001 has been inferred to have formed at 4.513 Ga based on the combined isotopic systematics of176Lu-176Hf,147Sm-143Nd, and146Sm-142Nd [Borg et al., 2003; Lapen et al., 2010]. However, this model is premised on the interpretation of >4 Ga Pb-Pb isochron ages as the crystallization age of the shergottites [Bouvier et al., 2005, 2008, 2009] is erroneous and that the younger ages (~170 Ma) obtained by all radiogenic isotopic systems (except for Pb-Pb isochrons) are their true crystallization ages [e.g., Borg et al., 2003; Nyquist et al., 1995; Moser et al., 2013; Symes et al., 2008]. This study aims to constrain

Journal of Geophysical Research: Planets

RESEARCH ARTICLE

10.1002/2015JE004809

Key Points:

• Initial Pb model ages confirm young ages of shergottites

• Enriched shergottites and ALH84001 linked to a similar reservoir • Core formation had little to no effect

on U-Pb ratio of Mars

Correspondence to:

J. J. Bellucci,

jeremy.bellucci@gmail.com

Citation:

Bellucci, J. J., A. A. Nemchin, M. J. Whitehouse, J. F. Snape, P. Bland, and G. K. Benedix (2015), The Pb isotopic evolution of the Martian mantle constrained by initial Pb in Martian meteorites, J. Geophys.

Res. Planets, 120, 2224–2240,

doi:10.1002/2015JE004809. Received 21 FEB 2015 Accepted 19 NOV 2015

Accepted article online 21 NOV 2015 Published online 21 DEC 2015

©2015. American Geophysical Union. All Rights Reserved.

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the isotopic composition of initial Pb (present at the time of crystallization) in ALH84001 and three enriched shergottites and subsequently construct a coherent, time-integrated Pb isotopic model of the Martian mantle that is in agreement and consistent with the majority of the other radiogenic isotopic information available for these meteorites.

1.1. Pb Isotopic System

The element Pb has four isotopes:208Pb,207Pb,206Pb, and204Pb, of which thefirst three are daughter products of long-lived isotopes232Th,235U, and238U, respectively. At any given time, the Pb isotopic composition of a rock or mineral is a combination of its initial Pb isotopic composition, which is dependent on its previous U-Th-Pb history, and subsequent in-growth of Pb from the decay of U and Th. The coupled U decay schemes mean that the system can be utilized by measurement of Pb isotopes alone, yielding both chronological and geochemical information. Pb-Pb geochronology is explicitly dependent on the following assumptions: (1) individual minerals within a given sample began with an identical Pb isotopic composition; (2) the system has been closed since the time of crystallization or has been disturbed very recently; and (3) if multiple samples are used, they must be cogenetic and abide by thefirst two assump-tions. If these assumptions are met, radiogenic Pb can be used to define a Pb-Pb isochron corresponding to the age of crystallization for a rock or suite of cogenetic samples. Initial Pb, which is present at the time of crystallization and preserved in minerals that contain no U or Th and, hence, have no radiogenic Pb in growth, can be used to define time-integrated chemical parameters (μ =238U/204Pb andκ =232Th/238U) of a sample’s source, based on a model of Pb growth in the host planetary body, and its crystallization age. Terrestrial Pb isotopic models have historically fallen into three categories: single stage, multiple stage, and those incorporating mass transport.

1.2. Single-Stage Pb Isotope Models and Model Ages

A single-stage Pb isotopic model starts at the beginning of the solar system, defined here by the oldest preserved solid materials in the solar system at 4.567 Ga [Connelly et al., 2012], with the Pb isotopic composi-tion of solar system initial (primordial) Pb in Canyon Diablo Troilite (CDT) [Chen and Wasserburg, 1983]. CDT is an FeS mineral found in the Canyon Diablo IAB iron meteorite, which has no U and therefore has accumulated no radiogenic Pb since solar system formation. Subsequently, the Pb isotopic composition evolves with a fixed μ value until a given rock/mineral crystallization age, time (t). The line that connects these two points is known as a single-stage geochron (for present day this is“The Geochron,” for any time in the past it is a paleo-geochron), on which all samples with the same crystallization age and differentiation history but differ-entμ values will plot [Holmes, 1946; Houtermans, 1946, Figure 1]. A paleo-geochron is, therefore, a prediction of the Pb isotopic composition at any time after 4.567 Ga, based on a single-stage growth for anyμ value anchored by CDT. This is in contrast to a Pb-Pb isochron used in Pb-Pb chronology following the assumptions listed earlier. Additionally, an isochron is not anchored to a specific primordial Pb isotopic composition (i.e., CDT). The crystallization age of a rock can be determined either by a Pb-Pb isochron or by independent radio-genic isotopic systems (e.g., U-Pb in zircon,87Rb-87Sr,147Sm-143Nd, and176Lu-176Hf). If the least radiogenic and here assumed“true” initial Pb in a sample lies on a paleo-geochron corresponding to its crystallization age, then the single-stageμ of its source can be calculated. Additionally, as a method to test whether a given Pb isotopic model is correct, the initial Pb can be used to calculate a Pb model age that should be in agreement with all other isotope systems. A similar approach can be taken in the 232Th-208Pb versus

238U-206Pb systems and yields aκ curve.

1.3. Multiple-Stage Pb Isotope Models and Model Ages

Whereas single-stage models can be used to explain very simple U-Th-Pb isotopic systems, they are generally inappropriate for use in planetary evolution, as U-Th-Pb systems on a planetary body have a complicated history. For example, initial Pb compositions determined for terrestrial samples and the best estimate for the Pb isotopic composition of the Earth plot to the more radiogenic side of the single-stage paleo-geochrons corresponding to their crystallization age and the age of the Earth, respectively. This shift away from a single-stage paleo-geochron for both a given sample and the bulk Earth results in inaccurate, younger initial Pb model ages and a problem commonly referred to as the“first terrestrial Pb paradox” [e.g., Allegre, 1969]. This explicitly requires at least a two-stage differentiation history and/or mixing with a more evolved end-member [e.g., Zartman and Haines, 1988]. The effect of a two-stage differentiation history is illustrated

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schematically in Figure 1, whereby the second-stage paleo-geochron and the second-stage growth curves are shifted to accommodate an episodic change inμ value at a given differentiation time [e.g., Gale and Mussett, 1973]. Multiple-stage models have been proposed for terrestrial samples to account for inadequacies of single-stage models and have been demonstrated to produce accurate initial Pb model ages [e.g., Stacey and Kramers, 1975].

In the context of planetary scale evolution of major Pb reservoirs, changes in U/Pb and/or Th/U ratios require special circumstances. Since the partition coefficients for U, Th, and Pb in a silicate-only melt are similar, silicate differentiation has little to no effect on their ratios [e.g., Gaffney et al., 2007; Hauri et al., 1994]. However, the partition coefficients for Pb in sulfides and/or liquid metal are significantly greater than that of U [Gaffney et al., 2007; Hauri et al., 1994; Malavergne et al., 2007] and fractionation of Pb from U and Th can occur during differentiation events that sequester Pb in precipitated sulfides and/or liquid metal [e.g., Allègre et al., 1982; Malavergne et al., 2007]. Therefore, core formation, in theory, should have the largest effect on a planetary body’s μ value [Allègre et al., 1982].

The initial Pb measured in any Martian meteorite has likely resided in at least three reservoirs, including (1) an undifferentiated, primitive reservoir before core formation; (2) a silicate mantle after the core was extracted; and (3) various subsequently differentiated Martian mantle silicate reservoirs, including the source(s) of the shergottites. To construct a coherent model for Martian Pb isotopic evolution, the timing of the differentia-tion events that formed these reservoirsfirst needs to be constrained to model μ and κ values relevant to each stage. In turn, this then facilitates an assessment of the nature of the differentiation process(es), such as silicate only, core segregation, and/or metal/sulfide/ silicate.

These types of multiple-stage models have limitations and are unlikely to be“physically” accurate, in that they reflect solely the behavior of U, Th, and Pb and cannot describe mass transfer. A basic assumption in most, nonmixing based, multistage models of Pb evolution [e.g., Stacey and Kramers, 1975] is a continuity of Pb isotopic composition between the stages; i.e., the Pb isotopic composition at the beginning at each stage is the same as it was at the end of the previous stage. This condition is satisfied if U and Th are the only elements migrating between the reservoirs and Pb is immobile. If the transition between the stages is associated with migration of Pb from reservoir A to reservoir B, then the multistage model will still be valid for reservoir A but not for reservoir B. On a planetary scale the assumption of continuity in Pb isotopic composition is highly unrealistic. For the Earth, Pb evolution models have been proposed that are based

Figure 1. Schematic diagram of single-stage and second-stage Pb paleo-geochrons that illustrates a paleo-geochron for a given time and the effect of the differentiation of a mantle with a higherμ value resulting in two-stage Pb growth.

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on physical, not purely mathematical, constraints involving mass transfer and continuousflow between disparate isotopic reservoirs [e.g., Zartman and Haines, 1988; Kramers and Tolstikhin, 1997]. However, the Martian sample set is currently limited to six different meteorite types and there is no unified theory to explain mass transfer on Mars (cf. plate tectonics on Earth). As such, a mathematically coherent model, while unlikely to be an accurate reflection of physical mechanisms, can provide a first-order understanding of Martian U-Th-Pb evolution.

1.4. Determining Initial Pb

Several studies have attempted to constrain the Pb isotopic systematics of Mars [e.g., Borg et al., 2005; Bouvier et al., 2008, 2009; Chen and Wasserburg, 1986; Gaffney et al., 2007; Jagoutz, 1991; Nakamura et al., 1982]. The focus of these studies has been to determine the Pb isotopic compositions of mineral separates in Martian meteorites via traditional dissolution and solution analysis procedures. However, solution analyses of minerals that contain heterogeneous Pb isotopic compositions (e.g., from U-bearing microin-clusions, different U concentrations, or contamination) will be homogenized and thus may not yield the true initial Pb. Feldspars, specifically, in this case plagioclase or shocked plagioclase (maskelynite), contain relatively abundant Pb and very little to no U [e.g., Gancarz and Wasserburg, 1977; Oversby, 1978]. Therefore, these grains should record the true initial Pb isotopic composition of each meteorite. Secondary ion mass spectrometry (SIMS) offers an in situ technique that minimizes crystal boundary/surface contamination by targeting the middle of crystals while avoiding U-bearing inclusions. Although individual measurements may be relatively imprecise, this method can produce large, statistically significant data sets. By utilizing SIMS in this study, the least radiogenic and true initial Pb has been obtained from maskelynite grains in ALH84001 and three enriched shergottites (Zagami, Roberts Massif (RBT) 04262, and Larkman Nunatuk (LAR) 12011).

2. Samples

Allan Hills (ALH) 84001 is a coarse-grained orthopyroxenite. It contains 97% orthopyroxene and minor chro-mite, maskelynite, phosphate, augite, olivine, pyrite, and secondary carbonate [Mittlefehldt, 1994]. ALH84001 has a crystallization age of 4.091 ± 0.030 Ga determined by176Lu-176Hf (2σ [Lapen et al., 2010]). Roberts Massif

Figure 2. Weighted average age determination for the enriched shergottites in this study. Circles indicate176Lu-176Hf isochron ages. Squares indicate87Rb-87Sr isochron ages. Triangles indicate147Sm-143Nd isochron ages. Diamonds indicate

238

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Table 1. Measur ed Maske lynite and Pyroxe ne Pb Isotopic Co mpositions a Sample Type Fall/Find Mineral 206 Pb/ 204 Pb 2σ 207 Pb/ 204 Pb 2σ 208 Pb/ 204 Pb 2σ 204 Pb/ 206 Pb 2σ 207 Pb/ 206 Pb 2σ 208 Pb/ 206 Pb 2σ U/Pb 2σ 208 Pb (c/s) ALH84001 Orthopyroxenite Find Maskelynite 9.51 0.77 10.96 0.88 28.1 2.21 0.105 0.009 1.15 0.04 2.95 0.08 13 Maskelynite 9.60 0.74 10.90 0.84 28.0 2.10 0.104 0.008 1.14 0.03 2.91 0.07 14 Maskelynite 9.61 0.92 11.23 1.07 28.5 2.65 0.104 0.010 1.17 0.04 2.96 0.09 10 Maskelynite 9.65 0.80 10.66 0.89 26.9 2.18 0.104 0.009 1.10 0.04 2.78 0.08 12 Maskelynite 9.70 0.79 10.85 0.88 27.9 2.20 0.103 0.008 1.12 0.04 2.87 0.08 13 Maskelynite 9.74 0.94 11.47 1.10 29.1 2.74 0.103 0.010 1.18 0.05 2.99 0.10 b.d. 10 Maskelynite 9.78 0.72 11.10 0.81 28.1 2.01 0.102 0.008 1.13 0.03 2.87 0.07 16 Maskelynite 9.97 0.73 11.60 0.84 29.4 2.09 0.100 0.007 1.16 0.03 2.94 0.07 17 Maskelynite 10.00 0.59 11.50 0.68 30.1 1.73 0.100 0.006 1.15 0.03 3.00 0.06 b.d. 27 Maskelynite 10.04 0.78 11.44 0.88 29.7 2.24 0.100 0.008 1.14 0.03 2.96 0.08 b.d. 16 Maskelynite 10.14 0.39 11.66 0.44 30.4 1.13 0.099 0.004 1.15 0.02 2.99 0.04 63 Maskelynite 10.16 0.35 11.51 0.40 30.1 1.02 0.098 0.003 1.13 0.02 2.96 0.03 76 Maskelynite 10.17 0.42 11.64 0.48 30.9 1.26 0.098 0.004 1.14 0.02 3.04 0.04 0.001 0.002 55 Maskelynite 10.29 0.63 11.53 0.70 29.4 1.76 0.097 0.006 1.12 0.03 2.86 0.06 24 Maskelynite 10.44 0.62 12.01 0.71 30.0 1.73 0.096 0.006 1.15 0.03 2.87 0.06 0.004 0.003 27 Maskelynite 10.54 0.96 11.70 1.06 30.0 2.65 0.095 0.009 1.11 0.04 2.84 0.08 11 Maskelynite 10.82 0.85 12.31 0.96 32.1 2.44 0.092 0.007 1.14 0.03 2.96 0.08 b.d. 16 Maskelynite 10.54 0.79 11.59 0.86 30.6 2.22 0.095 0.007 1.10 0.03 2.90 0.07 b.d. 17 Excluded × Maskelynite 10.83 0.96 12.11 1.07 31.3 2.70 0.092 0.008 1.12 0.04 2.89 0.08 12 Excluded × Maskelynite 10.93 0.89 12.12 0.98 32.4 2.57 0.091 0.007 1.11 0.03 2.96 0.08 15 Excluded × Maskelynite 10.97 0.88 12.46 0.99 32.3 2.52 0.091 0.007 1.13 0.03 2.94 0.07 15 Excluded × Maskelynite 11.58 0.52 12.51 0.56 31.4 1.39 0.086 0.004 1.08 0.02 2.71 0.04 47 Excluded × Maskelynite 12.33 0.96 13.28 1.03 35.5 2.71 0.081 0.006 1.08 0.03 2.87 0.07 18 Excluded × Maskelynite 12.84 1.26 13.24 1.30 33.5 3.24 0.078 0.008 1.03 0.04 2.61 0.08 11 Excluded × Maskelynite 13.38 1.29 13.16 1.27 32.9 3.11 0.075 0.007 0.98 0.03 2.46 0.07 11 Orthopyroxene 18.54 0.64 15.37 0.54 37.8 1.30 0.054 0.002 0.829 0.009 2.04 0.02 94 Orthopyroxene 18.57 0.47 15.21 0.39 37.6 0.95 0.054 0.001 0.819 0.007 2.03 0.01 174 Orthopyroxene 19.54 0.74 15.96 0.61 39.3 1.48 0.051 0.002 0.817 0.010 2.01 0.02 81 LAR12011 Enriched olivine-phyric shergottite Find Maskelynite 13.08 0.49 12.31 0.46 31.5 1.15 0.076 0.003 0.941 0.013 2.40 0.03 0.004 0.002 73 Maskelynite 13.24 0.65 12.63 0.62 31.7 1.52 0.076 0.004 0.954 0.017 2.39 0.04 43 Maskelynite 13.34 0.56 12.47 0.52 31.8 1.31 0.075 0.003 0.935 0.014 2.39 0.03 59 Maskelynite 13.35 0.60 12.69 0.57 32.1 1.41 0.075 0.003 0.950 0.016 2.40 0.03 0.05 0.008 52 Maskelynite 13.40 0.72 12.66 0.68 32.4 1.70 0.075 0.004 0.945 0.018 2.42 0.04 36 Maskelynite 13.48 0.68 12.97 0.65 32.8 1.61 0.074 0.004 0.962 0.017 2.43 0.04 b.d. 42 Maskelynite 13.50 0.32 12.81 0.31 32.6 0.76 0.074 0.002 0.949 0.008 2.41 0.02 186 Maskelynite 13.59 0.49 13.06 0.47 33.6 1.18 0.074 0.003 0.960 0.012 2.47 0.03 0.008 0.004 85 Maskelynite 13.59 0.46 12.73 0.43 32.9 1.10 0.074 0.003 0.936 0.012 2.42 0.02 0.005 0.002 92 Maskelynite 13.65 0.41 12.94 0.39 33.1 0.99 0.073 0.002 0.948 0.010 2.42 0.02 0.04 0.003 116 Maskelynite 13.67 0.56 13.12 0.54 33.2 1.33 0.073 0.003 0.960 0.014 2.42 0.03 b.d. 64 Maskelynite 13.70 0.43 12.95 0.41 32.8 1.01 0.073 0.002 0.945 0.011 2.39 0.02 108 Maskelynite 13.75 0.55 13.12 0.53 33.5 1.31 0.073 0.003 0.954 0.014 2.43 0.03 0.015 0.010 68 Maskelynite 13.77 0.90 13.21 0.86 32.6 2.09 0.073 0.005 0.959 0.023 2.36 0.05 25 Maskelynite 13.78 0.34 13.00 0.32 33.6 0.81 0.073 0.002 0.943 0.008 2.44 0.02 180 Maskelynite 13.79 0.49 12.76 0.45 33.0 1.14 0.073 0.003 0.925 0.012 2.39 0.03 0.05 0.017 86 Maskelynite 13.79 0.42 13.16 0.40 33.7 1.00 0.073 0.002 0.954 0.010 2.44 0.02 120 Maskelynite 13.80 0.44 13.06 0.42 33.5 1.05 0.072 0.002 0.946 0.011 2.43 0.02 b.d. 106 Maskelynite 13.81 0.26 13.06 0.25 33.2 0.61 0.072 0.001 0.946 0.006 2.40 0.01 0.0010 0.0004 305 Maskelynite 13.82 0.44 13.05 0.42 33.6 1.06 0.072 0.002 0.944 0.011 2.43 0.02 106 Maskelynite 13.82 0.31 13.19 0.30 33.6 0.74 0.072 0.002 0.954 0.008 2.43 0.02 0.0010 0.0003 213 Maskelynite 13.87 0.46 13.17 0.44 33.6 1.10 0.072 0.002 0.949 0.011 2.42 0.02 b.d. 97 Maskelynite 13.89 0.32 13.18 0.31 33.7 0.76 0.072 0.002 0.949 0.008 2.42 0.02 0.020 0.002 204

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Table 1. (continued) Sample Type Fall/Find Mineral 206 Pb/ 204 Pb 2σ 207 Pb/ 204 Pb 2σ 208 Pb/ 204 Pb 2σ 204 Pb/ 206 Pb 2σ 207 Pb/ 206 Pb 2σ 208 Pb/ 206 Pb 2σ U/Pb 2σ 208 Pb (c/s) Maskelynite 13.94 0.46 13.15 0.44 33.5 1.09 0.072 0.002 0.943 0.011 2.40 0.02 0.009 0.002 98 Maskelynite 13.98 0.85 13.27 0.81 33.5 2.01 0.072 0.004 0.949 0.021 2.39 0.04 29 Maskelynite 14.02 0.47 13.23 0.44 33.5 1.10 0.071 0.002 0.943 0.011 2.39 0.02 97 Maskelynite 14.02 0.44 13.41 0.42 34.2 1.06 0.071 0.002 0.956 0.011 2.44 0.02 0.0012 0.0004 110 Maskelynite 14.04 0.63 13.52 0.61 34.3 1.52 0.071 0.003 0.963 0.015 2.44 0.03 0.005 0.002 54 Maskelynite 14.09 0.50 13.38 0.47 34.3 1.19 0.071 0.003 0.950 0.012 2.43 0.03 88 Maskelynite 14.11 0.51 13.28 0.48 33.9 1.19 0.071 0.003 0.941 0.012 2.40 0.03 85 Maskelynite 14.15 0.51 13.42 0.49 34.6 1.23 0.071 0.003 0.948 0.012 2.44 0.03 b.d. 84 Maskelynite 14.17 0.62 13.60 0.59 34.3 1.47 0.071 0.003 0.960 0.015 2.42 0.03 0.003 0.001 58 Maskelynite 14.21 0.47 13.33 0.44 34.3 1.10 0.070 0.002 0.937 0.011 2.41 0.02 0.287 0.013 103 Maskelynite 14.26 0.55 13.48 0.52 34.3 1.29 0.070 0.003 0.945 0.013 2.41 0.03 75 Maskelynite 14.29 0.53 13.59 0.50 34.6 1.25 0.070 0.003 0.951 0.012 2.42 0.03 82 Maskelynite 14.34 0.66 13.75 0.64 34.9 1.59 0.070 0.003 0.959 0.016 2.43 0.03 52 Maskelynite 14.38 0.58 13.59 0.55 34.1 1.36 0.070 0.003 0.945 0.014 2.37 0.03 67 RBT04261 Enriched olivine-phyric shergotitte Find Maskelynite 13.87 0.73 13.10 0.69 33.1 1.70 0.072 0.004 0.944 0.018 2.39 0.04 b.d. 39 Maskelynite 13.66 0.75 12.73 0.70 32.8 1.76 0.073 0.004 0.932 0.018 2.40 0.04 b.d. 36 Maskelynite 13.75 0.49 13.06 0.47 33.1 1.17 0.073 0.003 0.950 0.012 2.41 0.03 0.0011 0.001 83 Maskelynite 13.76 0.51 13.13 0.49 33.3 1.22 0.073 0.003 0.954 0.013 2.42 0.03 b.d. 78 Maskelynite 12.84 1.64 12.35 1.58 31.2 3.91 0.078 0.010 0.961 0.045 2.43 0.10 b.d. 6 Maskelynite 14.25 1.40 13.99 1.38 34.4 3.33 0.070 0.007 0.981 0.034 2.41 0.07 b.d. 11 Maskelynite 13.70 1.06 13.05 1.01 32.7 2.48 0.073 0.006 0.952 0.026 2.39 0.06 0.0005 0.001 18 Maskelynite 13.67 1.08 12.95 1.02 32.8 2.53 0.073 0.006 0.947 0.027 2.40 0.06 b.d. 17 Maskelynite 13.78 1.36 13.48 1.33 33.5 3.25 0.073 0.007 0.978 0.034 2.43 0.07 b.d. 11 Maskelynite 14.79 1.46 13.71 1.36 34.6 3.36 0.068 0.007 0.927 0.032 2.34 0.07 b.d. 11 Maskelynite 13.68 0.89 13.24 0.86 33.5 2.14 0.073 0.005 0.967 0.023 2.45 0.05 b.d. 26 Maskelynite 13.71 0.79 12.92 0.74 33.0 1.86 0.073 0.004 0.943 0.020 2.41 0.04 b.d. 32 Maskelynite 14.84 1.40 14.24 1.34 35.5 3.28 0.067 0.006 0.959 0.031 2.39 0.06 b.d. 13 Maskelynite 14.62 0.67 13.72 0.63 35.1 1.59 0.068 0.003 0.938 0.015 2.40 0.03 b.d. 53 Zagami

Encriched basaltic shergottite

Fall Maskelynite 12.85 1.15 12.43 1.11 31.6 2.77 0.078 0.007 0.967 0.032 2.45 0.07 0.04 0.007 13 Maskelynite 12.86 0.67 12.72 0.67 31.7 1.63 0.078 0.004 0.989 0.019 2.46 0.04 b.d. 39 Maskelynite 12.90 0.55 12.22 0.52 31.4 1.31 0.078 0.003 0.947 0.015 2.43 0.03 58 Maskelynite 12.92 0.56 12.49 0.55 31.7 1.36 0.077 0.003 0.966 0.015 2.45 0.03 56 Maskelynite 13.07 1.04 12.55 1.00 32.2 2.51 0.077 0.006 0.960 0.028 2.46 0.06 17 Maskelynite 13.10 0.62 12.53 0.59 31.6 1.47 0.076 0.004 0.957 0.017 2.41 0.03 48 Maskelynite 13.13 1.00 12.54 0.96 32.3 2.42 0.076 0.006 0.955 0.027 2.46 0.06 19 Maskelynite 13.18 0.82 12.80 0.80 32.2 1.96 0.076 0.005 0.971 0.022 2.44 0.05 0.006 0.003 28 Maskelynite 13.31 0.98 12.81 0.95 32.2 2.33 0.075 0.006 0.962 0.026 2.42 0.05 0.0001 0.0001 20 Maskelynite 13.34 0.39 12.69 0.37 32.3 0.93 0.075 0.002 0.951 0.010 2.42 0.02 0.006 0.003 126 Maskelynite 13.43 0.62 12.95 0.60 32.8 1.49 0.074 0.003 0.964 0.016 2.44 0.03 52 Maskelynite 13.44 0.68 13.04 0.66 33.2 1.64 0.074 0.004 0.970 0.018 2.47 0.04 44 Maskelynite 13.46 0.77 13.02 0.75 33.2 1.87 0.074 0.004 0.967 0.020 2.47 0.04 34 Maskelynite 13.46 1.09 12.22 0.99 31.7 2.51 0.074 0.006 0.907 0.027 2.35 0.06 16 Maskelynite 13.48 0.27 13.01 0.26 33.1 0.65 0.074 0.001 0.965 0.007 2.46 0.01 279 Maskelynite 13.53 0.42 12.99 0.40 33.1 1.00 0.074 0.002 0.960 0.011 2.45 0.02 0.013 0.009 117 Maskelynite 13.55 0.78 13.16 0.76 33.3 1.89 0.074 0.004 0.970 0.020 2.46 0.04 34 Maskelynite 13.61 0.78 12.86 0.73 32.9 1.84 0.073 0.004 0.945 0.019 2.41 0.04 b.d. 34 Maskelynite 13.65 0.81 13.07 0.77 33.5 1.94 0.073 0.004 0.957 0.020 2.45 0.04 29 Maskelynite 13.79 1.07 13.25 1.03 33.9 2.59 0.073 0.006 0.960 0.027 2.46 0.06 17 Excluded × Maskelynite 13.81 0.12 13.25 0.11 33.8 0.28 0.072 0.001 0.959 0.003 2.45 0.01 1426

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(RBT) 04262 is an enriched lherzolitic shergottite. It is a coarse-grained rock with ~40% olivine, 30–40% pyroxene, and 15–20% maskelynite, with minor merrillite, chromite, and ilmenite, and has been extensively described by Usui et al. [2010]. The age of RBT04262 has been determined by

176Lu-176Hf,87Rb-87Sr, and147Sm-143Nd, which yield ages of

200 ± 21 Ma, 167 ± 6 Ma, and 174 ± 14 Ma, respectively (all errors 2σ [Lapen et al., 2008; Shih et al., 2009]). Additionally, baddeleyite in RBT04262 yields a U-Pb age of ~200 Ma [Niihara, 2011]. Zagami is an enriched basaltic shergottite that contains at least two different textural types (coarse-grained andfine-grained), with ~77% pyroxene, ~18% maskelynite, and 1–2% each of oxides and mesostasis [McCoy et al., 1992]. The crystallization age of Zagami has been determined by87Rb-87Sr,147Sm-143Nd, and238U-206Pb, which yield ages of 166 ± 6 Ma, 166 ± 12 Ma, and 156 ± 6 Ma, respectively (all errors 2σ [Borg et al., 2005]). Larkman Nunatuk (LAR) 12011 is an enriched olivine-phyric shergottite. Due to its recent discovery (2012–2013), there is limited information available for this sample. However, LAR12011 is most likely paired with LAR06319 [Benedix and Roberts, 2014; Liu et al., 2014; Righter et al., 2015]. LAR12011 contains millimeter-sized olivine crystals in a matrix of pyroxene and maskelynite. No isotopic information is available for LAR12011, but the age of paired meteorite LAR06319 has been determined by 147Sm-143Nd and176Lu-176Hf, which have yielded ages of 183 ± 12 and 197 ± 29 Ma, respectively (all errors 2σ [Shafer et al., 2010]). A weighted average for each meteorite yields an age of 170 ± 25 for RBT04262, 162 ± 15 for Zagami, and 185 ± 11 for LAR 12011. A summary for all of the enriched shergottite ages listed above is presented in Figure 2.

3. Analytical Methods

The Pb isotopic compositions of maskelynite were deter-mined using secondary ion mass spectrometry (SIMS) on polished epoxy mounts or thin sections of each sample. Before analysis, each sample was cleaned in alternating 1 min ultrasonic baths of water and ethanol. After thorough washing, a 30 nm coating of Au was applied to the surface. All measurements were conducted using a CAMECA IMS1280 instrument at the Swedish Museum of Natural History, Stockholm (NordSIM facility), using previously described experimental protocols [Bellucci et al., 2015; Whitehouse et al., 2005]. An area of 35 × 35μm was rastered for 70 s prior to Pb isotopic analysis to remove the gold and further minimize surface contamination. A 300μm aper-ture was used to project a 12–15 nA O2 primary beam with

a slightly elliptical 30μm (long axis) spot on the surface of the sample. All analyses were conducted in multicollector mode at a mass resolution of 4860 (M/ΔM), using an NMR field sensor in regulation mode to maintain the stability of the magneticfield. Lead isotopic ratios were measured in a multicollector array of low-noise (<0.03 c/s) ion-counting

Table 1. (continued) Sample Type Fall/Find Mineral 206 Pb/ 204 Pb 2σ 207 Pb/ 204 Pb 2σ 208 Pb/ 204 Pb 2σ 204 Pb/ 206 Pb 2σ 207 Pb/ 206 Pb 2σ 208 Pb/ 206 Pb 2σ U/Pb 2σ 208 Pb (c/s) Excluded × Maskelynite 13.81 0.34 13.22 0.33 33.6 0.81 0.072 0.002 0.957 0.008 2.43 0.02 169 Excluded × Maskelynite 13.85 0.53 13.10 0.50 33.5 1.26 0.072 0.003 0.946 0.013 2.42 0.03 76 Excluded × Maskelynite 13.85 0.48 13.39 0.47 33.9 1.16 0.072 0.003 0.967 0.012 2.44 0.03 94 Excluded × Maskelynite 13.89 1.07 12.99 1.00 33.4 2.52 0.072 0.006 0.935 0.026 2.40 0.05 19 Excluded × Maskelynite 13.92 0.17 13.30 0.16 34.2 0.41 0.072 0.001 0.955 0.004 2.46 0.01 692 Excluded × Maskelynite 14.01 0.59 13.32 0.56 34.3 1.42 0.071 0.003 0.950 0.014 2.45 0.03 59 Excluded × Maskelynite 14.07 0.65 13.30 0.62 33.7 1.54 0.071 0.003 0.945 0.016 2.40 0.03 53 Excluded × Maskelynite 14.09 1.55 13.80 1.52 35.0 3.78 0.071 0.008 0.979 0.038 2.48 0.08 9 Excluded × Maskelynite 14.12 0.48 13.39 0.46 34.2 1.15 0.071 0.002 0.948 0.011 2.42 0.02 89 Excluded × Maskelynite 14.19 0.20 13.26 0.18 34.1 0.47 0.070 0.001 0.934 0.005 2.40 0.01 537 Excluded × Maskelynite 14.23 1.24 13.61 1.19 34.1 2.93 0.070 0.006 0.956 0.029 2.39 0.06 14 a b.d. = below detection for U. De tection lim its ~2 ng/g in basa ltic glass BCR 2-g. Bold valu es in the table ar e indic ated for values out side of statistica lly dete rmin ed x-y weig hted means .

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electron multipliers for 160 cycles with a count time of 10 s, resulting in a total collection time of 1600 s. Isotopic ratios were calculated using integrated means for all analyses. Mass fractionation and gain calibrations between detectors were made by bracketing the unknowns with analyses of U.S. Geological Survey basal-tic glass reference material BCR-2G (~11μg/g Pb). The degree of mass fractionation in Pb isotopes measured by SIMS between different silicate glasses is negligible and unresolvable at <0.2 amu 1 [e.g., Belshaw et al., 1994; Shimizu and Hart, 1982] and is encompassed by the large, individual uncertainty estimates. The largest poten-tial source of inaccuracy is in the relative gain differences between ion counters, which is accounted for by bracketing unknown measurements with BCR-2G and correcting iso-topic measurements using the accepted values of this reference material [Woodhead and Hergt, 2000]. External reproducibility in208Pb/206Pb and207Pb/206Pb was 0.3% and in 208Pb/204Pb, 207Pb/204Pb, and 206Pb/204Pb was 0.9%, 0.7%, and 1.0%, respectively (all errors 2σ). Following Pb isotopic measurements, the238U/208Pb ratio of several analytical spots was measured by peak jumping between masses208Pb and238U. Ions were collected in the central electron multiplier, and measured 238U+/208Pb+ ratios were converted to approximate elemental U/Pb ratios using the relative sensitivity factors determined by measuring BCR-2G and the natural isotopic abundances of U and Pb, following previously established procedures [Bellucci et al., 2015].

3. Results

Results from Pb isotopic analyses of maskelynite and pyroxene grains are presented in Tables 1 and 2, and U/Pb elemental ratios for individual spots are provided in Table 1.

Initial Pb compositions were calculated by taking the least radiogenic Pb isotopic population measured in maskely-nite and performing an X-Y weighted mean to determine a statistically identical population (ISOPLOT 4.15 [Ludwig, 2012]) in207Pb/204Pb versus206Pb/204Pb and208Pb/204Pb versus206Pb/204Pb. To perform these calculations, ana-lyses that did notfit in the X-Y weighted mean and were outside of the statistical population of least radiogenic were rejected. Those analyses are indicated in Table 1. Results from the X-Y weighted calculations are presented in Table 2 and Figures 3 and 4. There was a larger spread in 208Pb/206Pb in ALH84001 and LAR12011, so an X-Y mean could not be calculated for 208Pb/206Pb versus

204Pb/206Pb. Instead, the208Pb/206Pb values

correspond-ing to the analyses used in the X-Y 207Pb/206Pb versus

204Pb/206Pb weighted mean calculations were used in a

one-dimensional weighted mean calculation (Figure 4).

Table 2. Initia l P b Comp osit ions of M artian Meteo rites and M odel Paramet ers Sampl e Type Fall/Fin d 206 Pb / 204 Pb 2σ 207 Pb/ 204 Pb 2σ 208 Pb/ 204 Pb 2σ 204 Pb/ 206 Pb 2σ 207 Pb / 206 Pb 2σ 208 Pb/ 206 Pb 2σ n ALH8400 1 O rthopyr oxenite Find 10.07 0.17 11.47 0.19 29 .7 0.6 0.0989 0.0016 1.138 0.00 7 2.949 0.030 1 8 LAR120 11 E nriched olivi ne-phy ric shergo ttite Find 13.81 0.10 13.09 0.09 33 .4 0.2 0.0723 0.0005 0.948 0.002 2.417 0.006 3 7 RBT0426 1 E nriched olivi ne-phy ric shergo titte Find 13.87 0.22 13.16 0.21 33 .4 0.5 0.0719 0.0013 0.949 0.006 2.405 0.011 1 4 Zagami Enc riched basaltic shergo ttite Fa ll 13.32 0.16 12.81 0.16 32 .6 0.4 0.07 49 0.0009 0.961 0.004 2.443 0.007 2 0 Canyon Diabl o T roilite at t =4 5 6 7 M a a,b 9.31 10.29 29.48 0.1074 1.106 3.167 Martian man tle at 4513 Ga c 9.33 10.33 29.49 0.1072 1.107 3.160 a Age of solar system de fi ned by Conn elly et al . [201 2]. b Values from Che n and W asserburg [1983] . c Age of ma ntle diff erentia tion of Lapen et al . [2010] .

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The initial Pb composition in ALH84001 is the least radiogenic of any Martian meteorite measured to date (204Pb/206Pb = 0.0989 ± 0.0016, 2σ, n = 18; all listed uncertainties for the reported measurements here are 2σ), significantly less radiogenic than that implied by the solution measurements of Bouvier et al. [2009], and lies within error of the 4.09 Ga single-stage paleo-geochron (Figure 5). The Pb isotopic composition of the orthopyroxene measured here is also significantly less radiogenic than the residue analyses of successive orthopyroxene leachates reported by Bouvier et al. [2009] (204Pb/206Pb of 0.053 versus 0.014). This implies that there is a U-bearing phase in ALH84001 that was incorporated into the dissolved orthopyroxene fraction of Bouvier et al. [2009]. This phase is likely phosphate, as phosphates have been documented in ALH84001 and contain relatively abundant U and Th.

The initial Pb isotopic compositions of the enriched shergottites are similar, with Zagami being the least radiogenic (204Pb/206Pb = 0.0749 ± 0.0009, n = 20) and LAR12011 and RBT04262 being indistinguishable (204Pb/206Pb of 0.0723 ± 0.0.0005, n = 37; 0.0719 ± 0.0.0013, n = 14, respectively). In comparison to solution measurements of separated maskelynite in the same meteorites, all of the corresponding data presented here are less radiogenic. For example, the204Pb/206Pb of maskelynite measured by Borg et al. [2005] for Zagami was 0.07399 ± 0.00008 2σ (the204Pb/206Pb of maskelynite measured by Bouvier et al. [2009] for RBT04262 was 0.07062 ± 0.00001 2σ). There are no literature Pb isotopic data available for LAR12011. These combined results confirm the ability of SIMS to target the least radiogenic parts of a mineral. Additionally, the U/Pb ratio measured in each sample does not correlate with Pb isotopic compositions, and in most cases U is effectively below detection limit (Table 1).

Figure 3. X-Y weighted means for207Pb/206Pb versus204Pb/206Pb for maskelynite measured via SIMS in this study. Individual error ellipses are 2σ. Colored ellipses represent 2σ weighted average.

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In addition to determining the initial Pb isotopic composition of ALH84001, orthopyroxene grains were also analyzed to determine the Pb-Pb crystallization age. These results yield a two-point isochron with an age of 4089 ± 73 Ma (2σ; Figure 5). This age is identical, within error to the 4091 ± 30 Ma determined by176Lu-176Hf [Lapen et al., 2010] and previously determined Pb-Pb isochron age of 4074 ± 99 Ma [Bouvier et al., 2009]. Pyroxenes were also measured in each of the enriched shergottites, but the uncertainties were too large for meaningful age calculations to be made.

4. Discussion

4.1. A Time-Integrated Forward Pb Isotopic Model, Initial Pb Model Ages, and Implications for Martian Core Formation

Despite being the least radiogenic Pb measured in each of the Martian meteorites, the initial Pb determined for each sample is slightly more radiogenic than predicted by the single-stage paleo-geochrons correspond-ing to the ages determined by147Sm-143Nd,87Rb-87Sr, and176Lu-176Hf for all samples (Figure 6). Lying on a single-stage paleo-geochron would be expected to intercept the initial Pb compositions if the source(s) for these rocks evolved as a closed system with afixed, single μ value since accretion and an assumed CDT initial composition. As such, these results explicitly imply at least a two-stage differentiation history. One approach to establishing a Pb isotopic model for these samples is to use previously established reservoir chronology as

Figure 4. X-Y weighted means and weighted averages for208Pb/206Pb versus204Pb/206Pb in Zagami and RBT04262. One-dimensional weighted averages were used in ALH84001 and LAR12011, where the X-Y spread in208Pb/206Pb was too great to define a single, statistical population of data; see text for further information. Individual error ellipses and error bars are 2σ. Colored ellipses represent 2σ weighted average.

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a framework for a time forward integrated Pb isotopic model. As noted earlier, core formation and sulfide precipitation during silicate differentiation should, in theory, have had the largest effect onμ values in the source reservoir, shown to be similar in Sm-Nd, Lu-Hf isotopic systems, for ALH84001 and the enriched sher-gottites. Therefore, there are two events that should be considered in the model presented here: core formation (1–15 Ma after calcium–aluminium-rich inclusion (CAI)) and the large silicate differentiation event (4513 Ma). The initial Pb isotopic composition of Mars for this model is assumed to be that of CDT. This is considered to be a reasonable proxy for the Martian mantle composition post core formation for the following reasons: (1) CDT is homogenous in all measured IAB meteorites [Göpel et al., 1985]; (2) the formation ages of the IAB core and Martian core are similar at 1.8+2.3/ 2.0Ma and 1–15 Ma after CAI, respectively [Schultz et al., 2009;

Dauphas and Pourmand, 2011; Foley et al., 2005; Kleine et al., 2004; Nimmo and Kleine, 2007]; and (3) it is a rea-sonable assumption that the materials that accreted to form both Mars and the IAB parent body had chon-dritic low-μ values such that radiogenic growth during the 1–15 Ma was extremely limited.

The formation age of the common source reservoir for ALH84001 and the enriched shergottites has been estimated at 4.513 Ga based on coupled Lu-Hf and Sm-Nd systematics that evolved from a chondritic reservoir [Borg et al., 2003; Lapen et al., 2010]. To test this forward model, an initial Pb model age can be calculated based on the slope between the measured initial Pb and the modeled Pb isotopic composition at 4.513 Ga based on a varying theμ1value (Table 3 and Figure 7). This is assuming that the time between

accretion and core formation (1–15 Ma after CAI) had no effect on the initial Martian mantle Pb isotopic composition, as stated previously. The correctness of this model can be evaluated by determining the differences between model ages and accepted crystallization ages inΔt (whereΔt= age of

crystalliza-tion initial Pb model age). Following the approach of Stacey and Kramers [1975], a μ1 value from

4.567 Ga to 4.513 Ga of 1.2 minimizes Δt (Δt= 0). This value results in Pb model ages for Zagami,

RBT04262, and LAR12011 of 170, 212, and 134 Ma, respectively. Thisμ value is not significantly elevated over ordinary and enstatite chondritic composition (~1 [e.g., Göpel et al., 1994; Tera, 1983]). Therefore, this μ value likely represents that of the material used to build Mars. Given the identical Cr isotopic composi-tions of ordinary and enstatite chondrites, and all other inner solar system objects, this value probably represents their primordialμ value [e.g., Trinquier et al., 2009].

Figure 5.207Pb/206Pb versus204Pb/206Pb isochron calculation for ALH84001 and 4.09 Ga single-stage paleo-geochron. Black ellipses represent maskelynite, and red ellipses represent orthopyroxene. All error ellipses are 2σ.

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While the initial Pb composition of ALH84001 plots within error of both the single-stage and second-stage 4.09 Ga paleo-geochron, the initial Pb model age is consistently younger (~3.9 Ga) than the accepted age for all μ1 values (Table 3). This discrepancy could be a function of one or both of the

following reasons: the measurements here do not represent true initial Pb and/or the source region for ALH84001 and the enriched shergottites cannot be linked to the same U-Pb reservoir, in contrast to inferences from the Lu-Hf and Sm-Nd isotopic systematics of these meteorites. Unfortunately, the aforementioned hypotheses are untestable, given the uncertainties in the SIMS measurements presented here. However, despite the model age for ALH84001 being 3.9 Ga instead of 4.09 Ga, it still is a reflection of the large absolute difference in ages needed to explain the initial Pb compositions of ALH84001 and the enriched shergottites.

Given that the best initial model ages require a chondritic-likeμ value from 4.567 Ga to 4.513 Ga and core formation happened at 1–15 Ma after CAI, core segregation likely had very little effect on the μ value of the Martian mantle, an observation that is in agreement with the partition coefficient experiments of Malavergne et al. [2007]. Afinal complication in Pb isotopic modeling of early planetary differentiation arises from recent experimental evidence suggesting that U may be sequestered into the core [e.g., Wohlers and Wood, 2015], challenging the long-held assumption that only Pb, in this isotopic system, partitions into the cores of planetary bodies [e.g., Allègre et al., 1982]. However, the determination of absolute concentrations during core segregation is far beyond the scope of this paper, as Pb isotopic information only yieldsμ and

Figure 6.207Pb/206Pb versus204Pb/206Pb diagrams for initial Pb for each meteorite measured here. Shown are 0.17 Ga/4.1 Ga single-stage paleo-geochrons and second-stage paleo-geochrons deriving from a Pb reservoir with a composition reflecting Pb growth from 4.567 Ga to 4.513 Ga with a μ value of 1.2. Included in appropriate figure panes are corre-sponding solution measurements of maskelynite in the same meteorites by Borg et al. [2003], Bouvier et al. [2005], and Bouvier et al. [2009].

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κ information. Furthermore, this problem is likely intractable given both the limited sample availability and current information on partition coefficient experiments.

4.2. Second-Stageμ Values in the Martian Mantle

Second-stage paleo-geochrons with slopes corresponding to the ages of each sample and the source reservoir Pb composition from 4.513 Ga (which evolved from 4.567 Ga with aμ value of 1.2) are shown in Figure 8. There is a significantly better fit than a single-stage paleo-geochron, which is in agreement with the model age calculations described above (Figures 6 and 8). As the measured initial compositions now lie on a second-stage paleo-geochron corresponding to their crystallization age, bothμ and κ values can be calculated (Figure 8). Zagami has a calculatedμ value of 4.1. RBT04262 and LAR12011 have almost identical initial Pb isotopic compositions with a calculated sourceμ value of 4.6. Both of these μ values have growth curves that intersect the initial Pb isotopic composition of ALH84001 (Figure 8). Aκ value of 3.4 is within error of all initial Pb measurements, although ALH84001 may have an initial Th/U ratio that is less than the enriched shergottites (Figure 8). Theseμ values are independent of all of the calculations for the model ages listed above. This increase inμ value from 1.2 to 4.6 could have occurred during a mantle differentiation event with the precipitation of sulfides [Gaffney et al., 2007]. It should be noted that this model is highly dependent on the Lu-Hf and Sm-Nd isotopic systematics put forth by previous workers and assumes that this silicate differentiation event also precipitated sulfides, thereby affect-ing all three isotopic systems. It would be valuable to independently confirm these results with the new Pb isotopic data presented here. Unfortunately, given the amount of variables and equations in the Pb isotopic system, independent determinations of time of mantle differentiation,μ1andμ2values are impossible with the current data

set. Hopefully, with the addition of more samples with significantly different crystallization ages, these calculations can be made.

4.3. Implications for Pb-Pb Isotopic Chronology

The Pb isotopic compositions of ALH84001 (initial Pb and pyroxene) and the enriched shergottites measured here can be used erro-neously to define an “isochron” of 4087 ± 93 (mean square weighted deviation (MSWD) = 4.6). This age is considered to have no signi fi-cance for the following reasons: (1) the samples are not cogenetic petrographically or geochemically as ALH84001 is an orthopyroxene cumulate, LAR12011 and RBT04262 are olivine-phyric shergottites, and Zagami is an enriched basaltic shergottite. As such, fitting them to an isochron violates the primary assumptions in Pb-Pb chron-ology listed above, and (2) the maskelynite grains analyzed in each sample have extremely low U/Pb ratios and in most cases undetect-able U (Tundetect-able 1) and, hence, have no ingrowth of radiogenic Pb. If there is no ingrowth of Pb in the maskelynite, these data are a re flec-tion of different Pb isotopic composiflec-tions at the time of crystallizaflec-tion. The goodness offit to the calculated isochron (MSWD = 4.6) between these data sets, thus, is coincidental and a reflection of placing very old (ALH84001 initial Pb and pyroxene) and young (shergottites) initial Pb, inappropriately, on the same isochron diagram.

Table 3. In itial Pb Mode l A g e s B a sed on Different m Values From 4.567 G a to 4 .513 Ga and D Valu es Crysta lization age a μ =0 μ =0 μ = 0.5 μ = 0.5 μ =1 μ =1 μ =1 .2 μ = 1.2 μ = 1.5 μ = 1.5 μ = 2.5 μ = 2.5 μ =3 μ =3 μ =4 μ =4 Sam ple Mode l age Δ b Model age Δ b Mode l age Δ b Mode l age Δ b Mode l age Δ b Mode l age Δ b Mode l age Δ b Model age Δ b ALH 84001 4090 3898 192 389 0 200 3880 210 3877 203 3 870 22 0 3850 24 0 3840 25 0 3 818 3880 RBT 04261 185 280 9 5 25 0 65 220 35 212 42 185 0 140 4 5 110 60 60 110 Zag ami 170 240 7 0 21 0 40 175 5 170 8 150 20 90 8 0 60 10 2 1 0 172 LAR 12011 162 200 3 8 18 0 18 145 17 134 51 110 52 60 10 2 2 0 1 6 5 4 0 225 Ave rage Δ t 68 41 8 0 24 7 6 10 9 169 2σ mean 33 27 30 54 30 3 3 61 66 a See Figu re 1 and discussion in the te xt. b W here D = crystallization age -calcula ted initial Pb mo del age, all ages in M a .

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As the major phase analyzed in this study is maskelynite, a shocked form of plagioclase, the potential effect of shock on the U-Th-Pb system needs to be assessed. Shock features permeate the shergottites [e.g., El Goresy et al., 2013, and references therein], and shock has had a well-documented profound effect on chronometers that have mobile daughter products (e.g., Ar-Ar [El Goresy et al., 2013; Jourdan et al., 2014]). El Goresy et al. [2013] further proposed that the radiogenic isotopic systems in the shergottites were likely partially or completely reset due to this shock.“Complete resetting” in isotopic systems homogenizes all individual minerals to a common isotopic composition, and this is kinetically impossible given the very short duration of heating and quenching. If the samples were“reset” during shock, all Pb would still be homogenous between mineral phases because there has not been sufficient time for new radiogenic Pb growth. Furthermore, shock has been documented to generate heterogeneous melt pockets with the isotopic composition of the surrounding minerals [e.g., Moser et al., 2013], not wholesale homogenization to a common value. Therefore, the homogeneity of our data (n> 14 for all samples on the scale of 30 μm), which cluster around the appropriate paleo-geochrons of each age, is more likely a reflection of true initial Pb instead of heterogeneously shock redistributed Pb. Additionally, the oldest cosmic exposure age, a proxy for ejection/impact age, for the samples studied here is 20 Ma and not 170 Ma [Nyquist et al., 1995, and references therein]. There are meteorites that have experienced multiple impacts, polymict impact breccias, which are no longer identified as basalts or cumulate rocks. To explain the Pb isotopic composi-tions measured here as shock resetting, a highly improbable and illogical set of events needed to have happened: (1) an impact at ~170 strong enough to convert plagioclase to maskelynite and reset the whole rock but not eject it from the surface and (2) a second impact at ~20 Ma strong enough to eject the rock from Mars but not reset the isotopic composition. Similarly, the whole rock Pb isotope composition for ALH84001 is significantly more radiogenic than that of the shergottites, which explicitly implies a whole rock μ value that is significantly higher [Bouvier et al., 2009]. If an impact strong enough to convert plagioclase to maskelynite reset the whole rock, the maskelynite composition for ALH84001 would be sig-nificantly more radiogenic than the shergottites, which is not the case. All of the samples have been ejected at around the same age, have been shocked, plagioclase was converted to maskelynite, and have vastly different initial Pb isotopic compositions. As such, the simplest explanation for this data set is that the initial Pb determined here in each meteorite is reflective of the Pb isotopic composition at the time of crystallization. Thus, the difference in initial Pb compositions determined for ALH84001 compared with those of the enriched shergottites requires ~4 Ga of Pb growth that could not have taken place in situ and

Figure 7. AverageΔt(crystallization age initial Pb model ages) for the enriched shergottites versusμ values from solar system initial (CDT) at 4.567 Ga to age of reservoir formation for ALH84001 and the enriched shergottites at 4.513 Ga [Borg et al., 2003; Lapen et al., 2010]. Solid line represents calculated averageΔtmodel age average for the enriched shergottites, and the dotted lines represent the 2σmeanfor the corresponding average. The minimization ofΔt(=0) yields a μ value of 1.2.

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therefore must reflect differences in crystallization ages of these meteorites. Lastly, as these are the least radiogenic Pb isotopic compositions measured for each Martian meteorite and the maskelynite contains no U, they are assumed to be the closest to“true initial Pb.” Therefore, the model presented here accounts for the 4 Ga difference in age necessary to explain the initial Pb compositions of ALH84001 and the enriched shergottites, a conclusion that is consistent with all other radiogenic isotope systems.

5. Conclusions

Initial Pb isotopic compositions preserved in maskelynite in ALH84001 and some enriched shergottites have been measured by SIMS and are vastly different. These differences are best explained by a ~4 Ga difference in their respective crystallization ages. These data are less radiogenic than corresponding solution measure-ments of identical comparable mineral separates, a result that highlights the unique ability of SIMS to mea-sure true initial Pb. A maskelynite-pyroxene Pb-Pb isochron for ALH84001 yields an age of 4089 ± 73 Ma (2σ). Using previously established reservoir chronology for the enriched shergottites and ALH84001, a forward model can be constructed to assess the Pb isotope composition of the Martian mantle after core formation and prior to large-scale silicate differentiation at 4.513 Ga, calculate initial Pb model ages, andμ values in the

Figure 8.207Pb/206Pb versus204Pb/206Pb and208Pb/206Pb versus204Pb/206Pb diagrams for the initial Pb determined for each of the meteorites studied here with second-stage paleo-geochrons from a source reservoir at 4.513 Ga that evolved from 4.567 Ga with aμ value of 1.2. Included are μ and κ values for each of the meteorites here.

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Martian mantle after 4.513 Ga. Assuming a CDT initial Martian mantle Pb isotopic composition and a modeled μ value of 1.2 between 4.567 and 4.513 Ga, the initial Pb model ages of the enriched shergottites are identical (~170 Ma) to those determined by Rb-Sr, U-Pb, and Lu-Hf. This model also indicates aμ value in the Martian mantle of 4.1–4.6 from 4.513 to 4.09 Ga and 0.17 Ga. The low-μ value postcore formation and before 4.513 Ga indicates that this event likely had little effect on theμ value of the Martian mantle and that silicate differen-tiation with sulfide precipitation is the largest contributor to the observed spread in Martian μ values.

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Acknowledgments

All data used here are available in Tables

1–3 and the peer-reviewed references

cited in the text andfigure captions. The

authors would like to acknowledge Marianne Ahlbom for access to the SEM at Stockholm University and Kerstin Lindén for preparing the samples. Ross Kielman is thanked for his help in preparing sample maps. The authors would like to thank Tom Lapen, James Connelly, two anonymous reviewers, and Editor David Baratoux whose comments and editorial handling

significantly improved the quality of

this manuscript. This work was funded by grants from the Knut and Alice Wallenberg Foundation (2012.0097) and the Swedish Research Council (VR 621-2012-4370) to M.J.W. and A.A.N. The Nordsim ion microprobe facility oper-ates as a Nordic infrastructure regulated by The Joint Committee of the Nordic Research Councils for Natural Sciences (NOS-N). This is Nordsim publication 428. P.A.B. acknowledges support from the Australian Research Council via their Australian Laureate Fellowship scheme. G.K.B. acknowledges support from Curtin University via their Research Fellowship scheme.

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