<<

and Planetary Science Letters 296 (2010) 67–77

Contents lists available at ScienceDirect

Earth and Planetary Science Letters

journal homepage: www.elsevier.com/locate/epsl

The evolution of a heterogeneous : Clues from K, P, Ti, Cr, and Ni variations in and shergottite meteorites

Mariek E. a,⁎, Timothy J. McCoy b a Dept. of Earth Sciences, Brock University, St. Catharines, ON, Canada L2S 3A1 b Dept. of Sciences, National Museum of Natural History, Smithsonian Institution, Washington, DC 20560-0119, USA article info abstract

Article history: Martian basalts represent samples of the interior of the , and their composition reflects their source at Received 10 December 2009 the time of extraction as well as later igneous processes that affected them. To better understand the Received in revised form 16 April 2010 composition and evolution of , we compare whole compositions of basaltic shergottitic meteorites Accepted 21 April 2010 and basaltic examined by the in Gusev Crater. Concentrations range from Available online 2 June 2010 K-poor (as low as 0.02 wt.% K2O) in the shergottites to K-rich (up to 1.2 wt.% K2O) in basalts from the Editor: R.W. Carlson (CH) of Gusev Crater; the basalts from the Gusev Plains have more intermediate concentrations of K2O (0.16 wt.% to below detection limit). The compositional dataset for the Gusev basalts is Keywords: more limited than for the shergottites, but it includes the minor elements K, P, Ti, Cr, and Ni, whose behavior Mars igneous processes during mantle melting varies from incompatible (prefers melt) to very compatible (remains in the shergottites residuum). The range in partitioning behavior of these elements provides leverage on interpreting igneous Gusev basalts processes. Models are presented that demonstrate how concentrations of these elements in Gusev basaltic Mars mantle would change by simple igneous processes (fractional crystallization, crustal contamination, and mantle ). The Gusev basalts may be related by two-stage batch melting of a primitive (K-rich) mantle source to first, generate the K-rich, Columbia Hills basalts and second, the lower K Adirondack basalts, leaving behind a K-poor residuum. The mantle source for the Gusev basalts is separate from the more depleted (K-poorer) source region of the shergottites. This indicates that primitive, K-rich mantle persisted until the Early during formation of the Columbia Hills basalts. We suggest that separate mantle reservoirs developed by inhomogeneous partial melting (at melt fractions greater than 0.05) to form the Martian in the first 1 Ga of the planet's history. © 2010 Elsevier B.V. All rights reserved.

1. Introduction For terrestrial like Mars, where we have few geophysical data, least-differentiated, unaltered basalts offer the only that In spite of the ancient, stable and absence of plate we have to its mantle composition. The composition of a basaltic melt on Mars, has likely persisted until the relatively is determined by its mantle source composition and by processes of recent past. Crystallization ages of the Martian shergottite–– formation and igneous evolution (e.g., partial melting, fractional chassignite (SNC) meteorites are late and as young as crystallization, and crustal contamination). Basalts among the Martian 173±3 Ma (Sm–Nd of ALHA77005; Borg et al., 2001). Crater SNC meteorites, the shergottites, commonly have cumulate textures densities suggest volcanic activity has occurred as recently as that indicate some degree of differentiation (fractional crystallization ∼10 Ma (Lucchitta, 1987; Berman and Hartmann, 2002; Hartmann, and/or crystal accumulation). Most shergottites tend to be depleted in 2005). In addition, significant uplift of the volcanic region is incompatible elements, particularly in alkalis (Fig. 1D) and light rare likely buoyed by both magmatic crustal thickening and hot mantle earth elements (REE; e.g., Treiman et al., 1986) and are likely derived upwelling (Harder and Christensen, 1996; Kiefer, 2003). Together, from a mantle that has been depleted in these elements for a long time these observations suggest that processing of the Martian mantle by based on Sm–Nd isotope systematics (Harper et al., 1995; Borg et al., convection and melt extraction is ongoing to the present day, 1997). A second, more enriched reservoir must also exist because although it is not nearly as vigorous as inside the Earth. some shergottites have flatter REE profiles (higher La/Lu ratios) and Sm–Nd systematics that indicate prolonged incompatible element enrichment. This enriched reservoir may reside in the crust (e.g., Borg ⁎ Corresponding author. et al., 2002) and/or in the mantle as primitive, undepleted domains E-mail address: [email protected] (M.E. Schmidt). (e.g., , 1989) or as the product of latest-stage melts from the

0012-821X/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2010.04.046 68 M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77

Fig. 1. SiO2 and MgO elemental variation diagrams for whole rock analyses of Martian shergottitic meteorites along with unaltered basalts examined by the Spirit rover at Gusev Crater, Mars. classes Adirondack, Backstay, Irvine, and HumboldtPeak are discussed in the text. Rock surface types analyzed by the Spirit APXS were RAT abraded (R), RAT brushed (B), or as is (U). The average undisturbed Gusev composition is also plotted to demonstrate how surface dust may affect Gusev rock surfaces. A) MgO vs. SiO2;B)Al2O3 vs. SiO2; C) FeO* vs. SiO2;D)K2O vs. MgO; E) P2O5 vs. MgO. The QUE 94201 analysis used in other plots (Warren et al., 1996) does not report P2O5 values, but the meteorite has high modal abundance of phosphates (∼6 vol.%; Mikouchi et al., 1998). QUE* shown is an analysis of fusion crust (Kring et al., 2003). F) TiO2 vs. MgO. Data are from Anand et al. (2008), Dreibus et al. (2000, 2002), Gellert et al. (2006), Gillet et al. (2005), Ikeda et al. (2006), Jambon et al. (2002), Ming et al. (2008), Neal et al. (2001), Rubin et al. (2000), Shirai and Ebihara (2004), Taylor et al. (2002), Treiman et al. (1986), Warren et al. (1996, 1999).

solidification of the primordial ocean (e.g., Borg and Draper, 2008; Ming et al., 2008) over narrow ranges in SiO2 and MgO con- 2003). centrations (Fig. 1). Although there is no absolute age information, Unaltered basalts encountered by the Mars Exploration Rover these diverse basalts are found in a Hesperian terrain that is estimated (MER) Spirit in Gusev Crater also give insight to the Martian mantle. In to have formed ∼3.7 billion years ago ( et al., 2005). Im- contrast to the shergottites, these basalts contain fewer conspicuous portantly, Gusev basalts are samples of the Martian interior at an crystals (McSween et al., 2004; Squyres et al., 2006) and more likely earlier time than the shergottites. represent primary melts (Monders et al., 2007) than accumulated For this paper, we focus on the Gusev basalts and their con- crystal mushes. Alkali abundances are variable in Gusev basalts (from centrations of minor elements as determined by Alpha Particle X-ray below detection limits to 1.2 wt.% K2O; McSween et al., 2006a, 2006b, Spectrometer (APXS), including K, P, Ti, Cr, and Ni (Gellert et al., 2006; M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77 69

Ming et al., 2008). The behavior of these elements during igneous must remain in the mantle (Kiefer, 2003). High pressure phase equi- processes, such as partial melting and fractional crystallization ranges libria experiments at to core–mantle boundary from very incompatible (partitions into the melt) to very compatible conditions (2 to 23.5 GPa) on WD composition mantle by Bertka (remains in the residuum). Concentrations of these minor elements and Fei (1997) indicate that the Martian upper mantle contains are modeled to determine what processes may have played a role in , orthopyroxene, clinopyroxene, and garnet, which transitions the formation and evolution of the basalts. We also compare the to majoritic garnet and γ-spinel in the . The heat- Gusev mantle source with that of the younger shergottites and producing elements K, U, and Th are not compatible in these investigate how mantle heterogeneities may have evolved over the and would readily fractionate out of the mantle during partial melting. history of the planet. Other geochemical data sets for Martian crustal Trace phases (e.g., Ca-perovskite and apatite) that might host heat- rocks that were analyzed by the Pathfinder mission (Wänke et al., producing elements were not part of the Bertka and Fei (1997) study, 2001; Foley et al., 2003), the Gamma-Ray Spectrometer (GRS) on the but would likely be the first minerals to melt. Therefore, some K-rich Mars Odyssey orbiter (Boynton et al., 2007) or indirectly by Thermal (potentially WD or LF) mantle likely persists in Mars today. Emission Spectrometer (TES) on the Mars Global Surveyor orbiter The Gusev basalts originate from a separate mantle reservoir than (Wyatt and McSween, 2001; Hamilton et al., 2001) indicate the basaltic shergottites based on high-pressure phase equilibria widespread surface alteration, sediment redistribution, and (or) studies on a synthetic analog to Adirondack, a Gusev plains basalt dust contamination and are not considered for this study. We also (Monders et al., 2007). Adirondack melt multiply saturates with do not consider the (clinopyroxenites) or chassignites olivine, orthopyroxene, and spinel near its liquidus at 1320 °C and (dunites) because they have cumulate textures and the compositions upper mantle pressures at 1.0 GPa (85 km depth). Garnet was not of their parental magmas are a source of debate (e.g., Goodrich et al., identified as a near-liquidus phase. Garnet makes up ∼10% of the 2010; Treiman, 2005). modal in the higher-pressure (2 GPa, 200 km depth) multianvil experiments on WD mantle by Bertka and Fei (1997). The 2. The Martian mantle source for Adirondack is different in major elements (higher FeO and

lower CaO/Al2O3) than the source for the olivine-phyric shergottite The composition of a basaltic melt is a function of its mantle Yamato 980459 (Y98; Agee and Draper, 2004). Because LREE source. Estimates for the composition of bulk (mantle+crust) silicate depletion of the shergottitic source occurred early (Borg et al., Mars (Morgan and Anders, 1979; Dreibus and Wänke, 1985; Wänke 1997), the two mantle domains co-existed at the time of the eruption and Dreibus, 1988; Lodders and Fegley, 1997) have relied on orbital of Gusev basalts (Monders et al., 2007). characteristics, gamma-ray spectroscopy, and compositions of the shergottite–nakhlite–chassignite (SNC) meteorites. The most widely 3. Martian basalts used model of Wänke and Dreibus (1988; here referred to as WD) for the bulk silicate Mars was determined by comparing ratios of 3.1. Shergottites elements that behave similarly during igneous processes (e.g., K/La) in SNC meteorites. Relative to bulk Earth, bulk Mars is richer in volatile Abundant geochemical data are available for the shergottites (as elements, although some alkalis may have fractionated to the surface summarized by McSween and Treiman, 1998; Meyer, 2009 http:// by vaporization during accretion and (or) aqueous alteration. De- curator.jsc.nasa.gov/antmet/mmc/index.cfm), including mineral com- pletion in chalcophile elements suggests that sulfide fractionation positions (e.g., Wadhwa et al., 1994), trace elements (e.g., Treiman et affected the Martian mantle uniformly and that Mars accreted al., 1986), and various isotopic ratios (e.g., Shih et al., 1982; Borg et al., homogeneously (Wänke and Dreibus, 1988). The Lodders and Fegley 2002). (1997; referred to here as LF) model composition for silicate Mars The three lithologic types of shergottites (basaltic, olivine-phyric links the isotopic composition of Martian meteorites to con- basaltic, and lherzolitic) share common compositional traits, such as tributions of different types of nebular materials. Relative to the WD super-chondritic CaO/Al2O3 ratios and low total alkali and Al2O3 model (Wänke and Dreibus, 1988), the LF model has more alkali and contents (Fig. 2). As a group, the shergottites are distinct from rocks halogen elements and lower abundances of siderophile and moder- examined by orbital and landed missions to Mars (McSween et al., ately siderophile elements (Ni and Cr). Thermal and crustal evolution 2009). Shergottites are typically coarse-grained and contain cumulate models for Mars by Hauck and Phillips (2002) indicate that K-rich LF , chromite, and olivine. Magnesium numbers (Mg#, defined mantle heat production would produce excessive crustal thicknesses, as the atomic ratio of 100×Mg/(Mg+Fe)) are an index of magma making this model less likely. differentiation and are variable for basaltic shergottites, ranging from After initial accretion, the Martian mantle was modified during 23.8 (Los Angeles; Rubin et al., 2000) to 56.2 (Zagami; Treiman et al., the early magma ocean by processes such as fractional crystallization 1986). A basaltic melt in equilibrium with the WD mantle olivine

(e.g., Borg and Draper, 2003), crust formation (e.g., Norman, 1999) (Fo76; Bertka and Fei, 1997) would have Mg#≈48, assuming Fe and and cumulate over-turn (Elkins-Tanton et al., 2003). The magma Mg partitioning (KD) of 0.30 (Roeder and Emslie, 1970). This supports ocean would have caused a homogeneous distribution of some textural interpretations that the basaltic shergottites are generally not element groups, such as siderophile elements (e.g., Righter, 2003), primary melts and have experienced some . In that fractionated during core formation. Heterogeneities originating lherzolitic and olivine-phyric shergottites, MgO increases to up to from the later stage crystallization of the magma ocean, including 33 wt.% (Mg#=74.5; ALHA 77005; Treiman et al., 1986) with differentiated melts, unmixed reservoirs, or refractory cumulates decreasing SiO2 concentration, indicating significant olivine accumu- may persist in today's Martian mantle. In addition, later partial lation (Fig. 1A). melting and melt extraction to form the crust would also have created heterogeneities in the distribution of elements in the 3.2. Gusev basalts Martian mantle. In spite of inherent heterogeneities of the Martian mantle, the Over the last 5 years in Gusev Crater, the Mars Exploration Rover models for bulk silicate Mars (WD and LF models) represent best Spirit has encountered mainly volcanic rocks, including diverse, estimates for the primary mantle prior to crust relatively unaltered basaltic lavas first, in the Gusev Plains and later, in formation. In order to sustain convection and volcanism in the the Columbia Hills (Gellert et al., 2006; McSween et al., 2004, 2006a, geologically recent past (e.g., Harder and Christensen, 1996; Lucchitta, 2006b; Ming et al., 2008). Major and minor element compositions of 1987), roughly half of all the heat-producing elements (K, U and Th) rock surfaces (abraded, brushed, and “as is”) were determined by 70 M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77

Fig. 2. A) Total alkali (K2O+Na2O) vs. SiO2 diagram for whole rock analyses of Martian shergottitic meteorites and unaltered basalts examined by the Spirit rover APXS at Gusev Crater,

Mars. This diagram is commonly used to classify and compare alkalinity of igneous rocks. Gusev basalts have higher total alkali concentrations than the shergottites. B) CaO/Al2O3 vs. MgO.

APXS (Table 1). Minerals at rock surfaces were identified remotely by rocks found on the flank of and the Inner Basin of the Miniature Thermal Emission Spectrometer (Mini-TES), and Fe- Columbia Hills (McSween et al., 2006a, b; Schmidt et al., 2008); and (4) bearing minerals were identified in situ by Mössbauer Spectrometer HumboldtPeak class, which is a single olivine-bearing rock that found at (Squyres et al., 2003). the end of in the Inner Basin (Ming et al., 2008; Morris All Gusev basalts are fine-grained with few visible in et al., 2008). Because clastic rocks are more susceptible to alteration, Microscopic Imager (MI) images. Four classes of Gusev basaltic rocks are such as the layered tephra deposit of Home Plate (Barnhill class; e.g., distinguished by elemental compositions and include: (1) Adirondack Schmidt et al., 2009) or the clastic Wishstone class (Hurowitz et al., class, which consists of widespread, massive, olivine- and pyroxene- 2006), they are not considered in this review. bearing blocks in the Gusev plains (McSween et al., 2004); (2) Backstay The diamond grinding teeth of the (RAT) dulled class, which includes a fine-grained olivine-, pyroxene-, magnetite-, and approximately 500 sols (Martian days) into the mission (Arvidson ilmenite-bearing float rock and other float identified by Mini-TES near et al., 2006). Of the basaltic classes, the RAT only abraded surfaces of the crest of Husband Hill (McSween et al., 2006b); (3) Irvine class, which Adirondack class rocks. The RAT brush removed dust from the surface includes massive to highly vesicular, pyroxene and magnetite-bearing of Backstay, but poor positioning of the rover or irregular rock surfaces prevented brushing of Irvine and HumboldtPeak class rocks. Although relatively dust-free surfaces were examined, surface contamination is possible for these rock classes. Even so, we do not normalize APXS Table 1 APXS analyses for type basalts of Gusev Crater. results of the basalts to relatively volatile-free compositions (0.3 wt.% SO3 and no Cl) as has been convention set by McSween et al. (2004; Target name Adirondack Backstay Irvine HumboldtPeak 2006a; 2006b) because all rocks have relatively low S and Cl con- Location Gusev Plains Husband Hill Husband Hill Home Plate centrations (b3 wt.% SO3). RATted rock interiors contain higher SO3 Target type R B U U (1–2 wt.%) than the renormalizing value of 0.3 wt.% and calculated S – # 34 511 600 1341 solubility (0.27 0.54 wt.% SO3; Righter et al., 2009). Some SO3 may have been added by alteration, but renormalization would change wt.% compositions by ∼1–3% with little effect on classifications or SiO2 45.7 (0.4) 49.5 (0.4) 47.0 (0.3) 46.6 (0.4) interpretations. Some amount of alteration has affected all the basaltic TiO2 0.48 (0.06) 0.93 (0.06) 1.06 (0.06) 0.84 (0.06)

Al2O3 10.87 (0.12) 13.25 (0.14) 8.29 (0.09) 10.11 (0.12) classes since amorphous nanophase ferric oxide (npOx) accounts for a FeO* 18.8 (0.12) 13.0 (0.08) 19.2 (0.08) 16.7 (0.10) portion of their Mössbauer spectra (4% in Irvine to up to 13% in MnO 0.41 (0.01) 0.24 (0.01) 0.36 (0.01) 0.39 (0.01) Backstay; Morris et al., 2006; 2008). The npOx alteration is likely a MgO 10.83 (0.12) 8.31 (0.10) 10.63 (0.12) 9.18 (0.13) CaO 7.75 (0.05) 6.04 (0.04) 6.03 (0.04) 6.39 (0.05) product of local low temperature alteration (Morris et al., 2006), fi Na2O 2.41 (0.20) 4.15 (0.21) 2.68 (0.26) 4.25 (0.28) possibly derived by the oxidation and devitri cation of volcanic glass K2O 0.07 (0.05) 1.07 (0.06) 0.68 (0.06) 1.21 (0.06) (Schmidt et al., 2009). P2O5 0.52 (0.07) 1.39 (0.08) 0.97 (0.07) 1.12 (0.07) The Gusev basalts encompass a narrow range in MgO (8.3 to Cr O 0.61 (0.03) 0.15 (0.03) 0.20 (0.03) 0.41 (0.03) 2 3 10.8 wt.%; Fig 1A) and have near primitive Mg#s (43–53), indicating SO3 1.23 (0.03) 1.52 (0.03) 2.37 (0.03) 2.13 (0.04) Cl 0.20 (0.01) 0.35 (0.01) 0.45 (0.01) 0.55 (0.01) they are more likely primary melts than most shergottites. Also, unlike the shergottites, Gusev basalts have subchondritic CaO/Al2O3 ratios ppm and notably higher total alkali abundances (Fig. 2). Backstay is a Ni 165 (39) 191 (34) 289 (37) 360 (39) notable outlier to the Gusev basalts with higher SiO2,Al2O3, and lower Zn 81 (11) 269 (11) 230 (11) 385 (13) Br 14 (15) 26 (14) 6 (13) 29 (14) FeO*. Other Gusev basalts contain similar total FeO* to the shergottites. In incompatible trace elements, the Gusev basalts are more variable. Target types are RATted (R), RAT brushed (B), and unbrushed (U). Total Fe is presented Concentrations of K O vary from 0.07–0.16 wt.% in RATted Adirondack as FeO*. Statistical errors are given in parentheses. Detailed summaries of the Spirit 2 APXS data are presented in Gellert et al. (2006) and Ming et al. (2008). Detection limits basalts to 1.21 wt.% in HumboldtPeak (Fig. 1D; note APXS detection ∼ are distance dependent and are at approximately 0.1 wt.% for K2O and TiO2. limits for K2O and TiO2 are at 0.1 wt.%; Gellert et al., 2006). M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77 71

4. Discussion and an absence of xenocrysts suggest that QUE 94201 is a bulk melt and the result of closed system fractional crystallization under relatively low

4.1. Comparing the two datasets fO2 conditions (0.3 to 1.0 log units below IW buffer; Mikouchi et al., 1998; Wadhwa, 2001; Kring et al., 2003). QUE 94201 is likely derived Trace and minor elements behave predictably in evolving melts from a mantle that has already undergone 4 to 5 episodes of partial and may be used to interpret igneous processes and source het- melting and extraction of an alkali and light REE-enriched melt (Borg et erogeneities. This is well illustrated in a comparison of the basaltic al., 1997). The source of the high P in QUE 94201 is unknown, but it likely shergottite meteorites, representing two extremes in the range in affects mineral stability fields (Kring et al., 2003). In contrast, Los trace elements QUE 94201 (Warren et al., 1996; Kring et al., 2003) and Angeles is the result of crystal fractionation, as indicated by its cumulate Los Angeles (Rubin et al., 2000). The two have generally similar major texture and depletion in the compatible element Cr. An alkali- and light element compositions (Fig. 1A–C) and among the lowest MgO and Ni REE-rich (possibly crustal) component was likely assimilated by the Los concentrations of the shergottites. But QUE 94201 is depleted in the Angeles magma (Rubin et al., 2000). large ion lithophile element (LILE) K while Los Angeles has the highest The APXS data of Gusev basalts include the minor elements K, P, Ti, K concentrations of the shergottites (Fig. 1D). Both meteorites are also Cr, and Ni and these elements vary in behavior from highly in- rich in Ti and QUE 94201 has notably high P (Fig. 1E–F; Rubin et al., compatible (partition coefficient, D≈0) to highly compatible (D≫1; 2000: Kring et al., 2003). Electronic Suppl. I) during igneous processing of basaltic melts, and In order to consider patterns in overall element concentrations, a span the full element abundance diagram (Fig. 3B). Although not all chondrite-normalized element abundance diagram (Fig. 3)lists elements are included, an overlay of the Gusev basalts onto the elements in order of increasing compatibility during (terrestrial) mantle shergottite element abundance diagram (Fig. 3B) allows us to make melting. The major elements Al, Ca, Fe, Si, and Mg are also included in comparisons. The Adirondack basalts have lower, but increasing from this diagram. All shergottites exhibit alkali depletions (Fig. 3A), but QUE left to right, normalized concentrations of the incompatible elements 94201 is also depleted in light REE, resulting in a steeper pattern. Los K, P, and Ti, resulting in a steep pattern. Basalts of the Columbia Hills Angeles has uniformly high concentrations of incompatible trace (CH; Backstay, HumboldtPeak, and Irvine) have higher normalized elements (except K) and has a flatter pattern. These variations suggest concentrations and a flatter pattern in incompatible elements (higher distinct igneous histories for the two meteorites. Mineral compositions K/Ti). Note that in major elements and Ni, the Gusev basalts are not distinguishable from one another in this diagram. Normalized concentrations of Cr span a wider range than the other more com- patible elements. While these patterns among the Gusev basalts are intriguing, surface contamination or alteration of these rocks must be ruled out before their igneous histories may be investigated. The average Gusev soil composition is too similar to the RATted, brushed, and unbrushed

basalt rock surfaces in SiO2, MgO, and Al2O3 to greatly affect rock compositions (Fig. 1A–B). In addition, the average soil has lower K2O and P2O5 than the most K-rich basalts, Backstay and HumboldtPeak (Fig. 1D–E), indicating that high concentrations of these elements do not reflect dust and soil contamination. Dust is most likely to affect concentrations of Ni because up to 3% meteoritic material has been identified in undisturbed (Yen et al., 2006). Ni concentrations in the Gusev basalts vary from 130 ppm in RATted Adirondack class basalts to 460 ppm in unbrushed Irvine basalts. Addition of 3% chondritic material (with 10,500 ppm Ni; Sun and McDonough, 1988) to unbrushed rock surfaces could contribute ∼315 ppm Ni and may account for some of the difference. But this maximum estimate is not significant in a chondrite-normalized diagram (Fig. 3B) and is unlikely because major elements do not suggest dust contamination. Low temperature alteration does not appear to have affected concentrations of K, P, and Ti in these basalts. NpOx, interpreted to be altered volcanic glass (Schmidt et al., 2009), is most abundant in the K-rich Backstay and HumboldtPeak classes (up to 13% of the Mössbauer spectral area), but is not abundant (4–6%) in the K-rich Irvine class rocks (Morris et al., 2008). Other rocks in Gusev Crater, such as those that make up Home Plate, have npOx contents that correlate with abundances of Cl or Zn (Schmidt et al., 2009), but this relation is not found in these basalts. Magnetite contents identified in the Mössbauer spectra (Morris et al., 2008) among the CH basalts are

anti-correlated with P2O5, perhaps suggesting combined phosphate dissolution and magnetite precipitation at rock surfaces. Phosphate dissolution, however cannot account for other trace element differences (K and Ti) between Adirondack class and the CH basalts. Fig. 3. Chondrite normalized element abundance diagrams for A) basaltic shergottites Also, the ratios of P/Ti of Gusev basalts (0.6 to 1.0) are grossly similar and B) Gusev basalts. The gray region in B corresponds to the range in shergottites in A. to estimates for bulk Mars (0.8; Wänke and Dreibus, 1988), indicating C1 chondrite element abundances are from Sun and McDonough (1989). Data are from no significant phosphate dissolution or addition such as has been Anand et al. (2008), Dreibus et al. (2000, 2002), Gellert et al. (2006), Gillet et al. (2005), suggested elsewhere in the Columbia Hills (Ming et al., 2006). Note Ikeda et al. (2006), Jambon et al. (2002), Ming et al. (2008), Neal et al. (2001), Rubin et fi al. (2000), Shirai and Ebihara (2004), Taylor et al. (2002), Treiman et al. (1986), Warren that variations among the CH basalts in P2O5 are not signi cant when et al. (1996, 1998).AsinFig 1E, the value for P for QUE 94201 is from Kring et al. (2003). chondrite normalized (Fig. 3B). 72 M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77

4.2. Relating Gusev basalts by igneous processes ation of a K-, P-, or Ti-bearing mineral, such as phlogopite, apatite, or ilmenite would affect these results. Ilmenite and magnetite have been Because the minor elements K, P, Ti, Cr, and Ni have known identified in the CH basalts Backstay and Irvine by the Mössbauer partitioning behavior in igneous systems (Electronic Suppl. I), we can spectrometer (Morris et al., 2008), but high Ti contents suggest that make inferences about the minerals and processes that may have little to no removal of these minerals has occurred. played a role during the formation of the Gusev basalts. Specifically, After 50% crystallization of an Adirondack composition magma we address what processes may have generated compositional (Fig. 4B), the increase in concentration of the most incompatible differences between incompatible element-poor Adirondack and the element K is too small (0.13 to 0.25 wt.% K2O) and the decrease in incompatible element-rich CH basalts (Backstay, Irvine, Humboldt- concentration of the most compatible element Ni is too great (149 to Peak). In this section, simple models of Rayleigh fractional crystalli- 5 ppm) for it to be parental to the CH basalts by Rayleigh fractional zation, crustal contamination, and partial melting of WD mantle are crystallization. Crystallization of 80 to 90% of Adirondack magma is presented (Fig. 4). Although we compare our models with variations required to attain the K concentrations of Irvine and Backstay. This among the shergottites, such calculations are not as straightforward scenario is unlikely because Mg and Ni concentrations are similar for these rocks because most do not represent true liquid composi- among all Gusev basalts (Fig. 1A and 3B). tions (e.g., McSween and Treiman, 1998). A notable exception is the Fractional crystallization of primary shergottitic magma may olivine-phyric shergottite Yamato 980459 (Y98; Figs 1 and 3A). generate major and REE variations among the older shergottites Thought to be a primary melt, Y98 contains magnesian olivine (474 Ma to 575 Ma; Dar al Gani 476, Sayh al Uhaymir 005/008, and megacrysts (Fo84–86) that appear to be in equilibrium with the Dhofar 019) as demonstrated by Symes et al (2008). Using the MELTS whole-rock composition (Greshake et al., 2004; McKay et al., 2004; algorithm (Ghiorso and Sack, 1995), these workers found that 43% Usui et al., 2008). crystallization of spinel, olivine, orthopyroxene, pigeonite, clinopy- Bulk partition coefficients (D) and molar proportions (X) used in roxene, and of Y98 magma can generate compositions these models are presented in the Electronic Supplement II. The bulk similar to the more evolved QUE94201. But they are quick to point out K distribution coefficient for K (Dbulk) will most likely always be ∼0 that it is unlikely these meteorites represent a single fractional because K is incompatible in the anhydrous mafic minerals considered crystallization path. More likely, the older shergottites are related Ni in this model. Likewise, the Dbulk will be most likely always be ≫1 magmas with a similar parent and similar differentiation histories. because Ni is very compatible in the refractory mafic minerals olivine This is unlikely to be the case for Gusev basalts because Rayleigh and pyroxene. The elements K and Ni are therefore the most robust fractional crystallization cannot replicate observed K and Ni concen- elements in the models presented. trations, and some other process(es) must have played a role in their

Although oxygen fugacity (fO2) greatly affects mineral stability petrogenesis. with implications for trace element partitioning (Herd et al., 2002), its effect is not considered in these models. We selected mineral 4.2.2. Crustal contamination assemblages on the basis of experiments near the QFM (– Trace element concentrations in a basaltic magma may also be fayalite–magnetite) buffer by Monders et al. (2007), and CIPW norms affected by assimilation of a partial melt of the crust. Assimilation of a (McSween et al., 2006a, 2006b) that are based on in situ measure- crustal component has been theorized to account for differences in REE ments of Fe3+/FeTotal (0.17–0.40; Morris et al., 2006, 2008. among the basaltic shergottites (e.g., EET A79001 vs. Zagami; Wadhwa et al., 1994). To test this for the Gusev basalts, we assume (K-poor) 4.2.1. Rayleigh fractional crystallization Adirondack magma (Table 1) intrudes and batch melts an Adirondack The instantaneous removal of crystals from a magma, or Rayleigh composition basaltic crust with either a) CIPW fractional crystallization causes steadily increasing concentrations of (McSween et al., 2006), reflecting a low pressure assemblage or b) incompatible elements and rapidly decreasing concentrations of plagioclase-free lherzolite with equal parts olivine and two compatible elements (Fig. 4B). Representing a limiting case, Rayleigh (Electronic Suppl. II). These models of crustal assimilation are in- fractional crystallization produces the greatest change in K/Ni ratios. tentionally simple. More complex processes of assimilation±fractional The equation for Rayleigh fractional crystallization is crystallization (e.g., AFC; dePaolo, 1981) or ± magma recharge are certainly possible, but would be poorly constrained because the i i ðÞD−1 ð Þ Cliq = C0F 1 composition of the crust is unknown. By considering this simplest case, we assume a homogeneous mantle that produced a single basalt i i where C0 is the initial concentration of element i in the magma, Cliq is composition (Adirondack) and a uniform basaltic crust, thereby the concentration of i in the liquid phase at fraction F after (1−F) has removing any other igneous processes from this calculation. been removed, and D is the bulk crystal-liquid distribution coefficient In order to calculate the concentration of element i in the partial i for i. melt (CPM), we use the batch melting equation McSween et al. (2006) suggested that Adirondack basalt is parental and forms the alkali-rich basalts Irvine and Backstay by fractional i i i i C = C = F + D −FD ð2Þ crystallization based on calculations using the MELTS algorithm at PM 0 bulk bulk various pressures (0.1 to 1 GPa) and oxygen fugacities (fO2; Ghiorso and Sack, 1995). In their model, McSween et al. (2006) consider only where F is the fraction of partial melt. The result of mixing Adirondack major element concentrations and end at 100% crystallization. To test with a 1% batch melt with residual CIPW normative mineralogy (CPM −1) this assertion with the Rayleigh fractional crystallization calculation, and with a 3% batch melt with lherzolite residual mineralogy (CPM −2) are we start with an Adirondack composition basalt (Table 1) and remove plotted in an unnormalized element abundance diagram in Fig. 4C(note an assemblage (plagioclase, olivine, two pyroxenes, and magnetite; linear scale). Also plotted are 20% incremental mixtures of partial melt and Electronic Suppl. II) chosen on the basis of calculated CIPW norms for Adirondack. The 1% batch melt for the CIPW normative residual

Adirondack (McSween et al., 2006a). For comparison, we also present mineralogy (CPM −1) was chosen to maximize K concentrations. Because olivine only fractional crystallization (Fig. 4B). The exact assemblage minor phases such as apatite and magnetite are present in the CIPW does not significantly affect the results because the partitioning normative mineral assemblage, concentrations of P and Ti are too low in behaviors of the more incompatible elements (K, P and Ti) are similar the batch melt to account for their high concentrations in the CH basalts for the minerals considered. Overlapping results of the normative like Backstay. The CPM −2 represents the partial melt with a residual assemblage and olivine only models illustrates this point. Fraction- assemblage that is free of plagioclase and minor phases. The 20% M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77 73

Fig. 4. Chondrite normalized K–P–Ti–Cr–Ni abundance diagrams for A) the Gusev basalts and B–H) modeled igneous processes that may have affected those basalts according to the partition coefficients and mineral modes in Electronic Supplements I and II. B) Rayleigh fractional crystallization (0 to 50%) of an Adirondack composition basalt (C0)at5% increments for a mineral assemblage and for olivine only. C) Weight% element abundance diagram of crustal contamination two-component mixing models that involve an

Adirondack composition magma and a batch melt of Adirondack with residual CIPW normative mineralogy (CPM −1) or residual plagioclase-free lherzolite (CPM −2). Incremental

20% mixtures with CPM −1 and CPM −2 are plotted as is Backstay, an alkali-rich Gusev basalt. D) Chondrite normalized element abundance diagram for a progressive batch melting model of a WD primitive Martian mantle (Cinit, Wänke and Dreibus, 1988). E) Chondrite normalized element abundance diagram for a progressive batch melting model of a LF primitive Martian mantle (Cinit, Lodders and Fegley, 1997) The composition of the resulting liquid (Cliq) and the residual mantle (Cres) at increments of 2% melt fractions (F) are shown for both batch melting models. F) Rayleigh partial melting of a WD* mantle at F=0.01 and 0.02. Two-stage batch melting of a WD* mantle at F=0.015 and 0.05 is presented in a chondrite normalized diagram, G, and an unnormalized linear abundance diagram, H. The results of the two-stage model are liquid partial melts (Cliq1 and Cliq2) and residual mantle

(Cres1 and Cres2). The Gusev basalts are plotted in H for comparison. 74 M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77

mixture of Adiroandack and CPM −2 is fairly similar to Backstay in K, P, and depleted in K to produce any of the Gusev basalts after b1% melting, Ti and may suggest that the Backstay magma interacted with a ruling out Rayleigh partial melting. plagioclase-free . The last case is two-stage batch melting of WD* mantle; the

But more importantly, the SiO2 contents of 1 or 3% partial melt of a primary reservoir melts to produce a (K-rich) magma like CH basalts, basaltic rock would be too high (N65%, e.g., Beard and Lofgren, 1991) leaving behind a depleted (K-poor) residuum that then melts a second to be a significant contribution to Adirondack (45.7 wt.% SiO2) to form time to produce a more depleted (K-poorer) magma like Adirondack. Backstay (49.5 wt.% SiO2, Table 1). In order for the difference in SiO2 Two-stage batch melts of WD* bulk mantle were calculated using i between Adirondack and Backstay magmas to be produced by Eq. (2), where the concentration of element i in the partial melt (CPM) i i assimilation of a ∼65% SiO2 melt, ∼5% addition of a melt with ∼17% correspond with Cliq1 and Cliq2 for the first and second batches at K2O is required. Such a K-rich melt is extremely unlikely to exist on F=0.015 and 0.05, respectively (Fig. 4G–H). After 1.5% melt is i Mars since terrestrial typically have ∼4 wt.% K2O(Irvine and extracted to generateCliq1, the residual mantle composition (Cres1)is i Barragar, 1971). melted another 5% to generate Cliq2, where the concentration of i i element i in the residuum (Cres1 and Cres2; Eq. (3)). Melt percents were 4.2.3. Partial melting of the Martian mantle chosen to best fit the observed variations in the Gusev basalt data. Bulk Melt generation processes and/or source heterogeneities may have distribution coefficients for the first batch melt are the same as those caused differences among the Gusev basalts. To test partial melting used in previous models (Electronic Suppl. II). For the second batch processes, we use the WD (Wänke and Dreibus, 1988) and LF melt, higher abundances of pyroxene and apatite were chosen in order Cr (Lodders and Fegley, 1997) estimates for bulk silicate Mars to to match the P and Ti of Adirondack. In addition, the Dbulk of the second represent primitive mantle. In the simplest case, batch melting batch melt reflects an absence of chromite and higher Cr concentra- (Eq. 2; Electronic Suppl. II) of the Martian mantle produces basaltic tions of Adirondack. Some chromite accumulation might account for magmas with decreasing enrichments in K and other incompatible differences in Cr concentrations between Adirondack and the modeled elements with increasing melt fractions (Fig. 4D–E). An incompatible melt compositions. Even so, modeled K, P, and Ti concentrations element depleted reservoir is left behind with slightly higher correspond very well with those observed in the Gusev basalts concentrations of compatible elements, where the residual concen- (Fig. 4H). It is unlikely that the same parcel of mantle melted to form all i tration of element i, Cres is known by: Gusev basalts. Instead, we suggest the CH basalts originated by melting a primary mantle reservoir, while Adirondack basalts came from a i i i ð Þ Cres = Dbulk × Cliq 3 source that had previously undergone melt extraction. These calculations reveal that the Gusev basalts can be related by Monders et al. (2007) favor batch-melting of WD-like mantle to partial melting of WD-like bulk mantle, either by simple batch or by produce Adirondack basalt, based on high-pressure (0.1 to 1.5 GPa) two-stage batch melting (Fig. 4D, F, G). Varying degrees of multiple phase equilibria experiments. Partition coefficients for our calculation stage partial melting has been suggested to explain major- and trace- reflect their garnet-free, near-liquidus assemblage (1.0 GPa), but with element variations among the nakhlites and shergottites (Longhi, a very minor amount of residual apatite (0.02%) in order to better fitP 1991). The primary mantle source for the Gusev basalts probably has concentrations. Batch melting of WD mantle (Fig. 4D) produces higher concentrations of K than was estimated by Wänke and Dreibus realistic variations in P, Ti, Cr, and Ni at melt fractions (F) of 0.33–0.35 (1988). This suggests that the Gusev basalts are partial melts of a and 0.01–0.005 for Adirondack and the CH basalts, respectively. The K discrete K-rich mantle domain that is possibly the product of late- concentrations in Backstay, however require an extremely low degree stage differentiation of the magma ocean or later metasomatism of the melt (F=0.005) in order for it to have been produced by melting of mantle. K the low-K WD mantle (Cinit =0.003 wt.%). Melt fractions of 0.05– Our partial melting models do not rule out later igneous processing 0.22% are required for melt segregation to occur (Philpotts, 1990), (i.e., fractional crystallization). Later processes would have a relatively making extraction of such low degree melts unlikely. minor effect on K concentrations as is demonstrated by the fractional Batch melting of the alkali-rich LF mantle (Fig. 4E) produces crystallization model (Fig 4B). On Earth, primary mantle melts are Backstay-like K concentrations at more realistic melt fractions (0.01– rare, but differences in incompatible elements of two important basalt 0.05). The P and Ti concentrations, however are too low and do not types, mid-ocean ridge low-K tholeiites and ocean island alkali match Gusev basalt compositions. A possible solution is that the basalts, are attributed to different degrees of partial melting and Gusev mantle source is like the WD model mantle in P, Ti, Cr, and Ni, melt extraction history of their source (e.g., Sun and McDonough, K but with higher K like the LF mantle (Cinit =0.01 wt.%). For sub- 1989). sequent mantle melting models, we use this revised (higher K) In the absence of more trace element data, it is difficult to dis- primary mantle composition that we call WD*. Note that because K is tinguish between the single- or two-stage batch melting, but our a heat-producing element, higher K in the Martian mantle would preference is the two-stage model for two reasons. First, multiple- cause more vigorous convection and an unrealistically thick crust stage batch melting could explain the formation of different mantle (Hauck and Phillips, 2002). The K-rich source for the Gusev basalts reservoirs with variable CaO/Al2O3 and concentrations of incompat- may represent a discrete domain in a heterogeneous mantle. ible elements, including volatiles (CO2, Cl and H2O). Abundant For comparison, Rayleigh partial melting or the instantaneous vesicular Irvine class basalts in the Inner Basin of the Columbia Hills removal of melts from WD* mantle generates trace element con- (Crumpler et al., 2007) suggest they are rich in volatiles and perhaps centrations by derived from a volatile-rich source. In addition, multiple stage melting would allow concentrations of K in the basaltic melt to be lower than ! the initial primary mantle composition, such as has been observed in i 1 −1 i C0 Di C = ×1ðÞ−F bulk ð4Þ some olivine-phyric shergottites (Y98 and Sayh al Uhaymir; Dreibus liq i Dbulk et al., 2000; Shirai and Ebihara, 2004). Also, variations in REE among the shergottites indicate different degrees of light REE depletion in Bulk distribution coefficients are the same as those used in the their mantle source, such as in the meteorites QUE 94201 and Los batch melting calculations (Electronic Suppl. II). Although it is difficult Angeles (Fig. 3; Borg et al., 1997; Rubin et al., 2000). For QUE 94201, for such low degree melts to migrate from their source, this Borg et al. (1997) suggest 4 to 5 episodes of partial melting of garnet- endmember case rapidly depletes the residuum (known by Eq. (3)) bearing cumulate formed during solidification of the Martian magma in least compatible elements (Fig. 4F). In fact, the residuum is too ocean. M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77 75

More importantly, two-stage batch melting allows for the in- we prefer a mean crustal thickness of 40–50 km based on Mars corporation of time, arguably the most significant variable in planetary Orbiter Laser Altimeter (MOLA) topography and gravity models evolution. Melt extraction and light REE depletion occurred very early (Neumann et al., 2004) and which is consistent with an estimate in the shergottitic source region and may have occurred within 33 Ma based on REE concentrations in shergottites (Norman, 1999). of the condensation of Mars based on Sm–Nd systematics of QUE Time in this model is not linear. Lack of absolute age or subsurface 94201 (Borg et al., 1997). Enriched 176Lu/177Hf in the orthopyroxenite information makes it impossible to constrain crustal growth rates. ALH84001 (crystallization age of ∼4.0Gyr)requirestheearly Instead, we assume the bulk of the Martian crust was built by the occurrence of a source that is enriched in Lu/Hf (Righter et al., (within first 1 Ga; Hauck and Phillips, 2002). In order to 2009). If CH basalts are partial melts of primary mantle as these determine the volume of Amazonian age crust, we use surface area models suggest, then primary K-rich mantle must have persisted until estimates of Amazonian volcanic rocks (Tanaka et al., 1992) and the Hesperian (∼3.7 Ga.). Because K and light REE behave similarly assume 5 km new crust underlies these deposits. While there are during igneous processes (Fig. 3), early depletion of these elements certainly Amazonian age surfaces with much greater intrusive (Fig. 3) in the shergottite mantle source region was not uniform. volumes, such as beneath , there must also be areas of Processing of the Martian mantle continues today, as indicated by relatively thin volcanic veneer, such as the relatively young crystallization ages (173 to 575 Ma) of the shergot- (Scott and Tanaka, 1986). The volumes of Noachian and Amazonian tites (Jones, 1985; Borg et al., 2001) and observations of geomorphi- crust are averaged to estimate the amount of crust added during the cally “fresh” flows (Lucchitta, 1987). Thus, the compositions of the Hesperian. In our model, the total crustal thickness added by partially Gusev basalts and shergottites are snapshots, tracking the evolution of melting of the mantle since the Noachian is b1 km. These estimates the Martian mantle. agree with thermal models that suggest the Martian mantle convected more vigorously early (e.g., Schubert et al., 1992; Sleep, 2000) and 4.3. The evolution of the Martian mantle with time crustal growth slowed over time (Hauck and Phillips, 2002). In addition, early light REE depletion of the shergottitic source (Harper To address how the composition of the Martian mantle varied with et al., 1995; Borg et al., 2002) indicates early partial melting and crustal time, we present a mass balance model of progressive crustal growth growth. and mantle depletion. Assuming no crustal recycling, our model Although our estimated crustal growth rates are not well con- involves three reservoirs that change in volume over time: the crust strained, we are able to evaluate how changes in F affects the relative and two mantle reservoirs — a K-rich, primitive mantle (PM) of WD volumes of PM and DM (Fig. 5). As the crust grows, the volume of PM composition and a K-depleted mantle (DM). The CH basalts represent decreases from 100%, and the volume of DM grows over time. At low partial melts of PM, while Adirondack and the low-K shergottites melt fractions (F=0.03) and crustal thicknesses N30 km, no PM is left (Fig 3A) are partial melts of DM. In this model, a volume of PM by the Noachian (Fig. 5). As demonstrated in the batch melting partially melts at melt fraction F (at F=0.03, 0.05, and 0.10) to models (Fig. 4D), even at low melt fractions, the residual mantle (DM) generate volumes of DM and crust, where the total volume of the crust would be depleted in K. The K-rich CH basalts suggest that some PM and the mantle is constant (Fig. 5). Crustal thickness is a significant persists at least until the Hesperian (∼3.7 Ga). Therefore, a higher unknown variable because it correlates with the volume of mantle degree of mantle partial melting (FN0.05, Fig. 5) is more likely. involved in crust formation. Estimates of the average thickness of the Thermal models by Kiefer (2003) suggest that in order for mantle Martian crust range from 30 to 100 km (e.g., Nimmo and Stevenson, convection to support the Tharsis bulge uplift today, ∼50% of ra- 2001; Wieczorek and Zuber, 2004, and Zuber, 2000). For our purposes, dioactive elements (including K) presently reside in the mantle. If all heat-producing elements are in PM domains, then PM makes up at least 10 to 15% of present total mantle volume. A possible solution is that the 40–50 km-thick Martian crust was created by ∼10% partial melting of the mantle (Fig. 5), leaving ∼20% PM still residing in the present Martian mantle. This model is overly simplistic since the Martian crust most likely formed by inhomogeneous, multi-stage partial melting of the mantle like we suggest for the Gusev basalts. However, it demonstrates that under certain circumstances of incomplete mantle melting, heterogeneities formed by melt extrac- tion, including some primary WD or more K-rich mantle domains, may persist over the history of Mars.

5. Conclusion

Variations in K, Ti, P, Cr, and Ni concentrations among the relatively unaltered basalts examined by the Spirit rover in Gusev Crater may be reconciled by two-stage partial melting of primary WD mantle. Domains of K-rich WD mantle likely persisted until at least the Early Hesperian (∼3.7 Ga.) during generation of the Gusev basalts. Alkali- and prolonged light REE-depletions that are characteristic of the shergottitic source (Borg et al., 1997; McSween and Treiman, 1998) were not ubiquitous to the Martian mantle. We infer that partial melting to form the crust created hetero- Fig. 5. Mass balance model of the evolution of the Martian crust and mantle since the geneities within the Martian mantle within the first billion years of the formation of the planet (T=0) to the present day. Volume curves are illustrated for the planet's history. Our models of crustal growth and mantle evolution 40–50 km-thick Martian crust (based on MOLA observations; Neumann et al., 2004), may be adjusted once additional mantle reservoirs are identified, and primitive mantle (PM) of WD composition (Wänke and Dreibus, 1988), and depleted we discover other unaltered basalts from the surface of Mars in mantle (DM). Modeled melt fractions (F) are F=0.03, 0.05, and 0.1. The nonlinearity of the timescale is by reason of a lack of absolute time constraints for Mars crustal meteorite collections or by future spacecraft missions. In particular, growth. isotopic data and absolute age constraints of a basalt sampled from the 76 M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77

Martian surface would allow us to consider isotopic reservoirs and Greeley, R., Foing, B.H., McSween, H.Y., , Pinet, P., van Kan, M., Werner, S.W., fl – – Williams, D.A., 2005. Fluid lava ows in Gusev Crater. J. Geophys. Res. 110, E05008. elemental partitioning (e.g., Rb Sr and Sm Nd) with a linear, rather doi:10.1029/2005JE002401. than relative timescale. Goodrich, C.A., Treiman, A.H., Filiberto, J., Jercinovic, M.J., 2010. The Nahkla parent magma: old problems, new approaches. Lunar. Planet. Sci. Con. 41 Abs#1387. Greshake, A., Fritz, J., Stöffler, D., 2004. and shock metamorphism of the Acknowledgements olivine-phyric shergottite Yamato 980459 Evidence for a two-stage cooling and a single-stage ejection history. Geochim. Cosmochim. Acta 68, 2359–2377. Hamilton, V.E., Wyatt, M.B., McSween, H.Y., Christensen, P.R., 2001. Analysis of This work benefited from the Mars Meteorite Compendium terrestrial and Martian volcanic compositions using thermal emission spectroscopy (Meyer, http://curator.jsc.nasa.gov/antmet/mmc/index.cfm) and was 2. Application to spectra from the Mars Global Surveyor Thermal – supported by a Mars Exploration Rover participating scientist award Emission Spectrometer. J. Geophys. Res. 106, 14,733 14,746. Harder, H., Christensen, U.R., 1996. A one-plume model of Martian mantle convection. to Tim McCoy. Careful and perceptive reviews by Kevin Righter and an 380, 507–509. anonymous reviewer greatly improved this manuscript. In addition, Harper, C.L., Nyquist, L.E., Bansal, B., Wiesmann, Shih, C.Y., 1995. Rapid accretion and 142 144 we acknowledge the engineers and scientists on the MER team for early differentiation of Mars indicated by Nd/ Nd in SNC meteorites. Science 267, 213–217. doi:10.1126/science.7809625. their contributions to the success of the mission. Hartmann, W.K., 2005. Martian cratering 8: isochron refinement and the chronology of Mars. Icarus 174, 294–320. doi:10.1016/j.icarus.2004.11.023. Hauck, S.A., Phillips, R.J., 2002. Thermal and crustal evolution of Mars. J. Geophys. Res. Appendix A. Supplementary data E7, 5052. doi:10.1029/2001JE001801. Herd, C.D.K., Borg, L.E., Jones, J.H., Papike, J.J., 2002. Oxygen fugacity and geochemical Supplementary data associated with this article can be found, in variations in the martian basalts: implications for martian basalt petrogenesis and oxidation state of the upper mantle of Mars. Geochim. Cosmochim. Acta 66, 2025–2036. the online version, at doi:10.1016/j.epsl.2010.04.046. Hurowitz, J.A., McLennan, S.M., McSween, H.Y., DeSouza, P.A., Klingelhöfer, G., 2006. Mixing relationships and the effects of secondary alteration in the Wishstone and Watchtower Classes of Husband Hill, Gusev Crater, Mars. J. Geophys. Res. 111, References E12S14. doi:10.1029/2006JE002795. Ikeda, Y., Kimura, M., Takeda, H., Shimoda, G., Kita, N.T., Morishita, Y., Suzuki, A., Jagoutz, Agee, C.B., Draper, D.S., 2004. Experimental constraints on the origin of Martian E., Dreibus, G., 2006. Petrology of a new basaltic shergottite: Dhofar 378. Antarct. meteorites and the composition of the Martian mantle. Earth Planet. Sci. Lett. 224, Meteorite Res. 19, 20–44. 415–429. Irvine, T.N., Barragar, W.R.A., 1971. A guide to the chemical classification of the common Anand, M., James, S., Greenwood, R.C., Johnson, D., Franchi, I.A., Grady, M.M., 2008. volcanic rocks. Can. J. Earth Sci. 8, 523–548. Mineralogy and of shergottite RBT 04262. Lunar Planet. Sci. XXXIX, Jambon, A., Barrat, J.A., Sautter, V., Gillet, Ph., Göpel, C., Javoy, M., Joron, J.L., Leseourd, M., 2173 abstact # 1391. 2002. The basaltic shergottites Northwest Africa 856: petrology and chemistry. Arvidson, R.E., et al., 2006. Overview of the Spirit Mars Exploration Rover mission to Meteorit. Planet. Sci. 37, 1147–1164. Gusev Crater: landing site to Backstay rock. J. Geophs. Res. 111E02S01. doi:10.1029/ Jones, J.H., 1985. A discussion of isotopic systematics and mineral zoning in the 2005JE002499. shergottites: evidence for a 180 m.y. igneous crystallization age. Geochim. Berman, D.C., Hartmann, W.K., 2002. Recent fluvial, volcanic, and tectonic activity on Cosmochim. Acta 50, 969–977. the Plains of Mars. Icarus 159, 1–17. Jones, J.H., 1989. Isotopic relationships among the shergottites, the nakhlites and Beard, J.S., Lofgren, G.E., 1991. Dehydration melting and -saturated melting of chassigny. Proc. 19th Lunar Planet Sci. Con. 465–474. basaltic and andesitic greenstones and amphibolites at 1, 3, and 6.9 kbar. J. Petrol. Kiefer, W.S., 2003. Melting in the Martian mantle: shergottite formation and 32, 365–401. implications for present-day mantle convection on Mars. Meteorit. Planet. Sci. 38, Bertka, C.M., Fei, Y., 1997. Mineralogy of the Martian interior up to the core–mantle 1815–1832. boundary pressures. J. Geophys. Res. 102, 5251–5264. Kring, D.A., Gleason, J.D., Swindle, T.D., Nishizumi, K., Caffee, M.W., Hill, D.H., Jull, A.J.T., Borg, L.E., Nyquist, L.E., Taylor, L.A., Weismann, H., Shih, C.-Y., 1997. Constraints on Boynton, W.V., 2003. Composition of the first bulk melt sample from a volcanic Martian differentiation process from Rb–Sr and Sm–Nd isotopic analyses of the region on Mars: Queen Alexandra Range 94201. Meteorit. Planet. Sci. 38, basaltic shergottite QUE 94201. Geochim. Cosmochim. Acta 61, 4915–4931. 1833–1848. Borg, L.E., Nyquist, L.E., Reese, Y., Wiesmann, H., Shih, C.-Y., Ivanova, M., Nazarov, M.A., Lodders, K., Fegley, B., 1997. An oxygen isotope model for the composition of Mars. Taylor, L.A., 2001. The age of Dhofar 019 and its relationship to the other Martian Icarus 126, 373–394. meteorites. Lunar Planet. Sci. XXXII abstract #1144. Longhi, J., 1991. Complex magmatic processes on Mars: inferences from the SNC Borg, L.E., Nyquist, L.E., Weismann, H., Reese, Y., 2002. Constraint on the petrogenesis of meteorites. Proc. Lunar. Planet. Sci. 21, 695–709. Martian meteorites from the Rb–Sr and Sm–Nd isotopic systematics of the lherzolitic Lucchitta, B.K., 1987. Recent mafic . Science 235, 565–567. shergottites ALH77005 and LEW88516. Geochim. Cosmochim. Acta 66, 2037–2053. McKay, G., Le, L., Schwandt, C., Mikouchi, T., Koizumi, E., Jones, J., 2004. Yamato 980459: Borg, L.E., Draper, D.S., 2003. A petrogenic model for origin and compositional variation the most primitive shergottite? Lunar Planet. Sci. XXXV abstract. # 1944. of the Martian basaltic meteorites. Meteorit. Planet. Sci. 38, 1713–1731. McSween, H.Y., Treiman, A.H., 1998. Martian meteorites. In: Papike, J.J. (Ed.), Planetary Boynton, W.V., et al., 2007. Concentration of H, Si, Cl, K, Fe, and Th in the low- and mid- Materials: Reviews in Mineralogy, 36, 6, pp. 1–53. latitude regions of Mars. J. Geophys. Res. 112, E12S99. doi:10.1029/2007JE002887. McSween, H.Y., et al., 2004. Basaltic rocks analyzed by the Spirit rover in Gusev Crater. Crumpler, L.S., McCoy, T., Schmidt, M.E., Cabrol, N., 2007. Physical Volcanology at Gusev Science 305, 842–845. doi:10.1126/science.3050842. Crater, Spirit Rover. 7th Internat. Conf. Mars Abstract #3385. McSween, H.Y., et al., 2006a. Characterization and petrologic interpretation of olivine- dePaolo, D.J., 1981. Trace element and isotopic effects of combined wallrock rich basalts at Gusev Crater, Mars. J. Geophys. Res. 111, E02S10. doi:10.1029/ assimilation and fractional crystallization. Earth Planet. Sci. Lett. 53, 189–202. 2005JE002477. Dreibus, G., Wänke, H., 1985. Mars, a volatile-rich planet. Meteoritics 20, 367–381. McSween, H.Y., Ruff, S.W., Morris, R.V., Bell III, J.F., Herkenhoff, K., Gellert, R., Stockstill, Dreibus, G., Spattel, B., Haubold, R., Jochum, K.P., Palme, H., Wolf, D., Zipfel, J., 2000. K.R., Tornabene, L.L., Squyres, S.W., Crisp, J.A., Christensen, P.R., McCoy, T.J., Chemistry of a new shergottite: Sayh al Uhaymir 005 [abstract]. Meteorit. Planet. Mittlefehldt, D.W., Schmidt, M., 2006b. Alkaline volcanic rocks from the Columbia Sci. 35 (Suppl), A39–A440. Hills, Gusev Crater, Mars. J. Geophys. Res. 111, E09S91. doi:10:1029/2006JE002698. Dreibus, G., Wlozka, F., Huisl, E., Jagoutz, E., Kubny, A., Spettel, B., 2002. Chemistry and McSween, H.Y., Ruff, S.W., Morris, R.V., Gellert, R., Klingelöfer, G., Christensen, P.R., petrology of the most feldspathic shergottite: Dhofar 378. Meteorit. Planet. Sci. 37, McCoy, T.J., Ghosh, A., Moersch, J.M., Cohen, B.A., Rogers, A.D., Schröder, C., Squyres, Suppl, A43. S.W., Crisp, J.A., Yen, A., 2008. Mineralogy of volcanic rocks in Gusev Crater, Mars: Elkins-Tanton, L.T., Parmentier, E.M., Hess, P.C., 2003. Magma ocean fractional crystalli- reconciling Mössbauer, Alpha Particle X-ray Spectrometer, and Miniature Thermal zation and cumulate overturn in terrestrial planets: implications for Mars. Meteorit. Emission Spectrometer spectra. J. Geophys. Res. 113, E06S04. doi:10.1029/ Planet Sci. 38, 1753–1771. 2007JE002970. Foley, C.N., Economou, T., Clayton, R.N., 2003. Final chemical results from the Mars McSween, H.Y., Taylor, G.J., Wyatt, M.B., 2009. Elemental composition of the Martian Pathfinder alpha proton X-ray spectrometer. J. Geophys. Res. 108 (E12), 8096. crust. Science 324, 736–739. doi:10.1029/2002JE002019. Meyer, C., 2009. The Mars meteorite compendium, astromaterials research and Gellert, R., Reider, R., Brückner, J., , B.C., Dreibus, G., Klingelhöfer, G., Lugmair, G., exploration science. http://curator.jsc.nasa.gov/antmet/mmc/index.cfm2009. Ming, D.W., Wänke, H., Yen, A., Zipfel, J., Squyres, S.W., 2006. Alpha Particle X-ray Mikouchi, T., , M., McKay, G.A., 1998. Mineralogy of Antarctic basaltic Spectrometer (APXS): results from Gusev Crater and calibration report. J. Geophys. shergottite Queen Alexandra Range 94201: similarities to Elephant Moraine Res. 111, E02S05. doi:10:1029/2005JE002555. A79001 (Lithology B) . Meteorit. Planet. Sci. 33, 181–189. Ghiorso, M.S., Sack, R.O., 1995. Chemical mass transfer in magmatic processes, IV. A Ming, D.W., Mittlefehldt, D.W., Morris, R.V., Golden, D.C., Gellert, R., Yen, A., Clark, B.C., revised and internally consistent thermodynamic model for the interpolation and Squyres, S.W., Farrand, W.H., Ruff, S.W., Arvidson, R.E., Klingelhöfer, G., McSween, H.Y., extrapolation of liquid–solid equilibria in magmatic systems at elevated tempera- Rodionov, D.S., Schröder, C., de Souza, P.A., Wang, A., 2006. Geochemical and tures and pressures. Contrib. Mineral. Petrol. 119, 197–212. mineralogical indicators for aqueous processes in the Columbia Hills of Gusev Crater, Gillet, P., Barrat, J.A., Beck, P., Marty, B., Greenwood, R.C., Franchi, I.A., 2005. Petrology, Mars. J. Geophys. Res. 111, E02S12. doi:10.1029/2005JE002560. geochemistry, and cosmic-ray exposure age of lherzolitic shergottite Northwest Ming, D.W., Gellert, R., Morris, R.V., Arvidson, R.E, Brückner, J., Clark, B.C., Cohen, B.A., Africa 1950. Meteorit. Planet. Sci. 40, 1175–1184. d'Uston, C., Economou, T., Fleischer, I., Klingelhöfer, G., McCoy, T.J., Mittlefehldt, D.W., M.E. Schmidt, T.J. McCoy / Earth and Planetary Science Letters 296 (2010) 67–77 77

Schmidt, M.E., Schröder, C., Squyres, S.W., Trégiur, E., Yen, A.S., Zipfel, J., 2008. Shirai, N., Ebihara, M., 2004. Chemical characteristics of a Martian meteorite, Yamato Geochemical properties of rocks and soils in Gusev Crater, Mars: results of the Alpha 980459. Antarct. Meteorite Res. 17, 55–67. Particle X-ray Spectrometer from Cumberland Ridge to Home Plate. J. Geophys. Res. Sleep, N.H., 2000. Evolution of the mode of convection within terrestrial planets. 113, E12S39. doi:10.1029/2008JE003195. J. Geophys. Res. 105, 17563–17578. Monders, A.G., Médard, E., Grove, T.L., 2007. Phase equilibrium investigations of the Squyres, S.W., Arvidson, R.E., Baumgartner, E.T., Bell III, J.F., Christensen, P.R., Gorevan, Adirondack class basalts from the Gusev plates, Gusev Crater. Mars Meteorit. S., Herkenhoff, K.E., Klingelhöfer, G., Madsen, M.B., Morris, R.V., Reider, R., Romero, Planet. Sci. 42, 131–148. R.A., 2003. The Athena science investigation. J. Geophys. Res. 108 (E12), Morgan, J.W., Anders, E., 1979. Chemical composition of Mars. Geochim. Cosmochim. 8062. doi:10.1029/2003JE002121. Acta 43, 1601–1610. Squyres, S.W., Arvidson, R.E., Blaney, D.L., Clark, B.C., Crumpler, L., Farrand, W.H., Morris, R.V., Klingelhöfer, G., Schröder, C., Fleischer, I., Ming, D.W., Yen, A.S., Gellert, R., Gorevan, S., Herkenhoff, K.E., Hurowitz, J., Kusack, A., McSween, H.Y., Ming, D.W., Arvidson, R.E., Rodionov, D.S., Crumpler, L.S., Clark, B.C., Cohen, B.A., McCoy, T.J., Morris, R.V., Ruff, S.W., Wang, A., Yen, A., 2006. Rocks of the Columbia Hills. Mittlefehldt, D.W., Schmidt, M.E., de Souza Jr., P.A., Squyres, S.W., 2006. Mössbauer J. Geophys. Res. 111, E02S11. doi:10.1029/2005JE002562. mineralogy of rock, soil, and dust at Gusev Crater, Mars: Spirit's journey through Sun, S.-s, McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: weakly altered olivine basalt on the plains and pervasively altered basalt in the implications for mantle composition and processes. Geol. Soc. London, Special Pub. Columbia Hills. J. Geophys. Res. 111, E02S13. doi:10.1029/2005JE002584. 42, 313–345. doi:10.1144/GSL.SP.1989.04.01.19. Morris, R.V., Klingelhöfer, G., Schröder, C., Fleischer, I., Ming, D.W., Yen, A.S., Gellert, R., Symes, S.J.K., Borg, L.E., Shearer, C.K., Irving, A.J., 2008. The age of the Martian meteorite Arvidson, R.E., Rodionov, D.S., Crumpler, L.S., Clark, B.C., Cohen, B.A., McCoy, T.J., Northwest Africa 1195 and the differentiation history of the shergottites. Geochim. Mittlefehldt, D.W., Schmidt, M.E., de Souza Jr., P.A., Squyres, S.W., 2008. Cosmochim. Acta 72, 1696–1710. doi:10.1016/j.gca.2007.12.022. mineralogy and aqueous alteration from Husband Hill through Home Plate at Tanaka, K.L., Scott, D.H., Greeley, R., Tanaka, K.L., Scott, D.H., Greeley, R., 1992. Global Gusev Crater, Mars: results from the Mössbauer instrument on the Spirit Mars Stratigraphy. Mars. University of Arizona Press, Tuscon, pp. 345–382. Exploration Rover. J. Geophys. Res. 113, E12S42. doi:10.1029/2008JE003201. Taylor, L.A., Nazorov, M.A., Shearer, C.K., McSween Jr., H.Y., Cahill, J., Neal, C.R., Ivanova, Neal, C.R., Taylor, L.A., Ely, J.C., Jain, J.C., Nazarov, M.A., 2001. Detailed geochemistry of M.A., Barsukova, L.D., Lentz, R.C., Clayton, R.N., Mayeda, T.K., 2002. Martian new shergottite, Dhofar 019. Lunar Planet. Sci. XXXII abstract #1671. meteorite Dhofar 019: a new shergottite. Meteorit. Planet. Sci. 37, 1107–1128. Neumann, G.A., Zuber, M.T., Wieczorek, M.A., McGovern, P.J., Lemoine, F.G., , D.E., Treiman, A.H., Drake, M.J., Janssens, M.-J., Wolf, R., Ebihara, M., 1986. Core formation in 2004. Crustal structure of Mars from gravity and topography. J. Geophys. Res. 109, the Earth and Shergottite Parent Body (SPB): chemical evidence from basalts. E08002. doi:10.1029/2004JE002262. Geochim. Cosmochim. Acta 50, 1071–1091. Nimmo, F., Stevenson, D.J., 2001. Estimates of Martian crustal thickness from viscous Treiman, A.H., 2005. The nakhlite meteorites: augite-rich igneous rocks from Mars. relaxation of topography. J. Geophys. Res. E3, 5085–5098. Chem. Erde 65, 203–270. Norman, M.D., 1999. The composition and thickness of the crust of Mars estimated from Usui, T., McSween, H.Y., Floss, C., 2008. Petrogenesis of olivine-phyric shergottites rare earth elements and neodymium isotopic compositions of Martian meteorites. Yamato 980459, revisited. Geochim. Cosmochim. Acta 72, 1711–1730. doi:10.1016/ Meteorit. Planet. Sci. 34, 439–449. j.gca.2008.01.011. Philpotts, A.R., 1990. Priciple of Igneous and Metamorphic Petrology. Prentice Hall, Wadhwa, M., 2001. Redox state of Mars upper mantle from Eu anomalies in shergottite Upper Saddle, N.J.. 498p. pyroxenes. Science 291, 1527–1530. Righter, K., 2003. Metal–silicate partitioning of siderophile elements and core Wadhwa, M., McSween, H.Y., Crozaz, G., 1994. Petrogenesis of shergottite meteorites formation in the Early Earth. An. Rev. Earth Planet. Sci. 31, 135–174. inferred from minor and trace element microdistributions. Geochim. Cosmochim. Righter, M., Lapen, T.J., Brandon, A.D., Beard, B.L., Shafer, J.T., Peslier, A.H., 2009. Lu–Hf Acta 58, 4213–4229. age and isotope systematics of ALH84001. Lunar. Planet Sci. Con. Abs# 2256. Wänke, H., Dreibus, G., 1988. Chemical composition and accretion history of terrestrial Roeder, P.L., Emslie, R.F., 1970. Olivine–liquid equilibrium. Contrib. Mineral. Petrol. 29, planets. Philos. Trans. Royal Soc. London 325, 545–557. 275–289. Wänke, H., Brückner, J., Dreibus, G., Reider, R., Ryabchikov, I., 2001. Chemical Rubin, A.E., Warren, P.H., Greenwood, J.P., Vervish, R.S., Leshin, L.A., Hervig, R.L., Clayton, composition of rocks and soils at the Pathfinder site. Space Science Rev. 96, R.N., Mayeda, T.K., 2000. Los Angeles: the most differentiated basaltic Martian 317–330. doi:10.1023/A:101196725645. meteorite. Geology 28, 1011–1014. doi:10.1130/0091-7613(2000) 28b1011: Warren, P.H., Kallemeyn, G.W., Arai, T., Kaneda, K., 1996. Compositional-petrologic LATMDBN2.0.CO;2. investigations of and the QUE 94201 shergottite. Proc. NIPR Symposium for Schmidt, M.E., Ruff, S.W., McCoy, T.J., Farrand, W.H., Johnson, J.R., Gellert, R., Ming, D.W., Anarctic Meteorites 21st. National Institute of Polar Research, Tokyo, pp. 195–197. Morris, R.V., Cabrol, N., Lewis, K.W., Schroeder, C., 2008. Hydrothermal origin of Warren, P.H., Kallemeyn, G.W., Kyte, F.T., 1999. Origin of planetary cores: evidence from halogens at Home Plate, Gusev Crater. J. Geophys. Res. 113, E 06S12. doi:10.1029/ highly siderophile elements in Martian meteorites. Geochim. Cosmochim. Acta 63, 2007JE003027. 2105–2122. doi:10.1016/S0016-7037(99)00156-8. Schmidt, M.E., Farrand, W.H., Johnson, J.R., Schröder, C., Hurowitz, J.A., McCoy, T.J., Ruff, Wieczorek, M.A., Zuber, M.T., 2004. Thickness of the Martian crust: improved constrains S.W., Arvidson, R.E., Des Marais, D.J., Lewis, K.W., Ming, D.W., Squyres, S.W., de for geoid-to-topography ratios. J. Geophys. Res. 109, E01009. doi:10.1029/ Souza Jr., P.A., 2009. Spectral, mineralogical, and geochemical variations across 2003JE002153. Home Plate, Gusev Crater, Mars indicate high and low temperature alteration. Earth Wyatt, M.B., McSween, H.Y., 2001. Spectral evidence for weathered basalt as an Planet. Sci. Lett. 281, 258–266. alternative to in the northern lowlands of Mars. Nature 417, 263–266. Schubert, G., Solomon, S.C., Turcotte, D.L., Drake, M.J., Sleep, N.H., 1992. Origin and Yen, A., Mittlefehldt, D.W., McLennan, S.M, Gellert, R., Bell III, J.F., McSween, H.Y., Ming, thermal evolution of Mars. Mars. University of Arizona Press, Tuscon, pp. 147–183. D.W., McCoy, T.J., Morris, R.V., Golombek, M., Economou, T., Madsen, M.B., Wdowiak, Scott, D.H., Tanaka, K.L. 1986, Geologic map of the western equatorial region of Mars, T., Clark, B.C., Joliff, B.L., Schröder, C., Brückner, J., Zipfel, J., Squyres, S.W., 2006. Ni on Map I-1802-A, 1L15,000,000, U.S. Geological Survey, Boulder, CO. Mars: constraints on meteoritic material at the surface. J. Geophys. Res. 111, E12S11. Shih, C.-Y., Nyquist, L.E., Bogard, D.D., McKay, G.A., Wooden, J.L., Bansal, B.M., doi:10.1029/2006JE002797. Weismann, H., 1982. Chronology and petrogenesis of young achondrites, Shergotty, Zuber, M.T., 2000. The crust and mantle of Mars. Nature 412, 220–227. Zagami, and ALH77005: late volcanism on a geologically active planet. Geochim. Cosmochim. Acta 46, 2323–2344.