# The Northwest Africa 8159 martian meteorite: Expanding the martian sample suite to the early Amazonian

Christopher D. K. Herd, Erin L. Walton, Carl B. Agee, Nele Muttik, Karen Ziegler, Charles K. Shearer, Aaron S. Bell, Alison R. Santos, Paul V. Burger, Justin I. Simon, et al.

► To cite this version:

Christopher D. K. Herd, Erin L. Walton, Carl B. Agee, Nele Muttik, Karen Ziegler, et al. The Northwest Africa 8159 martian meteorite: Expanding the martian sample suite to the early Amazonian. Geochimica et Cosmochimica Acta, 2017, 218, pp.1-26. 10.1016/j.gca.2017.08.037. hal-01765617

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# The Northwest Africa 8159 Martian Meteorite: Expanding the Martian Sample Suite to the Early Amazonian

C. D. K. Herd, E. L. Walton, C. B. Agee, N. Mutik, K. Ziegler, C. K. Shearer, A. S. Bell, A. R. Santos, P. V. Burger, J. I. Simon, M. J. Tappa, F. M. McCubbin, J. Gattacceca, F. Lagroix, M. E. Sanborn, Q. Z. Yin, W. S. Cassata, L. E. Borg, R. E. Lindvall, T. S. Kruijer, G. A. Brennecka, T. Kleine, K. Nishiizumi, M. W. Caffee

September 29, 2017

Geochimica et Cosmochimica Acta
This document was prepared as an account of work sponsored by an agency of the United States government. Neither the United States government nor Lawrence Livermore National Security, LLC, nor any of their employees makes any warranty, expressed or implied, or assumes any legal liability or responsibility for the accuracy, completeness, or usefulness of any information, apparatus, product, or process disclosed, or represents that its use would not infringe privately owned rights. Reference herein to any specific commercial product, process, or service by trade name, trademark, manufacturer, or otherwise does not necessarily constitute or imply its endorsement, recommendation, or favoring by the United States government or Lawrence Livermore National Security, LLC. The views and opinions of authors expressed herein do not necessarily state or reflect those of the United States government or Lawrence Livermore National Security, LLC, and shall not be used for advertising or product endorsement purposes.
The Northwest Africa 8159 Martian Meteorite:
Expanding the Martian Sample Suite to the Early Amazonian

Christopher D. K. Herd¹, Erin L. Walton¹,², Carl B. Agee³, Nele Muttik³, Karen Ziegler³, Charles

K. Shearer³, Aaron S. Bell³, Alison R. Santos³, Paul V. Burger³, Justin I. Simon⁴, Michael J.

Tappa⁵, Francis M. McCubbin⁴, Jérôme Gattacceca⁶, France Lagroix⁷, Matthew E. Sanborn⁸,

Qing-zhu Yin⁸, William S. Cassata⁹, Lars E. Borg⁹, Rachel E. Lindvall⁹, Thomas S. Kruijer¹⁰,

Gregory A. Brennecka¹⁰, Thorsten Kleine¹⁰, Kunihiko Nishiizumi¹¹, Marc W. Caffee¹²,¹³

¹Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, T6G 2E3, Canada

²Department of Physical Sciences, MacEwan University, Edmonton, Alberta, T5J 4S2, Canada

³Institute of Meteoritics, Department of Earth and Planetary Sciences, University of New Mexico, Albuquerque, NM 87131, USA

⁴NASA Johnson Space Center, Mailcode XI2, 2101 NASA Parkway, Houston, Texas 77058

⁵Jacobs Technology, NASA Johnson Space Center, Houston, TX, 77058, USA

⁶CNRS, Aix Marseille Univ, IRD, Coll France, CEREGE, Aix-en-Provence, France

⁷Institut de Physique due Globe de Paris, Paris, France

⁸Department Earth and Planetary Sciences, University of California-Davis, One Shields Avenue, Davis, CA 95616

⁹Lawrence Livermore National Laboratory, 7000 East Avenue (L-231), Livermore, CA 94550

¹⁰Institute for Planetology, Westfälische Wilhelms-Universität Münster, Wilhelm-Klemm-Str. 10, 48149 Münster, Germany

¹¹Space Sciences Laboratory, University of California Berkeley, Berkeley, CA, 94720-7450, USA

¹²Department of Physics and Astronomy, Purdue University, West Lafayette, IN 47907-2036, USA

¹³Department of Earth, Atmospheric, and Planetary Sciences, Purdue University, West Lafayette, IN 47907-2036, USA

Revised in response to AE comments

Geochimica et Cosmochimica Acta

Version of: August 22, 2017
40 Corresponding author:

41

42 Christopher Herd

43 Professor

44 Department of Earth and Atmospheric Sciences

45 1-26 Earth Sciences Building

46 University of Alberta

47 Edmonton, AB T6G 2E3

48 Phone: (780)492-5798

49 Fax: (780)492-2030

50 [herd@ualberta.ca](mailto:herd@ualberta.ca)

51

52
**Abstract**

Northwest Africa (NWA) 8159 is an augite-rich shergottite, with a mineralogy dominated by Ca-, Fe-rich pyroxene, plagioclase, olivine, and magnetite. NWA 8159 crystallized from an evolved melt of basaltic composition under relatively rapid conditions of cooling, likely in a surface lava flow or shallow sill. Redox conditions experienced by the melt shifted from relatively oxidizing (with respect to known Martian lithologies, ~QFM) on the liquidus to higher oxygen fugacity (~QFM+2) during crystallization of the groundmass, and under subsolidus conditions. This shift resulted in the production of orthopyroxene and magnetite replacing olivine phenocryst rims. NWA 8159 contains both crystalline and shock-amorphized plagioclase (An₅₀-₆₂), often observed within a single grain; based on known calibrations we bracket the peak shock pressure experienced by NWA 8159 to between 15 and 23 GPa. The bulk composition of NWA 8159 is depleted in LREE, as observed for Tissint and other depleted shergottites; however, NWA 8159 is distinct from all other martian lithologies in its bulk composition and oxygen fugacity. We obtain a Sm-Nd formation age of 2.37 ± 0.25 Ga for NWA 8159, which represents an interval in Mars geologic time which, until recently, was not represented in the other martian meteorite types. The bulk rock ¹⁴⁷Sm/¹⁴⁴Nd value of 0.37 ± 0.02 is consistent with it being derived directly from its source and the high initial ε¹⁴³Nd value indicates this source was geochemically highly depleted. Cr, Nd, and W isotopic compositions further support a unique mantle source. While the rock shares similarities with the 2.4-Ga NWA 7635 meteorite, there are notable distinctions between the two meteorites that suggest differences in mantle source compositions and conditions of crystallization. Nevertheless, the two samples may be launch-paired. NWA 8159 expands the known basalt types, ages and mantle sources within the Mars sample suite to include a second igneous unit from the early Amazonian.
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# 1. INTRODUCTION

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The suite of meteorites known to be derived from Mars are predominantly igneous, and consist of the shergottite (>80% of all known martian meteorites), nakhlite (~10%) and chassignite (~3%) groups, the names of which derive from the 'original three' meteorites (Shergotty, Nakhla and Chassigny). The shergottites may be subdivided on the basis of texture and mineralogy, giving rise to aphyric (also known as "basaltic" and containing pyroxene-plagioclase in subequal proportions), olivine-phyric (like aphyric but containing olivine phenocrysts), and poikilitic (also known as lherzolitic, containing oikocrysts of low-Ca pyroxene enclosing olivine chadocrysts), or variations thereof. Pyroxene in the basaltic shergottites varies in composition, although most commonly low- and high-Ca pyroxene are present in subequal proportions (Papike et al., 2009). The nakhlites are olivine clinopyroxenites and the chassignites are olivine-rich rocks (mostly dunites) with minor pyroxene and chromite (e.g., see reviews by Bridges and Warren, 2006; Papike et al., 2009; Treiman, 2005). Additional types are the orthopyroxenite Allan Hills (ALH) 84001 (Mittlefehldt, 1994), and the unique polymict regolith breccia (originally classified as a "basaltic breccia") Northwest Africa (NWA) 7034 (Agee et al., 2013) and its pairs (Humayun et al., 2013; Wittmann et al., 2015). One other martian meteorite, NWA 7635, is characterized by phenocrysts of plagioclase, augite, ferroan olivine and Ti-poor magnetite in a pyroxene-plagioclase groundmass (Lapen et al., 2017); its texture and mineralogy set it apart from the other known martian meteorites, although it shares some similarities to NWA 8159, described here.

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The martian meteorite suite is dominated by young (< 600 Ma) basalts (shergottites of various types) and the members of the nakhlite and chassignite groups, which date to ~1300 Ma (e.g., Nyquist et al., 2001; Papike et al., 2009). ALH 84001 represents the only known sample
100 from the early history of Mars, dating to 4100 Ma (Bellucci et al., 2015; Lapen et al., 2010).
101 NWA 7034 is a regolith breccia containing zircons dated to 4428 Ma with evidence for U-Pb
102 disturbance at ~1500-1700 Ma. The nature of the younger disturbance is the subject of continued
103 discussion, and may relate to reheating involving interaction with near-surface fluids (Cartwright
104 et al., 2014; McCubbin et al., 2016; Nemchin et al., 2014), possibly associated with the annealing
105 of the breccia. It is generally agreed that the age bias in the martian meteorite suite is the result of
106 the inherent bias in the impact lofting mechanism in their delivery to Earth, which favors more
107 competent igneous rocks (such rocks tend to be younger; e.g., Walton et al., 2008). The martian
108 meteorite suite notably lacks samples that formed in the early Amazonian (~3000 to 1500 Ma); a
109 recently-discovered exception is the 2400-Ma NWA 7635 shergottite (Lapen et al., 2017).
110 Therefore, the discovery of any new martian meteorite with a crystallization age within this
111 timeframe is of interest for expanding our understanding of the igneous history of Mars, as well
112 as providing insights into the potential compositional variability of the martian mantle.

113 Northwest Africa (NWA) 8159 is a single, 149.83 g stone that was purchased by Brahim
114 Tahiri from a Moroccan meteorite hunter and sent to his partner Sean Tutorow for classification
115 in 2013. The distinct mineralogy of the meteorite – in particular the dominance of augite over
116 other pyroxenes – was noted during classification, leading to its designation as a martian “augite
117 basalt” (Ruzicka et al., 2015). Here we describe the petrography, mineral chemistry, bulk
118 composition, and stable and radiogenic isotopic composition of this sample, for the purpose of
119 understanding its relationship to other samples in the Martian meteorite suite. We find that NWA
120 8159 has some characteristics in common with the shergottites, and is best termed an augite-rich
121 shergottite. However, NWA 8159 is distinct among the Martian samples available to date,
122 sharing some similarities (including possible launch pairing) only with the recently discovered
123 NWA 7635.

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## 2. METHODS

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### 2.1 Petrographic and mineral composition analysis

127 Several polished samples of NWA 8159 including thin sections and polished probe
128 mounts were investigated using transmitted and reflected light microscopy and backscattered
129 electron (BSE) imaging at the University of Alberta (UAlberta) and the University of New
130 Mexico (UNM). Two polished sections were initially investigated by transmitted (thin section
131 only) and reflected light microscopy to study shock deformation and transformation features in
132 the host rock and to locate shock veins. Detailed micro textures of plagioclase in these sections
133 were mapped as a function of distance from shock veins using a Zeiss EVO MA LaB₆ filament
134 SEM in BSE mode at UAlberta. BSE images were acquired using a Si diode detector under
135 conditions of 20 kV (accelerating voltage) and 8-10 mm working distance, and subsequently
136 imported into a commercial image analysis program to make accurate measurements of features
137 such as grain size and shock vein thickness. These measurements were made in the two
138 dimensions provided by the thin section and therefore represent apparent dimensions. Further
139 observations of the mineralogy and texture of NWA 8159 were carried out using the same
140 apparatus.

141 Major and minor element abundances were measured using wavelength dispersive
142 spectrometry (WDS) using a JEOL 8900 electron microprobe (EMP) at UAlberta or a JEOL
143 8200 EMP at UNM. In general, a focused (≤1 µm) spot size was used for all analyses except
144 feldspar and shock-produced glass (10 µm beam) to minimize alkali element mobilization and
damage to the sample. Analytical conditions at UNM included a 15 kV accelerating voltage and 145
20 nA beam current. Additional analytical details are provided in Electronic Annex 1.
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2.2 Raman spectroscopy
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Raman spectroscopy was employed to further investigate the mineralogy of NWA 8159
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using a Bruker SENTERRA instrument at MacEwan University. An Olympus BX51
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petrographic microscope was used to focus the excitation laser beam (532 nm line of Ar+ laser)
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and to collect the Raman signal in the backscattered direction. The 100X objective was used to
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achieve a 1 µm spot size. A sequence of two-10 s exposures was acquired using a laser power of
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10 mW. Due to the fine grained nature of some minerals the analyses were repeated to check
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reliability. These multiple exposures were then summed to achieve the final spectrum.
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Backgrounds of the spectra were graphically reduced using the OPUS v. 6.5 spectroscopy
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software.
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157 2.3 Trace element analysis of bulk sample

A 0.4 g interior fragment was used for bulk trace element analysis at UAlberta. The
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fragment was broken into smaller pieces, ground, dissolved, and analyzed using a Perkin Elmer
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Elan6000 quadrupole ICP-MS instrument at UAlberta. Additional digestion of a 10 mg bulk
160
sample was carried out and analyzed at Lawrence Livermore National Laboratory (LLNL), using
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a Thermo XSII quadrupole ICP-MS instrument. Details of sample preparation and analysis are
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provided in Electronic Annex 1.
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164 2.4 Magnetic properties

Hysteresis properties and their variation with temperature between 20 and 650 °C were
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measured with a Princeton Micromag Vibrating Sample Magnetometer on a 115 mg sample at
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CEREGE and on a 34 mg sample at IPGP, respectively. The S<sub>300</sub> ratio (ratio of saturation
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168 remanence superimposed with a back-field isothermal remanence magnetization at 300 mT over
169 saturation remanence) was determined for two ~100 mg samples. Magnetic susceptibility was
170 measured on a 2.75 g sample with a MFK1 susceptometer from Agico. The evolution of
171 magnetic susceptibility with temperature between -195°C and 650°C was measured on a 150 mg
172 sample with an Agico MFK1 apparatus equipped with a CS3 furnace and a CSL cryostat at
173 CEREGE.

## 2.5 Samarium-neodymium isotopic analysis

175 To limit the effects of desert weathering on radiogenic isotopic measurements, a ~1.35 g
176 cube was extracted from the center of NWA 8159. This material from the meteorite interior was
177 then gently crushed and processed according to existing JSC procedures (e.g., Nyquist et al.,
178 2016 and references therein). The fine-grained texture, zoned pyroxene, and shock melt veins in
179 this meteorite made clean mineral separation difficult. Nevertheless, bulk “whole” rock (WR),
180 concentrated mineral separations, and nearly pure plagioclase and pyroxene splits were
181 recovered, all of which have varying amounts of impurities. Each fraction was placed in an
182 ultrasonication bath and leached using HCl for 10 minutes. Various acid concentrations (1N to
183 6N HCl) were explored in an effort to reduce the effects of contamination due to weathering. The
184 leached component was removed from the residue via pipette, and the residue washed using
185 ultra-pure H₂O to ensure all dissolved components were removed. A combination of
186 concentrated acids (HF, HNO₃, HCl, HClO₄) was used for dissolution and visual inspection
187 ensured complete digestion of all phases. All 11 samples including WR, concentrated mineral
188 separates, and acid leachates were thoroughly homogenized with the JSC mixed ¹⁴⁹Sm-¹⁵⁰Nd
189 spike during the decomposition, which included multiple steps of sample dry-down and acid
190 fluxing. Chemical separation was achieved following the method detailed by Shih et al. (1999),
which uses cation exchange resin AG 50WX8 200-400 mesh in NH₄⁺ form and eluted with 0.2M α-hydroxyisobutyric acid for the separation of Sm and Nd. Following separation, fractions were passed through AG 50WX12 200-400 resin to remove the α-hydroxyisobutyric acid. The purified Nd and Sm fractions were analyzed on the Thermo-Finnigan Triton thermal ionization mass spectrometer (TIMS) housed in the Center for Isotope Cosmochemistry and Geochronology at NASA-JSC. Additional analytical details are provided in Electronic Annex 1. Additional fractions of NWA 8159, including the purest plagioclase and pyroxene separates, remain to be measured.

## 2.6 Tungsten and neodymium isotopic analysis of bulk sample

The W and Nd isotope compositions of NWA 8159 were determined following the methods documented in Kruijer et al. (2015) and Borg et al. (2016). Two samples of NWA 8159 (Wr-1 and Wr-2) weighing ~0.5 g and ~1 g were powdered in an agate mortar and then digested in concentrated HF-HNO₃ (2:1). Small (~4%) solution aliquots were taken for determination of Sm and Nd concentrations by isotope dilution. A mixed <sup>149</sup>Sm-<sup>150</sup>Nd tracer was added to the samples and they were dried. Tungsten from the un-spiked samples was separated from the sample matrix using two-stage anion exchange chromatography (Kruijer et al., 2015), and the W isotope composition measured on a ThermoScientific® Neptune Plus multicollector ICP-MS at the University of Münster (Kruijer et al., 2015). Neodymium isotopic compositions were determined from the washes of the W separations at LLNL using a ThermoScientific Triton TIMS. Additional analytical details are provided in Electronic Annex 1.

## 2.7 Noble Gas and Cosmogenic Radionuclide Mass Spectrometry

Noble gas extractions of whole-rock and fragments visibly enriched in shock-produced glass of NWA 8159 were conducted in the Livermore Noble Gas Lab at LLNL following
procedures described in Cassata and Borg (2016), after sample preparation methods outlined in Electronic Annex 1. Ages were calculated using the decay constants and standard calibration of Renne et al. (2011). He, Ne, and Ar exposure ages obtained from un-irradiated aliquots were calculated from the elemental production rates given in Eugster and Michel (1995) using the bulk chemical compositions obtained by other methods (Table 3). For <sup>38</sup>Ar exposure age calculations, the relative contributions of trapped and cosmogenic Ar were determined by a two-component deconvolution assuming a trapped component with <sup>38</sup>Ar/<sup>36</sup>Ar = 0.189. All <sup>21</sup>Ne was assumed to be cosmogenic. Cosmogenic radionuclide concentrations of <sup>10</sup>Be (t<sub>1/2</sub> = 1.36 Myr), <sup>26</sup>Al (0.705 Myr), and <sup>36</sup>Cl (0.30 Myr) were measured by accelerator mass spectrometry at Purdue University (Sharma et al., 2000), after sample preparation methods outlined in Electronic Annex 1. The reported cosmic ray exposure age is based on the measured chemical composition (wt %) of the sample (2.31% Mg, 5.42% Al, 7.69% Ca, 3130 ppm Mn, and 16.1% Fe) and modeled production rates (Masarik and Reedy, 1994).

## 2.8 Oxygen isotopic analysis of bulk sample

Oxygen triple isotope analyses of several subsamples of the meteorite were performed by laser fluorination at UNM (Sharp, 1990), using a Thermo Finnigan MAT 253 dual inlet isotope ratio mass spectrometer, after sample preparation methods outlined in Electronic Annex 1. Oxygen isotopic ratios were calculated using the following procedure: The δ<sup>18</sup>O values refer to the per-mil deviation in a sample (‰O/‰O) from SMOW, expressed as δ<sup>18</sup>O = [‰O/‰O]<sub>sample</sub>/(‰O/‰O)<sub>SMOW-1</sub> x 10<sup>3</sup>. The delta-values were converted to linearized values by calculating: δ<sup>18</sup>/<sup>17</sup>O' = ln[(δ<sup>18</sup>/<sup>17</sup>O + 10<sup>3</sup>)/10<sup>3</sup>] x 10<sup>3</sup> in order to create straight-line mass-fractionation curves. The Δ<sup>17</sup>O' values were obtained from the linear δ-values by the following relationship: Δ<sup>17</sup>O' = δ<sup>17</sup>O' - 0.528 * δ<sup>18</sup>O'; Δ<sup>17</sup>O' values of zero define the terrestrial mass-
fractionation line. $\Delta^{17}$O' values lying on any line with a slope of ~0.52-0.53 are due to mass-dependent processes. Typical analytical precision of the laser-fluorination technique is better than ± 0.01‰ for $\Delta^{17}$O'. Additional analytical details are provided in Electronic Annex 1.

## 2.9 Chromium isotopic analysis of bulk sample

A 42.36 mg interior chip of NWA 8159 was powdered using an agate mortar and pestle, and dissolved using Parr bomb methods. Chromium isotope ratios were measured using a Thermo *Triton Plus* TIMS at the University of California, Davis. Additional analytical details are provided in Electronic Annex 1.

## 3. RESULTS

The specimen has a weathered, yellow-brown patina exterior, light colored desert soil coating on one side, and lacks fusion crust (Figure 1). A saw cut reveals a green-gray interior with aphanitic texture crosscut by black to gray colored, glassy veins (up to 1 mm thick) of shock-melted material (Fig. 1b). The outer rind of the stone shows a light-colored weathering band ~3 mm thick. A few light-colored weathering veins penetrate ~0.5 cm into the stone.

### 3.1 Petrography

Based on petrographic observations, electron microprobe mass balance calculations and magnetic properties (discussed below), NWA 8159 has an intergranular texture, consisting of augite (~48 modal %), both birefringent and amorphous plagioclase (~37%), olivine (~5%), magnetite (~5%), maghemite (~0.5%) and orthopyroxene (~1%). Augite, plagioclase, olivine, and magnetite are approximately equigranular, although the distribution of grain size appears to be bimodal; in some areas, the average grain size is ~200 µm, whereas in others it is <100 µm based on visual estimation (e.g., Fig. 2a). Minor amounts of ilmenite, merrillite, chlorapatite, Cr-
spinel, and secondary alteration minerals (see Discussion) are located throughout the meteorite.
Sulfides were sought but not found. There are also domains of dark, largely opaque veins of
shock-melted and quenched material, and associated mineral transformations such as plagioclase
amorphization and formation of tissintite. Representative mineral compositions are reported in
Tables 1 and 2; the full dataset is reported in Electronic Annex 2.

## 3.2 Magnetic properties

NWA 8159 is the second most magnetic among analyzed martian meteorites, with a
saturation remanence of 0.83 Am²/kg, about one third of that of NWA 7034 (Gattacceca et al.,
2014) and twice the most magnetic nakhlite or shergottite (Rochette et al., 2005). This value is
well above the limit value of 0.29 Am²/kg that is required for Martian lithologies to account for
the intense magnetic anomalies on Mars (Gattacceca et al., 2014). Similarly magnetic
susceptibility is 1.69 x 10⁻⁵ m³/kg, three times higher than the most magnetic nakhlites. Pure
magnetite (based on Curie point at 580°C and Verwey transition) dominates, but a detectable
amount of maghemite (like in NWA7034; e.g., Muttik et al., 2014) is also identified by the
irreversible drop in susceptibility at 300 and 350 °C (Figure 4). The associated drop in saturation
magnetization indicates a maghemite content of 0.24 wt%. Slight oxidation of the magnetite is
indicated by the Verwey transition temperature, which is lower than the canonical temperature of
120 K (King and Williams, 2000). Pyrrhotite is absent as indicated by the S₃₀₀ ratio of -0.99.
Saturation magnetization of 3.28 Am²/kg indicates a magnetite plus maghemite content of about
3.5 wt%.

280

## 3.3 Geochemistry
The compositions of quenched melt produced by shock in veins or pockets are reflective of bulk rock composition, in cases where the scale of melting is much larger than the grain size of igneous minerals (Walton et al., 2012). We document several mm-scale veins of shock-melted material in NWA 8159. Defocused beam electron microprobe analysis of one vein yields a bulk major and minor element composition that is similar to those obtained from bulk ICP-MS analysis (Table 3). Bulk trace elements are also provided in Table 3, and CI-normalized REE plotted in Figure 8, along with data from selected martian lithologies for comparison.

The Sm-Nd isotopic data of the leached residual materials define a six point internal isochron that yields a formation age for NWA 8159 of 2.37 ± 0.25 billion years (95% conf., MSWD=59), an initial $\epsilon^{143}$Nd value of +46 ± 11, and a calculated source $^{147}$Sm/$^{144}$Nd ratio of 0.36 ± 0.04 (or 0.40 ± 0.04 depending on the choice of modeled source age, i.e., 4.5 Ga, e.g., Nyquist et al. 2001, or 4.1 Ga, e.g., Lapen et al. 2010, respectively), shown in Figure 9. Lower Sm/Nd values come from fragments selected because they were lighter in color (i.e., dominated by maskelynite/plagioclase) and the cluster of fragments with higher Sm/Nd values were selected because of their darker color (likely containing more pyroxene). The WR samples were those that looked to generally represent the meteorite mineralogy (not dominated by any specific phase). Leaching experiments were undertaken in order to effectively remove minerals from the rock fragments that are susceptible to secondary Sm and Nd mobilization (e.g., phosphate). Collectively, these leachates and the WR samples (see Figures S1, S2, in Electronic Annex 1) help define a secondary Nd reservoir, which is generally less radiogenic than the meteorite phases, but still highly radiogenic compared to any likely terrestrial Nd sources and thus likely originates from the martian crust. Further radiogenic isotopic work is required to test this hypothesis. Nd isotopic, Sm, and Nd abundance data collected at NASA-JSC are provided in
Table 4. Uncertainties are 2σ and reflect both measurement standard error and reproducibility of the Nd and Sm isotope standards.

The measured ε<sup>142</sup>Nd, ε<sup>143</sup>Nd, and <sup>147</sup>Sm/<sup>144</sup>Nd values for the 0.5 g bulk fragment of NWA 8159 measured at LLNL are +0.88 ± 0.04, +99.62 ± 0.03, and 0.3689 ± 0.0004, respectively (Table 4). Of note, a separate digestion of 0.999 g of NWA 8159 yields somewhat lower but more precise ε<sup>142</sup>Nd, ε<sup>143</sup>Nd, and <sup>147</sup>Sm/<sup>144</sup>Nd values of +0.77 ± 0.02, +78.34 ± 0.01, and 0.3178 ± 0.0003, respectively. For this fragment the W isotopic composition was also measured resulting in an ε<sup>182</sup>W value of +1.13 ± 0.10 (2s.d.) (Table 5). The differences in the values obtained at the two laboratories are likely due to differences in modal proportions of minerals in each sample and whether leaching was used in the procedure.

Results of <sup>40</sup>Ar/<sup>39</sup>Ar investigations are shown in Figure 10, and the complete isotopic dataset is given in Electronic Annex 2. Data obtained from the incremental heating of ~1.4 mg of glass extracted from a mm-sized shock vein of NWA 8159 define an isochron with an age of 2.15 ± 0.10 Ga (2σ; MSWD = 6.6), statistically indistinguishable from the Sm-Nd age of 2.37 ± 0.25 Ga (Fig. 10a). The <sup>40</sup>Ar/<sup>36</sup>Ar ratio of the trapped component (296 ± 23) is consistent with terrestrial contamination and was likely acquired during desert weathering. An age spectrum obtained from the incremental heating of a whole-rock fragment of NWA 8159 is indicative of significant loss of radiogenic Ar (Fig. 10b). Maximum ages do not exceed ~1.2 Ga, which post-dates crystallization by ~1 Ga. Given this evidence for a thermal disturbance to the K-Ar system, it is possible that the <sup>40</sup>Ar/<sup>39</sup>Ar isochron age obtained from the shock-produced glass is systematically biased toward a younger age than the Sm-Nd system due to partial resetting, despite the statistically equivalent ages.
Cosmic ray exposure ages were determined from ³He, ²¹Ne, and ³⁸Ar measurements of whole-rock and shock melt glass fragments (Table 6). Shock melt glass exposure ages, calculated using the average chemical composition of EPMA spot analyses of melt veins (Table 3) are concordant at 1.4-1.5 Ma. Whole rock ³He and ²¹Ne exposure ages, calculated using the whole-rock chemical composition determined by ICP-MS (Table 3), are slightly younger at 0.9-1.2 Ma. Deviations in the chemistry of the mg-sized aliquots that were analyzed from the compositions listed in Table 3 may explain some of the spread in apparent exposure ages. The average exposure age from ³He, ²¹Ne, and ³⁸Ar is 1.29 ± 0.25 Ma (1σ). The measured activities of ¹⁰Be, ²⁶Al, and ³⁶Cl are 8.75±0.16 dpm/kg, 57.2±1.5 dpm/kg, and 9.23±0.19 dpm/kg, respectively, yielding a cosmic ray exposure age of 1.0 ± 0.1 Ma. Although the recovered mass is small, the pre-atmospheric radius, based on our cosmogenic nuclide measurements, is estimated to be 30-60 cm. Using the ³⁶Cl concentration we obtain a terrestrial age of 0.23 ± 0.06 Ma. The ejection age from Mars for NWA 8159 is thus 1.2 ± 0.1 Ma, the sum of the two, which is in agreement with LLNL results from ³He, ²¹Ne, and ³⁸Ar, and similar to the ²⁶Al and ¹⁰Be exposure age of 1.0 ± 0.1 Ma determined on NWA 7635 (Lapen et al., 2017). NWA 8159 may therefore belong to a group of ~12 depleted shergottites having ~1.1 Ma ejection ages (e.g., Lapen et al., 2017; Wieler et al., 2016).

Oxygen isotopic analysis of bulk meteorite samples yield $\Delta^{17}O' = 0.247, 0.321, 0.336, 0.316, 0.280$ (all linearized, ‰), with $\delta^{18}O'$ ranging from 3.328 to 4.197 ‰ (Table 7). These values fall within the range of typical martian meteorites such as those belonging to the traditional SNC group (Figure 11) and ALH 84001, indicating that the oxygen within NWA 8159 is from the same martian oxygen reservoir as these other martian samples (e.g., Agee et al., 2013 and references therein; Clayton and Mayeda, 1983; Franchi et al., 1999). Evidence for an
alternate martian reservoir represented by portions of the martian breccia NWA 7034 (Agee et al., 2013) were sought but not found.

The chromium isotopic composition (ε<sup>53</sup>Cr and ε<sup>54</sup>Cr) of NWA 8159 is provided in Table 5. As shown in Figure 12, NWA 8159 plots within the same general region in Δ<sup>17</sup>O-ε<sup>54</sup>Cr isotope space as other martian (SNC) meteorites. In terms of ε<sup>54</sup>Cr, NWA 8159 is not resolvable from either Shergotty or Nakhla; it is however, resolvable from Chassigny. The resolved nature of NWA 8159 from Chassigny can also be seen in a plot of ε<sup>53</sup>Cr versus ε<sup>54</sup>Cr (Fig. 12c).

## 4. DISCUSSION

### 4.1 Texture, mineral chemistry, and conditions of crystallization and subsolidus equilibration

The aphanitic texture of NWA 8159 is similar to that of basaltic rocks that have crystallized rapidly, for example, as an eruptive flow or shallow intrusion. Dark glassy veins crosscut the igneous texture. These features are attributed to a post-crystallization impact event, which also deformed host rock minerals and are associated with mineral transformations (see section 4.2 for details). Secondary alteration is observed throughout the meteorite, and includes preferential replacement of glassy shock veins, (mostly Ca-)carbonate and barite filling of fractures, and replacement of magnetite and olivine with Fe-oxides/hydroxides. Fe-oxide replacement of olivine and magnetite tends to be spatially associated with fractures or veins, consistent with these fractures acting as preferential conduits for terrestrial fluids. Olivine phenocrysts have been affected to variable degrees by low-temperature alteration, such that some are enriched in P and Al and consist of a fine-grained mixture of phases (Shearer et al., 2015). The alteration phases are consistent with formation under low-temperature conditions in a desert
environment; considering also that the main mass lacks a fusion crust and has a visible weathering rind (Fig. 1) we conclude that the bulk of the alteration is likely terrestrial, and that this meteorite had a significant terrestrial residence time. Nevertheless, some near-surface modification on Mars is suggested by the Nd isotopic compositions of the leachates (see Section 4.3.2).

Electron microprobe analyses of NWA 8159 augite grains show that they are chemically zoned, becoming increasingly ferroan and calcium-poor from core to rim (Figure 3) over a range of $Fs_{25}En_{37}Wo_{38}$ to $Fs_{66}En_{7}Wo_{27}$. Olivine, ranging in composition from $Fo_{41}$ to $Fo_{20}$ (average of $Fo_{33}$; Fig. 3), occurs as anhedral grains mantled by low-Ca pyroxene and magnetite (Fig. 2 c-e). Individual olivine grains display normal igneous zoning that extends across most of this compositional range. Low-Ca pyroxene is only found in this mineral assemblage, and ranges in composition from $Fs_{45}En_{53}Wo_2$ to $Fs_{73}En_{26}Wo_1$; as with olivine, individual low-Ca pyroxene grains display zoning over most of this composition, with the most ferroan compositions at the margins with the groundmass (Fig. 2e). $Wo$ contents, which are rarely > 2.5 mol% even for the most ferroan compositions, are consistent with orthopyroxene. Fe/Mn ratios in both augite ($Fe/Mn = 35 \pm 3$; n = 201) and olivine ($Fe/Mn = 47 \pm 7$; n = 43) are consistent with a martian origin (e.g., Papike et al., 2009; Treiman et al., 2000). In contrast, Fe/Mn ratios of orthopyroxene ($Fe/Mn = 22 \pm 4$; n = 15) reflect a post-crystallization reaction involving olivine, discussed below. Plagioclase occurs as both crystalline and amorphous phases, but we avoid the use of the term maskelynite to describe the amorphous plagioclase as the maskelynite typical to shergottites has different properties from those of the amorphous phase observed here (i.e., maskelynite in shergottites is typically non-stoichiometric, has degraded igneous zoning, and K enrichment near shock veins; Chen and El Goresy, 2000). Plagioclase exhibits the following range in
composition: An₅₀-6₂Ab₃₈-4₉Or₀.₁-₀.₅, and has an average composition of An₅₇Ab₄₂Or₀.₃. Both the crystalline and amorphous plagioclase share this compositional range. In contrast to other martian meteorites, the dominant oxide mineral in NWA 8159 is magnetite of near-endmember composition (Mt₉₆.₀-9₉.₅Sp₀.₁-₀.₆Usp₀.₂-₃.₅). Magnetite occurs with low-Ca pyroxene in olivine rims (Fig. 2d-e), as < 20 µm grains in the groundmass, and also as ~ 150 µm phenocrysts (Fig. 2b). Some of the latter are characterized by a core consisting of a symplectic intergrowth of magnetite and at least one additional phase (Fig. 2b inset); quartz was identified by a sharp peak in the Raman spectrum near 464 cm⁻¹. Many of the non-magnetite grains in the symplectite are replaced by fine-grained terrestrial alteration, or are void spaces. All magnetite phenocrysts analyzed have TiO₂ below detection (0.1 wt%); the maximum Ti contents (up to ~0.6 wt%) were observed in grains within the groundmass. Ilmenite is less abundant than the magnetite, typically occurring as lath-shaped grains adjacent to, or in contact with magnetite and associated with mesostasis and silica grains. Ilmenite ranges in composition from Ilm₆₅Hem₃₀Gk₂.₃Py₂.₂ to Ilm₅₀Hem₄₆Gk₂Py₂. The most oxidized ilmenite, which contains up to 46% of the hematite endmember, is the highest known among martian meteorites and is indicative of highly oxidizing conditions of crystallization.

Pyroxene compositions in NWA 8159 are distinct from those known from all other Martian meteorites, with the possible exception of NWA 7635 (Figure 3). Nakhlite pyroxenes show some similarities, e.g., in the range of molar Fs; however, augite in the nakhlites is generally more calcic. Furthermore, whereas nakhlite augites are relatively coarse-grained (up to 1 cm long) cumulates, the augites in NWA 8159 occur as fine-grained (<200 µm) crystals intergrown with plagioclase and magnetite as part of a presumably rapidly cooled basaltic magma. NWA 8159 pyroxene compositions do overlap with high-Ca pyroxene in shergottites
(Figure 3); however, NWA 8159 lacks the low-Ca pyroxene (pigeonite) having similar Fe/Mg ratio which typifies most shergottites (Papike et al., 2009). Despite these differences, NWA 8159 is most similar to shergottites in its pyroxene compositions and presence of other typical shergottite minerals (e.g., plagioclase, oxides, and phosphates). For these reasons, NWA 8159 is best classified as an augite-rich shergottite. A full range of pyroxene compositions in NWA 7635 has not been reported (Lapen et al., 2017); nevertheless, representative magnesian and ferroan compositions overlap with their counterparts in NWA 8159 (Figure 3). Therefore NWA 7635 could also be classified as an augite-rich shergottite based on the composition of its pyroxene and its reported mineralogy.

The most magnesian pyroxene in even the most evolved shergottites (e.g., Los Angeles or QUE 94201, ~En₅₀) is more magnesian than augites in NWA 8159 (~En₄₀; Figure 3) – along with bulk Mg# (Table 3), this observation indicates that NWA 8159 crystallized from a relatively evolved melt. Notably, augite rim compositions in NWA 8159 cross into the 1-bar “forbidden region” of pyroxene stability (Figure 3; see also Lindsley, 1983), although augite core compositions fall outside of this field. While it is not uncommon to observe these metastable compositions – or their symplectitic breakdown assemblages – in shergottites (e.g., Aramovich et al., 2002), the metastable pyroxene compositions in NWA 8159 are much more calcic. The presence of metastable pyroxene compositions is consistent with a relatively rapid cooling rate, as indicated by the relatively fine-grained texture of the rock. Calculation of Fe³⁺ contents of pyroxenes based on stoichiometry results in Fe³⁺/(Fe³⁺+Fe²⁺) < 2% for orthopyroxene, while that for clinopyroxene (Wo₄₀En₃₅Fs₂₅ to Wo₂₂En₁₈Fs₆₀) is only slightly higher and more variable (> 3%). These relatively low Fe³⁺ contents are consistent with conditions approximately at, or 2 log units below the Quartz-Fayalite-Magnetite (QFM) buffer (ΔQFM to ΔQFM-2) based on
experimental augite-glass pairs using compositions relevant to martian meteorites (McCanta et al., 2004), assuming the Fe³⁺ contents are the result of mineral-melt partitioning.

The presence of near endmember magnetite and hematite-rich ilmenite in NWA 8159 is distinct from the shergottites and nakhlites, in which the magnetite component in titanomagnetite is rarely > 40% and in which hematite rarely exceeds 5% in ilmenite (e.g., Herd et al., 2001; Treiman, 2005; Walton et al., 2012). However, it is similar to oxides within some of the igneous clasts within martian regolith breccia NWA 7034, which have near endmember magnetite and ilmenite with up to 12% hematite component (Santos et al., 2015). Endmember magnetite is only stable in Fe-rich martian magma compositions either at conditions much more oxidizing than QFM, or under subsolidus conditions (Papike et al., 2015; Righter et al., 2013). Examination of the magnetite and ilmenite compositions in NWA 8159 shows that they are not in Mn-Mg exchange equilibrium, falling well outside the Bacon-Hirschmann line (Bacon and Hirschmann, 1988; Ghiorso and Evans, 2008). Applications of the Ghiorso-Evans (Ghiorso and Evans, 2008) and Ca-QUIF (Andersen et al., 1993) models to co-existing oxides corroborate this inference: the Ca-QUIF model results in unrealistic fO₂ conditions (≥ 8 log units above QFM) and spuriously low temperatures (250-420 °C), whereas the Ghiorso-Evans model yields ΔQFM+4 to ΔQFM+6 and similar temperatures (250-365 °C). Notably, the Ghiorso-Evans model calculates the liquid activity of TiO₂ = 2-4.4, a result that indicates that the composition of the magnetite is too Ti-poor to be in equilibrium with the ilmenite used in the calculation. Otherwise put, the ilmenite observed in NWA 8159 would be in equilibrium with a (titano)magnetite containing appreciably more TiO₂; our calculations suggest that such a titanomagnetite would contain TiO₂ ≥ ~2.5 wt% and would be in equilibrium at ~ ΔQFM+2. In contrast, the most Ti-rich magnetite grains we observe contain < 0.6 wt% TiO₂. It is likely that ilmenite formed during the very last
stage of crystallization, when the Ti was sufficiently enriched in the melt to stabilize ilmenite.
465

The observed occurrence of ilmenite most commonly as part of the mesostasis assemblage is
466

consistent with this inference.
467

Application of the Ca-QUIIF model to typical magnetite (either phenocrystic or
468
groundmass), the most Mg-rich olivine (Fo₄₀), and augite (Fs₄₀En₃₀Wo₃₀) and orthopyroxene
469
(Fs₅₉.₅En₃₉.₅Wo₁) compositions in equilibrium with the olivine yields ΔQFM+2 with high
470
uncertainties (±2 log units, reported by the Ca-QUIIF program) and temperatures (430-510 °C)
471
indicative of some part of the assemblage being out of equilibrium. Inclusion of ilmenite
472
compositions in the calculation results in convergence errors. Nevertheless the assemblage – as
473
well as the presence of maghemite (assuming it is magmatic) – is indicative of fO₂ above the
474
QFM buffer for the latter stages of crystallization, conditions that are higher than most
475
shergottites (ΔQFM-4 to -1; Herd, 2003; Herd et al., 2002), and similar to the most oxidized
476
lithologies within the shergottites (~ΔQFM+0.5; Herd, 2006; Peslier et al., 2010), nakhlites (up
477
to ΔQFM+2; McCubbin et al., 2013; Treiman and Irving, 2008), and igneous clasts in NWA
478
7034 (ΔQFM+2 to +4; Santos et al., 2015).
479

Olivine core compositions have similar Fe/Mg ratios to augite cores, indicating that
480
olivine and augite are approximately in Fe-Mg equilibrium (Figure 3). This observation suggests
481
that the olivine phenocrysts are cognate with the melt represented by the groundmass
482
(alternatively, the olivine phenocrysts represent xenocrysts included in the melt that had
483
fortuitously similar Fe/Mg ratios). All olivine phenocrysts are replaced at their margins by an
484
assemblage of orthopyroxene and near-end-member magnetite. Magnetite-orthopyroxene
485
intergrowths such as those observed in NWA 8159 have been attributed to both magmatic (e.g.,
486
Kuno, 1950; Presnall, 1966) and very late-stage magmatic to subsolidus (e.g., Morse, 1969;
487
488 Yoder and Tilley, 1962) processes that are highly dependent on oxygen fugacity ($fO_2$). The
489 relevant equilibria are:

$490 \quad 6 \mathrm{Fe}_2\mathrm{SiO}_4(\mathrm{ol}) + \mathrm{melt} = 2 \mathrm{Fe}_3\mathrm{O}_4(\mathrm{mt}) + 6 \mathrm{FeSiO}_3(\mathrm{opx}) \quad [1]$

491 or

$492 \quad 6 \mathrm{Fe}_2\mathrm{SiO}_4(\mathrm{ol}) + \mathrm{O}_2 = 2 \mathrm{Fe}_3\mathrm{O}_4(\mathrm{mt}) + 6 \mathrm{FeSiO}_3(\mathrm{opx}) \quad [2]$

493 Equation [1] describes a peritectic reaction in which olivine is destabilized in the presence of the
494 melt (i.e., supersolidus), under high $fO_2$ conditions (Presnall, 1966). Equation [2] is an oxidation
495 reaction, which may effectively occur at any stage during the evolution of the system; however,
496 it typically pertains to lower-temperature (e.g., subsolidus) conditions in which olivine is out of
497 equilibrium with the prevailing $fO_2$, as a result of a relative $fO_2$ increase.

498 Several observations suggest that equation [2] is most relevant to the olivine-opx-
499 magnetite assemblages in NWA 8159. Texturally, olivine phenocrysts have broadly euhedral
500 outlines, with margins embayed by orthopyroxene and magnetite products (Figure 2). This
501 texture is more consistent with subsolidus re-equilibration than the peritectic reaction; if a
502 peritectic reaction were the cause of the olivine embayment, then one would expect the olivine
503 phenocrysts to have irregular outlines consistent with resorption in the melt, and for
504 orthopyroxene (and magnetite) to be present as discrete grains within the groundmass. The
505 presence of magnetite, and lack of orthopyroxene, in the groundmass is not consistent with a
506 peritectic reaction. Furthermore, orthopyroxene grains are compositionally zoned, with
507 decreasing incompatible elements (e.g., Fe and Mn) as the olivine-orthopyroxene boundary is
508 approached (Figure 2d,e). This compositional zoning is the opposite of what is expected if
509 orthopyroxene nucleated after resorption of olivine margins (i.e., in the case of a peritectic
510 reaction). Instead, the compositional zoning in the orthopyroxene appears to continue the trend in
Fe and Mn observed in the olivine phenocrysts; MnO contents of olivine range from 0.8 to 1.6 wt% (average 1.1 wt%), increasing towards the rims, whereas orthopyroxene MnO ranges from 1.1 to 2.0 wt% (average 1.5 wt%). Lastly, additional support that olivine was transformed to orthopyroxene (+ magnetite) in a subsolidus reaction (equation [2]) is provided by the Fe/Mn ratios of the orthopyroxene (~22), which deviate significantly from what is expected (~35) based on crystallization from the melt, as an indicator of planetary parentage (Papike et al., 2009).

A similar assemblage of olivine, low-Ca pyroxene, and magnetite is known from the NWA 998 nakhlite, and is attributable to a subsolidus reaction under fO₂ conditions slightly above QFM (Treiman and Irving, 2008). Notably, in that case the oxide phase is a titanomagnetite (Mt₃₈Sp₂Usp₅₈Chr₂; endmembers recalculated in the same way as in our study), and the highest fO₂ estimates result from oxides in a magmatic inclusion that nevertheless contain appreciable Ti (Mt₅₃Sp₆Usp₃₈Chr₃ and Ilm₇₉Hem₁₆Gk₃.₅Py₁.₅). In the case of NWA 8159, fO₂ conditions may have been similar (~ΔQFM+2); however, in this case the dominant oxide is magnetite, as a result of the very low Ti content of the system, relative to other martian lithologies (Table 3).

The crystallization history of NWA 8159 that we infer from our observations is as follows:

1) Crystallization of the parental melt at fO₂ < QFM. The first phases to crystallize are olivine of Fo ~41 and augite with a composition of ~ Fs₄₀En₃₀Wo₃₀. Although the initial fO₂ of crystallization is poorly constrained, it must have been sufficiently low for olivine to be stable (i.e., below QFM). Plagioclase crystallization initiated shortly after olivine and augite, based on textural relationships.
2) Increase in the fO₂ of the melt, causing destabilization of the olivine phenocrysts and replacement of olivine phenocryst rims with orthopyroxene and magnetite. The metastability of olivine facilitated the subsolidus reaction according to equation [2], preferentially producing opx from olivine; notably, all of the opx that formed in this process is stable at low P (i.e., outside of the 'forbidden zone' of pyroxene crystallization at 1 bar; Figure 3). Some of the opx is more magnesian than would be in equilibrium with the most Mg-rich olivine; this opx is likely the result of the preferential partitioning of Fe in the olivine into the magnetite as a result of the oxidation involved in the reaction, leaving the opx relatively Fe-depleted. The fO₂ during this stage is estimated to be at least QFM+2, assuming magnetite is in equilibrium; the relatively low Fe³⁺ contents of opx and augite may be the result of preferential partitioning of Fe³⁺ into magnetite during the subsolidus reaction. Phenocrysts of magnetite likely grew at this stage; the presence of symplectitic cores in some magnetite phenocrysts (e.g., Figure 2b) may indicate the former presence of a phenocryst phase that was destabilized due to the increase in fO₂. The exact identity of this phase is unknown, although the presence of quartz+magnetite in the symplectite indicates the involvement of an Fe-bearing silicate. The increase in oxidation may have been caused by the introduction of an oxidizing fluid.

3) Crystallization of the groundmass, consisting of augite+plagioclase+magnetite occurs at the higher (≥QFM+2) fO₂ conditions. Ilmenite crystallizes once Ti has been sufficiently enriched in the remaining melt.

It should be noted that application of thermodynamic models (e.g., Ca-QUIF) to NWA 8159 are limited because the models are designed primarily to assess minerals that have attained
equilibrium whether during co-crystallization, subsolidus re-equilibration, or metamorphism (Andersen et al., 1993). It is clear from textural evidence that the phases in NWA 8159 are not all in equilibrium, and that reactions (e.g., equation [2]) have not gone to completion. Accordingly, $fO_2$ conditions for NWA 8159 are poorly constrained. Nevertheless, we conclude that the NWA 8159 melt experienced a significant (minimum 2 log unit) increase in $fO_2$ between the liquidus and the solidus that resulted in the observed texture and mineralogy.

## 4.2 Conditions of shock metamorphism

The igneous texture is cut across by a network of black, glassy veins visible in hand specimen and thin section, which range in thickness from 1 µm to 1 mm (Fig. 1b, 2a, 5). The boundary between the veins and host rock minerals is sharp, with micrometer-scale offset and displacement of olivine, pyroxene and plagioclase along vein margins (Fig. 5). These petrographic characteristics, and the association with high pressure minerals (Sharp et al., 2015), constrain the veins to have formed and evolved during the shock-compression regime, and they will hereafter be referred to as “shock veins”.

Plagioclase (An$_{50.5-62.3}$Ab$_{37.6-49.1}$Or$_{0.1-0.5}$; Tables 2, S4) in the host rock, far from local zones of shock melt, is birefringent ($\delta_{max}$ = 0.011) with typical albite polysynthetic twinning and carlsbad twins. All igneous plagioclase grains are pervasively fractured and exhibit undulose extinction. The fractures are predominately irregular, cutting across the entire grain or restricted to small areas inside the crystal. No evidence for incipient development of shock-induced planar fractures was found, despite a systematic search of over 100 grains. These may be distinguished from inherent cleavage planes, which are observed in some grains, by their close spacing and crystallographic orientation. Raman spectra exhibit sharp peaks characteristic of crystalline
material with positions at 178, 477, 509 and 660 cm⁻¹ (Fig. 6) matching the expected vibrational modes for labradorite (Mernagh, 1991).

Within a zone ~20-300 µm distant from shock veins, plagioclase is amorphous with original grain boundaries, igneous zoning profiles and twins preserved (Fig. 3a). Amorphous plagioclase is stoichiometric (4.98–5.00 cations based on 8 oxygen), with a composition (An₅₈.₄-60.8Ab₃₉.₁-4₁.₄Or₀.₁-0.₂) overlapping that of crystalline plagioclase (Table 2). The transition from crystalline to amorphous plagioclase is observed using a polarizing light microscope by an abrupt change from anisotropic to isotropic crystals. In BSE images this appears as a change from highly fractured (crystalline) to smooth, unfractured (amorphous) grains. Larger plagioclase crystals are partially transformed, with the amorphous portion closest to the shock vein, and the highly-fractured portion of the same crystal furthest from the vein (Fig. 7b). Raman spectra show a broad hump centered over 509 cm⁻¹, consistent with an amorphous state (Fig. 6). In direct contact with the shock vein, amorphous plagioclase has transformed to tissintite (Fig. 7c), a dense, vacancy-rich (Ca,Na,□)₂AlSi₂O₆ pyroxene first documented within and surrounding quenched shock melt in the depleted olivine-phyric shergottite Tissint (Ma et al., 2015). The mean composition of this mineral in NWA 8159 is (Na₀.₃Ca₀.₄₄□₀.₂₆)(Al₀.₉₇Fe₀.₀₃)(Si₁.₈₆Al₀.₁₄)O₆ (Table 2). In BSE images, tissintite appears brighter compared to neighboring plagioclase as a result of the higher density of this phase and the slightly higher Fe-concentration (0.74–1.05 wt% FeO versus 0.40–0.85 wt% FeO). Peak positions in the Raman spectrum collected on tissintite in the type specimen, as reported by Walton et al. (2014), are identical to those spectra collected on NWA 8159 in this study with peaks near 378 cm⁻¹, 520 cm⁻¹, 697 cm⁻¹ and 1008 cm⁻¹ (Fig. 6).

The direct relationship between plagioclase transformation and shock veins is perhaps the most significant shock metamorphic finding in NWA 8159. Such an association has been
documented from only a handful of shock-veined rocks (e.g., Walton et al., 2016). Insight into
the relationship between plagioclase transformation and shock veins is gained from *in situ* high-
pressure X-ray diffraction experiments, in which the amorphization conditions of albite
(Ab<sub>98.0</sub>An<sub>0.4</sub>Or<sub>1.6</sub>) and labradorite (An<sub>45.0</sub>Ab<sub>551.8</sub>Or<sub>3.2</sub>) are investigated (Kubo et al., 2010). These
experiments show that the pressure at which plagioclase becomes amorphous decreases with
increasing temperature and is compositionally dependent. For temperatures ≤ 200 °C, complete
amorphization of labradorite at 23 GPa was not achieved, and complete amorphization of albite
(Ab<sub>98.0</sub>An<sub>0.4</sub>Or<sub>1.6</sub>) occurred at static pressures of 37 GPa. In contrast, labradorite, appropriate for
plagioclase in NWA 8159, at elevated temperatures (>200 °C) undergoes complete
amorphization at 19 GPa with transformation to jadeite occurring at the same pressure but under
temperatures ≥800 °C. The compression experiments of Daniel et al. (1997) also shed light on
plagioclase transformation because they demonstrate that the high-pressure behavior of anorthite
is not only temperature-dependent, but is also very sensitive to deviatoric stresses, such as those
conditions experienced by minerals in direct contact with shock veins. In these experiments the
onset of amorphization is reduced from 16 GPa to 11 GPa, compared to anorthite compressed
under hydrostatic conditions. The results of Daniel et al. (1997) and Kubo et al. (2010) account
for the observed distribution of plagioclase transformation in NWA 8159, showing that this
distribution is a result of temperature gradients and deviatoric pressure experienced by igneous
plagioclase along shock vein margins (Fig. 5). The temperature gradients are established because
shock melt is at ~2000-2500 °C compared to the surrounding host rock, where the temperature
increase by shock compression and the nonadiabatic deposition of heat following decompression,
is limited to a few hundred degrees (Sharp and DeCarli, 2006). The shock veins cool by
conduction of heat to the host rock, with the hottest portion of the grain in direct contact with
shock melt (Shaw and Walton, 2013); it is within this zone that the conditions for tissintite formation and plagioclase amorphization were realized.

Partial transformation of labradorite in NWA 8159 constrains the shock conditions experienced by this meteorite to be >15 GPa and <23 GPa. This is in good agreement with the results of shock recovery experiments, in which plagioclase of An₅₀.₄-63.₅ composition remains completely crystalline up to a maximum shock pressure of ~26 GPa (Stöffler et al., 1986). Shock deformation in non-veined regions of the host rock is consistent with low to moderate degrees of shock metamorphism. Olivine and pyroxene exhibit irregular fractures and undulose extinction to weak mosaicism, with minor development of planar fractures in olivine. Polysynthetic mechanical twin lamellae are present in some pyroxene grains. Raman spectra from olivine show sharp peaks at 816 cm⁻¹ and 846 cm⁻¹, characteristic of crystalline (Mg,Fe)₂SiO₄ olivine. Additional features in Raman spectra documented from olivine in strongly shocked martian meteorites NWA 2737 and NWA 1950 are notably absent in NWA 8159. These include a broadening of peaks characteristic of highly disordered olivine crystals; the presence of defects within the structure indicated by a weak peak at 760 cm⁻¹; and additional broad bands appearing at 650 and 686 cm⁻¹ that are consistent with the presence of a shock-produced metastable olivine polymorph (ζ-(Mg,Fe)₂SiO₄; Van De Moortèle et al., 2007).

Free SiO₂ in the host rock occurs as small crystals (10-25 µm) with hexagonal basal sections. Raman spectra collected on these grains show a strong peak at 464 cm⁻¹, with several weaker peaks at 127, 204 and 355 cm⁻¹ that are consistent with reported Raman spectra from unshocked quartz, and samples shocked to 10.7 GPa and 14.5 GPa (McMillan et al., 1992) (Fig. 6). Quartz grains in NWA 8159 show no evidence for phase transformation such as the
646 formation of seifertite, which is found in the silica-rich mesostasis of strongly shocked
647 shergottites Shergotty, Zagami and Dhofar 378 (Blass, 2013; El Goresy et al., 2008).

648 Our estimate of ~15-23 GPa for the shock pressure experienced by NWA 8159, although
649 lower than most shergottites, is within the range obtained for other martian meteorites, which
650 have experienced shock pressures of at least 5-14 GPa (Fritz et al., 2005). These results also
651 support predictions from numerical modeling that material cannot be ejected from the surface of
652 Mars without experiencing a shock pressure of at least 10 GPa (Artemieva and Ivanov, 2004).
653 NWA 8159 joins a small, but increasing number of shergottite-like martian meteorites which
654 contain precursor igneous plagioclase preserved throughout the host rock; others include NWA
655 4480 and NWA 10416 (Irving et al., 2016; Walton et al., 2016).

656

## 4.3 Geochemistry, geochronology, and mantle source characteristics

### 4.3.1 Parental melt characteristics

659 The bulk composition of NWA 8159, determined by EPMA of shock veins and solution
660 ICP-MS of bulk sample, is given in Table 3 with the compositions of NWA 7635, MIL 03346
661 intercumulus melt, NWA 5790 mesostasis, QUE 94201, and Los Angeles for comparison. REE
662 patterns are plotted in Figure 8. Terrestrial alteration has only minimally affected the bulk
663 composition of NWA 8159 – neither the anomalous Ce nor the LREE enrichment expected from
664 hot desert meteorites (e.g., Crozaz and Wadhwa, 2001) are observed (Table 3, Figure 8),
665 although enrichment in Sr, Ba, and Pb are evident (Table 3; cf. NWA 8159 and Los Angeles,
666 QUE 94201), consistent with desert alteration (Barrat et al., 1999; Stelzner et al., 1999) including
667 the precipitation of barite. The NWA 8159 bulk composition is very similar to that of the
668 recently discovered NWA 7635 meteorite (Fig. 8; Lapen et al., 2017). The major and minor
element composition of NWA 8159 is most similar to that of Los Angeles, with the exception of lower TiO₂ and P₂O₅; however, these meteorites contain different mineralogies, possibly the result of differing redox conditions during crystallization. Despite the mineralogical similarities shown by the experimental crystallization of nakhlite mesostasis (Herd et al., 2014; Herd and Walton, 2008), significant differences of some elements in the bulk NWA 8159 major and minor element composition and REE patterns (Fig. 8) indicate there is no direct relationship between this meteorite and nakhlites through simple igneous differentiation processes. While the bulk rock REE patterns for NWA 8159 and NWA 7635 show LREE depletion similar to that of depleted shergottites such as Tissint and Yamato 980459 (Figure 8), these meteorites are distinguished from all other martian lithologies by the slope of the HREE, best quantified by CI-normalized Dy/Lu. The augite-rich shergottites NWA 8159 and NWA 7635 have Dy/Lu ~0.84, whereas other martian lithologies have Dy/Lu > 1 (Table 3). Also of note is the low concentration of Ni and Co in NWA 8159. Bulk Co is higher than Ni by approximately a factor of 3; this is most similar to the mesostasis of the nakhlite NWA 5790, but in contrast to Los Angeles in which Ni and Co are subequal in concentration (Table 3). It is evident from major, minor and trace element concentrations that NWA 8159 and NWA 7635 are derived from parental melts which are distinct from those of other martian meteorites.

Shergottites are characterized by a wide range of time-integrated incompatible element enrichment, clustering into three distinct geochemical groups – depleted, intermediate and enriched – based on their Rb-Sr and Sm-Nd isotopic compositions and REE patterns (e.g., McCoy et al., 2011; Symes et al., 2008). REE patterns range from LREE-depleted to relatively enriched (Figure 8). Early and rapid differentiation of the silicate portion of Mars is called upon to generate enriched and depleted reservoirs by 4.5 Ga (e.g., Borg et al., 2016), either through
692 formation of primitive martian crust by partial melting of the mantle, or crystallization of a
693 magma ocean (Blinova and Herd, 2009; Borg et al., 1997). In the latter model, the last dregs of
694 magma ocean crystallization comprise the enriched portion, and the cumulate mantle is the
695 depleted counterpart (Blinova and Herd, 2009; Borg and Draper, 2003); whether a martian
696 magma is depleted, intermediate, or enriched depends on the mix of depleted or enriched
697 components in its mantle source, which was set during crystallization at 4.5 Ga. A correlation
698 between long-term incompatible element enrichment and oxygen fugacity is also apparent
699 among the shergottites (Herd et al., 2002; Wadhwa, 2001), which can be explained by
700 differences in the redox states of the mantle sources (more incompatible element depleted
701 sources are less oxidized, while enriched sources are more oxidized; Herd, 2003). This
702 correlation is illustrated in Figure 13. Notably, the LREE depleted but high fO₂ nature of NWA
703 8159 distinguishes it from the other shergottites, and strongly suggests that NWA 8159
704 originates from a distinct mantle source and that its petrogenesis involves a unique melt
705 composition and conditions of crystallization. Increases in fO₂ during crystallization on the order
706 of ~2 log units are known from the olivine-phyric shergottites (e.g., Gross et al., 2013; Herd,
707 2006; Peslier et al., 2010), yet in all of those cases, the initial (liquidus) fO₂ is much lower,
708 ~ΔQFM-3 to ΔQFM-2. Our observations of NWA 8159 suggest that it formed from a melt with
709 an initial fO₂ of < QFM; the rapid increase in fO₂ evidenced by the olivine phenocryst rims and
710 groundmass assemblages may be explained by interaction with an oxidizing fluid during the
711 latter stages of crystallization.

712

713 *4.3.2 Formation Age*
At face value, the Sm-Nd isotopic data exhibit significant scatter on an isochron diagram indicating disturbance (Figure S1). However, when residual leached materials are separated from unleached and leachate materials, a more meaningful isochron is revealed (Figure 9). Distinct trends among the Sm-Nd data can also be distinguished in a plot of Nd isotopes vs. 1/Nd where it becomes clear that the leaching procedures removed a surficial Nd component presumably leaving only the indigenous or magmatic Nd in the residual materials (Figure S2). Although clearly off the isochron, the secondary Nd source contaminant is highly radiogenic and not consistent with any known terrestrial composition. Once the measurements that contain the effects of this reservoir have been removed the 6 remaining residual data points define the 2.37 ± 0.25 Ga isochron age reported herein. It is notable that the Nd isotopic compositions of potential contaminants such as terrestrially-sourced sedimentary materials from northwestern Africa (e.g., Borg and Banner, 1996) are less radiogenic and lie far from a line extrapolated from the Sm-Nd isochron, indicating that terrestrial disturbances are unlikely to have significantly affected the Sm-Nd age.

The <sup>40</sup>Ar/<sup>39</sup>Ar isochron age obtained from glass produced by quenching of shock melt confirms an early Amazonian age for NWA 8159. Given the distinct chemical behavior of the Sm-Nd and K-Ar isotope systems to secondary processes such as impact heating and weathering, the concordance of these two chronometers attests to the significance of the age. It thus appears that the glass in NWA 8159 shock veins formed shortly after the rock crystalized (within 0.22 ± 0.35 Ga) or formed without significant loss of radiogenic Ar. Although the rock experienced one or more thermal events in its history that resulted in significant loss of radiogenic Ar and He from the fine-grained groundmass (Cassata and Borg, 2016), the shock-produced glass appears to have retained >95% of its radiogenic Ar during these thermal events. The high retention of Ar
and He may be explained by the cooling time of the shock-produced melt, which depends on the conductive cooling time from the initial temperature of the shock melt (~2500 °C) to the background temperature of the host rock. The quench time will be most efficient for shock melt in rocks experiencing lower shock levels because shock temperature increases with shock pressure. We conclude therefore, that the lower levels of shock experienced by NWA 8159 (15-23 GPa; this study) helped to retain radiogenic Ar and He, and that the shock veins likely formed at the time of ejection from Mars.

The bulk rock ¹⁴⁷Sm/¹⁴⁴Nd values of NWA 8159 range from 0.318 to 0.381 and are consistent with derivation of this sample directly from its mantle source. This inference is based on the observation that the ¹⁴⁷Sm/¹⁴⁴Nd ratio of the NWA 8159 source region, calculated from the Nd isotopic systematics of Wr-1 and Wr-2 and the crystallization age of 2.37 Ga, is estimated to be 0.31 to 0.33 and approximates the bulk rock values. This is in contrast to most other martian meteorites which show significant deviations between the ¹⁴⁷Sm/¹⁴⁴Nd measured for the bulk rocks and the ¹⁴⁷Sm/¹⁴⁴Nd calculated for their source regions using ¹⁴³Nd/¹⁴⁴Nd ratios (e.g., Borg et al., 2016; Borg et al., 2003; Debaille et al., 2007); the exception is NWA 7635, which has a bulk rock ¹⁴⁷Sm/¹⁴⁴Nd value of 0.314 and a source ¹⁴⁷Sm/¹⁴⁴Nd value of 0.303 (Lapen et al., 2017). The high initial ε¹⁴³Nd value of NWA 8159 indicates its source was depleted in incompatible elements – probably similar to that defined by the most depleted meteorites from Mars – and is consistent with its bulk REE composition.

### 4.33. Mantle source characteristics

Prior studies of the classical martian ("SNC") meteorites observed a homogenous Cr isotopic composition among the limited number of samples analyzed, attributable to Cr isotopic
homogeneity in the martian mantle (e.g., Trinquier et al., 2007). NWA 8159 is indistinguishable
from Shergotty and Nakhla in either ε<sup>53</sup>Cr or ε<sup>54</sup>Cr isotopic composition, within analytical
uncertainty (Table 5, Figure 12). However, a resolvable difference exists between NWA 8159
and Chassigny, in particular in ε<sup>54</sup>Cr. This difference may be attributable to the presence of two
distinct mantle sources from which the melts that generated NWA 8159 and Chassigny
originated. Since Shergotty and Nakhla overlap both Chassigny and NWA 8159 in Cr isotopic
composition, the compositional relationship of these samples to either source reservoir remains
unclear based on Cr alone. However, our results for NWA 8159 strongly suggest that Cr isotopic
heterogeneity is present within the martian mantle, and this sample is the first to be recognized as
having been derived from a distinct mantle Cr isotopic reservoir.

The idea that NWA 8159 derives from a distinct martian mantle reservoir is also
supported by its W and Nd isotope compositions. In particular, the ε<sup>182</sup>W and ε<sup>142</sup>Nd determined
here are distinct from any of the known groups of shergottites as well as the nakhlites (e.g., Borg
et al., 2016; Debaille et al., 2009; Debaille et al., 2007; Foley et al., 2005; Kleine et al., 2004),
with the exception of NWA 7635 (Table 8). Our observations of NWA 8159 are generally
consistent with derivation from a mantle source that is somewhat more incompatible element
depleted than that of the depleted shergottites.

5. CONCLUSIONS

The petrogenetic history, bulk elemental and isotopic composition, and age of NWA
8159 distinguishes this meteorite from others in the martian meteorite suite: the NWA 8159
parent melt is relatively evolved, Ca-rich, Ti-poor and LREE-depleted, has Dy/Lu ~ 0.8, and
crystallized an assemblage dominated by augite, olivine, plagioclase and magnetite under
relatively oxidized conditions. The presence of magnetite as the dominant oxide establishes
784 NWA 8159 as another candidate lithology to account for the strong crustal magnetization in
785 certain parts of the martian crust. The rock experienced distinct conditions of shock – generally
786 lower than most shergottites – which preserves the partial transformation of plagioclase to
787 maskelynite. The high-$fO_2$, low-La/Yb ratio and bulk isotopic composition of NWA 8159
788 strongly suggest that it is derived from a previously unsampled mantle source. Its ~2.4 Ga age is
789 only matched by that of NWA 7635, with which it also shares similarities in REE pattern (as
790 quantified by La/Yb and Dy/Lu ratios) and ejection age (Table 8), the latter suggesting the two
791 rocks were launched from Mars in the same impact event. Although the datasets required to fully
792 compare the two samples are incomplete, differences in mineralogy, texture, and Nd and W
793 isotopic composition between NWA 8159 and NWA 7635 (Table 8) suggest different mantle
794 sources and conditions of crystallization for these two rocks. Therefore, NWA 8159 represents a
795 distinct sampling of an early Amazonian igneous unit on Mars.

796
797
5. ACKNOWLEDGEMENTS

798 We thank Nathan Gerein (UAlberta) for assistance with SEM analysis, Andrew Locock
799 (UAlberta) and Mike Spilde (UNM) for assistance with EPMA, Guangcheng Chen (UAlberta)
800 for ICP-MS analysis for trace elements, and Timothy J. Peters for help with mineral separations
801 for Sm-Nd isotopic work at NASA-JSC. Funding for this study was provided by Natural
802 Sciences and Engineering Research Council (NSERC) Grant 261740 to CDKH, and NSERC
803 Grant RES00007057 to ELW. Partial funding to support the isotopic work at NASA-JSC was
804 provided by NASA LASER and Cosmochemistry Programs (grants 10-LASER10-0077 & 11-
805 COS11-0066 to JIS). Partial funding to support noble gas analyses at LLNL was provided by the
806 NASA Mars Fundamental Research Program (grant NNH14AX56I to WSC). FMM also
807 acknowledges partial support from the NASA Mars Fundamental Research Program (grant
808 NNX13AG44G to FMM). Partial funding to support Cr isotope work at UC Davis is provided by

809 NASA Cosmochemistry (NNX14AM62G) and Emerging Worlds (NNX16AD34G) Programs to

810 QZY. KN and MWC acknowledge the NASA Cosmochemistry program (NNX14AK62G) for

811 funding to support cosmogenic radionuclide work. All samples were provided by the UNM

812 Meteorite Museum. Constructive reviews by Tony Irving and two anonymous reviewers, and

813 editorial handling by Munir Humayun resulted in changes that significantly improved the

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1163

1164
1165 Tables

1166

<caption>Table 1. Representative compositions of NWA 8159 olivine, pyroxene, magnetite and ilmenite</caption>

1168

<table><thead><tr><th></th><th>olivine</th><th>olivine</th><th>augite</th><th>augite</th><th>opx</th><th>opx</th><th>magnetite</th><th>ilmenite</th><th>ilmenite</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>32.40</td><td>30.95</td><td>51.73</td><td>45.47</td><td>52.11</td><td>48.66</td><td>0.12</td><td>0.16</td><td>0.30</td></tr><tr><td>TiO<sub>2</sub></td><td>b.d.</td><td>b.d.</td><td>0.25</td><td>0.99</td><td>b.d.</td><td>b.d.</td><td>0.36</td><td>35.87</td><td>27.29</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>b.d.</td><td>b.d.</td><td>1.81</td><td>0.73</td><td>0.10</td><td>0.09</td><td>0.13</td><td>0.77</td><td>0.87</td></tr><tr><td>FeO</td><td>49.38</td><td>58.04</td><td>14.68</td><td>36.64</td><td>27.82</td><td>40.70</td><td>32.05</td><td>32.31</td><td>25.38</td></tr><tr><td>Fe<sub>2</sub>O<sub>3</sub></td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td><td>67.15</td><td>28.17</td><td>44.07</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.12</td><td>b.d.</td><td>0.97</td><td>b.d.</td><td>0.13</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>0.14</td></tr><tr><td>MnO</td><td>1.04</td><td>1.49</td><td>0.42</td><td>0.76</td><td>1.56</td><td>1.91</td><td>0.12</td><td>0.99</td><td>0.88</td></tr><tr><td>MgO</td><td>15.72</td><td>8.24</td><td>12.22</td><td>2.22</td><td>18.37</td><td>9.23</td><td>0.16</td><td>0.60</td><td>0.47</td></tr><tr><td>CaO</td><td>0.60</td><td>0.66</td><td>17.91</td><td>11.98</td><td>0.75</td><td>0.17</td><td>0.12</td><td>0.47</td><td>0.39</td></tr><tr><td>Na<sub>2</sub>O</td><td>b.d.</td><td>b.d.</td><td>0.18</td><td>0.11</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td></tr><tr><td>K<sub>2</sub>O</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td></tr><tr><td>Total</td><td>99.26</td><td>99.38</td><td>100.16</td><td>98.90</td><td>100.84</td><td>100.75</td><td>100.21</td><td>99.34</td><td>99.81</td></tr><tr><td>Fe/Mn</td><td>46.7</td><td>38.6</td><td>34.5</td><td>47.9</td><td>17.6</td><td>21.0</td><td></td><td></td><td></td></tr><tr><td>Fs (mol%) or<br/>Fa (Xmol)</td><td>0.64</td><td>0.80</td><td>24.7</td><td>65.5</td><td>45.2</td><td>70.9</td><td></td><td></td><td></td></tr><tr><td>En (mol%) or<br/>Fo (Xmol)</td><td>0.36</td><td>0.20</td><td>36.7</td><td>7.1</td><td>53.2</td><td>28.7</td><td></td><td></td><td></td></tr><tr><td>Wo (mol%)</td><td>-</td><td>-</td><td>38.6</td><td>27.4</td><td>1.6</td><td>0.4</td><td></td><td></td><td></td></tr><tr><td>Mt or Hem<br/>(mol%)</td><td></td><td></td><td></td><td></td><td></td><td></td><td>97.6</td><td>29.9</td><td>46.2</td></tr><tr><td>Usp or Ilm<br/>(mol%)</td><td></td><td></td><td></td><td></td><td></td><td></td><td>2.1</td><td>65.6</td><td>50.0</td></tr></tbody></table>

b.d. = below detection

- = not determined

1169

1170

1171

1172

1173
<caption>Table 2. Representative compositions of tissintite, plagioclase and plagioclase glass in NWA</caption>

1174
1175 8159

<table><thead><tr><th></th><th>tissintite</th><th>tissintite</th><th>tissintite</th><th>tissintite</th><th>glass</th><th>glass</th><th>glass</th><th>glass</th><th>plag</th><th>plag</th><th>plag</th><th>plag</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>55.07</td><td>54.99</td><td>54.93</td><td>54.61</td><td>53.48</td><td>54.38</td><td>53.62</td><td>54.88</td><td>53.26</td><td>55.72</td><td>54.17</td><td>54.90</td></tr><tr><td>TiO<sub>2</sub></td><td>0.02</td><td>0.03</td><td>0.01</td><td>0.05</td><td>0.01</td><td>0.05</td><td>0.02</td><td>0.01</td><td>b.d.</td><td>b.d.</td><td>0.03</td><td>0.03</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>28.10</td><td>27.88</td><td>27.69</td><td>27.20</td><td>29.48</td><td>28.68</td><td>29.13</td><td>28.07</td><td>29.20</td><td>27.62</td><td>27.89</td><td>27.87</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.04</td><td>0.00</td><td>0.04</td><td>0.01</td><td>0.01</td><td>0.02</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>b.d.</td><td>0.01</td><td>b.d.</td></tr><tr><td>FeO</td><td>0.89</td><td>0.74</td><td>0.82</td><td>1.05</td><td>0.60</td><td>0.65</td><td>0.67</td><td>0.83</td><td>0.49</td><td>0.65</td><td>0.85</td><td>0.66</td></tr><tr><td>MnO</td><td>0.02</td><td>0.03</td><td>b.d.</td><td>b.d.</td><td>0.03</td><td>b.d.</td><td>0.02</td><td>0.03</td><td>0.02</td><td>0.02</td><td>0.02</td><td>0.03</td></tr><tr><td>MgO</td><td>0.05</td><td>0.03</td><td>0.04</td><td>b.d.</td><td>0.03</td><td>0.04</td><td>0.04</td><td>0.00</td><td>0.00</td><td>0.00</td><td>0.00</td><td>b.d.</td></tr><tr><td>CaO</td><td>11.40</td><td>12.04</td><td>12.00</td><td>12.25</td><td>12.27</td><td>11.92</td><td>12.26</td><td>11.52</td><td>12.50</td><td>10.40</td><td>12.07</td><td>11.26</td></tr><tr><td>Na<sub>2</sub>O</td><td>4.61</td><td>4.43</td><td>4.40</td><td>4.73</td><td>4.36</td><td>4.42</td><td>4.37</td><td>4.51</td><td>4.24</td><td>5.58</td><td>4.64</td><td>4.88</td></tr><tr><td>K<sub>2</sub>O</td><td>0.01</td><td>0.04</td><td>0.01</td><td>0.04</td><td>0.02</td><td>0.04</td><td>0.02</td><td>0.02</td><td>0.03</td><td>0.07</td><td>0.04</td><td>0.01</td></tr><tr><td>Total</td><td>100.19</td><td>100.19</td><td>99.94</td><td>99.93</td><td>100.28</td><td>100.20</td><td>100.15</td><td>99.88</td><td>99.74</td><td>100.06</td><td>99.73</td><td>99.65</td></tr><tr><td colspan="14">Number of cations based on 6 oxygens for tissintite and 8 oxygens for plagioclase</td></tr><tr><td>Si</td><td>1.8541</td><td>1.8608</td><td>1.8697</td><td>1.8609</td><td>2.4245</td><td>2.4537</td><td>2.4259</td><td>2.4814</td><td>2.4122</td><td>2.4997</td><td>2.4560</td><td>2.4859</td></tr><tr><td>Ti</td><td>0.0005</td><td>0.0007</td><td>0.0003</td><td>0.0013</td><td>0.0004</td><td>0.0017</td><td>0.0007</td><td>0.0004</td><td>0.0000</td><td>0.0000</td><td>0.0012</td><td>0.0011</td></tr><tr><td>Al</td><td>1.1270</td><td>1.1120</td><td>1.1008</td><td>1.0923</td><td>1.5603</td><td>1.5252</td><td>1.5536</td><td>1.4960</td><td>1.5588</td><td>1.4605</td><td>1.4905</td><td>1.4875</td></tr><tr><td>Cr</td><td>0.0010</td><td>0.0000</td><td>0.0010</td><td>0.0003</td><td>0.0003</td><td>0.0009</td><td>0.0000</td><td>0.0000</td><td>0.0000</td><td>0.0000</td><td>0.0005</td><td>0.0000</td></tr><tr><td>Fe</td><td>0.0249</td><td>0.0209</td><td>0.0233</td><td>0.0299</td><td>0.0225</td><td>0.0246</td><td>0.0255</td><td>0.0313</td><td>0.0186</td><td>0.0244</td><td>0.0323</td><td>0.0251</td></tr><tr><td>Mn</td><td>0.0005</td><td>0.0008</td><td>0.0001</td><td>0.0000</td><td>0.0011</td><td>0.0000</td><td>0.0006</td><td>0.0013</td><td>0.0007</td><td>0.0007</td><td>0.0008</td><td>0.0012</td></tr><tr><td>Mg</td><td>0.0027</td><td>0.0014</td><td>0.0018</td><td>0.0001</td><td>0.0017</td><td>0.0027</td><td>0.0027</td><td>0.0000</td><td>0.0000</td><td>0.0000</td><td>0.0000</td><td>0.0000</td></tr><tr><td>Ca</td><td>0.4111</td><td>0.4365</td><td>0.4359</td><td>0.4472</td><td>0.5904</td><td>0.5763</td><td>0.5942</td><td>0.5579</td><td>0.6066</td><td>0.4999</td><td>0.5864</td><td>0.5463</td></tr><tr><td>Na</td><td>0.3138</td><td>0.2904</td><td>0.2910</td><td>0.3127</td><td>0.3795</td><td>0.3868</td><td>0.3832</td><td>0.3956</td><td>0.3724</td><td>0.4854</td><td>0.4079</td><td>0.4285</td></tr><tr><td>K</td><td>0.0005</td><td>0.0019</td><td>0.0006</td><td>0.0017</td><td>0.0012</td><td>0.0022</td><td>0.0012</td><td>0.0013</td><td>0.0018</td><td>0.0038</td><td>0.0025</td><td>0.0008</td></tr><tr><td>Total</td><td>3.7360</td><td>3.7253</td><td>3.7244</td><td>3.7463</td><td>4.9820</td><td>4.9740</td><td>4.9876</td><td>4.9652</td><td>4.9710</td><td>4.9744</td><td>4.9780</td><td>4.9763</td></tr><tr><td>An</td><td></td><td></td><td></td><td></td><td>60.8</td><td>59.7</td><td>60.7</td><td>58.4</td><td>61.9</td><td>50.5</td><td>50.5</td><td>56.0</td></tr><tr><td>Ab</td><td></td><td></td><td></td><td></td><td>39.1</td><td>40.1</td><td>39.2</td><td>41.4</td><td>38.0</td><td>49.1</td><td>36.4</td><td>43.9</td></tr><tr><td>Or</td><td></td><td></td><td></td><td></td><td>0.1</td><td>0.2</td><td>0.1</td><td>0.1</td><td>0.2</td><td>0.4</td><td>0.1</td><td>0.1</td></tr></tbody></table>

b.d. = below detection
<table><caption>**Table 3.** Bulk composition of NWA 8159 compared to other martian lithologies</caption><thead><tr><th>Sample</th><th>NWA<br>8159<br>(UAlberta)</th><th>NWA<br>8159<br>(LLNL)</th><th>NWA<br>8159</th><th>NWA<br>7635</th><th>MIL<br>03346</th><th>NWA<br>5790</th><th>QUE 94201</th><th>Los<br>Angeles</th></tr><tr><th>Major</th><th>ICP-MS<sup>a</sup></th><th></th><th>EMPA<sup>b</sup></th><th>XRF<sup>c</sup></th><th>EMPA<sup>d</sup></th><th>EMPA<sup>f</sup></th><th>INAA<sup>g</sup></th><th>INAA<sup>h</sup></th></tr><tr><th>Trace</th><th>ICP-MS<sup>a</sup></th><th>ICP-MS</th><th></th><th>ICP-MS<sup>c</sup></th><th>ICP-MS<sup>e</sup></th><th>ICP-MS<sup>f</sup></th><th>INAA/RNAA<sup>g</sup></th><th>ICP-MS<sup>i</sup></th></tr><tr><th colspan="9">wt%</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>48</td><td></td><td>47(1)</td><td>47.60</td><td>48.5</td><td>51.75</td><td>47.90</td><td>48.60</td></tr><tr><td>TiO<sub>2</sub></td><td>0.65</td><td></td><td>0.61(8)</td><td>0.61</td><td>1.39</td><td>1.03</td><td>1.98</td><td>1.43</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>10.74</td><td></td><td>11.2(9)</td><td>11.16</td><td>8.17</td><td>13.19</td><td>9.82</td><td>10.40</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.14</td><td></td><td>0.12(9)</td><td>0.15</td><td>-</td><td>0.00</td><td>0.15</td><td>0.02</td></tr><tr><td>MgO</td><td>4.35</td><td></td><td>3.9(6)</td><td>4.11</td><td>3.92</td><td>0.58</td><td>6.30</td><td>3.74</td></tr><tr><td>CaO</td><td>11.98</td><td></td><td>9.1(6)</td><td>9.77</td><td>9.51</td><td>4.68</td><td>11.47</td><td>9.89</td></tr><tr><td>MnO</td><td>0.46</td><td></td><td>0.46(5)</td><td>0.47</td><td>0.53</td><td>0.32</td><td>0.47</td><td>0.47</td></tr><tr><td>FeO</td><td>20.86</td><td></td><td>24(1)</td><td>22.50</td><td>24.4</td><td>16.67</td><td>19.17</td><td>21.40</td></tr><tr><td>Na<sub>2</sub>O</td><td>2.46</td><td></td><td>2.1(2)</td><td>2.18</td><td>2.38</td><td>5.30</td><td>1.37</td><td>2.13</td></tr><tr><td>K<sub>2</sub>O</td><td>0.10</td><td></td><td>0.10(4)</td><td>0.07</td><td>0.71</td><td>1.62</td><td>0.04</td><td>0.31</td></tr><tr><td>P<sub>2</sub>O<sub>5</sub></td><td>0.26</td><td></td><td>0.29(3)</td><td>0.28</td><td>0.70</td><td>1.04</td><td></td><td>1.49</td></tr><tr><td>Cl</td><td></td><td></td><td>0.07(2)</td><td></td><td></td><td>-</td><td></td><td></td></tr><tr><td>Total</td><td>100.0</td><td></td><td>98(1)</td><td>98.9</td><td>100.2</td><td>96.2</td><td>98.7</td><td>99.9</td></tr><tr><td>Mg#</td><td>0.27</td><td></td><td>0.22</td><td>0.25</td><td>0.22</td><td>0.06</td><td>0.37</td><td>0.24</td></tr><tr><td>Dy/Lu</td><td>0.87</td><td>0.80</td><td></td><td>0.86</td><td>1.28</td><td>1.47</td><td>1.13</td><td>1.52</td></tr><tr><th colspan="9">ppm</th></tr><tr><td>Li</td><td>8.30</td><td></td><td></td><td></td><td></td><td></td><td></td><td>5.03</td></tr><tr><td>Be</td><td>b.d.</td><td>0.012(3)</td><td></td><td></td><td></td><td></td><td></td><td>0.54</td></tr><tr><td>B</td><td>2.90</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>V</td><td>44.3</td><td>35(1)</td><td></td><td>40.4</td><td>210</td><td></td><td>124</td><td></td></tr><tr><td>Co</td><td>14.8</td><td>9.7(3)</td><td></td><td></td><td>38.7</td><td>11.6</td><td>24.4</td><td>29.2</td></tr><tr><td>Ni</td><td>3.33</td><td>2.55(7)</td><td></td><td>7.7</td><td>58.1</td><td>3.3</td><td>6.6</td><td>32</td></tr><tr><td>Cu</td><td>5.30</td><td>5.2(2)</td><td></td><td></td><td>13.3</td><td></td><td></td><td>22.6</td></tr><tr><td>Zn</td><td>95.3</td><td>98(4)</td><td></td><td></td><td>61.5</td><td></td><td>90</td><td>62.3</td></tr><tr><td>Ga</td><td>22.1</td><td>14.9(5)</td><td></td><td></td><td>6.77</td><td></td><td>25.9</td><td>22.09</td></tr><tr><td>Ge</td><td>0.73</td><td></td><td></td><td></td><td></td><td></td><td>1.95</td><td></td></tr><tr><td>As</td><td>1.62</td><td>0.31(4)</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Se</td><td>0.72</td><td>1.0(3)</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Rb</td><td>0.40</td><td>0.28(2)</td><td></td><td>0.40</td><td>4.41</td><td>14.66</td><td></td><td>12</td></tr><tr><td>Sr</td><td>163</td><td>120(4)</td><td></td><td>69.2</td><td>121.3</td><td>282.3</td><td>59</td><td>81</td></tr></tbody></table>
<table><tr><td>Y</td><td>15.6</td><td>12.8(8)</td><td></td><td>8.5</td><td>14.1</td><td></td><td>29.39</td></tr><tr><td>Zr</td><td>11.1</td><td>9.1(2)</td><td>9.8</td><td>21.2</td><td>63.6</td><td>94</td><td>79.6</td></tr><tr><td>Nb</td><td>0.29</td><td>0.055(9)</td><td></td><td>3.65</td><td></td><td></td><td>4.99</td></tr><tr><td>Mo</td><td>0.16</td><td>0.070(4)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Ru</td><td>b.d.</td><td>b.d.</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Rh</td><td>b.d.</td><td>0.0022(2)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pd</td><td>0.65</td><td>0.41(2)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Ag</td><td>0.33</td><td>0.0376(8)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Cd</td><td>0.37</td><td>0.236(5)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Sn</td><td>0.89</td><td>0.442(9)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Sb</td><td>0.06</td><td>0.0023(5)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Te</td><td>b.d.</td><td>b.d.</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Cs</td><td>b.d.</td><td>0.0134(3)</td><td></td><td>0.27</td><td></td><td></td><td>0.88</td></tr><tr><td>Ba</td><td>65.2</td><td>140(4)</td><td>9.9</td><td>56.9</td><td>422</td><td></td><td>46.8</td></tr><tr><td>La</td><td>0.24</td><td>0.139(2)</td><td>0.19</td><td>3.89</td><td>13.16</td><td>0.44</td><td>3.97</td></tr><tr><td>Ce</td><td>0.85</td><td>0.535(6)</td><td>0.83</td><td>11.3</td><td>33.66</td><td>1.63</td><td>9.84</td></tr><tr><td>Pr</td><td>0.19</td><td>0.122(5)</td><td>0.196</td><td>1.78</td><td>4.19</td><td></td><td>1.43</td></tr><tr><td>Nd</td><td>1.38</td><td>0.92(1)</td><td>1.38</td><td>8.04</td><td>17.9</td><td>2.36</td><td>7.07</td></tr><tr><td>Sm</td><td>0.84</td><td>0.557(1)</td><td>0.85</td><td>1.83</td><td>3.51</td><td>2.55</td><td>2.64</td></tr><tr><td>Eu</td><td>0.39</td><td>0.223(3)</td><td>0.33</td><td>0.52</td><td>1.02</td><td>1.09</td><td>1.02</td></tr><tr><td>Gd</td><td>1.64</td><td>1.02(3)</td><td>1.65</td><td>1.86</td><td>2.89</td><td></td><td>4.28</td></tr><tr><td>Tb</td><td>0.34</td><td>0.24(1)</td><td>0.36</td><td>0.30</td><td></td><td>0.93</td><td>0.792</td></tr><tr><td>Dy</td><td>2.61</td><td>1.73(8)</td><td>2.57</td><td>1.66</td><td>2.64</td><td>6.1</td><td>5.04</td></tr><tr><td>Ho</td><td>0.58</td><td>0.402(4)</td><td>0.62</td><td>0.32</td><td></td><td>1.32</td><td>1.03</td></tr><tr><td>Er</td><td>1.83</td><td>1.24(1)</td><td>1.91</td><td>0.84</td><td>1.44</td><td></td><td>2.76</td></tr><tr><td>Tm</td><td>0.32</td><td>0.19(1)</td><td>0.289</td><td>0.13</td><td></td><td>0.51</td><td></td></tr><tr><td>Yb</td><td>1.95</td><td>1.28(5)</td><td>1.99</td><td>0.80</td><td>1.28</td><td>3.5</td><td>2.35</td></tr><tr><td>Lu</td><td>0.30</td><td>0.216(7)</td><td>0.30</td><td>0.13</td><td>0.18</td><td>0.54</td><td>0.331</td></tr><tr><td>Hf</td><td>0.49</td><td>0.339(9)</td><td>0.30</td><td>0.66</td><td>1.62</td><td>3.4</td><td>2.19</td></tr><tr><td>Ta</td><td>n.r.<sup>1</sup></td><td>b.d.</td><td></td><td>0.21</td><td></td><td></td><td>0.28</td></tr><tr><td>W</td><td>n.r.<sup>1</sup></td><td>0.0157(3)</td><td></td><td></td><td></td><td></td><td>0.66</td></tr><tr><td>Re</td><td>b.d.</td><td>0.0022(3)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Ir</td><td>b.d.</td><td>0.0016(2)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pt</td><td>b.d.</td><td>0.0044(6)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Tl</td><td>b.d.</td><td>0.0584(9)</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pb</td><td>2.06</td><td>1.26(2)</td><td>0.93</td><td>0.87</td><td></td><td></td><td>1.12</td></tr><tr><td>Th</td><td>b.d.</td><td>0.0083(3)</td><td></td><td>0.43</td><td>1.44</td><td></td><td>0.57</td></tr><tr><td>U</td><td>b.d.</td><td>0.0452(9)</td><td></td><td>0.11</td><td>0.38</td><td></td><td>0.12</td></tr></table>

1179

1180 b.d. = below detection

1181 <sup>a</sup>This study - represents average of two replicate analyses, except for Hf, in which one of the replicate

1182 data was excluded; major and minor elements assuming SiO2 = 48 wt% and normalized to 100%
1183 $^b$Based on defocused beam analysis of shock melt vein (n = 49); numbers in parentheses are 1σ standard

1184 deviation of mean

1185 $^c$After Irving et al. (2013), in which the same sample preparation methods and ICP-MS instrumentation

1186 were used

1187 $^d$Intercumulus melt composition of Hammer (2009)

1188 $^e$After Day et al. (2006)

1189 $^f$Mesostasis composition of Jambon et al. (2016)

1190 $^g$After Warren et al. (1999)

1191 $^h$After Rubin et al. (2000)

1192 $^i$After Jambon et al. (2002)

1193 $^j$n.r. = not reported, on suspicion of contamination

1194

1195
<table><caption>**Table 4.** Sm-Nd isotopic data for NWA 8159</caption><thead><tr><th rowspan="2">Sample</th><th rowspan="2">Weight<br>(mg)</th><th rowspan="2">Sm<br>(ppm)</th><th rowspan="2">Nd<br>(ppm)</th><th colspan="2"><sup>147</sup>Sm/<sup>144</sup>Nd</th><th colspan="2"><sup>143</sup>Nd/<sup>144</sup>Nd</th><th colspan="6"><sup>142</sup>Nd/<sup>144</sup>Nd</th></tr><tr><th>present</th><th>±2se</th><th>present</th><th>±2se</th><th>±2se</th><th>ε<sup>142</sup>Nd</th><th>±2se</th><th>ε<sup>143</sup>Nd</th><th>±2se</th></tr></thead><tbody><tr><td colspan="14"><strong>NASA-JSC<sup>a</sup></strong></td></tr><tr><td>WR</td><td>3.72</td><td>0.841</td><td>1.334</td><td>0.3813</td><td>0.0004</td><td>0.516477</td><td>0.000020</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>WR Residue I</td><td>3.60</td><td>0.503</td><td>0.805</td><td>0.3778</td><td>0.0004</td><td>0.516982</td><td>0.000017</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>WR Leach I</td><td>0.73</td><td>2.507</td><td>3.931</td><td>0.3857</td><td>0.0004</td><td>0.515969</td><td>0.000019</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Residue II</td><td>4.02</td><td>0.528</td><td>0.857</td><td>0.3727</td><td>0.0005</td><td>0.516924</td><td>0.000015</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Leach II</td><td>0.53</td><td>3.501</td><td>5.371</td><td>0.3942</td><td>0.0012</td><td>0.515909</td><td>0.000031</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Residue III</td><td>3.98</td><td>0.487</td><td>1.144</td><td>0.2575</td><td>0.0003</td><td>0.515125</td><td>0.000015</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Leach III</td><td>0.64</td><td>3.405</td><td>6.587</td><td>0.3127</td><td>0.0003</td><td>0.515272</td><td>0.000015</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Residue IV</td><td>3.35</td><td>0.442</td><td>0.801</td><td>0.3335</td><td>0.0004</td><td>0.516174</td><td>0.000020</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Leach IV</td><td>0.80</td><td>2.590</td><td>4.025</td><td>0.3891</td><td>0.0215</td><td>0.515956</td><td>0.000042</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Residue V</td><td>3.78</td><td>0.467</td><td>1.109</td><td>0.2548</td><td>0.0004</td><td>0.515228</td><td>0.000050</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FR Residue VI</td><td>4.61</td><td>0.423</td><td>0.689</td><td>0.3717</td><td>0.0005</td><td>0.516959</td><td>0.000015</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td colspan="14"><strong>LLNL<sup>b</sup></strong></td></tr><tr><td>WR-1</td><td>557</td><td>0.852</td><td>1.396</td><td>0.3689</td><td>0.0004</td><td>0.517153</td><td>0.000002</td><td>1.141939</td><td>0.000004</td><td>0.88</td><td>0.04</td><td>99.62</td><td>0.03</td></tr><tr><td>WR-2</td><td>999</td><td>1.034</td><td>1.968</td><td>0.3178</td><td>0.0003</td><td>0.516113</td><td>0.000001</td><td>1.141926</td><td>0.000002</td><td>0.77</td><td>0.02</td><td>78.34</td><td>0.01</td></tr><tr><td>JNdi (N=12)</td><td></td><td></td><td></td><td></td><td></td><td>0.511201</td><td>0.000006</td><td>1.141838</td><td>0.000008</td><td></td><td></td><td></td><td></td></tr></tbody></table>

<footnote><sup>a</sup>NASA-JSC data normalized to <sup>146</sup>Nd/<sup>144</sup>Nd = 0.724487 and adjusted to 143Nd/144Nd = 0.511138 for the Ames Nd Standard (Wasserburg et al., 1981). Nd analyzed as NdO<sup>+</sup> on Triton TIMS at NASA-JSC. Sample signal size bracketed by those run for Ames standard (0.2 to 1.4 V for <sup>144</sup>NdO<sup>+</sup>). Error limits account for uncertainty of each measurement and reproducibility of standard over a 3-year period (±19 ppm 2 SD, n=58).
WR is bulk “whole” rock and FR is rock fragment of NWA 8159. The NASA-JSC WR sample listed first is unleached.</footnote>
<footnote>bLLNL data normalized to ¹⁴⁶Nd / ¹⁴⁴Nd = 0.7219. Long-term average of JNdi represents N=68 mass spectrometer runs and uncertainty is 2x standard deviation of the 68 runs. Uncertainty on samples are 2 standard errors of 540 (Wr-1) and 2700 (Wr-2) cycles of mass spectrometer data.</footnote>
<table><caption>**Table 5.** Cr and W isotopic composition of NWA 8159 and other martian (SNC) meteorites</caption><thead><tr><th>Sample</th><th>Type</th><th>ε<sup>53</sup>Cr</th><th>ε<sup>54</sup>Cr</th><th>ε<sup>182</sup>W</th></tr></thead><tbody><tr><td>NWA 8159</td><td>augite basalt</td><td>0.25 ± 0.03</td><td>-0.04 ± 0.07</td><td>1.13 ± 0.10<sup>a</sup></td></tr><tr><td>Shergotty</td><td>shergottite</td><td>0.22 ± 0.08<sup>b</sup></td><td>-0.18 ± 0.17<sup>b</sup></td><td>0.27 ± 0.13<sup>c</sup></td></tr><tr><td>Nakhla</td><td>nakhlite</td><td>0.21 ± 0.06<sup>b</sup></td><td>-0.4 ± 0.08<sup>b</sup></td><td>2.99 ± 0.20<sup>d</sup></td></tr><tr><td>Chassigny</td><td>chassignite</td><td>0.17 ± 0.07<sup>b</sup></td><td>-0.20 ± 0.05<sup>b</sup></td><td>2.14 ± 0.62<sup>d</sup></td></tr></tbody></table>

<table><caption>**Table 6.** Cosmic ray exposure ages for NWA 8159</caption><thead><tr><th rowspan="2">Method</th><th colspan="2">Shock melt glass</th><th colspan="2">Whole rock</th></tr><tr><th>Age (Ma)</th><th>± (1σ)</th><th>Age (Ma)</th><th>± (1σ)</th></tr></thead><tbody><tr><td><sup>38</sup>Ar<sup>a</sup></td><td>1.5</td><td>0.1</td><td>-</td><td>-</td></tr><tr><td><sup>21</sup>Ne</td><td>1.4</td><td>0.1</td><td>0.9</td><td>0.1</td></tr><tr><td><sup>3</sup>He</td><td>1.4</td><td>0.1</td><td>1.2</td><td>0.1</td></tr><tr><td><sup>10</sup>Be, <sup>26</sup>Al, <sup>36</sup>Cl</td><td></td><td></td><td>1.0</td><td>0.1</td></tr></tbody></table>

<table><caption>**Table 7.** Oxygen isotopic composition of NWA 8159</caption><thead><tr><th>Sample size (mg)</th><th>δ<sup>17</sup>O'</th><th>δ<sup>18</sup>O'</th><th>Δ<sup>17</sup>O'</th></tr></thead><tbody><tr><td>1.2</td><td>2.406</td><td>4.089</td><td>0.247</td></tr><tr><td>1.2</td><td>2.405</td><td>3.947</td><td>0.321</td></tr><tr><td>1.8</td><td>2.093</td><td>3.328</td><td>0.336</td></tr><tr><td>2.0</td><td>2.532</td><td>4.197</td><td>0.316</td></tr><tr><td>1.0</td><td>2.329</td><td>3.880</td><td>0.280</td></tr><tr><td>average</td><td>2.353</td><td>3.888</td><td>0.300</td></tr></tbody></table>

<footnote><sup>a</sup>This study, Münster data</footnote>

<footnote><sup>b</sup>Trinquier et al. (2007), Trinquier et al. (2008)</footnote>

<footnote><sup>c</sup>Foley et al. (2005)</footnote>

<footnote><sup>d</sup>Lee and Halliday (1997)</footnote>

<footnote><sup>a</sup><sup>38</sup>Ar age calculated assuming (<sup>38</sup>Ar/<sup>36</sup>Ar)<sub>trap</sub> is 0.188</footnote>
<table><caption>**Table 8.** Comparison of the augite-rich shergottites NWA 8159 and NWA 7635</caption><thead><tr><th>Observation type</th><th>NWA 8159</th><th>NWA 7635<sup>a</sup></th></tr></thead><tbody><tr><td>Texture and mineralogy<sup>b</sup></td><td>intergranular; ol+mt<br>phenocrysts with rims<br>replaced by opx+mt in<br>plag+px groundmass<br>lacking sf</td><td>porphyritic;<br>pl+aug+ol+mt<br>phenocrysts in<br>plag+px+sf<br>groundmass<br>not reported</td></tr><tr><td>Peak shock pressure</td><td>15-23 GPa</td><td></td></tr><tr><td>Crystallization Age</td><td>2370 ± 250 Ma</td><td>2403 ± 140 Ma</td></tr><tr><td>Ejection Age</td><td>1.2 ± 0.1 Ma</td><td>1.0 ± 0.1 Ma</td></tr><tr><td>Terrestrial Age</td><td>230 ± 60 ka</td><td>2.3 ± 1.3 ka</td></tr><tr><td>ε<sup>143</sup>Nd</td><td>46 ± 11</td><td>66.3 ± 0.8<sup>c</sup></td></tr><tr><td>ε<sup>142</sup>Nd</td><td>0.77 ± 0.02</td><td>0.92 ± 0.08</td></tr><tr><td>ε<sup>182</sup>W</td><td>1.13 ± 0.10</td><td>1.80 ± 0.13<sup>d</sup></td></tr><tr><td>La/Yb</td><td>0.08</td><td>0.06</td></tr><tr><td>Dy/Lu</td><td>0.87</td><td>0.86</td></tr></tbody></table>

<footnote><sup>a</sup>Data/observations from Lapen et al. (2017) unless otherwise indicated</footnote>

<footnote><sup>b</sup>Mineral abbreviations: ol=olivine; mt=magnetite; opx=orthopyroxene; pl=plagioclase; aug=augite; px=pyroxene; sf=sulfide</footnote>

<footnote><sup>c</sup>Calculated from the <sup>143</sup>Nd/<sup>144</sup>Nd value reported by Lapen et al. (2017)</footnote>

<footnote><sup>d</sup>After Kruijer et al. (2017)</footnote>
**Figure Captions**

**Figure 1.** Photograph of the main mass of NWA 8159 (a) after cutting to reveal a green-grey, fine-grained interior texture cross-cut by several distinct shock veins (b). Cube is 1 cm.

**Figure 2.** Textures observed in NWA 8159. (a) backscattered electron (BSE) image showing typical texture of intergrowth augite (medium grey) and plagioclase (dark grey) with olivine (light grey) and magnetite (white) phenocrysts. (b) BSE image of a magnetite phenocryst with a core consisting of a symplectic intergrowth of magnetite and quartz (inset image). c,d) BSE images showing olivine phenocrysts with assemblages of orthopyroxene and magnetite at their margins. In (c) the olivine cores are affected by secondary alteration, visible as mottled areas and veins of Fe-oxides which continue into the augite-plagioclase groundmass. In (d) details of the orthopyroxene-magnetite (opx, mt) assemblage replacing rims of olivine (ol) are provided. The presumed original outline of the olivine phenocryst is shown. Pyroxene in the groundmass is augite (cpx). (e) Fe X-ray map of the same area shown in (d); scale bar same as in (d).

**Figure 3.** Compositions of pyroxene and olivine in NWA 8159. Also shown are representative compositions of pyroxene in the nakhlites (MIL 03346 and Yamato (Y) 000593, after Treiman (2005)), and shergottites (Z = Zagami, after McCoy et al. (1999); LA = Los Angeles, after Warren et al. (2004), with high-Wo metastable compositions excluded), and NWA 7635 (from Lapen et al. 2017). Shown by dashed line is the approximate boundary of the 'forbidden region' of pyroxene crystallization at 1 bar pressure, after Lindsley (1983).

**Figure 4.** Magnetic susceptibility as a function of temperature for NWA 8159. The experiment starts at low temperature and includes one temperature cycle between 450 and 20 °C to search for maghemite; red and blue lines represent warming and cooling phases respectively.

**Figure 5.** Offset and displacement of igneous minerals along the margins of a thin shock vein. Aug = augite, plag = amorphous plagioclase, terr alteration = cracks and fractures cutting across the igneous portion of the rock as well as shock features, infilling by secondary alteration products (calcite).

**Figure 6.** Representative Raman spectra for crystalline plagioclase in the host rock, tissintite and plagioclase glass associated with shock veins.

**Figure 7.** BSE images of plagioclase and its transformation products in NWA 8159. (a) Smooth, unfractured amorphous plagioclase with oscillatory igneous zoning profiles preserved (contrasting greyscale in BSE image). (b) The amorphization of plagioclase is clearly associated with shock veins, with amorphous plagioclase occurring in those portions closest to the shock vein and highly fractured crystalline plagioclase in those portions of the grains distant from the vein. (c) In direct contact with the shock melt plagioclase has transformed to tissintite, the Ca-analogue of jadeite.

**Figure 8.** Chondrite-normalized REE patterns of NWA 8159 and selected martian meteorite lithologies for comparison. Data from sources listed in Table 3, with the exception of Yamato
(Y)980459, for which data are from Shirai and Ebihara (2004), and Tissint, for which data are
from Aoudjehane et al. (2012).

**Figure 9.** Sm-Nd isochron diagram for NWA 8159. The age of 2.37 ±0.25 Ga (2σ) is defined by measurements of five “residual” leached rock fragments containing maskelynite+plagioclase+pyroxene in varying proportions and a minimally leached whole rock sample (WR). See text for additional details.

**Figure 10.** A. $^{40}$Ar/$^{39}$Ar isochron diagram obtained from NWA 8159 shock melt glass. B. $^{40}$Ar/$^{39}$Ar age spectrum obtained from a whole-rock fragment of NWA 8159.

**Figure 11.** Oxygen isotopic compositions of NWA 8159 samples, with data from other SNCs (Agee et al., 2013, and references therein) for comparison.

**Figure 12.** (a) Comparison of the Cr and O isotopic composition of NWA 8159 with other martian (SNC) meteorites and other achondrite and chondrite groups. (b) Magnified view of the martian meteorite Cr-O field denoted by solid square in panel (a). (c) Comparison of the $\varepsilon^{53}\text{Cr}$ and $\varepsilon^{4}\text{Cr}$ composition for NWA 8159 and other martian meteorites; literature data from Ueda et al. (2006), Shukolyukov and Lugmair (2006), Trinquier et al. (2007), Trinquier et al. (2008), Jenniskens et al. (2012), Popova et al. (2013), and Jenniskens et al. (2014). Literature data for O isotopic composition are from Clayton et al. (1991; 1984), Clayton and Mayeda (1996; 1999), Franchi et al. (1999), Scott et al. (2009), Jenniskens et al. (2012), Popova et al. (2013), and Jenniskens et al. (2014).

**Figure 13.** Oxygen fugacity (in log units, relative to the QFM buffer) of select shergottites and NWA 8159 vs. whole-rock La/Yb (C1-normalized). NWA 8159 data point shows estimated liquidus $fO_2$ (see text for details). Oxygen fugacity data sources for shergottites from Herd et al. (2013) for Tissint; Peslier et al. (2010) for LAR 06319; Walton et al. (2012) for average poikilitic (“Ave. Poik.”; a.k.a. lherzolitic), Gross et al. (2011) for NWA 5789 (5789), Herd (2004) for Yamato 980459 (Y98), Herd (2003) for Sayh al Uhaymir 005 (SaU) and Gross et al. (2013) for NWA 6234; other shergottite data, including Zagami (Z), Shergotty (S), Los Angeles (LA), EETA79001A and B (EA, EB) from Herd et al. (2001). La/Yb data from various sources, found in Meyer (2016).
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Herd et al. Figure 12

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