В€Їpetrology of the Enriched Poikilitic Shergottite Northwest Africa 10169

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Geochimica et Cosmochimica Acta 266 (2019) 435–462 www.elsevier.com/locate/gca

Petrology of the enriched poikilitic shergottite Northwest Africa 10169: Insight into the martian interior

Logan M. Combs a, Arya Udry a,⇑, Geoﬀrey H. Howarth b, Minako Righter c, Thomas J. Lapen c, Juliane Gross d,e, Daniel K. Ross f, Rachel R. Rahib a, James M.D. Day g

a Department of Geoscience, University of Nevada Las Vegas, Las Vegas, NV 89154, USA b Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa c Department of Earth and Atmospheric Sciences, University of Houston, Houston, TX 77204, USA d Department of Earth and Planetary Sciences, Rutgers University, Piscataway, NJ 08854, USA e Department of Earth and Planetary Sciences, The American Museum of Natural History, New York, NY 10024, USA f Jacobs-JETS/NASA Johnson Space Center, Houston, TX 77058, USA g Scripps Institution of Oceanography, University of California San Diego, La Jolla, CA 92093, USA

Received 1 June 2018; accepted in revised form 1 July 2019; available online 9 July 2019

Abstract

The martian meteorite Northwest Africa (NWA) 10169 is classiﬁed as a member of the geochemically enriched poikilitic shergottites, based on mineral composition, Lu-Hf and Sm-Nd isotope systematics, and rare earth element (REE) concentrations. Similar to other enriched and intermediate poikilitic shergottites, NWA 10169 is a cumulate rock that exhibits a bimodal texture characterized by large pyroxene oikocrysts (poikilitic texture) surrounded by olivine-rich interstitial material (non-poikilitic texture). Olivine chadacrysts and pyroxene oikocrysts have higher Mg#s (molar Mg/Mg + Fe) than those in the interstitial areas, suggesting that the poikilitic texture represents early-stage crystallization and accumulation, as opposed to late-stage non-poikilitic (i.e., interstitial material) crystallization. Calculated oxygen fugacity values are more reduced (FMQ À2.3 ± 0.2) within the poikilitic regions, and more oxidized (FMQ À1.1 ± 0.1) within the interstitial areas, likely representing auto-oxidation and degassing during magma crystallization. Calculated parental melt compositions using olivine- hosted melt inclusions display a dichotomy between K-poor and K-rich melts, thus possibly indicating mixing of parental melt with K-rich melt. The 176Lu-176Hf crystallization age for NWA 10169 is 167 ± 31 Ma, consistent with the ages reported for other enriched shergottites. Based on the isochron initial 176Hf/177Hf value, the modeled source 176Lu/177Hf composition for NWA 10169 is 0.02748 ± 0.00037, identical within uncertainty to the source compositions of the enriched shergottites Sher- gotty, Zagami, LAR 06319, NWA 4468, and Roberts Massif (RBT) 04262, suggesting a shared, long-lived geochemical source, and distinct from the source of other enriched shergottites Los Angeles, NWA 856, and NWA 7320. This study reveals that at least two sources are responsible for the enriched shergottites, and that the martian mantle is more heterogeneous than previously thought. Additionally, the enriched shergottites, which share a source with NWA 10169, have consistent crystallization ages and magmatic histories, indicating that a common magmatic system on Mars is likely responsible for the formation of this group. Ó 2019 Elsevier Ltd. All rights reserved.

Keywords: Martian meteorites; Shergottites; Martian magmatism

⇑ Corresponding author at: Department of Geoscience, University of Nevada Las Vegas, Las Vegas, NV 89154, USA. E-mail address: [email protected] (A. Udry). https://doi.org/10.1016/j.gca.2019.07.001 0016-7037/Ó 2019 Elsevier Ltd. All rights reserved. 436 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

1. INTRODUCTION intermediate between the depleted and enriched groups of shergottites (e.g., Debaille et al., 2008). Poikilitic shergot- Shergottites are the most abundant of the main three tites are represented by two of the three geochemical groups: groups of martian meteorites (shergottites, nakhlites, and the intermediate and the enriched (McSween et al., 1979; chassignites), accounting for 90% of known martian sam- Harvey et al., 1993; Mikouchi and Miyamoto, 1996; ples. Currently, martian meteorites represent the only phys- Gleason et al., 1997; Ikeda, 1997; Goodrich, 2003; Hsu ical samples of Mars available for direct study, and as such, et al., 2004; Gillet et al., 2005; Lin et al., 2005). Recent stud- these materials are crucial for advancing our understanding ies show that some enriched poikilitic shergottites have sim- of the geology and magmatic processes of Mars. Four ilar petrologic characteristics to the intermediate poikilitic groups can be identified within the shergottites based on shergottites, yet have distinct geochemical characteristics texture and mineralogy: The basaltic, olivine-phyric, poiki- derived from the enriched source reservoir (e.g., Usui litic, and gabbroic shergottites (e.g., McSween and et al., 2010; Jiang and Hsu, 2012; Lin et al., 2013; Treiman, 1998; Goodrich, 2002; Bridges and Warren, Howarth et al., 2014; Howarth et al., 2015). 2006; Filiberto et al., 2014; Udry et al., 2017; Filiberto Crystallization and cosmic ray exposure (CRE) ages et al., 2018). The basaltic and olivine-phyric shergottites suggest that there are two ejection age clusters within the are thought to represent hypabyssal to extrusive rocks, enriched poikilitic shergottites; one group related to the while the poikilitic and gabbroic shergottites are thought enriched olivine-phyric and basaltic shergottites (3 Ma), to represent intrusive rocks emplaced at depth within the and the other, related to the intermediate poikilitic shergot- martian crust (e.g., Goodrich, 2002; Mikouchi and tites (4.5 Ma) (Nyquist et al., 2001; Christen et al., 2005; Kurihara, 2008; Usui et al., 2010; Howarth et al., 2014; Borg et al., 2008; Lapen et al., 2008; Nishiizumi and Howarth et al., 2015; Udry et al., 2017; Filiberto et al., Caffee, 2010; Usui et al., 2010; Niihara, 2011; Wieler 2018). Poikilitic shergottites, formerly known as ‘‘lher- et al., 2016). Robert Massif (RBT) 04261/2 and Northwest zolitic” shergottites, comprise >20% of the martian mete- Africa (NWA) 4468, the only enriched poikilitic shergot- orite collection; to date, 27 have been recovered. They are tites to have both their CRE and crystallization ages deter- characterized by a bimodal texture, consisting of a poikilitic mined, yielded crystallization ages of 225 ± 21 Ma and 150 texture and an interstitial non-poikilitic texture. The poiki- ± 29 Ma, respectively, and ejection ages of 3Ma(Borg litic shergottites have been formed through polybaric crys- et al., 2008; Lapen et al., 2008; Nishiizumi and Caffee, tallization, as suggested by Howarth et al. (2014), involving 2010; Niihara, 2011; Wieler et al., 2016). These ejection formation and accumulation of pyroxene oikocrysts within and crystallization ages, along with the crystallization ages a magma staging chamber, and subsequent formation of of the intermediate poikilitic shergottites (180 Ma), over- the interstitial non-poikilitic texture following the rise of lap with the ejection and crystallization ages of the enriched the magma towards the surface. This unique texture and basaltic shergottites, of 3 Ma and 165 to 177 Ma, crystallization process amongst shergottites allow us to respectively (Nyquist et al., 2001; Christen et al., 2005; investigate their petrogenesis from deep crustal chambers Usui et al., 2010). Two other enriched poikilitic shergot- to late-stage evolution in the upper crust on Mars. tites, NWA 6342 and NWA 7397, yielded CRE ages of In addition to mineralogical groups, shergottites are 4.5 Ma, coinciding with the ejection ages of the intermedi- divided into three geochemical groups (enriched, intermediate poikilitic shergottite suite (4 to 5 Ma) (Nyquist et al., ate, and depleted), based on their relative rare earth element 2001; Christen et al., 2005; Wieler et al., 2016). Thus, (REE) abundances, and their Sr, Nd, and Hf isotope abun- although intermediate and enriched poikilitic shergottites dances (e.g., Bridges and Warren, 2006; Debaille et al., 2008; have similar textures, some enriched poikilitic shergottites McSween, 2015). These geochemical variations within the may be spatially related (as evidenced by similar CRE ejec- shergottite suite have been interpreted to reflect mixtures tion ages) to basaltic shergottites. of at least two different mantle source compositions (e.g., Here we present bulk and mineral major and trace ele- Borg et al., 2002; Borg et al., 2003; Brandon et al., 2012; ment compositions, bulk melt composition from melt inclu- Lapen et al., 2010; Lapen et al., 2017). Generally, it is sions, a Lu-Hf age, and Sm-Nd and Lu-Hf isotopic believed that their source reservoirs formed early in Mars’ compositions for the poikilitic shergottite NWA 10169. history (Borg et al., 2016), within 100 million years after This shergottite is a 22 g meteorite that was found in the solar system formation (Debaille et al., 2007). One source Laayoune region of Western Sahara in January 2015, and reservoir is relatively reduced and depleted in incompatible classified as a poikilitic shergottite in July 2015 (Hewins trace elements, and is characterized by a relative depletion and Zanda, 2015). The geochemistry, mineralogy, and tex- in light rare earth elements (LREE) (low La/Yb); low Rb/ ture of this meteorite indicate that NWA 10169 is a member Sr ratios resulting in relatively low 87Sr/86Sr initial isotope of the geochemically enriched poikilitic shergottites. By ratios; and high Sm/Nd and Lu/Hf ratios that result in rel- conducting a comprehensive petrologic study of NWA atively high 143Nd/144Nd and 176Hf/177Hf initial isotope 10169, we aim to gain a further understanding of the source ratios (high e143Nd and e176Hf values). The shergottites of the enriched shergottites and the formation of poikilitic derived from the enriched source reservoir exhibit the oppo- shergottites. This study has implications for constraining site geochemical characteristics, whereas the intermediate mantle reservoirs on Mars and the petrological and geo- shergottites are derived from source compositions that are chemical link between different shergottite groups. L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 437

2. ANALYTICAL METHODS of 20 nA, and a take-off angle of 34.1°. X-ray maps were acquired using an SDD-EDS detector, with Thermo 2.1. Electron microprobe analyses Electron software. Melt inclusions were digitized in that software, and all X-ray counts from the digitized area were 2.1.1. In situ major and minor element analyses gathered into a single X-ray spectrum for each melt inclu- In situ and textural analyses were performed on two thin sion. The resulting extracted X-ray spectrum was quantified sections (T1 and T2), cut from a chip of NWA 10169 by reference to natural and synthetic glass and mineral (Fig. 1). Major and minor element compositions of the min- standards, with Phi-Rho-Z matrix corrections. erals of NWA 10169 were measured using the JEOL JXA- Using the JEOL JXA-8900 EMP at UNLV, in situ WDS 8900 Superprobe electron microprobe (EMP) analyzer at analyses of major and minor elements were taken for the the University of Nevada Las Vegas (UNLV). Analyses pyroxene, high-Si glass, and Fe-Ti-Cr oxides within the were performed with an accelerating voltage of 15 kV, a melt inclusions, as well as the surrounding host olivine. beam current of 20 nA, a beam diameter of 5 mm, and stan- These analyses were conducted with an accelerating voltage dard PAP corrections for all phases but maskelynite and of 15 kV, a beam current of 10 nA (20 nA for host olivine), glass. The beam current was decreased to 10 nA for maske- a spot size of 1 mm(2mm for host olivine), and standard lynite and glass analyses to prevent volatilization of Na. PRZ corrections. Peak count times were 30 s for all ele- Peak count times were 10 s for Na and F; and 30 s for ments measured, excluding Na (10 s). The phosphates, Ca, Mg, Al, Fe, Si, Mn, Ti, P, Cr, Ni, S, and K. Both nat- ilmenite, and sulfides present within the melt inclusions ural and synthetic standards were used for calibration. Ele- were too small (<1 mm) to obtain reliable analyses. mental X-ray maps of both thin sections were obtained for Elemental X-ray maps of P, Cr, and Al within the host Fe, Mg, Ca, and Si, with an accelerating voltage of 15 kV, a olivine were taken using the JEOL JXA-8200 Superprobe at beam current of 25 nA, and a pixel dwell time of 9 ms. Rutgers University with an accelerating voltage of 15 kV, Modal abundances of mineral phases were calculated from beam current of 300 nA, pixel dwell time of 500 ms, and a the resulting X-ray maps using pixel counting techniques step size of 2 mm. Phosphorus ka was measured on three within the ImageJ software. spectrometers simultaneously [PET, TAP, and large PET (LPET)]. 2.1.2. Melt inclusion analyses Olivine-hosted polymineralic melt inclusions in both 2.2. Bulk composition calculation: poikilitic and non-poikilitic textures were analyzed in order to calculate the primary trapped liquid composition. Six To derive a bulk chemical composition, the modal abun- melt inclusions were analyzed in both thin sections of dances of mineral phases were calculated and averaged NWA 10169. Prior to EMP analyses, backscatter electron from Fe, Mg, Ca, and Si X-ray maps of both thin sections (BSE) images and qualitative energy-dispersive spectral of NWA 10169 using the software ImageJ, and following a (EDS) analyses of each melt inclusion were taken on the method similar to Maloy and Treiman (2007). It was JEOL JSM-5610 scanning electron microscope at UNLV. assumed that the analyzed area proportions (and thus the At the ARES laboratory of NASA Johnson Space Center estimated area modal abundances) are equal to volume pro- (JSC), using a JEOL 8530F field emission electron probe, portions. Volume proportions of minerals were converted elemental X-ray maps of Fe, Mg, Ca, Si, Ti, Al, Cr, K, to mass proportions using mineral densities interpolated Na, O, P, Mn, S, and Cl were measured for each of the melt from literature values similar to the method presented in inclusions with an accelerating voltage of 15 kV, a current Gross and Treiman (2011). The bulk composition was then

Fig. 1. Plane polarized light photomicrographs of (a) slide 1 and (b) slide 2 of NWA 10169. Yellow dashed lines outline pyroxene oikocrysts. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.) 438 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 calculated as the sum of average mineral compositions 2.5. 176Lu/177Hf and 147Sm/143Nd isotopic analyses weighted by their mass proportions. A separate 1 g bulk sample of NWA 10169 was crushed 2.3. Laser Ablation-Inductively Coupled Plasma-Mass in an agate mortar used exclusively for martian meteorites. Spectrometry (LA-ICP-MS) The resulting powder was sieved into >100 mm, 75–100 mm, 45–75 mm, and < 45 mm size fractions. Pyroxene and In situ trace element compositions of the major phases of maskelynite cuts were hand-picked from the >100 mm size NWA 10169 were analyzed using the New Wave Research fraction. Maskelynite was also separated from the 75– UP213 (213 nm) solid-state laser ablation system coupled 100 mm size fraction using lithium metatungstate to a ThermoScientific iCAP Q ICP-MS at the Scripps Insti- (2.75 g/cc). Fe-Cr-Ti oxides were isolated from tute of Oceanography (SIO) at the University of California, both >100 mm and 75–100 mm size fractions using thallium San Diego (UCSD). Low-Ca pyroxene, high-Ca pyroxene, malonate (4.36 g/cc). For both heavy liquid separations, olivine, maskelynite, and merrillite were measured using a powder and liquid were added to a 15 mL conical test tube 100 mm laser spot size for the silicate analyses and 50 mm and spun in a centrifuge at 2500 rpm for 30 minutes. The for merrillite analyses. Material was ablated by the laser heavy separate at the bottom of the test tube was frozen beam in a helium atmosphere, then transported to the with liquid nitrogen and the light fraction was poured off. ICP-MS with an argon carrier gas. The laser was main- The separates were rinsed 10 times with pure water, and tained at a frequency of 5 Hz, and a fluence of 3 J/cm2 washed in pure water, in an ultrasonic bath for four 15- throughout all analyses. The measurements were performed minute sessions. The final aliquots: <45 mm bulk with a background measurement time of 20 s, an ablation (100 mg), >100 mm bulk minus mineral separates (100 mg, time of 40 s, and a washout time of 120 s. NIST 610, corresponding to the bulk powder after picking out the BHVO-2g, and BCR-2g were used as reference standards, other phases), pyroxene (100 mg), maskelynite (15 mg), and the order of points analyzed were as follows: 6 standard and Fe-Cr-Ti oxides (<1 mg), were taken to the University points (2 for each standard), 8 unknowns (NWA 10169), of Houston for chemical processing and mass spectrometry. and 6 additional standard points. Blanks were measured Mineral fractions were spiked with a 176Lu-178Hf tracer during the laser warm-up time. Isotopes analyzed were and initially digested in 4 mL of concentrated HF and 7 24 29 44 45 48 51 52 59 60 Li, Mg, Si, Ca, Sc, Ti, V, Cr, Co, Ni, 0.5 mL of concentrated HNO3 at 110 °C for 24 h. The stan- 66Zn, 71Ga, 85Rb, 88Sr, 89Y, 90Zr, 93Nb, 133Cs, 137Ba, dards BHVO-2, BCR-2, and a blank, underwent the same 139La, 140Ce, 141Pr, 146Nd, 147Sm, 153Eu, 157Gd, 159Tb, procedures as S1-5. A whole rock aliquot was spiked with 163Dy, 165Ho, 166Er, 169Tm, 172Yb, 175Lu, 178Hf, 181Ta, a 149Sm-150Nd isotope tracer along with USGS basalt stan- 204Pb, 206Pb, 207Pb, 232Th, and 238U. Laser traces were dard BCR-2 and sample digestion follows the same proce- examined to ensure no major penetration into underly- dure applied to the 5 aliquots. The samples were dried ing phases using trace element signal profiles. Reproducibil- down and completely digested with a Milestone Ultra- ity for unknown standard measurements of BHVO-2 WAVE microwave digestion system in 2 mL concentrated (n = 12) and BCR-2 (n = 12) were generally better HNO3, 1 mL concentrated HF, and 1 mL 8 M HCl. The than 6% for all elements. Accuracy is generally within samples were then dried down and dissolved once in 8 M 1SD of recommended values. Data reduction was HNO3, and twice more in 8 M HCl to ensure complete dis- performed using an in-house Excel spreadsheet developed solution of any salts. at SIO. Chemical separations of Lu and Hf follow the procedures of Mu¨nker et al. (2001) and Lapen et al. (2004). 2.4. Inductively Coupled Plasma-Mass Spectrometry (ICP- Chemical separations of Sm and Nd follow the procedures MS) outlined in Lapen et al. (2004, 2005, 2017). The 176Lu/177Hf, 176Hf/177Hf, 147Sm/144Nd, and 143Nd/144Nd A 0.5 g chip of bulk NWA 10169, which included both isotope ratios were analyzed using a Nu Instruments poikilitic and non-poikilitic textures, was crushed in an NuPlasma II MC-ICP-MS at the University of Houston, agate mortar used exclusively for martian meteorites. The following the spike subtraction and internal fractionation powder was digested (along with standards BHVO-2, correction procedures of Lapen et al. (2004). Analyses of AGV-1, and BCR-2) in 4 mL of concentrated HF and Hf standards, blanks, and aliquots were performed at 1 mL of concentrated HNO3 for over 72 hours at approxi- 40 ppb (15 ppb for pyroxene, maskelynite, and oxide ali- mately 150 °C. The NWA 10169 solution and the standards quots), and introduced to the ICP-MS through an Aridus were analyzed for bulk major, minor, and trace element II desolvating nebulizer fitted with a 50 ml/min self- compositions, using a ThermoScientific iCAP Q ICP-MS, aspirating nebulizer. Analyses of UH JMC 475 Hf were used exclusively for analyses of solutions, at SIO. The anal- performed, bracketing each analysis of unknown, to correct yses were standardized against BHVO-2. Reproducibility for instrument fractionation. Additionally, analyses of for unknown standard measurements of BHVO-2 (n = 3) USGS JMC 475 Hf and UH AMES Hf were performed and BCR-2 (n = 3) were better than 3% for elements, except to monitor data quality. The average measured 176Hf/177Hf B, Zn, Zr, Pb at better than 5%, and Sn and W at better ratios for USGS JMC-475 Hf, BHVO-2, and BCR-2 stan- than 18%. Accuracy is generally within 1SD of recom- dards are 0.282160 ± 17 (n = 6), 0.283100 ± 10 (n = 3), mended values. This analytical technique follows the proce- and 0.282869 ± 18 (n = 3), respectively (all errors are 2r). dure outlined in Day et al. (2015). Analyses of Lu standards, blanks, and aliquots were L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 439 performed using the methods outlined in Lapen et al. (2004) (approximately 500 mm) high-Ca pyroxene rims (Fig. 2). and Vervoort et al. (2004). External reproducibility for Olivine chadacrysts range in diameter from 100 to 800 mm 176Lu/177Hf and 176Hf/177Hf ratios are 0.2% and 0.005%, and are rounded and anhedral (Fig. 3). Rare olivine chada- respectively. The Sm-Nd isotope data was analyzed at crysts contain polymineralic melt inclusions, which are gen- 80 ppb and JNdi Nd was used to correct for instrumental erally between 100 mm and 400 mm in diameter, and are mass fractionation. Note that Hf and Nd concentrations composed of low and high-Ca pyroxene, Si-rich glass, sul- were chosen based on those yielded from the sample mate- fides, Fe-Cr-Ti oxides, ilmenite, apatite, and merrillite rials. USGS standard BCR-2 was analyzed and yielded a (Fig. 4). Fe-Cr-Ti oxide chadacrysts are anhedral to subhe- 143Nd/144Nd ratio of 0.512639 ± 15 (n = 3). Long-term dral and range from 25 to 100 mm in diameter. procedural blanks are <50 pg, <15 pg, <40 pg, <17 pg, for The non-poikilitic texture is interstitial to the pyroxene Nd, Sm, Hf, and Lu, respectively. oikocrysts (Figs. 1 and 2). In contrast to pyroxene oikocrysts, pyroxenes in the non-poikilitic texture have a less 3. RESULTS regular distribution of low-Ca and high-Ca pyroxenes. Non-poikilitic pyroxenes are subhedral and approximately 3.1. Petrography 500–1000 mm in length. Olivines within the non-poikilitic texture have average size of 800 mm, but can be as large Both thin sections of NWA 10169 exhibit distinct bimo- as 3 mm in length, and are generally larger in diameter than dal textures, comprised of a poikilitic and an interstitial the olivine chadacrysts. Non-poikilitic olivine crystals are non-poikilitic texture (Fig. 1). The poikilitic texture is char- subhedral to anhedral, and enclose a greater abundance acterized by pyroxene oikocrysts that both fully and par- of polymineralic melt inclusions relative to the poikilitic oli- tially enclose olivine and Fe-Cr-Ti oxide chadacrysts. The vine chadacrysts. These melt inclusions are similar in size, pyroxene oikocrysts are anhedral and significantly larger texture, and mineralogy to those observed within the olivine than any other minerals in the meteorite (up to 4 mm in chadacrysts, and are discussed further in Section 3.3. Non- diameter). The pyroxene oikocrysts have low-Ca cores poikilitic Fe-Cr-Ti oxides are generally partially or fully and display a sharp compositional boundary to thick enclosed by olivine, anhedral to subhedral, and are slightly

Fig. 2. Ca ka and Mg ka X-ray maps of both the poikilitic and non-poikilitic textures. (a) Mg X-ray map and (b) Ca X-ray map of a typical zoned (low-Ca to high-Ca) pyroxene oikocryst with olivine chadacrysts decreasing in forsterite content towards the oikocryst rims. (c) Mg X- ray map and (d) Ca X-ray map of a typical region of non-poikilitic texture. Lcp = low-Ca pyroxene; hcp = high-Ca pyroxene; ol = olivine; cr = chromite; ulv = ulvo¨spinel; msk = maskelynite; mer = merrillite; MI = melt inclusion. 440 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

Fig. 3. Backscatter electron (BSE) images of different regions of interest within NWA 10169. (a) An area of non-poikilitic texture showing euhedral olivine shapes and olivine-hosted melt inclusions. (b) Shock melt pocket in the non-poikilitic texture. (c) One of the smaller pyroxene oikocrysts with olivine chadacrysts ranging from subhedral to anhedral, showing characteristics of adsorption. (d) Group of late stage minerals in the non-poikilitic texture, with a large, highly fractured merrillite grain and an elongate ilmenite grain. Pyx = pyroxene; ol = olivine; cr = chromite; tr = troilite; ulv = ulvo¨spinel; msk = maskelynite; ilm = ilmenite; mer = merrillite; MI = melt inclusion. larger than the oxide chadacrysts (up to 300 mm in diame- 12–13 vol.% augite (high-Ca pyroxene), 13–17 vol.% ter). Maskelynite grains within the non-poikilitic textures maskelynite, 1–2 vol.% Fe-Cr-Ti oxides, and trace amounts are 500–1000 mm in diameter, and are interpreted to repre- of phosphates, pyrrhotite, and ilmenite. sent completely shocked and pseudomorphed plagioclase. Pyrrhotite and ilmenite are fine-grained (<100 mm), and in 3.2. Major, minor, and trace element mineral compositions the case of ilmenite, are elongated. Phosphates within NWA 10169 are predominantly merrillite with minor apa- 3.2.1. Pyroxene tite. They exist solely within the non-poikilitic components The pyroxene oikocrysts of the poikilitic textures have and are generally fine-grained (100–200 mm) elongate average core compositions of Wo6En68Fs26, becoming pro- crystal laths, however, there are a small number of larger gressively more calcium-rich (Wo7-15En60-67Fs24-26) with (400 mm) crystals. Apatite was too fine-grained for reli- increasing distance from the oikocryst cores (Fig. 5; Tables able EMP analyses. 1 and S1). The rims of the pyroxene oikocrysts have an Shock features in NWA 10169 include: (1) plagioclase average composition of Wo35En45Fs18, marking a sharp being completely converted to maskelynite; (2) darkening compositional boundary between the pigeonite cores and of olivine crystals within both thin sections, with some crys- augite rims (Figs. 2 and 5; Table 1). Pyroxenes within the tals darkened to near-opaque; and (3) one (shock) melt non-poikilitic textures, interstitial to the pyroxene pocket >400 mm in size, with abundant rounded <1 mm oikocrysts, have compositions ranging from pigeonite oxide and sulfide inclusions (Fig. 3b), located within the (Wo10En58Fs31) to augite (Wo33En47Fs20)(Fig. 5; Table 1). non-poikilitic texture of thin Section 1. Non-poikilitic pyroxene does not exhibit the sharp transi- Using the ImageJ software, the area modal abundance tion between low-Ca cores and high-Ca rims like the pyrox- ranges from both thin sections are estimated to be: 33– ene oikocrysts, but instead shows a patchy, random 40 vol.% olivine, 28–40 vol.% pigeonite (low-Ca pyroxene), distribution of pigeonite and augite (Fig. 2). The Mg#s L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 441

Fig. 4. Images of melt inclusions and their respective host olivine. (a) Schematic of a typical melt inclusion in NWA 10169, with labels for commonly observed phases; pyx = pyroxene; cr = chromite; ulv = ulvo¨spinel; modiﬁed from Peslier et al. (2010). (b) BSE image and (c) Mg,

Ca, Al composite X-ray map of MI1 within a poikilitic Fo66 olivine chadacryst, the most primitive olivine in NWA 10169 with a measurable melt inclusion. (d) Composite P X-ray map (P measured on PET, TAP, and LIF crystals) of MI1’s host olivine, showing the melt inclusion in the core of the host olivine, continuous zonation in the interior of the olivine, and discontinuous zonation near the rim. (e) BSE image and (f) a

Mg, Ca, Al composite X-ray map of MI6 within a non-poikilitic Fo59 olivine. (g) Composite P X-ray map of MI6’s host olivine, showing MI6 in the low-P core of its host olivine, and dislocations of the P zonation from shock. (h) BSE image and (i) a Mg, Ca, Al composite X-ray map of MI10 within a non-poikilitic Fo61 olivine. (j) Composite P X-ray map (from PET and TAP crystals) of MI10’s host olivine, showing MI10 trapped near the rim of the olivine, and multiple nucleation sites of olivine with low to high-P cores.

[molar 100 * Mg/(Mg + Fe)] of pyroxene oikocrysts are while Cr2O3 shows no relationship. Within the non- higher (>70) than those of the non-poikilitic pigeonite poikilitic pigeonite, TiO2 continues to increase, while (<67). A similar relationship is observed between the Al2O3 and Cr2O3 begin to decrease with decreasing Mg#. high-Ca oikocryst rims and non-poikilitic augite, albeit Augite within NWA 10169 shows decreasing Cr2O3, with greater degrees of overlap (Fig. 6a). Within the poiki- increasing TiO2, and no signiﬁcant relationship with litic pigeonite, as Mg# decreases, TiO2 and Al2O3 increase, Al2O3, with decreasing Mg# (Fig. 6). The Ti/Al ratios of 442 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

Di Ha

Intermediate poikilitic Enriched poikilitic Non-poikilitic Poikilitic

En Fs

Fo90 80 70 60 50 40 30 20 10 Fa

Fig. 5. Pyroxene quadrilateral and olivine major element compositions for NWA 10169. Compositions overlap with the enriched poikilitic shergottites. Compositional ﬁelds for the intermediate poikilitic shergottites are from Mikouchi (2005) and Mikouchi and Kurihara (2008). Compositional ﬁelds for the enriched poikilitic shergottites are from Bingkui et al. (2004), Usui et al. (2010), and Howarth et al. (2014).

NWA 10169 pyroxene are relatively uniform within the Fe-rich (Fo65-61)(Tables 1 and S1). The range of poikilitic pyroxene oikocrysts with Ti increasing with Al in poikilitic olivine chadacryst compositions is greater than the pyroxene and no observable trends for non-poikilitic observed compositional range of the non-poikilitic olivine pyroxene (Fig. 7). (Fo61-57)(Fig. 5). Major element compositions of individual The trace elements Y and Zr have positive correlations grains of olivine in NWA 10169 are relatively homoge- with Ti in NWA 10169 pyroxene, while Li has a slight neg- neous, with no evidence for zonation. ative correlation with Ti in NWA 10169 pigeonite and no Olivine shows oscillatory phosphorus zonation, with relationship in the augite (Fig. 7; Table S1). The trace P2O5 contents from <0.05 wt.% to 0.4 wt.% (Fig. 4). element contents vary between the poikilitic and non- These zonations are preserved as phosphorus diffuses very poikilitic textures, with Li concentrations slightly decreas- slowly in olivine (as opposed to Fe and Mg zonations; ing, and Ti, Y, and Zr showing higher abundances in e.g., Milman-Barris et al., 2008; Watson et al., 2015) non-poikilitic pyroxene compared to poikilitic pyroxene, (Fig. 4). In both poikilitic and non-poikilitic textural com- displaying an expected increase in incompatible elements ponents, irregular P-poor regions surround melt inclusions across the textures (Fig. 7). As observed with major element within the mapped olivine crystals. These P-poor regions trends, there is also a greater degree of overlap between the often have rounded boundaries, in contrast to the sharp, poikilitic and non-poikilitic augite than pigeonite. angular compositional boundaries observed in the more Pyroxene REE contents in NWA 10169 show LREE- P-rich zoned regions of the olivine (Fig. 4d, g, and j). depleted profiles, with the greatest degree of LREE- Two of the three mapped olivine grains contain melt inclu- depletion seen in the cores of pyroxene oikocrysts [(La/ sions near their cores, thus obscuring P-zonation in the Yb)CI = 0.06]. Pigeonite within the non-poikilitic textures cores of these crystals. The mapped poikilitic olivine shows have higher overall REE abundances than the pyroxene a large (150 mm thick) homogeneous, relatively P-poor oikocrysts, and similar degrees of LREE-depletion [(La/ region surrounded by continuous oscillatory zonings Yb)CI = 0.07] (Fig. 8a). Augite of the oikocryst rims and (<10 mm thick) becoming more discontinuous towards the within the non-poikilitic textures have overlapping REE rims. The cores of the non-poikilitic olivine crystals are P- profiles with overall REE greater abundances, and lower poor and surrounded by euhedral, continuous oscillatory degrees of LREE-depletion [(La/Yb)CI = 0.10] than the zonings (<10 mm thick) of P-rich and P-poor olivine. Closer pigeonite of NWA 10169 (Fig. 8b). The non-poikilitic to the rims, the P zonation in olivine becomes progressively pyroxenes, both pigeonite and augite, have slight negative more discontinuous, with alternating P-rich bands follow- 1/2 Eu anomalies [Eu/Eu* = EuCI/(SmCI *GdCI) ] of 0.74 ing euhedral crystal shapes that abruptly terminate at the and 0.76, respectively. We use here the geometric mean to continuous zonation of the interior crystal (Fig. 4d and j). calculate the Eu anomalies has been more commonly used Phosphorus zonations reveal agglomerated olivine crystals in the martian literature than the arithmetic mean (e.g., in what otherwise appears as a single olivine crystal in Day et al., 2006; Hui et al., 2011). BSE imagery (Fig. 4d and j). Divalent cations and slow dif- fusing elements in olivine, such as Al and Cr, do not show 3.2.2. Olivine any correlation with the P zonings. Poikilitic olivine chadacrysts located within the cores of Trace element concentrations in NWA 10169 olivine are pyroxene oikocrysts show the most primitive olivine com- similar in both the poikilitic olivine chadacrysts and non- positions in NWA 10169 (Fo68), while olivine chadacrysts poikilitic olivine, with no significant trace element trends located closer to the oikocryst rims are progressively more observed between the two groups of olivine (Tables 2 and Table 1 Representative major element analyses of olivine and pyroxene in NWA 10169. Olivine Pigeonite Augite Poikilitic Non poik. Poikilitic Non poik. Poikilitic Non poik. oxide wt.% Olivine in oik core ———— –> Olivine in oik rim Oikocryst core Oikocryst mantle 443 435–462 (2019) 266 Acta Cosmochimica et Geochimica / al. et Combs L.M.

SiO2 37.6 36.9 36.2 36.2 54.4 52.8 52.2 52.1 51.9 TiO2 – <0.03 <0.03 – 0.07 0.19 0.35 0.26 0.34 Al2O3 <0.03 – <0.03 – 0.44 0.86 1.00 1.60 1.90 Cr2O3 <0.05 <0.05 <0.05 <0.05 0.33 0.43 0.37 0.85 0.71 FeOT 28.1 31.8 34.0 35.7 17.0 16.8 19.4 10.5 11.8 MnO 0.50 0.66 0.72 0.67 0.43 0.44 0.57 0.32 0.36 MgO 34.5 31.3 29.6 28.7 24.9 22.4 20.2 15.8 16.2 CaO 0.11 0.13 0.16 0.12 2.53 5.06 5.30 16.8 15.4 Na2O <0.03 – – <0.03 0.08 0.08 0.09 0.23 0.20 P2O5 <0.03 – – – – – – <0.03 – NiO 0.09 <0.03 0.06 0.03 <0.03 0.06 – <0.03 0.06 Total 100.9 100.9 100.8 101.4 100.2 99.0 99.5 98.5 98.8 Cations per 4 oxygen Cations per 6 oxygen Si 1.00 1.00 0.99 0.99 1.98 1.96 1.96 1.96 1.95 Ti – <0.01 <0.01 – <0.01 0.01 0.01 0.01 0.01 Al <0.01 – <0.01 – 0.02 0.04 0.04 0.07 0.08 Cr <0.01 <0.01 <0.01 <0.01 0.01 0.01 0.01 0.03 0.02 Fe2+ 0.62 0.72 0.78 0.82 0.52 0.52 0.61 0.33 0.37 Mn 0.01 0.02 0.02 0.02 0.01 0.01 0.02 0.01 0.01 Mg 1.36 1.26 1.21 1.17 1.35 1.24 1.13 0.89 0.91 Ca <0.01 <0.01 <0.01 <0.01 0.10 0.20 0.21 0.68 0.62 Na <0.01 – – <0.01 0.01 0.01 0.01 0.02 0.01 P– – – – – – ––– Ni <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 – <0.01 <0.01 Total 3.00 3.00 3.01 3.01 4.00 4.01 4.00 3.99 3.99 Fo 68.6 63.7 60.8 58.9 Wo 5.0 10.3 10.9 35.7 32.6 En 68.6 63.2 57.9 46.8 47.8 Fs 26.3 26.6 31.2 17.5 19.6

FeOT = total iron, Fo = forsterite, Wo = wollastonite, En = enstatite, Fs = ferrosillite, – = zero. 444 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

a) b) 20 1.2 p np p np 18 Augite 1.0 Augite 16 Pigeonite Pigeonite 14 0.8 12

(wt%) 0.6

10 2 8

TiO 0.4

CaO (wt%) CaO 6 4 0.2 2 0 0 60 62 64 66 68 70 72 74 76 60 62 64 66 68 70 72 74 76 Mg# Mg# c) d) 0.9 2.5 p np p np 0.8 Augite Augite 0.7 Pigeonite 2.0 Pigeonite 0.6 0.5 1.5 (wt%) (wt%) 3 0.4 3 1.0 O O 2 0.3 2 Cr 0.2 Al 0.5 0.1 0 0 60 62 64 66 68 70 72 74 76 60 62 64 66 68 70 72 74 76 Mg# Mg#

Fig. 6. Variation of (a) CaO, (b) TiO2, (c) Cr2O3, and (d) Al2O3 with Mg# in NWA 10169 pyroxene. Blue points represent pyroxene oikocrysts (poikilitic areas), and orange points represent interstitial, non-poikilitic pyroxene. (For interpretation of the references to colorin this ﬁgure legend, the reader is referred to the web version of this article.)

S1). Olivine within NWA 10169 have LREE-depleted pro- The chromite chadacrysts within pyroxene oikocrysts and files and REE concentrations lower than the cores of the olivine chadacrysts range in composition from Chr59 (chro- pyroxene oikocrysts; low enough that LREE concentra- mite) Mgn7 (magnetite) Spn21 (spinel) Ulv13 (ulvo¨spinel) to tions in olivine are often below the detection limit of the Chr80Mgn3Spn12Ulv4. One chadacryst located near the rim LA-ICP-MS (0.003 ppm; Table 2). of a pyroxene oikocryst yielded a composition of Chr33- Mgn10Spn10Ulv47 (Tables 3 and S1). The oxides interstitial 3.2.3. Maskelynite to the pyroxene oikocrysts, within the non-poikilitic texture Maskelynite occurs exclusively interstitial to the pyrox- range in composition from Chr42Mgn10Spn11Ulv37 to ene oikocrysts within the non-poikilitic textural compo- Chr16Mgn14Spn5Ulv66 (Fig. S2). nent. The compositions range from An34 to An56 with an average of An49, with no evidence for compositional zoning 3.2.5. Phosphates (Fig. S1; Tables 3 and S1). The majority of maskelynite is Phosphates are solely located within the non-poikilitic K-poor, with an average composition of Or3, increasing textures and are predominantly merrillite, a water-poor from Or2 to Or13 with decreasing An content. The most Mg-rich Ca phosphate with 45 wt.% P2O5, 46 wt.% CaO, K-rich maskelynite present is Or13, located near a late- 3 wt.% MgO, 2 wt.% Na2O, and 1 wt.% FeO. Merrillite stage phosphate grain (Fig. S1). The REE contents of compositions within NWA 10169 are summarized in Tables maskelynite in NWA 10169 yield LREE-enriched profiles 3 and S1. Apatite is present and mainly located within the [(La/Yb)CI = 1.18], with HREE concentrations often below glassy interiors of olivine-hosted polymineralic melt inclu- the detection limit of the LA-ICP-MS (Table 2; Fig. 8c). sions. The diameters of the apatite crystals are commonly The average REE profile shows a positive Eu anomaly <1 mm, below the size that can be reliably analyzed with [Eu/Eu* = 14]; less than the anomalies observed in the an electron microprobe. Merrillite is enriched in the REE enriched shergottites RBT 04261/2, Shergotty, and Zagami with La 555 Â CI. The average REE profile of merrillite (33, 40, and 35, respectively) (Wadhwa et al., 1994; Usui is LREE enriched [(La/Yb)CI = 1.45] with a slight negative et al., 2010). Eu anomaly [Eu/Eu* = 0.78] (Fig. 8d).

3.2.4. Fe-Cr-Ti oxides 3.3. Olivine-hosted polymineralic melt inclusions The Fe-Cr-Ti oxides are more Cr-rich (chromite) within the pyroxene oikocrysts, relative to more Ti-rich oxides Both the olivine chadacrysts in the poikilitic texture and (ulvo¨spinel) within the interstitial non-poikilitic textures. the interstitial olivine in the non-poikilitic texture, are hosts L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 445

a) b) 7 18 6 16 14 5 12 4 10 3 8

Zr (ppm) 6 Li (ppm) 2 p np 4 p np 1 Augite Augite Pigeonite 2 Pigeonite 0 0 0 500 1000 1500 2000 2500 0 500 1000 1500 2000 2500 Ti (ppm) Ti (ppm) c) d) 0.040 16 p np 0.035 Augite surface/hypabyssal 14 Pigeonite 12 0.030 upper crust 10 0.025 8 0.020 lower crust 6 (pfu) Ti

Y (ppm) 0.015 4 p np Augite 0.010 2 upper mantle Pigeonite 0.005 0 0 500 1000 1500 2000 2500 0 Ti (ppm) 0 0.02 0.04 0.06 0.08 0.10 0.12 0.14 Al (pfu)

Fig. 7. Variation of the trace elements (a) Li, (b) Zr, and (c) Y with Ti in NWA 10169 pyroxene. Blue points represent pyroxene oikocrysts (poikilitic areas), and orange points represent interstitial, non-poikilitic pyroxene. The circles represent low-Ca pyroxene analyses, while the squares represent high-Ca pyroxene analyses. (d) Ti versus Al contents in NWA 10169 pyroxene, plotted on experimentally calibrated isobars from Nekvasil et al. (2004) and Filiberto et al. (2010). (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

to abundant melt inclusions, with a greater proportion (>20 mm) Fe-Cr-Ti rich oxides are observed. Their compo- being hosted by non-poikilitic olivine (Fig. 4; Table S1). sitions trend from chromite to u¨lvospinel with decreasing Melt inclusions are present within olivine across the range Fo content of the host olivine (Fig. 4). The major phase of observed compositions (Fo58-68), in both poikilitic and within the interior of the melt inclusions is Si-Al-Ca rich non-poikilitic olivine, and are generally circular to elongate glass, with feldspathic compositions. The glass has average in cross-section. They vary in size from as large as 650 mm SiO2 contents of 73 wt.%, Al2O3 contents of 17 wt.%, up to to less than 100 mm. The majority of melt inclusions are 3 wt.% CaO, and up to 1 wt.% FeO. Glasses in inclusions polymineralic, though there are a number of small hosted by olivine (Fo<63) are enriched in K2O (up to 4 wt. (100 mm) homogenous inclusions of Si and Al-rich glass, %). Blebs of Si-rich (SiO2 > 85 wt.%) glass pockets are as well as some melt inclusions that are nearly monominer- observed within the feldspathic glass of the interior alic (commonly containing high-Ca pyroxene). The (Fig. 4). Maskelynite grains are occasionally observed polymineralic melt inclusions show a relatively concentric within the Si- and Al-rich glass, with compositions of succession of minerals from the rim to core (Fig. 4a). The 48 wt.% SiO2, 34 wt.% Al2O3, 16 wt.% CaO, and 1 wt.% edges of the melt inclusions trapped in more primitive oli- Na2O. In addition, ﬁne-grained phosphates (<5 mm) are vine (Fo65-69) have thin (<5 mm) rims of low-Ca pyroxene present within the interior of the melt inclusions. They of compositions (Wo2En68Fs30 to Wo4En65Fs31), are too small to be reliably analyzed, though EDS analyses with a sharp compositional boundary to high-Ca rims often yield F and Cl peaks, indicating that they are apatite. (Wo43En41Fs16 –Wo46En38Fs16) towards the interior of Less commonly, Mg peaks are observed, possibly indicating the inclusion. Melt inclusions within the more evolved, the presence of merrillite. Needles (<5 mm) of troilite and non-poikilitic olivine (Fo<63) show a similar relationship, ilmenite are also present within the interior of the melt though occasionally the low-Ca pyroxene rim is absent. inclusions, although are less abundant than the ﬁne- The rims of high-Ca pyroxene in these inclusions have com- grained phosphates. positions of Wo44En35Fs21 to Wo52En31Fs18. Both low-Ca and high-Ca pyroxene are present as discrete subhedral to 3.4. Bulk rock composition of NWA 10169 anhedral grains within the interior of the inclusions and are rich in Al2O3 (up to 12 wt.%). Occasionally at the edges The poikilitic shergottites are cumulate rocks, as of the inclusions, small troilite (<10 mm) grains and large shown by their texture, and thus, their measured bulk 446 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

a) b) 100 100 Low-Ca pyroxene High-Ca pyroxene

10 10

1.0 1.0

0.1 0.1 non-poikilitic non-poikilitic poikilitic poikilitic Sample/CI-chondrite Sample/CI-chondrite 0.01 0.01 La Ce PrNd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce PrNd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu c) d) 100 10000 Maskelynite Merrillite

10 1000

1.0

100 0.1 Sample/CI-chondrite Sample/CI-chondrite 0.01 10 La Ce PrNd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce PrNd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu e) 1000

100

1.0 np p np merrillite 0.1 p maskelynite bulk (meas) high-Ca pyx Sample/CI-chondrite bulk (calc) low-Ca pyx 0.01 La Ce PrNd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Fig. 8. Rare earth element (REE) compositions of NWA 10169 minerals normalized to CI chondrites (McDonough and Sun, 1995). (a) low- Ca pyroxene, (b) high-Ca pyroxene, (c) maskelynite and (d) merrillite. Grey envelopes correspond to the REE profiles of minerals from enriched poikilitic shergottites NWA 7755, NWA 7397, RBT 04262, and GRV 020090 (Usui et al., 2010; Jiang and Hsu, 2012; Howarth et al., 2014; Howarth at al., 2015). P = poikilitic, np = non-poikilitic. (e) Average chondrite-normalized REE profiles for pyroxene, maskelynite, and merrillite, as well as the calculated (black) and measured (pink) bulk REE profiles for NWA 10169. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) compositions are not entirely representative of their true calculated as the sum of average mineral REE compositions parent magma compositions (e.g., Stolper and McSween, for olivine, pigeonite, augite, maskelynite, and merrillite, 1979; McSween and Treiman, 1998; Goodrich, 2002; according to their respective modal abundances in NWA Papike et al., 2009; Filiberto et al., 2010; Filiberto and 10169, weighted by their mass proportions. The calculated Dasgupta, 2011). For this reason, we both measured and REE profiles from both thin sections were then averaged. calculated the bulk REE compositions of NWA 10169 Though making up 2% of NWA 10169, Fe-Cr-Ti oxides through the analysis of a bulk powder and in situ laser abla- were not included in the calculation, as their REE abun- tion analysis (Table S1). The calculated bulk REE compo- dances are extremely low, near the detection limit of the sition was inferred using a method similar to Maloy and LA-ICP-MS. The calculated bulk REE profile is shown in Treiman (2007) and Gross and Treiman (2011), and was Fig. 8e. We estimate an uncertainty of 5% for the modal Table 2 Representative trace element analyses of olivine, pyroxene, maskelynite, and merrillite in NWA 10169. Olivine Pigeonite Augite Maskelynite Merrillite Poikilitic Non poik. Poikilitic Non poik. Poikilitic Non poik. Non poik. ppm Olivine in oik core Olivine in oik rim Oikocryst core Oikocryt rim Low-An High-An Li 2.70 3.43 2.70 3.89 2.72 6.13 4.51 5.48 3.76 1.65 Sc 10.2 9.29 10.5 34.1 53.6 101 98.8 7.30 7.79 33.9 Ti 36.1 40.8 47.6 368 941 1154 1466 566 447 438 781 435–462 (2019) 266 Acta Cosmochimica et Geochimica / al. et Combs L.M. V 6.20 5.15 2.73 172 256 591 491 1.49 4.67 7.39 Cr 123 55.6 40.8 2279 2384 6431 4440 1.14 1.13 27.7 Co 79.6 68.6 73.3 46.2 47.0 39.5 28.6 0.289 0.334 1.89 Ni 383 318 261 165 127 155 99.6 0.415 0.801 4.15 Zn 56.7 55.1 77.4 54.1 69.9 39.4 33.6 3.55 4.14 3.34 Ga 0.135 0.149 0.205 1.64 3.98 4.91 6.65 97.1 55.8 1.79 Rb – 0.013 0.026 0.005 0.013 0.163 0.202 22.7 5.25 0.918 Sr 0.006 0.025 0.307 0.249 0.589 9.29 12.3 227 193 115 Y 0.091 0.141 0.092 1.31 4.30 6.70 10.6 0.513 0.161 759 Zr 0.116 0.331 1.07 0.341 2.01 3.57 8.65 3.34 6.98 19.2 Nb – 0.020 0.015 0.004 0.020 0.016 0.063 0.029 0.166 0.131 Cs – 0.009 0.007 0.005 0.009 0.020 0.012 0.041 0.043 0.060 Ba – 0.069 3.13 0.169 1.29 95.5 76.7 302.8 60.8 12.5 La 0.011 0.010 0.011 0.003 0.010 0.093 0.183 0.141 0.317 132.5 Ce 0.003 0.023 0.006 0.023 0.068 0.485 0.834 0.220 0.552 324.5 Pr – 0.018 0.016 0.007 0.018 0.092 0.182 0.052 0.062 44.2 Nd – <0.003 0.008 0.044 0.134 0.696 1.13 0.049 0.300 202.4 Sm 0.010 0.006 0.003 0.027 0.152 0.400 0.748 – 0.043 73.0 Eu 0.003 0.029 0.036 0.007 0.052 0.178 0.262 0.929 0.663 20.7 Gd 0.019 0.005 0.046 0.098 0.353 0.708 1.35 0.015 0.048 110.3 Tb <0.001 0.005 0.002 0.019 0.080 0.176 0.282 0.019 0.008 20.4 Dy 0.006 0.052 0.008 0.197 0.649 1.24 2.07 0.006 0.151 135.8 Ho 0.003 0.023 0.017 0.058 0.172 0.291 0.451 0.002 0.006 28.9 Er 0.020 0.041 0.021 0.131 0.510 0.764 1.21 – <0.004 74.4 Tm 0.009 0.026 0.008 0.026 0.079 0.115 0.170 – 0.003 10.1 Yb 0.050 0.098 0.078 0.195 0.619 0.757 1.04 0.009 0.030 63.5 Lu 0.011 0.026 0.016 0.021 0.093 0.111 0.149 0.003 0.004 8.31 Hf 0.002 0.022 0.041 0.021 0.093 0.185 0.459 0.105 0.167 0.331 Ta 0.004 0.012 0.002 0.001 <0.001 – 0.001 0.093 0.030 0.024 Th 0.008 0.010 0.005 – <0.001 0.001 0.006 0.015 0.027 18.8 U <0.001 0.020 0.02 <0.001 0.001 0.006 0.007 0.033 0.020 3.29 oik = oikocryst, An = anorthite, – = zero. 448 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

Table 3 Representative major element analyses of maskelynite, Fe-Cr-Ti oxides, and phosphates in NWA 10169. oxide wt.% Maskelynite Fe-Cr-Ti oxide Merrillite Low-An High-An High-Or Chr-rich Spn-rich Ulv-rich

SiO2 57.9 54.6 58.8 SiO2 0.14 0.06 <0.03 SiO2 0.19 0.18 TiO2 0.05 0.05 0.05 TiO2 0.72 2.61 17.9 Ce2O3 0.05 0.04 Al2O3 26.1 28.6 25.1 Al2O3 5.99 10.9 3.52 FeOT 1.46 1.12 FeOT 0.41 0.59 0.41 Cr2O3 58.4 46.3 18.5 MgO 2.99 3.29 MnO – <0.05 – Fe2O3 2.65 5.88 13.2 CaO 46.7 46.7 MgO 0.07 0.08 0.08 FeO 29.0 30.9 44.0 Na2O 2.14 1.95 CaO 7.66 10.9 6.94 MnO 0.71 0.61 0.69 K2O 0.08 0.06 Na2O 6.49 4.99 5.93 MgO 3.12 3.30 2.64 P2O5 46.4 46.5 K2O 0.88 0.37 2.14 NiO <0.03 0.03 <0.03 Cl <0.03 – P2O5 0.11 <0.03 0.13 Na2O <0.03 – – Total 100.1 99.8 Total 99.7 100.1 99.6 Total 100.7 100.6 100.4 Cations per 8 oxygen Cations per 4 oxygen Cations per 28 oxygen Si 2.61 2.46 2.65 Si <0.01 <0.01 <0.01 Si 0.03 0.03 Ti <0.01 <0.01 <0.01 Ti 0.02 0.07 0.48 Ce <0.01 <0.01 Al 1.38 1.52 1.33 Al 0.25 0.44 0.15 Fe 0.22 0.17 Fe2+ 0.02 0.02 0.02 Cr 1.62 1.26 0.53 Mg 0.80 0.88 Mn – <0.01 – Fe3+ 0.07 0.15 0.36 Ca 8.96 8.95 Mg <0.01 0.01 0.01 Fe2+ 0.85 0.89 1.32 Na 0.74 0.67 Ca 0.37 0.53 0.34 Mn 0.02 0.02 0.02 K 0.02 0.01 Na 0.57 0.44 0.52 Mg 0.16 0.17 0.14 P 7.03 7.04 K 0.05 0.02 0.12 Ni <0.01 <0.01 <0.01 Cl <0.01 – P <0.01 <0.01 0.01 Na <0.01 – – Total 17.80 17.75 Total 5.00 5.00 4.99 Total 3.00 3.00 3.00 Or 5.1 2.2 12.6 Chr 85.0 68.6 32.0 An 37.5 53.5 34.4 Ulv 2.0 7.4 58.9 Ab 57.4 44.4 53.0 Spn 13.0 24.0 9.1

FeOT = total iron, Or = orthoclase, An = anorthite, Ab = albite, Chr = chromite, Ulv = ulvo¨spinel, Spn = spinel, – = zero.

abundances based on the method of Maloy and Treiman compositions are presented in Table 4. The 176Lu/177Hf (2007). versus 176Hf/177Hf 5-point isochron yields a crystallization The calculated bulk rock REE pattern is consistent with age of 167 ± 31 Ma (2r, MSWD = 1.4), and an initial the REE profile of the measured bulk rock, although over- 176Hf/177Hf ratio of 0.282186 ± 0.000011 (Fig. 9). From this all less enriched in REE, indicating that either method is age, an initial 143Nd/144Nd ratio of 0.512032 ± 0.000009 is sufficient. Both show a relatively flat LREE-enriched pro- derived (Table 4). The Sm-Nd isochron was not obtained file, although the measured bulk REE profile shows a for this sample, but the 176Lu/177Hf age is used to calculate slightly greater enrichment of La (but not the other LREE), the initial 143Nd/144Nd ratio from a bulk rock aliquot. Using including La/YbCI = 1.18, than that of the calculated bulk the CHUR parameters of Bouvier et al. (2008b), the calcu- REE profile (La/YbCI = 0.78) (Fig. 8e) with slight positive lated initial eHf(CHUR) and eNd(CHUR) values are À17.5 Eu anomalies [Eu/Eu* = 1.14 (meas), 1.32 (calc)] (Fig. 8e). and À7.6, respectively, consistent with the enriched shergot- Since the measured REE composition was obtained from a tites (e.g., Bouvier et al., 2005; Debaille et al., 2008; Shafer larger and thus, more representative sample and includes et al., 2010). In order to calculate the source 176Lu/177Hf both poikilitic and non-poikilitic textures, we used the mea- and 147Sm/144Nd ratios, the methods of Borg et al. (1997, sured REE bulk composition for the discussion. Depending 2003) were used, on the basis of the following assumptions: on where the thin section was cut, the ratio of poikilitic/ a bulk chondritic Mars (for the REE and Hf) that formed at non-poikilitic textures can vary drastically within one thin 4.567 Ga (T0); an age of 4.504 Ga (T1) for the differentiation section. Thus, using measured REE bulk composition that of source reservoirs (Borg et al., 2016); and CHUR param- 176 177 176 177 all include both poikilitic and non-poikilitic textures, miti- eters of Lu/ Hf = 0.0336, Hf/ Hf4.567 Ga = gates this issue. 0.279825, k-176Lu = 1.865 Â 10À11 yÀ1, 147Sm/144Nd = 143 144 147 0.1967, Nd/ Nd4.567 Ga = 0.506674, and k- Sm = 3.5. Crystallization age and Lu-Hf and Sm-Nd isotope 6.54 Â 10À12 yÀ1 (Bouvier et al., 2008b). The calculation systematics was performed from the isochron-derived initial 176Hf/177Hf 143 144 and Nd/ Nd ratios at T2 (crystallization age). The The Lu-Hf isotopic compositions of the five aliquots ana- model source 176Lu/177Hf composition is 0.02789 lyzed, crystallization age, isochron initial isotopic composi- ± 0.00017 (2r), and the model source 147Sm/144Nd compo- tion, and model 176Lu/177Hf and 147Sm/144Nd source sition is 0.18317 ± 0.00060 (2r). ..Cmse l eciiae omciiaAt 6 21)4542449 435–462 (2019) 266 Acta Cosmochimica et Geochimica / al. et Combs L.M.

Table 4 Lu-Hf and Sm-Nd isotopic data for NWA 10169. 176 176 176 176 176 176 a 176 176 b Sample Lu/ Hf % unc Hf/ Hf ±1 SD Hf/ Hfi ±2 SD eHfi Isochron age ±2 SD Lu/ Hfsource ±2 SD S1-WR 0.0175 0.2 0.282246 0.000004 0.282186 0.000011 À17.5 167 31 0.02789 0.00017 S2-WR* 0.0129 0.2 0.282214 0.000004 S3-pyx 0.0337 0.2 0.282293 0.000011 S4-mask 0.0106 0.2 0.282225 0.000009 S5-oxide 0.0007 0.5 0.282191 0.000018

147 144 143 144 143 144 a 147 144 b Sample Sm/ Nd % unc Nd/ Nd ±1 SD Nd/ Ndi ±2 SD eNdi Isochron age ±2 SD Sm/ Ndsource ±2 SD S1-WR 0.21384 0.2 0.512266 0.000010 0.512032 0.000009 À7.6 – – 0.1832 0.0006 WR = whole rock, WR* = whole rock minus separates, pyx = pyroxene, mask = maskelynite. a The initial eHf value was calculated using the CHUR parameters of Bouvier et al. (2008b). b Source isotopic composition was calculated using the source diﬀerentiation age of Borg et al. (2016), and the CHUR parameters ofBouvier et al. (2008b). 450 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

0.28234 data-point error ellipses are 2 (Figs. 7 and 8a), in comparison to the non-poikilitic pyroxene. In contrast, oikocryst augite rims have trace element Pyroxene and REE compositions that overlap with non-poikilitic 0.28230 augite, indicating crystallization from a similar stage of parent magma evolution (Figs. 7 and 8b). These major and

WR trace element compositions of pigeonite and augite are akin 0.28236 to those observed in the enriched poikilitic shergottites Hf Mask-rich RBT 04261/2, NWA 7397, NWA 7755, and Grove Moun- 177 tain (GRV) 020090 (Usui et al., 2010; Jiang and Hsu, 2012; Hf/ 0.28232 Howarth et al., 2014; Howarth et al., 2015). Additionally, 176 maskelynite is only observed within the non-poikilitic textu- WR* ral component, suggesting crystallization from interstitial 0.28218 Age = 167 31 Ma melt, following oikocryst formation. The transition from Chromite Initial 176Hf/177Hf = 0.282186 0.000011 MSWD = 1.4 chromite-spinel to chromite-u¨lvospinel substitution in the poikilitic to non-poikilitic regions, respectively, reflects the 0.28214 0.00 0.01 0.02 0.03 0.04 removal of Al from the interstitial melt following the initi- 176Lu/177Hf ation of plagioclase growth (Fig. S2), which was also observed within the enriched poikilitic shergottites RBT Fig. 9. Five-point Lu-Hf isochron calculated with Isoplot. 04261/2 and NWA 7397, and the intermediate poikilitic MSWD = mean standard weighted deviation, mask = maskelynite, shergottite NWA 4797 (Usui et al., 2010; Walton et al., WR = whole rock, WR* = whole rock minus mineral separates. 2012; Howarth et al., 2014). As the interstitial melt continues to crystallize, pockets of highly fractionated melt are isolated, and late stage phases such as merrillite, apatite, 4. DISCUSSION sulfides, and ilmenite crystallize. Phosphorus zonation in olivine arises through rapid 4.1. Conditions of formation growth rate changes leading to disequilibrium incorporation of phosphorus along the advancing crystallization 4.1.1. Crystallization sequence front, providing a record of olivine crystallization in The significant differences in textures and mineral major NWA 10169 from the poikilitic to non-poikilitic textures and trace element compositions observed across the bimo- (Milman-Barris et al., 2008). The core to rim transition dal texture of NWA 10169 provide the basis for defining from homogenous P content to oscillatory P zonation, its crystallization history (Fig. 10). The poikilitic textural observed in the olivine chadacryst of Fig. 4d, likely repre- component is inferred to have crystallized from a more sents a transition from equilibrium to disequilibrium P primitive melt, at the early-stages of crystallization, prior incorporation upon a cooling rate increase associated with to the non-poikilitic component, as also observed in the transport towards the surface (e.g., Milman-Barris et al., other poikilitic shergottites (Usui et al., 2010; Jiang and 2008; Shearer et al., 2013). In contrast, the two mapped Hsu, 2012; Walton et al., 2012; Howarth et al., 2014; non-poikilitic olivine (Fig. 4g and j) show P-poor cores that Howarth et al., 2015). This is evidenced by the higher sharply transition into continuous thin (<5 mm) oscillatory Mg#s of the olivine chadacrysts and pyroxene oikocrysts P zonations. These olivine likely crystallized after the per- (Fig. 5), as well as lower incompatible element abundances iod of slower crystallization recorded in the olivine chada- and greater LREE-depletion of pigeonite oikocrysts cryst, at high cooling rates during transport towards the

Poikilitic component Chr Chr Ulv Spn 85 2 13 Ol Fo 68 Pyx core Wo En Fs 6 68 26 Non-poikilitic component

Wo En Fs mantle 10 63 27 Ulv

Chr58Ulv26Spn16 Ol Pyx Chr69Ulv7Spn24 Fo61 Plag Phos Fo61 Wo10En58Fs31 An56

Cooling Tr rim Wo35En47Fs18 Ilm

Chr30Ulv62Spn9 Fo57 Wo33En47Fs20 An34

Fig. 10. Crystallization sequence of NWA 10169 estimated based on textural relationships, and major and trace element compositional trends. Chr = chromite; ol = olivine; pyx = pyroxene; ulv = u¨lvospinel; plag = plagioclase; phos = phosphate; tr = troilite; ilm = ilmenite. L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 451 surface, consistent with the generation of continuous oscil- tions used for calculations of fO2 within the enriched poiki- latory P zonation (e.g., Milman-Barris et al., 2008; Shearer litic and olivine-phyric shergottites (e.g., Herd, 2003; Usui et al., 2013). Additionally, the non-poikilitic olivine in et al., 2010; Howarth et al., 2014). The subsolidus equilibra- Fig. 4j shows discontinuous P zonation towards its rim, tion temperature and fO2 estimates for the poikilitic com- representing varying growth rates across the crystal face, ponent are 1065 ± 26 °C and 2.30 ± 0.2 log units below suggesting final growth in a highly crystalline, static mag- the fayalite-magnetite-quartz (FMQ) buffer, respectively. matic environment following emplacement (Shearer et al., For the non-poikilitic component, the subsolidus equilibra- 2013). tion temperature and fO2 estimates are 800 ± 61 °C and The apparent crystallization sequence of NWA 10169 is FMQ À1.07 ± 0.14, respectively. This late-stage fO2 value summarized in Fig. 10. Similar crystallization sequences is over one log unit more oxidizing than the estimate for have been proposed for other members of the enriched poi- the early stage poikilitic texture. The only other poikilitic kilitic shergottites (Usui et al., 2010; Howarth et al., 2014; shergottites to have fO2 estimated for both the poikilitic Howarth et al., 2015). Note that the crystallization and non-poikilitic textures are NWA 7397 and NWA sequences observed in the intermediate poikilitic shergot- 7755, and both show late-stage oxidation events of 1 log tites are similar to NWA 10169 and the other enriched poi- unit of magnitude (Howarth et al., 2014; Howarth et al., kilitic shergottites (e.g., Mikouchi and Kurihara, 2008). 2015). Through thermodynamic MELTS modeling, Castle However, mineral compositions in intermediate poikilitic and Herd (2017) showed that fO2 variation greater than shergottites are more primitive with pyroxene and olivine 1 log unit cannot be accounted for solely through auto- more Mg-rich (e.g., GRV 99027, Lewis Cliffs – LEW – oxidation. The 1.2 log unit difference between poikilitic 88516, and NWA 1950; Harvey et al., 1993; Hsu et al., and non-poikilitic fO2 of NWA 10169 may be accounted 2004; Mikouchi, 2005). for through auto-oxidation during continuous crystallization, with the possibility for degassing, as observed in 4.1.2. Temperature, pressure, and oxygen fugacity NWA 7397 and NWA 7755 (Howarth et al., 2014; To constrain the pressure during the crystallization of Howarth et al., 2015). NWA 10169, Ti/Al ratios within pyroxene were determined A polybaric formation model has been proposed for the from both the pyroxene oikocrysts and non-poikilitic enriched poikilitic shergottites, in which the chromite, oli- pyroxene (Fig. 7d). Pyroxene Ti/Al ratios decrease with vine, and enclosing pyroxene oikocrysts crystallized at increasing pressure. This relationship has been calibrated depth in a magma staging chamber, followed by the subse- for the bulk composition of the enriched olivine-phyric quent rise of the parent magma towards the surface and NWA 1068, which has a similar bulk composition and crystallization of the interstitial melt at lower pressure Ti/Al ratio to NWA 10169 (Barrat et al., 2002; Filiberto (Howarth et al., 2014). This hypothesis is supported by et al., 2010). The Ti/Al ratios of pyroxene oikocrysts are the dynamic pressures, temperatures, and oxygen fugacities relatively constant, and have values consistent with crystal- present during the formation of NWA 10169, and indicates lization at constant pressure, and at depths at or near the a petrogenetic link with the enriched poikilitic shergottite base of the martian crust (10 kbars) (e.g., Wieczorek NWA 7397 (Howarth et al., 2014). and Zuber, 2004)(Fig. 7d). As plagioclase crystallization was initiated following oikocryst formation, Ti/Al cannot 4.1.3. Relationship of oxygen fugacity to geochemical sources be used to estimate crystallization pressures of non- It was previously proposed that the degree of LREE- poikilitic pyroxene (McSween et al., 1996). However, the enrichment of the shergottites, represented by their bulk subequigranular, finer grained texture of the non-poikilitic La/Yb ratio, is correlated with the oxygen fugacity condi- regions suggests crystallization under faster cooling rates tions present at their respective formations, suggesting that than the coarse pyroxene oikocrysts, likely at lower pres- the sources of the geochemically enriched and depleted sher- sures closer to the surface. Additionally, pressures consis- gottites are oxidized and reduced, respectively (e.g., Herd tent with near surface conditions (1 bar) were estimated et al., 2002; Herd, 2003)(Fig. 11). However, as more sher- to have been present during late-stage crystallization for gottites are studied, the correlation between oxygen fugacity the poikilitic shergottites Y-000027 (and pairs) and RBT and LREE-enrichment becomes less pronounced, as also 04262 (Mikouchi and Kurihara, 2008; Usui et al., 2010). suggested by Castle and Herd (2017) and Howarth et al. To constrain the oxygen fugacity and temperature con- (2018). The early-stage oxygen fugacity of NWA 10169, ditions during various stages of crystallization of NWA more reduced than other enriched shergottites, overlaps 10169, the olivine-spinel geothermometer and the olivine- with the intermediate shergottites (e.g., Goodrich et al., pyroxene-spinel oxybarometer were used on mineral assem- 2003; Herd, 2003; Lin et al., 2005). Several of the olivine- blages in both poikilitic and non-poikilitic textures (Sack phyric shergottites (Tissint, NWA 6234, NWA 10170; and Ghiorso 1989, 1991a, 1991b, 1994a, 1994b, 1994c; NWA 1068, and LAR 06319) show evidence for a late stage Ballhaus et al., 1991; Wood, 1991; http://melts.ofm- oxidation event, leading to an up to 4 log unit difference research.org/CORBA_CTserver/Olv_Spn_Opx/index.php). between early-stage and late-stage fO2 (Herd, 2006; Peslier As oxygen fugacity is pressure dependent, the approximate et al., 2010; Gross et al., 2013; Castle and Herd, 2017; 10 kbars pressure estimate from pyroxene Ti/Al ratios Howarth and Udry, 2017). The enriched gabbroic shergot- was used for the subsolidus equilibration fO2 calculations tite, NWA 7320, formed under an oxygen fugacity of for the poikilitic textures. Pressures of 1 bar were assumed FMQ, but has a low bulk La/Yb ratio (0.55) for the non-poikilitic calculations, consistent with assump- (Udry et al., 2017). Finally, the geochemically depleted, 452 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

1.4

Poikilitic NWA 10169 Enriched Olivine-phyric 1.2 np Basaltic NWA 1068 p NWA 7397 Gabbroic Los Angeles 1.0 LAR 06319 GRV 020090 Shergotty

0.8 RBT 04262 Zagami NWA 8657 NWA 4797 NWA 5298

0.6 Intermediate NWA 7320 NWA 6234 GRV 99027 ALHA 77005 0.4 NWA 10170 LEW 88516 EETA 79001B Depleted EETA 79001A 0.2 Tissint SaU 005 DaG 476 La/Yb (CI-normalized) La/Yb to +2 QUE 94201 NWA 8159 0 -5 -4 -3 -2 -1 0 1 f log O2

Fig. 11. Oxygen fugacities relative to the FMQ buffer, for NWA 10169 and other shergottites, versus chondrite-normalized La/Yb ratio as a proxy for LREE enrichment. Note a common oxidation event recorded in the olivine-phyric and poikilitic shergottites. Data for the basaltic shergottites are from Lodders (1998), Rubin et al. (2000), Herd et al. (2001), Hui et al. (2011), and Howarth et al. (2018). Data for the olivine- phyric shergottites are from Dreibus et al. (2000), Zipfel et al. (2000), Barrat et al. (2002), Herd (2003; 2006), Peslier et al. (2010), Gross et al. (2013), Castle and Herd (2017), and Howarth and Udry (2017). Data for the poikilitic shergottites are from Goodrich et al. (2003), Lin et al. (2005), Usui et al. (2010), Jiang and Hsu (2012), Walton et al. (2012), and Howarth et al. (2014). Data for NWA 7320 and NWA 8159 are from Udry et al. (2017) and Herd et al. (2017), respectively. Modified from Howarth and Udry (2017). augite-rich, basaltic shergottite NWA 8159 experienced 300 ppm, respectively, suggesting some degree of terres- highly oxidizing conditions (FMQ + 2) (Herd et al., 2017). trial alteration. Similar enrichments have been observed in This overlap of the oxygen fugacities of the shergottites sug- the other hot desert enriched poikilitic shergottite NWA gests that the martian mantle is perhaps more heterogeneous 7397 (Howarth et al., 2014). In addition, although the cal- than previously thought, including reservoirs and sources of culated bulk rock REE pattern is subparallel with the REE varying oxygen fugacity. Processes such as degassing and profile of the measured bulk rock, the measured bulk com- crustal contamination through fluids could also explain position is overall an order of magnitude more enriched in the large ranges of fO2 derived from individual shergottites. REE than the calculated bulk REE compositions of NWA 10169 (Fig. 8e). This could be explained by higher modal 4.2. Geochemical classification and the REE budget content of merrillite in the measured chip than in the thin sections. Merrillite, the main carrier of the REE in NWA We both measured and calculated the bulk REE compo- 10169 (341–555 Â CI), is solely present within the non- sitions of NWA 10169. The first discrepancy between mea- poikilitic textures, thus the overall abundance of merrillite sured and calculated bulk compositions is the greater will change with the relative proportions of interstitial degree of La enrichment in the measured bulk REE profile, and oikocrystic material (e.g., Hsu et al., 2004; Shearer which may have arisen through hot desert alteration that et al., 2015). In addition, the bimodal nature of the poikili- has been shown to mobilize and elevate La, Ce, Sr, and tic shergottites makes these rocks inherently heterogeneous Ba contents (e.g., Crozaz et al., 2003). Cracks filled with ter- with regards to pyroxene oikocryst distribution (e.g., restrial calcite are present within NWA 10169, and suggest McSween and Treiman, 1998; Bridges and Warren, 2006). that alteration had occurred, although the extent of alter- Due to the discrepancies between measured and calculated ation is not clear. Efforts to avoid these cracks were made bulk rock compositions, we used the measured REE bulk during LA-ICP-MS analysis, though it was impossible to composition for the remaining discussion. avoid the cracks while crushing a sample of NWA 10169 The measured CI-normalized bulk REE profile of NWA for bulk analysis. The bulk composition of NWA 10169 10169, normalized to CI, is relatively flat and LREE- shows enrichments in Sr and Ba, 40 ppm and enriched (La/YbCI = 1.18), consistent with the REE L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 453

Los Angeles

Shergotty Zagami LAR 06319

4.0 NWA 1950 LEW 88516

NWA 10169 ALH 77005 RBT 04262 NWA 7397 GRV 99027 GRV 020090 Sample/CI-chondrite Enriched Shergottites 0.4 La Ce Pr Nmd P Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Fig. 12. Chondrite-normalized REE profiles for NWA 10169, intermediate poikilitic shergottites (black), and other enriched shergottites (gray). The measured REE profile for NWA 10169 is consistent with the other enriched shergottites. Data for the enriched poikilitic shergottites are from Usui et al. (2010), Jiang and Hsu (2012), and Howarth et al. (2014). Data for the intermediate poikilitic shergottites are from Lodders (1998), Gillet at al. (2005), and Lin et al. (2005). Data for Zagami and Shergotty are from Lodders (1998), LAR 06319 is from Basu Sarbadhikari et al. (2009), and Los Angeles is from Jambon et al. (2002). profiles of the enriched shergottites Zagami, Shergotty, the fact that some of the daughter phases in the melt inclu- LAR 06319, RBT 04262, NWA 7397, and GRV 0200090 sions were too fine grained (e.g. sulfides, phosphates) to (Lodders, 1998; Basu Sarbadhikari et al., 2009; Usui reliably measure with standard WDS techniques, we were et al., 2010; Jiang and Hsu, 2012; Howarth et al., 2014) unable to use the modal recombination technique outlined (Fig. 12). This La/YbCI value is higher than any recorded by Goodrich et al. (2013) to obtain the present bulk compo- within the shergottite suite, though with some possible con- sition (PBC) of the six selected melt inclusions. Instead, to tribution from hot desert alteration (Fig. 11). The REE measure the PBC, we employed a new method in which ele- composition, textural characteristics, and mineral major mental X-ray maps were mapped by ED-SDD (energy element compositions of NWA 10169 are consistent with dispersive-silicon drift detector) on the JEOL 8530F at those of the enriched poikilitic shergottites, allowing for NASA-JSC (Ross and Simon, 2019). The ThermoFisher its classification as the sixth member of this group of mete- NSS software was used to digitize the areal extent of the orites (Figs. 5, S1, and 9). melt inclusions within the X-ray maps, and subsequent amalgamation of X-ray counts into a single X-ray spectrum 4.3. Parental melt of NWA 10169 for each melt inclusion. The resulting X-ray spectra were then compared to standard X-ray spectra and quantified As NWA 10169 has a cumulate texture, its bulk compo- into bulk compositions through the phi-rho-Z (PRZ) sition does not represent a magmatic composition. To esti- matrix correction routine. Our results using this method mate a parental melt composition, olivine-hosted, are similar to studies of melt inclusions in olivine-phyric polymineralic melt inclusions, located in the primitive oli- shergottites that obtained PBCs through modal recombina- vine chadacrysts and in the later-stage interstitial olivine, tion and defocused beam techniques (e.g., Peslier et al., were analyzed in NWA 10169. Melt inclusions have been 2010; Goodrich et al., 2013). used to estimate a parental melt composition for a number We measured six melt inclusions in NWA 10169: One of olivine-phyric shergottites and nakhlites (e.g., Peslier (MI1) trapped in a relatively primitive olivine chadacryst et al., 2010; Goodrich et al., 2013; Potter et al., 2015; (Fo66), assumed to represent the closest approximation Sonzogni and Treiman, 2015). These melt inclusions repre- to the parental melt of NWA 10169, and five others sent portions of trapped melt that have been enclosed by (MI5, MI6, MI10, MI15, and MI19) trapped in non- growing olivine crystals at a specific point in time. poikilitic olivine (Fo63-59), assumed to represent melt Several different techniques and methods have been used trapped at later stages of parent melt evolution (Table 5). in the past to estimate the primary trapped liquid (PTL): To ensure that the melt inclusions can yield primary analytical, in which melt inclusions are studied in thin sec- trapped liquid (PTL) compositions that are representative tion, and experimental, in which melt inclusions are heated of the magma composition extant during the trapping of and rehomogenized. The analytical technique can further the melt, we followed methods similar to those outlined be subdivided into ‘‘spread beam analyses”, ‘‘grid analy- by Peslier et al. (2010). To establish that the analyzed melt ses”, and ‘‘modal recombination”. For a detailed descrip- inclusions cross-sections do not represent off-center cuts, we tion of each technique and for a detailed comparison have only studied melt inclusions containing late-stage between techniques, see Goodrich et al. (2013). Due to phases (e.g., ilmenite, sulfide, high-Si glass) in their centers. 454 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

Table 5 NWA 10169 parent melt and bulk compositions estimated from melt inclusions and pyroxene compositions. oxide wt.% Melt inclusion PBCb MI1a MI5 MI10 MI15 MI19 MI6

SiO2 61.8 54.1 63.6 50.9 61.8 61.8 TiO2 1.00 1.20 0.90 1.91 0.85 1.28 Al2O3 13.2 8.73 13.4 12.9 13.4 14.1 Cr2O3 0.08 0.05 0.06 1.37 <0.05 <0.05 FeOT 5.97 11.1 6.83 8.83 7.69 7.06 MnO 0.21 0.43 0.23 0.33 0.25 0.10 MgO 6.15 11.8 5.68 3.87 6.39 2.37 CaO 10.1 8.46 7.79 5.99 7.87 9.14

Na2O 0.18 0.11 0.20 3.69 0.17 0.16 K2O 0.21 1.59 0.24 5.01 0.16 1.30 P2O5 0.87 1.86 0.52 4.12 0.79 1.16 SO3 – 0.12 0.14 0.13 0.25 0.27 Cl – – <0.03 0.21 - 0.06 Total 99.7 99.5 99.6 99.3 99.7 98.8 Mg# 64.7 65.5 59.7 43.9 59.7 37.4 Melt inclusion PTLc Fe-Mg corrected NWA 10169 WR LAR 06319 MId MI1a MI5 MI10 MI15 MI19 MI6 host olivine Fo66 Fo63 Fo61 Fo61 Fo61 Fo59 n/a Fo66 FeOT melt 19.4 18.5 18.4 18.4 18.4 18.4 n/a 19.3 SiO2 53.7 52.7 56.8 47.5 56.1 55.1 48.0 52.3 TiO2 0.72 1.18 0.71 1.55 0.69 0.97 0.54 0.99 Al2O3 9.54 8.61 10.5 10.5 10.9 10.6 4.57 9.96 Cr2O3 0.06 0.05 0.05 1.11 <0.05 <0.05 n/a 0.16 Fe2O3 0.81 1.75 1.33 2.57 1.32 1.53 n/a 2.14 FeO or FeOT 18.7 16.9 17.2 16.3 17.2 17.1 21.6 17.3 MnO 0.15 0.42 0.18 0.27 0.20 0.08 0.50 0.19 MgO 8.15 6.49 6.32 4.98 6.27 5.78 19.34 7.09 CaO 7.28 8.34 6.12 4.86 6.39 6.89 3.97 5.24

Na2O 0.13 0.11 0.16 3.00 0.14 0.12 0.96 1.04 K2O 0.15 1.57 0.19 4.07 0.13 0.98 0.11 2.91 P2O5 0.63 1.83 0.41 3.35 0.64 0.88 0.40 0.67 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Mg# 43.8 40.6 39.6 35.3 39.3 37.6 61.5 42.2

K2O/Na2O 1.15 14.27 1.19 1.36 0.93 8.17 0.11 2.80

FeOT = total iron, Fo = forsterite, MI = melt inclusion, n/a = not analyzed, – = zero. a MI1 is entrapped within a poikilitic olivine chadacryst, and represents the estimated parent melt of NWA 10169. b The present bulk compositions (PBC) of melt inclusions were calculated through X-ray maps (Fig. 4). c The primary trapped liquid (PTL) compositions of melt inclusions were calculated by correcting for Fe-Mg exchange with the host olivine in PETROLOG3, assuming original melt FeOT content (Danyushevsky et al., 2002; Danyushevsky and Plechov, 2011). d The PTL compositions of melt inclusions hosted by Fo66 olivine of LAR 06319 (Peslier et al., 2010).

Melt inclusions that formed during rapid growth of skeletal Phosphorus maps of the host olivines reveal rounded zones or dendritic crystals can exhibit stronger boundary layer of low-P olivine that surround melt inclusions and termi- effects (i.e., compositional differences from the parent nate intersecting P zonations, a characteristic interpreted magma due to slow diffusion of incompatible elements to represent olivine crystallization from the trapped melt excluded from the crystal structure) than those formed dur- onto the host crystal (Fig. 4d, g, and j) (Milman-Barris ing slow growth of their host crystal (Faure and Schiano, et al., 2008; Peslier et al., 2010). The new, more Fe-rich oli- 2005), and thus the compositions of the trapped melt would vine, crystallized from the melt inclusion, drives the initia- not represent the original PTL. However, due to its slow tion of re-equilibration between the melt inclusion and diffusion rate, zoning profiles of P in olivine grains can be the host olivine, in which Fe from the melt inclusion diffuses used to determine whether the analyzed inclusions were into the host olivine, and Mg diffuses into the melt inclusion trapped during a period of rapid crystal growth, and thus, (e.g., Danyushevsky et al., 2000; Goodrich et al., 2013). if extensive boundary layer effects should be expected in the The PBC for each melt inclusion is not equivalent analyzed melt inclusions (Milman-Barris et al., 2008). In to the PTL composition, as the PBC Mg#s, using the Fe/Mg this study, we measured P X-ray maps for three of the host 0.35 KDolivine/melt of Filiberto and Dasgupta (2011), do not olivine to ensure that no disruptions of P-zonation patterns yield corresponding olivine compositions consistent with link the melt inclusions to their host olivine rims (Fig. 4). their respective host olivine Fo contents. This is likely due L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 455 to the ‘‘Fe-loss” process described by Danyushevsky et al. p np Fo>72 - 48 Fo>75 - 52 (2000). Following melt entrapment, new olivine, which is RBT 04262 LAR 06319a LAR 06319b indistinguishable from the host olivine, begins to crystallize p np p np np NWA 10169 NWA 7397a NWA 7397b along the wall of the host olivine, becoming more Fe-rich 70 as crystallization continues (Goodrich et al., 2013). To

account for the wall olivine crystallization and the Fe-Mg 65 re-equilibration with the host olivine, we used the PETRO- LOG3 software with the olivine/melt model of Ford et al. 60 (wt%)

(1983) (Danyushevsky et al., 2002; Danyushevsky and 2 Plechov, 2011). The main assumption of this model is that 55 the host olivine composition has not signiﬁcantly changed 50 since melt entrapment, and the algorithm requires the user input of the host olivine forsterite content and the FeO con- 45 T 25 tent of the melt upon entrapment (FeO*) (Danyushevsky and Plechov, 2011). As the FeO* value is unknown, we per- 20 formed the calculations using a variety of FeOT values from 15 the following compositions of meteorites within the enriched (wt%) shergottite suite: Bulk LAR 06319 (Basu Sarbadhikari et al., 3 O

2 10 2009), PTL of LAR 06319 melt inclusions in Fo>72 olivine (Peslier et al., 2010), bulk RBT 04262 (Anand et al., 2008), 5 bulk Zagami, and bulk Shergotty (Lodders, 1998). We did not use the bulk composition of NWA 10169 as it does not 0 20 represent a liquid composition. Once the PTL composition was calculated for the melt 15 inclusion (MI1) trapped within the poikilitic, Fo66 olivine, we performed isobaric (1 kbar) fractional crystallization modeling by running the alphaMELTS algorithm 10 (Ghiorso and Sack, 1995; Asimow and Ghiorso, 1998; Smith and Asimow, 2005). We ran the algorithm with an CaO (wt%)5 Al initial temperature of 1600 °C to ensure modeling started 0 above the liquidus, and an initial oxygen fugacity of 5 FMQ-2, consistent with values determined through oxybarometry (see Section 4.1). Through this method, FeO* 4 values for the melt inclusions trapped within the non- 3 poikilitic Fo63-59 olivine were estimated by isolating melt (wt%) SiO compositions in equilibrium with the respective olivine O 2 2

compositions, thus, enabling the calculation of the PTL K compositions of the remaining melt inclusions. 1 K-rich The PBC and PTL compositions of the six melt inclu- K-poor sions studied within NWA 10169 are presented in Table 5. 0 20 30 40 50 60 The presented MI1 PTL composition is calculated using the FeO* value from the bulk Shergotty composition of Mg# Lodders (1998), as this value is nearly identical to the Fig. 13. Mg# of calculated primary trapped liquid compositions of

FeO content of melt trapped in similar olivine in LAR NWA 10169 melt inclusions, plotted against SiO2,Al2O3, and CaO 06319 (Peslier et al., 2010). Additionally, it provided results wt.%. Lighter colors correspond to lower host olivine Fo-content. the most consistent with prior melt inclusion studies within NWA 10169 has similar parent melt evolution to LAR 06319, the enriched shergottite suite (e.g., Peslier et al., 2010; Basu though less evolved trapped melt compositions than RBT 04262 Sarbadhikari et al., 2011). The calculated PTL composi- and NWA 7397. Data for the enriched poikilitic shergottites are RBT 04262 (Potter et al., 2015), NWA 7397a (Ferdous et al., 2018), tions of the NWA 10169 melt inclusions show decreasing b Mg# with decreasing host olivine forsterite content, indi- and NWA 7397 (He and Xiao, 2014), respectively. Data for the enriched olivine-phyric shergottites are LAR 0619a (Peslier et al., cating that MI1, with a host olivine of Fo , was likely 66 2010) and LAR 06319b (Basu Sarbadhikari et al., 2011), respec- trapped at the early stages of crystallization; thus, its PTL tively. P = poikilitic, np = non-poikilitic. (For interpretation of the composition may be used as a proxy for the parental melt references to color in this figure legend, the reader is referred to the composition of NWA 10169 (Table 5 and Fig. 13). The esti- web version of this article.) mated parent melt of NWA 10169 is similar to the PTL of a melt inclusion trapped within a Fo66 olivine megacryst of sions of LAR 06319 (Peslier et al., 2010; Basu LAR 06319, though with less K2O and Na2O(Peslier Sarbadhikari et al., 2011). Both NWA 10169 and LAR et al., 2010). The trends observed from MI1 to the melt 06319 show constant Al2O3 content with decreasing Mg#, inclusions trapped at later stages of magma evolution are indicating melt entrapment before fractionation of plagio- broadly consistent with those observed in the melt inclu- clase (Fig. 13). There is a slight increase of SiO2 observed 456 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 with decreasing Mg# of melt inclusions in NWA 10169, in inclusions in NWA 10169, trapped in the non-poikilitic oli- contrast with the slight decrease of SiO2 observed in the vine, possibly represent evidence for the addition of metaso- melt inclusions of LAR 06319 (Fig. 13). The low SiO2 con- matized, K-rich material following transport towards the tents of MI15 are likely a result of there being a large surface and initiation of non-poikilitic olivine crystalliza- (>10 mm) Fe-Cr-Ti oxide grain within the melt inclusion tion; an event that perhaps also occurred during the forma- that may have been primary (i.e., present prior to melt tion of NWA 7397 and LAR 06319. An open-system process entrapment) (Table 5). The CaO contents of NWA 10169 is favored, as the difference in K2O/Na2O in the melt inclu- melt inclusions decrease with decreasing Mg#, a trend that sion cannot be explained by compositional change in the extends into the trend observed in preliminary MI data for residual liquid. Major element compositions of these melt RBT 04262, consistent with fractionation of a Ca-rich inclusions do not suggest bulk assimilation and could be phase, likely pyroxene in the case of NWA 10169 (Potter through selective addition of low melting point, K-rich min- et al., 2015)(Fig. 13). Preliminary PTL compositions of erals, but addition of K-rich material could have also melt inclusions within RBT 04262 and NWA 7397 have occurred through magma mixing, similar to processes inter- more evolved Mg# (as low as 20) than those observed preted to have occurred during the crystallization of ALHA in NWA 10169 and LAR 06319, suggesting that the melt 77005 (Ikeda, 1998). inclusions within these meteorites may have been trapped at a later stage than those of NWA 10169 and LAR 4.4. Shergottite crystallization ages and model Lu-Hf and 06319 (Peslier et al., 2010; Basu Sarbadhikari et al., 2011; Sm-Nd source composition He and Xiao, 2014; Potter et al., 2015; Ferdous et al., 2018)(Fig. 13). Note that melt inclusion compositions in The crystallization ages of the shergottites, dated using NWA 7397 measured by He and Xiao (2014) are very close the Lu-Hf, Sm-Nd, and Rb-Sr isotope systems, range from to the more evolved melt inclusions of NWA 10169. The 150 Ma to 574 Ma (e.g., Nyquist et al., 2001; Shafer et al., compositional trends observed in the PTL compositions 2010; Brennecka et al., 2014), with two meteorites, NWA of NWA 10169 melt inclusions show similar magma evolu- 7635 and NWA 8159, as old as 2.4 Ga, extending shergot- tion with LAR 06319 and potentially RBT 04262. tite magmatism to the early Amazonian (Herd et al., 2017; The K2O contents and K2O/Na2O of melt inclusions in Lapen et al., 2017). The 167 ± 31 Ma age measured in this NWA 10169 are highly variable and show a dichotomy study of NWA 10169 (Fig. 9) is concordant with the 150– between K-poor melts (<0.20 wt.%) and K-rich melts (up 225 Ma Lu-Hf, Sm-Nd, and Rb-Sr ages determined for the to 4 wt.%) with K2O/Na2O between 0.93 and 14.5 (Fig. 13 other enriched shergottites (e.g., NWA 4468, Shergotty, and Table 5). This dichotomy is also observed in the melt NWA 856, RBT 04262, LAR 03619) (e.g., Nyquist et al., inclusions of LAR 06319 and NWA 7397, where MI15 of 2001; Borg et al., 2008; Lapen et al., 2009; Shafer et al., NWA 10169 has K2O contents nearly identical to the K- 2010; Ferdous et al., 2017). The age of NWA 10169 is also rich melt of NWA 7397 (Peslier et al., 2010; Basu similar to the 170–212 Ma Lu-Hf, Sm-Nd, and Rb-Sr Sarbadhikari et al., 2011; He and Xiao, 2014)(Fig. 13). ages of the intermediate shergottites, though younger than The observed K-rich melt inclusions in NWA 10169 are the 327–574 Ma ages of the depleted shergottites (e.g., not the result of boundary layer effects because the magni- Nyquist et al., 2001; Shih et al., 2005; Misawa et al., tude of observed enrichment is too large, and there is no cor- 2006; Shih et al., 2011; Brennecka et al., 2014). relation of increasing K-enrichment with decreasing melt The modeled source 176Lu/177Hf composition of NWA inclusion size, as observed by Goodrich et al. (2013). Fur- 10169 is 0.02789 ± 0.00017 and is identical, within uncer- ther evidence that the observed variation of K2O is represen- tainty, to the modeled source values (0.02763 ± 0.00013 tative of the trapped melt, comes from K X-ray maps of the to 0.02783 ± 0.00014) of the enriched shergottites LAR two most K-enriched melt inclusions (MI5 and MI15), 06319, NWA 4468, Shergotty, Zagami, and RBT 04262, showing ubiquitous homogeneous K-rich glass, over an indicating a likely shared geochemical source between order of magnitude more K-rich than the glass of the K- NWA 10169 and this group of enriched shergottites poor melt inclusions, indicating that the variation did not (Bouvier et al., 2005, 2008a; Lapen et al., 2008, 2009; arise through variable exposure of K-rich phases at the thin Shafer et al., 2010)(Fig. 14 and Table 4). The crystalliza- section surface (Table S1). The K-rich melt inclusions MI5 tion ages within this group of shergottites span from 150 and MI15 have variable K2O/Na2O, 14.5 and 1.4, respec- ± 29 Ma to 225 ± 21 Ma for NWA 4468 and RBT 04262, tively. The large K2O/Na2O ratio reported for MI5 is likely respectively, with the crystallization age of NWA 10169 a result of the volatile loss of Na2O from the unstable Si-Al falling within this range (Borg et al., 2008; Lapen et al., rich glass during peak shock pressures (e.g., Goodrich, 2003; 2008). The range of crystallization ages indicates that this Basu Sarbadhikari et al., 2011). Furthermore, the K/Ti shared geochemical source was persistent in the interior ratios of the K-rich melt inclusions of NWA 10169 are con- of Mars for at least 75 Ma. Another group of enriched sher- sistent with those of the nakhlite parent melt, Gale crater gottites, Los Angeles, NWA 856, and NWA 7320, crystal- rocks, and some Gusev crater rocks, and are suggestive of lized contemporaneously with the first group, though metasomatism; meanwhile, the K/Ti ratios of the K-poor apparently sampled a source with distinct modeled melt inclusions, are consistent with igneous fractionation 176Lu/177Hf compositions (0.027834 ± 0.00015) (Nyquist (Treiman and Filiberto, 2015; Filiberto, 2017). However, et al., 2001; Brandon et al., 2004; Debaille et al., 2008; although the K-rich melt inclusions in the nakhlites suggest Lapen et al., 2010; Udry et al., 2017)(Fig. 14). Both groups, sampling of a metasomatized mantle source, the K-rich melt defined by Lu-Hf isotopic compositions, have similar mod- L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 457

0.1855 members of the enriched shergottites. A summary of this RBT 04262 Los Angeles petrogenetic model is presented in Fig. 15. 0.1850 Shergotty (A) Based on the REE contents and the measured 0.1845 enriched eHfi of À17.5 and the calculated eNdi of À7.6 for

source Zagami

LAR 12011 NWA 10169, this rock was derived from a geochemically

Nd 0.1840 176 177 NWA 4468 enriched source. The modeled source Lu/ Hf composi- 143 tion, as discussed in Section 4.4, indicates the source for 0.1835 NWA 856 NWA 10169 was similar to those for the enriched shergot- Sm/ 0.1830 tites Shergotty, Zagami, LAR 06319, NWA 4468, and

147 poikilitic LAR 06319 RBT 04262. Based on poikilitic olivine, low-Ca pyroxene, 0.1825 olivine-phyric NWA 10169 basaltic and spinel assemblages in NWA 10169, the fO2 of this source 0.1820 is likely at least two log units below the FMQ buffer; a value 0.0274 0.0276 0.0278 0.0280 0.0282 0.0284 0.0286 that overlaps with some intermediate shergottites and shows 176 177 Lu/ Hf that there may not be as strong of a correlation between fO2 source and geochemical source as was previously suggested. Fig. 14. 176Lu/177Hf and 147Sm/144Nd source compositions for (B) After melting of this enriched source, initial crystal- enriched shergottites. The source compositions are calculated based lization of olivine and chromite chadacrysts, followed by on crystallization age and isotope data from LAR 06319 (Shafer pyroxene oikocrysts, occurred at constant pressures in a et al., 2010), LAR 12011 (Righter et al., 2015), NWA 4468 (Lapen magma chamber near the base of the martian crust (e.g., et al., 2009; Borg et al., 2016), RBT 04262 (Lapen et al., 2008; Shih Wieczorek and Zuber, 2004). Olivine growth rates were rel- et al., 2009), Shergotty (Nyquist et al., 1979; Debaille et al., 2008), atively slow at this stage based on relatively homogenous P Zagami (Nyquist et al., 1995; Debaille et al., 2008), Los Angeles content in the interior of an olivine chadacryst (Fig. 4d). (Bouvier et al., 2008a; Debaille et al., 2008) and NWA 856 Melt inclusions trapped in olivine at this stage are assumed (Debaille et al., 2007; Borg et al., 2016). NWA 7320 has been modeled to share source compositions with Los Angeles and NWA to represent an approximation of the parental melt compo- 856 (Udry et al., 2017). sition. The estimated parent melt composition for NWA 10169 is similar to melt inclusions in LAR 06319 eled source 147Sm/143Nd compositions ranging from (Lodders, 1998; Peslier et al., 2010). 0.18317 ± 0.00060 to 0.18480 ± 0.00028. Based on (C) Following crystallization in a magmatic chamber at 176Lu/177Hf and 147Sm/144Nd source mixing relationships depth, oikocrysts were entrained in a rising magma and of intermediate and enriched shergottites (e.g., Borg et al., transported close to the surface, then were likely resorbed 2003; Debaille et al., 2008; Lapen et al., 2010, 2017), the during decompression, as described by Howarth et al. 176Lu/177Hf ratio is more sensitive to variations in enriched (2014). During magma ascent and eventual transport of and depleted end-member components than the oikocrysts to the near-surface, cooling rates increased while 147Sm/144Nd ratio. Thus, the Lu-Hf isotope system is able pressure decreased, and the more equigranular non- to distinguish source composition variations within this poikilitic olivine, pyroxene, and plagioclase began to crys- group of enriched shergottites. Although we did not define tallize, along with the augitic oikocryst rims. In contrast ejection age analyses in this study, the similar crystallization to the mapped olivine chadacryst, non-poikilitic olivine age and source compositions of NWA 10169, RBT 04262, show continuous oscillatory P zonations surrounding P- NWA 4468, LAR 06319, Zagami, and Shergotty indicate poor cores, indicative of fast olivine growth in the sur- that these shergottites must have sampled the same enriched rounding melt, likely during magma ascent. Melt inclusions source in the martian mantle, leading to the formation of a trapped during this stage show normal magmatic evolution possible shared magmatic system forming these at the same similar to those trapped in LAR 06319, but less evolved location on Mars. The source Lu-Hf isotopic compositions than those trapped in RBT 04262 and NWA 7397, indicat- and contemporaneous crystallization ages of both of these ing that these melts were trapped at a later stage of magma shergottite groups suggest that there must have been at least evolution (Peslier et al., 2010; Basu Sarbadhikari et al., two enriched sources in the interior of Mars that were sam- 2011; Potter et al., 2015; Ferdous et al., 2018). pled by the existing suite of enriched shergottites. The ori- (D) Potassium-rich melt inclusions in NWA 10169, per- gin of these two geochemical sources remains enigmatic, haps provide evidence for the addition of K-rich material at although their signatures in the enriched shergottites imply this stage, a process that likely occurred during the forma- a greater degree of heterogeneity in the martian mantle and tion of NWA 7397 and LAR 06319 as well (Peslier et al., provides evidence of multiple shergottite magmatic systems 2010; He and Xiao, 2014). Addition of K-rich material being contemporaneously active during the Late Amazo- was likely not significant, however, as NWA 10169 and nian Epoch on Mars. other members of the enriched shergottites retain their mantle-derived Sm-Nd and Lu-Hf isotope compositions. 4.5. Petrogenetic model for NWA 10169 and the enriched (E) From the initial magma chamber at depth to shergottites emplacement in the near-surface, auto-oxidation and possible degassing contributed to a more oxidizing environment Here we present a petrogenetic model for NWA 10169, in the NWA 10169 parental magma, with an fO2 of around from its source extraction to its emplacement in the crust. one log unit below the FMQ buffer. Upon emplacement at We compare this model with those developed for other hypabyssal depths, late stage merrillite, ilmenite, and sul- 458 L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462

Fig. 15. Schematic of a petrogenetic model for NWA 10169 and the enriched shergottites. (A) Partial melting of two distinct enriched sources generating two diﬀerent magma systems. (B) Ponding of magma near the base of the martian crust at 10 kbar, followed by crystallization and accumulation of pyroxene oikocrysts and olivine, and entrapment of MI1. (C) Transport towards the surface, initiation of non-poikilitic crystallization, progressive auto-oxidation, and possible degassing. (D) Assimilation of K-rich, metasomatized material, and entrapment of MI15. (E) Emplacement in the crust and ﬁnal crystallization of the late-stage non-poikilitic material of the enriched poikilitic shergottites. (F) Ejection of the two groups of enriched shergottites from the martian atmosphere based on overlapping CRE ages (e.g., Nyquist et al., 2001; Wieler et al., 2016). See main text for details.

fides crystallize, and the rims of olivine show discontinuous suggesting that, although the intermediate and enriched oscillatory P zonations, consistent with growth in a highly poikilitic shergottites were likely emplaced in a similar man- crystalline, static magmatic environment (e.g., Shearer ner in the martian crust, they were likely not all formed in et al., 2013). This three-stage, polybaric, petrogenetic model the same location on Mars. Additionally, though no iso- for the formation of NWA 10169 is likely common for the topic data is available, the comparisons discussed in this enriched poikilitic shergottites. paper between NWA 10169 and the enriched poikilitic sher- The crystallization age for NWA 10169 is 167 ± 31 Ma, gottite NWA 7397 point towards this meteorite likely also falling within the range of the enriched shergottites Sher- being a member of this enriched magmatic system. Mean- gotty, Zagami, LAR 06319, NWA 4468, and RBT 04262 while, the enriched shergottites Los Angeles, NWA 7320, (150 Ma to 225 Ma) (e.g., Nyquist et al., 2001; Borg and NWA 856, yield distinct 176Lu/177Hf source composi- et al., 2008; Lapen et al., 2008). These meteorites also have tions than those of NWA 10169, RBT 04262, NWA 4468, identical model 176Lu/177Hf and 147Sm/144Nd source com- LAR 06319, Shergotty, and Zagami (e.g., Udry et al., positions to NWA 10169. Thus, a long-lived (at least 2017), indicating two groups of enriched shergottites are 75 Ma) magmatic system, originating from a similar source, represented in these meteorites. existed on Mars and is likely responsible for the formation (F) The crystallization and CRE ages of these two of this group of shergottites. This model links shergottites groups overlap, indicating that two spatially correlated with different textures (e.g., poikilitic, olivine-phyric, and magmatic systems, sampling two distinct enriched sources, basaltic) to the same long-lived magmatic system, while also were active contemporaneously within the interior of Mars, L.M. Combs et al. / Geochimica et Cosmochimica Acta 266 (2019) 435–462 459 providing implications for the heterogeneity of the martian inclusions in Larkman Nunatak 06319. Geochim. Cosmochim. mantle (e.g., Nyquist et al., 2001; Nishiizumi and Caffee, Acta 75, 6803–6820. 2010; Wieler et al., 2016). Bingkui M., Ziyuan O., Daode W., Yitai J., Guiqin W. and Yangting L. (2004) A new martian meteorite from Antarctica: 5. CONCLUSIONS grove mountains (GRV) 020090. Acta Geol. Sin. 78, 1034–1041. Borg L. E., Nyquist L. E., Taylor L. A., Wiesmann H. and Shih C.- Y. (1997) Constraints on Martian differentiation processes from The work presented in this study indicates not only a Rb-Sr and Sm-Nd isotopic analyses of the basaltic shergottite shared, long-lived geochemical source tapped by a subset QUE 94201. Geochim. Cosmochim. Acta 61, 4915–4931. of the enriched shergottites, but also a shared petrogenesis Borg L. E., Nyquist L. E., Wiesmann H. and Reese Y. (2002) and magmatic history that reveals a tantalizing glimpse into Constraints on the petrogenesis of Martian meteorites from the a potential common magmatic system for NWA 10169 and Rb-Sr and Sm-Nd isotopic systematics of the lherzolitic this group of enriched shergottites. This mantle source shergottites ALH77005 and LEW88516. Geochim. Cosmochim. tapped by NWA 10169 is distinct from other enriched sher- Acta 66, 2037–2053. gottites, indicating further heterogeneity of martian mantle Borg L. E., Nyquist L. E., Wiesmann H., Shih C.-Y. and Reese Y. with at least two enriched geochemical sources sampled (2003) The age of Dar al Gani 476 and the differentiation history of the martian meteorites inferred from their radiogenic across the same time span. isotopic systematics. Geochim. Cosmochim. Acta 67, 3519–3536. Borg L. E., Gaffney A. M. and Depaolo D. J. (2008) Preliminary ACKNOWLEDGMENTS Age of Martian Meteorite Northwest Africa 4468 and Its Relationship to the Other Incompatible-Element-enriched We thank Minghua Ren for his assistance with EMP and SEM Shergottites. LPSC XXXIX, abstract #1391. analyses. We also thank Stephanie Suarez for providing assistance Borg L. E., Brennecka G. A. and Symes S. J. (2016) Accretion with column chemistry and Racheal Johnson for her help with timescale and impact history of Mars deduced from the isotopic heavy liquid separation. We thank Tomohiro Usui and two anony- systematics of martian meteorites. Geochim. Cosmochim. Acta mous reviewers for their helpful comments on this manuscript. This 175, 150–167. work was funded in part by the National Aeronautics and Space Bouvier A., Blichert-Toft J., Vervoort J. D. and Albare´de F. (2005) Administration under Grant EPSCoR No. NNX15AIO2H, The age of SNC meteorites and the antiquity of the Martian awarded to L.M. Combs, the NASA Solar System Workings grant surface. Earth Planet. Sci. Lett. 240, 221–233. No, to A. Udry, the NASA Solar System Workings grant No. Bouvier A., Blichert-Toft J., Vervoort J. D., Gillet P. and Albare´de NNX16AR95G to J.M.D. Day, and NASA Solar System Work- F. (2008a) The case for old basaltic shergottites. Earth Planet. ings grant No. 80NSSC18K0312 to T.J. Lapen. Sci. Lett. 266, 105–124. Bouvier A., Vervoort J. D. and Patchett P. J. (2008b) The Lu-Hf and Sm-Nd isotopic composition of CHUR: Constraints from APPENDIX A. SUPPLEMENTARY MATERIAL unequilibrated chondrites and implications for the bulk compositions of terrestrial planets. 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