Silicate mineralogy of martian meteorites

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Geochimica et Cosmochimica Acta 73 (2009) 7443–7485 www.elsevier.com/locate/gca

Invited Review

Silicate mineralogy of martian meteorites J.J. Papike, J.M. Karner, C.K. Shearer, P.V. Burger * Institute of Meteoritics, Department of Earth and Planetary Sciences, 1 University of New Mexico, Albuquerque, NM 87131, USA Received 19 December 2008; accepted in revised form 2 September 2009; available online 17 September 2009

Abstract Basalts and basaltic cumulates from Mars (delivered to Earth as meteorites) carry a record of the history of that planet – from accretion to initial differentiation and subsequent volcanism, up to recent times. We provide new microprobe data for plagioclase, olivine, and pyroxene from 19 of the martian meteorites that are representative of the six types of martian rocks. We also provide a comprehensive WDS map dataset for each sample studied, collected at a common magnification for easy comparison of composition and texture. The silicate data shows that plagioclase from each of the rock types shares similar trends in Ca–Na–K, and that K2O/Na2O wt% of plagioclase multiplied by the Al content of the bulk rock can be used to determine whether a rock is “enriched” or “depleted” in nature. Olivine data show that meteorite Y 980459 is a primitive melt from the martian mantle as its olivine crystals are in equilibrium with its bulk rock composition; all other olivine-bearing Shergottites have been affected by fractional crystallization. Pyroxene quadrilateral compositions can be used to isolate the type of melt from which the grains crystallized, and minor element concentrations in pyroxene can lend insight into parent melt compositions. In a comparative planetary mineralogy context, plagioclase from Mars is richer in Na than terrestrial and lunar plagioclase. The two most important factors contributing to this are the low activity of Al in martian melts and the resulting delayed nucleation of plagioclase in the crystallizing rock. Olivine from martian rocks shows distinct trends in Ni–Co and Cr systematics compared with olivine from Earth and Moon. The trends are due to several factors including oxygen fugacity, melt compositions and melt structures, properties which show variability among the planets. Finally, Fe–Mn ratios in both olivine and pyroxene can be used as a fingerprint of planetary parentage, where minerals show distinct planetary trends that may have been set at the time of planetary accretion. Although the silicate mineralogical data alone cannot support one specific model of martian magmatism over another, the data does support the basic igneous reservoirs proposed for Mars, and may also be used to constrain some aspects of specific petrogenetic models. Examples include enriched and depleted reservoirs that can be identified by plagioclase K, Na and Al composition, multivalent element partitioning in olivine and pyroxene (V, Cr) elucidates oxygen fugacity conditions of the reservoirs, and minor element concentrations (i.e., Cr in pyx) show that proposed fractional crystallization models linking Y 980459 to QUE 94201 will not work. Ó 2009 Elsevier Ltd. All rights reserved.

1. INTRODUCTION 1.1. Preamble Over the last 10 years there has been an amazing amount of activity related to the scientific study and exploration of * Corresponding author. Tel.: +1 505 277 8327; fax: +1 505 277 3577. E-mail address: [email protected] (P.V. Burger).

0016-7037/$ - see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2009.09.008

Mars. NASA, together with international partners, has initiated several orbital missions. Likewise, landed missions with rovers (e.g., Spirit and Opportunity) have explored Mars from 2004 to the present and have shed great light on the mineralogy and geology of the planet’s surface. In turn, data collected from these missions has generated intense work on theoretical calculations and experiments concerning the stability fields of Mars relevant alteration phases, which include sulfates and sheet silicates (phyllosilicates, clay minerals).

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In addition to this wealth of new information from the flight missions (and theoretical and experimental studies that have followed) there has been concentrated activity in the study of martian meteorites. These samples are basalts and basaltic cumulates that hold a record of volcanism on Mars over most of its history since accretion and differentiation, to recent times. More importantly, the martian meteorites provide not only “ground truth” for the surface (i.e., primary, igneous) mineralogy of Mars but also a starting point for which the secondary surface mineralogy of Mars (i.e., alteration phases) can be derived. These meteorites from Mars are the focus of this paper. 1.2. Objectives of paper There have been literally thousands of papers published on martian meteorites (Meyer, 2008) and their appearance in scientific journals and even the discovery of new martian meteorites is difficult to keep up with. In fact, the pace of activity is so intense that there has been little time to review and synthesize what we now know about these meteorites. The McSween and Treiman (1998) review in the book “Planetary Materials” included description and discussion of a dozen then identified martian meteorites, but since then there have been 40 more martian meteorites discovered! The focus of this review is on the major silicate minerals (plagioclase, olivine, and pyroxene) found in martian meteorites. These three phases also most likely dominate the primary surface mineralogy of Mars. Several studies have shown that the silicate minerals in basalts reflect the differing chemical and physical conditions from which they crystallized (e.g., Papike, 1981, 1998; Papike et al., 2003, 2005; Karner et al., 2006), and therefore planetary mineralogy studies are a viable means to compare planetary bodies and processes. The focus on individual minerals, rather than rocks, was prompted by early studies of the lunar regolith, that is largely composed of individual minerals rather than representative rock samples (Papike et al., 1982). It is likely that the martian regolith will also be dominated by individual minerals (derived from basalts) rather than representative rocks. These individual basaltic minerals will give clues to their parental melts and perhaps even possible source localities on Mars. Therefore, the objectives of this paper are to (1) introduce the different martian meteorite groups and several specific meteorites that are representative of these rock types, (2) review what we know about the representative meteorites we have chosen for this study, and (3) examine the compositional attributes, i.e., major, minor and select trace element chemistry (Co, Ni in olivine) and zoning trends of the silicate minerals in these martian meteorites that reflect the basaltic melts from which they came. The silicate data in turn will allow us to (4) compare and contrast the silicate mineralogy of martian rocks with those from the Earth, Moon, and 4 Vesta. 2. INTRODUCTION TO MARTIAN METEORITES As of the writing of this paper (May, 2009), there are 53 unique/unpaired martian meteorites in the world’s collec-

tions (see http://www.imca.cc/mars/martian-meteorites.htm for an up to date list). The first martian meteorites discovered (Shergotty, Nakhla, and Chassigny), were all observed falls, coming to Earth in 1865, 1911, and 1815, respectively. Since these falls, 49 more martian meteorites have been found, while only one more sample (Zagami, in 1962) has been seen to fall. Of the total finds, 16 have been found on the Antarctic continent since 1977 as a result of organized meteorite recovery expeditions; 8 by the US, 6 by the Japanese, and 2 by the Chinese. The vast majority of the rest of the finds (30) have been discovered in the hot deserts of northwest Africa and the Saudi Arabian peninsula, at a rate of about 2 per year since 1997 (Meyer, 2008). Martian meteorites were originally known as Shergottites, Nakhlites, and Chassignites, i.e., SNCs, an acronym referring to the first three martian meteorites discovered. Shergottites are basalts and basaltic cumulates composed of pyroxene and plagioclase plus or minus olivine, showing igneous textures; Nakhlites are cumulates of calcium-rich (clino-) pyroxene with minor olivine; and Chassignites are Dunites composed of greater than 90% olivine. Today we recognize additional rock types from Mars including Lherzolites, coarse grained ultramafic rocks of olivine and two pyroxenes and similar to martian basalts, and one Orthopyroxenite (cumulate rock consisting of almost all orthopyroxene). For the purpose of this paper we try and avoid antiquated nomenclature and divide the meteorites into six groups based on rock type, which are Pyroxene-phyric basalts, Olivine-phyric basalts, Lherzolites (note that historically all three are subtypes of the Shergottites), Clinopyroxenites (traditionally Nakhlites), Orthopyroxenites, and Dunites (traditionally Chassignites). A more detailed introduction to the six rock groups and individual meteorites used in this study is given in Section 4. Fig. 1 provides a comparison of the average modal proportions of the silicate phases that makeup each of the six rock types. Pyroxene analyses were divided into low- and high-Ca groups by using both Ca X-ray maps and quantitative electron microprobe (EMP) point analyses. HighCa pyroxenes are those with >30% Wo (augite) and lowCa pyroxenes (orthopyroxene plus pigeonite) are those that contain <20% Wo. It is important to note that out of all the martian meteorites only two, Y 980459 (an Olivine-phyric basalt) and QUE 94201 (a Pyroxene-phyric basalt), have been proven to be liquids when crystallization commenced. This means that the huge majority of the martian meteorites have been affected by crystal accumulation or subtraction and their modal abundance of phases reflect these processes. It should also be realized that we almost certainly do not have a representative sample of the igneous rocks that occur near the surface of Mars. Fig. 2 plots the age distributions for several martian meteorites compared with the timing of other events in martian history. The figure clearly shows three distinct age groupings for martian meteorites that include all rock types: Shergotty, along with the other martian basalts and the Lherzolites, crystallized from 165 to 475 Ma ago; Nakhla, other Clinopyroxenites, and Chassigny (Dunite) all crystallized around 1300 Ma ago; and the oldest martian meteorite, ALH 84001 (Orthopyroxenite),

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100% 90% 80%

Modal mineralogy

70% 60%

Plagioclase Hi-Ca pyx

50%

Low-Ca pyx 40%

Olivine

30% 20% 10% 0% Ol- phyric

Lherz.

Pyx- phyric

Clinopyx.

Orthopyx.

Dunite

Martian suite

Fig. 1. Average modal mineralogy for the meteorites in this study (normalized to 100% for the phases noted) for the six rock groups. Complete modes are given in the petrography section and appropriate references are given there. CIA (i.e., solar system) formation

Shergotty (165 Ma)

Nakhla (1260 Ma)

Chassigny (1320 Ma)

ALH 84001 (4500 Ma) Mars silicate differentiation Mars core formation

0

1

2

3

such as Y 980459 and QUE 94201 (Fig. 3c). These depleted rocks are characterized by low LREE concentrations and a very steep slope in a chondrite normalized REE pattern defined by La/Yb < 0.3. Depleted rocks also equilibrated at relatively low fO2 conditions (IW to IW + 1) (Wadhwa, 2001; Herd, 2003, 2006; Shearer et al., 2006; Karner et al., 2007a). At the other end of the continuum are the “enriched” rocks such as Shergotty, NWA 1110/1068, and Los Angeles (Fig. 3a). These rocks have flat REE patterns 10X chondrite, and crystallized at relatively higher fO2 conditions, most near FMQ (see Herd, 2006 for discussion and references). Therefore, the martian basalts and Lherzolites experienced a significant range of fO2 from  IW to FMQ but none are as reduced as basalts from the Moon and 4 Vesta, that crystallized at fO2’s  IW 1.

4

Age (Ga)

Fig. 2. Absolute ages of select martian meteorites and events (modified from Borg and Drake (2005)). Shergotty and other martian basalts and Lherzolites are in red, Nakhla and other Clinopyroxenites are in blue, Chassigny (Dunite) is in green, and ALH 84001 (Orthopyroxenite) is in black.

crystallized 4500 Ma ago. Mars has certainly experienced more than three distinct volcanic episodes in its history, and thus this figure serves to illustrate that our sample suite is biased (non-representative), most likely by the impact events that delivered these samples to Earth. Further information pertinent to this introduction is that the martian basalts (Olivine-phyric and Pyroxene-phyric) and Lherzolites display a significant range of trace element and isotopic characteristics (McSween, 1994; Borg et al., 2002). Fig. 3 illustrates this range for bulk rock REE patterns for representative martian basalts and Lherzolites. At one end of a trace element and isotopic continuum are the “depleted” rocks, represented by meteorites

3. ANALYTICAL APPROACH The following silicate data is new, all collected at the Institute of Meteoritics’ microbeam facilities using the EMP and Secondary Ion Mass Spectrometer (SIMS). Plagioclase, olivine and pyroxene grains in thin sections from 19 samples (Table 1) were analyzed to determine their major and minor element compositions. Analyses were made using a JEOL 733 Superprobe equipped with a back-scattered electron detector, a thin-window energy dispersive spectrometer, and five wavelength dispersive spectrometers, all controlled by an Oxford eXL II analyzer system. Plagioclase analysis conditions were an accelerating voltage of 15 kV, a beam current of 20 nA, and a beam diameter of 10 lm. A beam diameter of 10 lm minimized loss of volatile elements such as Na and K during feldspar analyses. Olivine and pyroxene grains were analyzed under the same conditions as plagioclase, except that the beam size was 1 lm. Wavelength dispersive spectrometer counting times of 20 s were used for major elements and 30–40 s for minor

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La/Yb> 0.7

Concentration normalized to CI

(a) Enriched

10.0

1.0

NWA 1110/1068 Los Angeles Shergotty RBT 04262 0.1 100.0

Concentration normalized to CI

(b) Intermediate

La/Yb= 0.3 to 0.7

10.0

1.0

ALH 77005 LEW 88516 0.1 100.0

Concentration normalized to CI

(c) Depleted

La/Yb< 0.3

10.0

1.0

Y 980459 DaG 476 SAU 005 QUE 94201 0.1

La

Ce

Pr Nd Sm Eu Gd Tb Dy

Ho

Er Tm Yb Lu

Fig. 3. REE patterns (normalized to CI-chondrite) for select meteorites from the (a) enriched, (b) intermediate, and (c) depleted groupings. REE bulk rock concentrations taken from Lodders (1998) (Shergotty, QUE 94201, LEW 88516, ALH 77005); Shirai and Ebihara (2004) (Y 980459); Dreibus et al. (2003) (NWA 1110, SAU 005, DaG 476); Jambon et al. (2002) (LA) and Anand et al. (2008) (RBT 04262).

elements in all three silicate phases. All data were reduced using a ZAF correction program. The detection limits reported are defined as three times the counting statistic standard deviation.

In order to ensure high-quality EMP data, each analysis for the three silicate phases had to meet several criteria: Plagioclase analyses were acceptable if (1) the total oxide sum was equal 100 ± 2 wt%, (2) the sum of the tetrahedral site

Silicate mineralogy of martian meteorites Table 1 Martian meteorite reference suite with number of new, highquality, EMP analyses from this study. ol

pyx

plag

Ol-phyric Y 980459 SAU 005 NWA 1110 DaG 476 NWA 1195 NWA 2046 NWA 2626

45 79 35 45 42 31 46

47 54 67 75 40 49 44

NA 42 40 46 49 42 39

Lherzolitic ALH 77005 LEW 88516 RBT 04262

47 37 88

75 39 70

27 47 49

Pyx-phyric Los Angeles NWA 3171 QUE 94201 Shergotty

NA NA NA NA

48 94 65 74

60 47 31 46

49 19 45

44 52 44

28 NA 14

NA

38

20

36

17

13

644

1036

640

Clinopyroxenites Nakhla MIL 03346 Governador Valadares Orthopyroxenites ALH 84001 Dunites Chassigny Total new analyses

cations, which includes Si, IVAl, and Fe3+ was equal to 4.00 ± 0.03 per 8 O atoms, and (3) the sum of the cations in the A-site (Ca, Na, and K) was equal to 1.00 ± 0.02. Likewise, pyroxene analyses were accepted if (1) the oxide total was between 98 and 102 wt%, (2) the sum of the tetrahedral site cations, which includes Si and IVAl, equaled 2 ± 0.02 atoms per six oxygens, and (3) octahedral cations in the M1 and M2 sites (Mn, Fe2+, Fe3+, Mg, Ti, Cr, VIAl, Ca, Na) summed to 2.0 ± 0.02, and (4) the charge balance equation {VIAl + VIFe3+ + VICr3+ + 2Ti4+ = IVAl + M2Na} balanced to within 0.03 charge. Ferric iron in pyroxene was estimated for each analysis using the method described by Droop (1987). Lastly, olivine analyses were considered superior if (1) the oxide total was 100 ± 2 wt%, (2) the sum of the tetrahedral site (only Si) was 1 ± 0.02, and (3) the total of all cations was 3 ± 0.03. Trace element analysis of olivine for Ni and Co was performed using a Cameca IMS 4f SIMS by focusing a primary beam of O ions with an accelerating voltage of 10 kV onto the sample. A beam size of 15 lm was obtained with a current of 15 nA. Sputtered secondary ions were energy filtered using an offset voltage of 105 V, and an energy window of ±25 V in order to reduce isobaric interferences. The analytical procedure involved repeated cycles of peak counting on the trace elements 30Si, 58Mn, 59 Co, and 60Ni. Absolute concentrations of the trace elements were calculated using the relationship between measured peak/30Si ratios normalized to known SiO2 content

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and elemental abundance in the olivine standards. Three olivine standards and one pyroxene standard were used to produce the calibration curves. The olivine standards were Marjalahti (Mason and Graham, 1970, additional analyses from Washington University INAA lab in St. Louis); San Carlos, and DW/D1 (INAA, Max-Planck-Institute, Mainz, Germany). The pyroxene standard was KAUG augite (Mason and Allen, 1973). Wavelength dispersive (WDS) X-ray maps were collected for Mg and Ca using an accelerating voltage of 15 kV and a beam current of 100 nA. Analytical map areas were determined by BSE imaging and illustrate the representative mineralogy of each sample. An equivalent scale for the images was maintained throughout the entire dataset. In addition, the contrast of each image (i.e., the relationship between the measured intensity and the color range in the resultant image) was shifted to maintain consistency from one image to another. This was accomplished by shifting the color spectrum so that equivalent mineral phases were the same color in all images within a particular rock suite (e.g., Olivine-phyric basalts). The focus on consistency from one meteorite map image to another was intended to provide a basis for comparison among all of the meteorites in this study. 4. PETROGRAPHY AND PREVIOUS WORK 4.1. Introduction For the purpose and scope of this study we focus on 19 of the 53 known martian meteorites. These 19 were chosen because they are representative of one of the six martian rock groupings. The following sections elaborate on the six rock groupings and the individual meteorites used in this study. Each section gives some basic background and history followed by a brief petrographic description of each sample with the emphasis on the silicate mineralogy. Meteorite geochemistry and geochronology is also discussed where relevant, along with specific research that may be notable or unique to a particular meteorite. We note that our review of any one meteorite is not comprehensive. We also note that all of the meteorites discussed have been “confirmed” to be of martian origin by their distinct oxygen isotopic signatures (Clayton and Mayeda, 1983), a topic which will not be examined further here. Those interested in the oxygen isotope data, along with a more complete cataloguing of martian meteorites and the work that has been done on each, should consult the outstanding Mars Meteorite Compendium (Meyer, 2008), online at http://wwwcurator.jsc.nasa.gov/antmet/mmc/index.cfm. 4.2. Olivine-phyric basalts Olivine-phyric basalts share several petrographic features including olivine–porphyritic textures, presence of chromite in addition to other Fe–Ti oxides, and low augite contents, which distinguish them from the other martian basalt groups (Goodrich, 2002). Sixteen meteorites fit into this grouping and we include 7 of the 16 in our reference suite. They are Y 980459, SAU 005, NWA 1110/1068, DaG 476, NWA 1195, NWA 2046 and NWA 2626.

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Y 980459 was found by the Japanese in 1998, near the Minami-Yamato Nunataks in northern Antarctica (Meyer, 2008) and is described as a mafic, Olivine-phyric basalt that contains no plagioclase or maskelynite (just quenched glass). The sample is very important because it has been proven to be a primitive “melt” (Koizumi et al., 2004; Usui et al., 2008a) from the martian mantle. In other words, the rock has experienced little accumulation or subtraction of phases before eruption, and more importantly, the liquidus minerals formed in equilibrium with the bulk composition of the rock. Mikouchi et al. (2004) report the modal mineralogy as 26% olivine, 48% pyroxene, and 25% glassy mesostasis. Fig. 4a shows a euhedral olivine megacryst within a matrix of smaller subhedral olivines, low-Ca pyroxenes and quench-textured glass. WDS maps suggest the Mg composition of megacryst rims is similar to that of the later-formed, small olivine grains. Both generations of olivine are rimmed by a narrow band, enriched in Fe relative to the remainder of the grain. Low-Ca pyroxene is most magnesian at the core of the grains, zoning toward greater Ca enrichment at the rims. Much as with olivine, a narrow band of Mg depletion (and Ca enrichment) occurs at the rim of the pyroxene grains. The Sm–Nd age of Y 980459 has been determined to be 472 Ma, while an Rb–Sr age has been difficult to obtain because of the absence of plagioclase and possible terrestrial contamination in the sample (Shih et al., 2005). By examining V partitioning in olivine megacrysts, Shearer et al. (2006) determined the fO2 of the sample to be IW + 1. Pieces of SAU 005 were first found in the deserts of Oman in 1999. Nine pieces appear to be paired and part of an apparent strewn field that is 2.5 km long; 11 kg have been collected. Fig. 4b shows porphyritic olivine (subhedral) set in a ground mass of pyroxene (mostly pigeonite), plagioclase (maskelynite), and later stage olivine. Porphyritic olivine grains show growth zoning from Mg-rich cores to more Feenriched rims. Like Y 980459, the Mg# of rims from the olivine phenocrysts seems to match the composition of smaller, later olivine crystals (Fig. 4b), but the narrow zone of Feenrichment seen in Y 980459 olivines does not occur in SAU olivines. Pyroxene is somewhat coarser grained relative to Y 980459, and inspection of the Ca WDS map reveals the maskelynite is somewhat heterogeneous. The sample mode is 25% olivine, 48% pyroxene, 15% plagioclase and trace amounts of opaques (Zipfel, 2000). The rock is severely shocked as evidenced by the total conversion of plagioclase to maskelynite, pyroxene twinning and fracturing, and olivine mosaicism (Zipfel, 2000). Three studies, using independent methods to calculate fO2, all indicate this meteorite crystallized in a very reducing environment, IW (Goodrich et al., 2003; Herd, 2003; Shearer et al., 2006). An age determination has yet to be made for this sample. NWA 1110/1068 are presumed to be paired meteorites found in the Saharan Desert in 2001. Twenty-two pieces (700 g total) of this lithology have been recovered from an undocumented strewn field. The sample examined in this study (NWA 1110) is characterized by phenocrysts of olivine (up to 2 mm in length) set in a basaltic matrix of low and highCa pyroxene and plagioclase (maskelynite) (Fig. 4c). Olivine grains show normal zoning from Mg-rich cores to Fe-rich rims. Olivine megacrysts often display corroded grain

boundaries and frequently occur in clusters that appear to be broken apart and intruded by groundmass material (Fig. 4c). These characteristics suggest the olivine grains may be xenocrysts (Barrat et al., 2002; Shearer et al., 2008). The matrix, consisting of low- and high-Ca pyroxenes has a finer grain size than SAU 005, possibly the result of quick cooling during the later stages of crystal growth. Barrat et al. (2002) observed a mineral mode of 21% olivine, 52% pyroxene, 22% plagioclase, 2% phosphate, 2% opaques, and 1% mesostasis. The apparent crystallization age of the meteorite is 185 Ma which is typical of martian basalts (Shih et al., 2003). Oxygen fugacity work by Goodrich et al. (2003) and Herd (2006) suggests the rock may have experienced a dramatic increase in fO2 during crystallization (from QFM 1.7 to QFM), indicating a non-buffered system. The Dar Al Gani meteorites (DaG) were found in the northwest African Sahara desert in 1997 and 1998. These meteorites were the first from a then fairly untapped collection area (i.e., hot deserts) that have since yielded thousands of meteorites. Our suite sample DaG 476 (2015 g) is paired with several fragments from an assumed strewn field (total mass 6 kg) (Meyer, 2008). DaG 476 consists of olivine phenocrysts/megacrysts (up to 5 mm) set in a groundmass of smaller pyroxene, plagioclase (maskelynite) and mesostasis (Fig. 4d). Olivine grains are euhedral to subhedral, the rims are devoid of any significant resorption texture, and there is no zone of extreme Fe-enrichment at the boundary of the grains, as was noticed in Y 980459 olivines. Olivine phenocrysts do show subtle growth zoning from a higher Mg core to a more Fe-enriched rim. Pyroxene grains are subhedral to anhedral, with a texture and grain size similar to SAU 005 and NWA 1195. Pyroxene crystallization appears to have begun with the nucleation of orthopyroxene, followed by a period of pigeonite crystallization. Pyroxenes of augitic composition occur as well, likely crystallizing later in the sequence. A mode by Zipfel et al. (1999) reports 14% olivine, 52% pigeonite, 3% augite, 3% opx, 17% plagioclase, 4% opaques, and 7% others which includes carbonate, impact glass, and phosphate. Both the mineral chemistry and bulk rock chemistry demonstrate that this is an ultramafic rock with Mg#68 (Zipfel, 2000). Borg et al. (2003) determined the Sm– Nd age of DaG 476 to be 474 Ma, which is younger than a previous determination by Jagoutz et al. (1999) of 703 Ma. The discrepancy could be due to the extreme desert weathering this meteorite has experienced. Oxygen fugacity studies indicate the meteorite crystallized approximately 1–2 log units below the QFM buffer (Herd et al., 2001; Herd, 2003). Irving et al. (2002) reported a 50 g stone was obtained near Safsaf, Morroco and was given the name NWA 1195. The rock has subhedral olivine megacrysts (4 mm) set in a matrix of low-Ca pyroxene, maskelynite, with minor Ti-chromite, ilmenite, phosphate, and sulfide (Fig. 4e). Olivine megacrysts show normal zoning and the textures seen in the sample are very similar to those of SAU 005 and NWA 1110. Some low-Ca pyroxene grains have orthopyroxene core compositions and these grains are reported to be aligned parallel to the long dimension of the stone (Irving et al., 2004a). Similar to other Olivine-phyric samples, there is a debate as to whether olivine megacrysts are phenocrysts or xenocrysts (see Shearer

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Fig. 4. Back-scattered electron (BSE) image followed (from left to right) by Mg and Ca X-ray maps of the 19 different meteorites in this study. Abbreviations are as follows: Pyx, pyroxene; Pig, pigeonite; Aug, augite; Opx, orthopyroxene; Plag, plagioclase; Ol, olivine; and Spin, spinel. In the X-ray maps, warmer colors denote higher concentrations; note the color code bar and scale bar to the right of the images.

et al., 2008). We determined the mode for this sample as 67% pyroxene, 13% olivine, 13% plagioclase, 6% oxides, and 1% sulfides and others. Recently, Symes et al. (2008) obtained an age of 347 ± 13 Ma using Sm–Nd techniques;

they also noted that the Rb–Sr system is disturbed and does not yield an isochron. NWA 2046 was purchased in Morroco in 2003 and has a mass of 63 g (Russell et al., 2004). The sample is de-

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Fig. 4 (continued)

scribed by Irving et al. (2004a) as olivine megacrysts in a basaltic matrix of low-Ca pyroxene, maskelynite, and minor spinel, oxides, phosphates, and fayalite. Fig. 4f shows the texture of the sample with two larger phenocrysts of low-Ca (orthopyroxene) pyroxene cores zoning to pigeon-

ite. Olivine megacrysts are generally anhedral, and often surrounded by pyroxene grains. Irving et al. (2004a) note a preferred orientation of both olivine and orthopyroxene phenocrysts, suggesting magmatic flow and/or crystal accumulation. No other data, such as mineral modes, bulk

Silicate mineralogy of martian meteorites

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Fig. 4 (continued)

chemistry, or age determinations have been reported in the literature. The last Olivine-phyric basalt in our sample suite is NWA 2626, which is a 31 g stone purchased in Morroco in 2004 (Irving et al., 2005). The sample is very similar pet-

rographically to NWA 1195 and NWA 2046 in that it features olivine megacrysts set in a matrix of smaller low-Ca pyroxene and plagioclase (maskelynite) grains, along with the usual accessory spinels, oxides, phosphates, sulfides, and fayalite (Fig. 4g). Olivines show typical growth zoning

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Fig. 4 (continued)

from Mg-rich cores to Fe-enrichment at the periphery of the grains. The boundaries of both the olivine and pyroxene grains are irregular which suggests resorption. Crosscutting glass-rich veinlets and pockets are also noted and are assumed to be shock-induced. Overall, NWA 1195,

2046, and 2626 are very similar, though Irving et al. (2005) suggested the mineral compositions in each are dissimilar enough to classify them as unpaired. As of the writing of this paper, no further data is available for this meteorite.

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Fig. 4 (continued)

4.3. Lherzolites The second group we introduce are the Lherzolitic rocks, which take their name from terrestrial, plutonic, ultramafic rocks which contain at least 40% modal olivine along with two pyroxenes: orthopyroxene and a high-Ca pyroxene. The martian variety of Lherzolite contains from 31 to 60 modal percent olivine that is often poikilitically enclosed by pyroxene, along with minor modal plagioclase (7–15%). To date there are 12 meteorites that fit into this group and here we concentrate on three, namely ALH 77005, LEW 88516 and RBT 04262. The type specimen of the Lherzolites, ALH 77005, was found in the Allan Hills area of Antarctica in late December of 1977. The stone weighed 482 g and was nearly devoid of fusion crust (Meyer, 2008). A first mode of the sample shows it consists of approximately 55% olivine, 35% pyroxene, 8% plagioclase (maskelynite) and 2% opaques, that include chro-

mite, ilmenite and sulfides (Mason, 1981). Fig. 4h displays the rock’s coarse grained cumulate texture with anhedral to subhedral olivine grains poikilitically enclosed in pyroxene (mostly augite), with minor plagioclase and opaques interstitial to the mafic silicates. Whereas the Olivine-phyric basalts have a bimodal grain size distribution (olivine megacrysts, smaller pyroxene and/or maskelynite), the grain size of silicate minerals in ALH 77005 is more comparable. Olivine grains tend to be more equant and somewhat larger, while pyroxene and maskelynite grains are more elongate. WDS maps (Fig. 4h), reveal that neither the olivines nor the pyroxenes show appreciable growth zoning and both appear relatively homogenous. Maskelynite also shows limited chemical variability in the WDS Ca map. ALH 77005 is heavily shocked as evidenced by numerous pockets of melt glass, planar elements in olivine, twinning and mosaicism in pyroxene, and olivines with a distinct brown color, possibly caused by shock-induced oxidation of Fe2+ to Fe3+ in the crystal (McS-

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ween and Sto¨ffler, 1980; Ostertag et al., 1984; Sto¨ffler et al., 1986; Treiman et al., 1994). Age dating of ALH 77005 and the other Lherzolites suffer from the same difficulties as dating the other martian basalts, but a crystallization age of 180 Ma is generally accepted at this time (Borg et al., 2002). The second martian Lherzolite (LEW 88516) to be discovered was found in the Lewis Cliff region of Antarctica in 1988, and is petrographically quite similar to ALH 77005 (Treiman et al., 1994). The sample is very small, about 2 cm in diameter and only 13.2 g (Meyer, 2008). Fig. 4i shows anhedral olivine and chromite grains enclosed in anhedral pyroxene grains (subequal pigeonite and augite) with plagioclase (maskelynite) in the interstices. As opposed to ALH 77005, some growth zoning can be seen in the WDS maps for LEW 88516. Specifically, pigeonite grains have their highest Mg concentration at their cores and trend toward greater Fe and Ca enrichment with crystallization. Treiman et al. (1994) give the mode as 46% olivine, 37% pyroxene, 7% plagioclase, 8% glass, and less than 1% of chromite, phosphate, ilmenite and sulfides. A generally accepted crystallization age determined by Rb–Sr is 183 Ga, which is very similar to the age of ALH 77005 (Borg et al., 2002). Treiman et al. (1994) summarize that LEW 88516 is very similar to ALH 77005 in its mineralogy, geochemistry, shock features, and noble gas contents, but is not paired with that meteorite based on different terrestrial residence ages and slight differences in mineral chemistry. The last Lherzolite in our suite is RBT 04261 which was found in the Robert’s Massif area of Antarctica on Christmas Day of 2004. A paired stone (RBT 04262) was found days later in the same area and the masses of each are 79 and 205 g. A mineral mode of RBT 04262 shows 43% low-Ca pyroxene, 30% olivine, 13% plagioclase (maskelynite) 10% augite, 2% Cr-spinel, 1% phophates, and 1% others (Mikouchi et al., 2008). Like ALH 77005, most silicate grains in RBT 04262 have an anhedral morphology. Olivine grains are somewhat equant, while pyroxene grains are slightly more elongate, with maskelynite filling in the interstices of the earlier formed grains. Subtle growth zoning is seen in the olivine grains (Fig. 4j), while pyroxene grains commonly zone from augite cores to pigeonite rims. The Ca map shows the transition between these two pyroxene phases is quite abrupt. The sample has been dated by the Lu–Hf technique and shows an age of 225 ± 21 Ma. (Lapen et al., 2008). The RBT meteorite was originally classified as an Olivine-phyric basalt, but the occurrence of mm-sized olivine and subhedral chromite grains enclosed in low-Ca pyroxene (see Fig. 4j) has prompted a suggestion to re-classify the rock as a Lherzolite (Mikouchi et al., 2008). Other workers have documented that the REE concentrations in RBT minerals are most akin to the other martian basalts (Sanborn et al., 2008; Usui et al., 2008b), thus the classification is still under dispute. We believe the sample is most similar to the other Lherzolitic meteorites (at least in its texture and mineralogy) and therefore we include it here. 4.4. Pyroxene-phyric basalts The third group of meteorites is the Pyroxene-phyric basalts. These rocks are basically composed of pyroxene

(augite and pigeonite) and plagioclase with no (early crystallizing) olivine or chromite. The absence of olivine in these rocks, their low Mg#, and crystallization experiments on their bulk compositions suggests these rocks crystallized from fractionated magmas (Goodrich, 2002). There are 15 meteorites classified as Pyroxene-phyric basalts and here we concentrate on four of them: Los Angeles, NWA 3171, QUE 94201, and Shergotty. The Los Angeles (LA) meteorites (2 paired stones, total 698 g) were discovered in a private collection in 1999 and verified to be of martian origin at UCLA; they may have been collected in the Mojave desert years before (Meyer, 2008). LA is a coarse-grained basalt mainly composed of subequal amounts of pyroxene and plagioclase 2–4 mm long (Fig. 4k). Rubin et al. (2000) describe the mode as 43% plagioclase (maskelynite), 38% pyroxene, 5% silica, 4% fayalite, 3% Ca-phosphate, 3% oxides, and 2% others. Like Shergotty, LA contains both pigeonite and augite that zones from Mgrich to Fe-rich compositions, although LA pyroxenes are more complex due to prevalent exsolution lamellae (Rubin et al., 2000). Similar to Shergotty, there is preserved heterogeneity in the maskelynite (Fig. 4k). Additionally, 5–10% of the rock contains 50–200-lm sized patches of a fine-grained vermicular symplectites of fayalite, hedenbergite, and silica, which is postulated to be a possible breakdown of pyroxferroite (Rubin et al., 2000; Aramovich et al., 2002; Xirouchakis et al., 2002). Ages determined by Rb–Sr and Sm–Nd are 165 and 173 Ma, respectively (Nyquist et al., 2001), and the fO2 of this sample is reported to be QFM 1, which is roughly equal to that of Shergotty (Herd et al., 2001). Pyroxene-phyric basalt NWA 3171 (506 g), was found in western Algeria (Meyer, 2008). The sample is mostly pyroxene (pigeonite and augite) and plagioclase (maskelynite) with accessory phases of ilmenite, ulvo¨spinel, feldspathic glass, phosphates, and silica (Irving et al., 2004b). Fig. 4l shows the basaltic texture and also reveals the sample is much finer-grained than the other Pyroxene-phyric basalts. We have calculated the mode to be 70% pyroxene, 22% plagioclase, 6% oxides, and 2% others. Park and Bogard (2007) report an Ar–Ar age of 225 Ma for the sample, while Irving et al. (2004b) calculate an fO2 from the compositions of the oxides to be QFM 1.4. QUE 94201 (12 g) was found among the rocks in a glacial moraine in the Queen Alexandra Range area of Antarctica in 1994 (Meyer, 2008). The basalt contains equal amounts of pyroxene and plagioclase with a bulk composition rich in iron (Fe#75) and phosphorous (2.5 wt% P2O3). The mode for the sample is 43% pyroxene, 42% plagioclase (maskelynite) 4% opaques, 6% phophates, and 5% mesostasis (Mikouchi et al., 1996). Fig. 4m shows the coarse grained fabric of the rock and also the complex zoning sequence of Mg-pigeonite to Mg-augite to Fe-pigeonite. Pyroxene grains within this sample are subhedral to anhedral, and the Ca map shows distinct compositional variability within maskelynite. The most important characteristic of this meteorite is that it has been proven to be a melt, although it is much more evolved than the other defined martian melt (Y 980459) in that it contains no olivine and has a much greater Fe# (Kring et al., 2003; McKay et al., 2003). Dating by several methods proves to be difficult, and crystallization ages range from 327 Ma to

Silicate mineralogy of martian meteorites

1.3 Ga (see Meyer, 2008). The composition of Fe–Ti oxides (Herd et al., 2001) and the size of Eu anomalies in augite (Wadhwa, 2001) show that QUE is one of the most reduced martian basalts. Later fO2 studies using V and Cr partitioning in QUE 94201 pigeonite (Karner et al., 2007a) place its fO2 between IW and IW + 1 (similar to Y 980459). The Shergotty meteorite fell in 1865 near a town called Shergahti, in Bihar State, India, and a 5 kg stone was recovered (Meyer, 2008). Shergotty is the type specimen for all the martian basalts (originally Shergottite basalts), and as such its petrologic characteristics were key in determining that these rocks (i.e., Shergottites) came from Mars. Shergotty is severely shocked; the rock is considered the “type locality” for maskelynite, as this was the first discovered occurrence of this shocked plagioclase phase (Tschermak, 1872). Modal mineralogy for the rock is 71% pyroxene (equal pigeonite and augite), 24% plagioclase (maskelynite), 3% mesostasis, and 4% others which includes ilmenite, magnetite, sulfide and phosphate (Stolper and McSween, 1979). Fig. 4n displays the rock’s basaltic texture of elongate pyroxene and plagioclase. Note the Mg and Ca maps which display the discrete grains of pigeonite and augite, both zoning from Mg-rich cores to Fe-rich rims. The discrete grains of augite and pigeonite suggest the two phases were growing concurrently, sometimes coalescing into aggregate grains. The Ca map (Fig. 4n) also illustrates the heterogeneity within maskelynite regions. The Ca concentration of these regions appears to be highest toward the center with lower Ca near the edges and surrounding pyroxenes. It is possible that this heterogeneity is the result of growth zoning which existed in plagioclase prior to the shock event which caused maskelytization. Accessory magnetite in the sample implies relatively oxidizing crystallization conditions (Tschermak, 1872), and subsequently Shergotty is still considered the most oxidized martian basalt we have QFM 1 (e.g., Herd et al., 2001; Wadhwa, 2001; McCanta et al., 2004). Again, the ages of martian basalts have been very difficult to interpret, and dating of Shergotty by several systems (K–Ar, Rb–Sr, Ar–Ar, U–Pb) has yielded discordant results (see Jones, 1986, 1989). At present, a generally accepted crystallization age for Shergotty is 165 Ma (Nyquist et al., 2001). 4.5. Clinopyroxenites The fourth group of meteorites is the Clinopyroxenites, which are igneous cumulate rocks composed of 70–80% augite, 10% olivine, with minor amounts of other intercumulus minerals. The martian Clinopyroxenites share the same ages, approximately 1.3 Ga, and differ from the other martian rock groups in that they show little evidence of shock. Our suite consists of the meteorites Nakhla, Governador Valaderes, and MIL 033346, which are three of the seven known martian Clinopyroxenites. Numerous pieces of the Nakhla meteorite fell as a shower near the small villages of El-Nakhla, El-Bahariya in Egypt in 1911 (Meyer, 2008). Approximately 40 pieces of the meteorite were recovered with a total mass of 10 kg. The rock is classified as an olivine-bearing Clinopyroxenite and an average mode for the sample is 80% augite, 11% olivine, and 9% mesostasis (Lentz et al., 1999). Fig. 4o

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shows subhedral augite crystals set in a fine-grained crystalline matrix. Pyroxene grains show subtle zoning, with the highest Mg and Ca contents in the cores. Minor, anhedral olivine crystals are noted along with plagioclase (un-maskelynitized) that occurs interstitially along with Fe–Ti oxides (bright white grains in BSE image). Olivine grains are fairly Fe-rich, suggesting these are late-stage crystallization products. Large melt inclusions (200 lm) are common in Nakhla olivine grains and these inclusions have been employed to calculate the composition of the original magma (e.g., Treiman, 1986; Harvey and McSween, 1992; Treiman and Goodrich, 2001). Several chronometers including K–Ar, Ar–Ar, Rb–Sr, and Sm–Nd show that Nakhla is 1.2–1.3 billion years old (Stauffer, 1962; Podosek, 1973; Nakamura et al., 1977; Papanastassiou and Wasserburg, 1974), which is quite consistent with all the other martian Clinopyroxenites. Smith et al. (1983) first noted high Ni contents in both Nakhla olivine and pyroxene grains and postulated this required relatively oxidizing conditions, near the QFM buffer (e.g., see Righter et al., 2008). MIL 03346 was found in the Miller Range area of Antarctica in December 2003. The rock was partially covered by a black fusion crust and weighed 715 g (Meyer, 2008). Modal analysis of the sample shows 78% augite, 19% mesostasis, 2% olivine, and about 1% “alteration” materialfound in voids and along grain boundaries (Stopar et al., 2005). Fig. 4p shows characteristic sharp euhedral grains of augite and rare, decomposing olivine set in a fine-grained matrix. Olivine grains in MIL 03346 are more Mg-rich and show more heterogeneity than those in the other Clinopyroxenites. The matrix in MIL 03346 is more abundant and finer-grained than in either Nakhla and Governador Valadares, and contains skeletal Fe–Ti oxide grains, assorted feathery silicate minerals, and feldspathic glass (Day et al., 2006). No plagioclase/maskelynite has been found in this sample. Similar to the other Clinopyroxenites, MIL 03346 shows a crystallization age of 1.2–1.4 Ga (Bogard and Garrison, 2006; Shih et al., 2006), and formed under relatively oxidizing conditions (Hammer and Rutherford, 2008; Righter et al., 2008). The Governador Valadares meteorite was found in 1958 near the city of its namesake in southeastern Brazil. The single stone weighs 158 g and is thought to be well preserved because of its nearly complete, glassy fusion crust (Meyer, 2008). A mode for the meteorite is 74% augite, 13% olivine, and 13% mesostasis (Lentz et al., 1999). The sample is classified as an olivine-bearing Clinopyroxenite very similar to Nakhla in its mineralogy and geochemistry (e.g., Burragato et al., 1975). Fig. 4q shows the sample consists of weakly aligned augite grains, minor olivine, and mixed interstitial matter composed of crystalline plagioclase grains, Fe–Ti oxides, sulfides, and silica-rich glass. The WDS maps from Nakhla and Governador Valadares show the two meteorites are nearly identical. Similar to Nakhla, augite grains in Governador Valadares are slightly zoned and olivine grains are Ferich with abundant magmatic melt inclusions (Harvey and McSween, 1992). The occurrence of magnetite in the meteorite suggests similarly oxidizing conditions as was for Nakhla. Age determinations from several chronometers confirm that Governador Valadares crystallized at basically the same time

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as Nakhla, at 1.3 Ga (Bogard and Husain, 1977; Wooden et al., 1979; Shih et al., 1999). 4.6. Orthopyroxenite The fifth group of martian igneous rocks consists of only one sample, ALH 84001. The ALH 84001 meteorite (1.9 kg) was found in the Allan Hills region of Antarctic in 1984, and is perhaps the most intriguing martian meteorite of all. First, it is an Orthopyroxenite; a unique composition among the martian meteorites. Second, ALH 84001 is very old, 4.5 Ga (Nyquist et al., 1995), which is 3–4 billion years older than any other martian meteorite and thus records very early igneous activity on Mars. Finally, carbonate globules within this rock were purported to contain evidence of past life on Mars (McKay et al., 1996)! ALH 84001 consists of 97% coarse grained, cataclastic orthopyroxene, with 2% chromite, 1% plagioclase (maskelynite), and approximately 0.15% phosphate (Mason et al., 1992). The sample also has minor augite, olivine, sulfides and secondary Fe–Mg–Mn–Ca carbonate. Fig. 4r shows the sample is mainly composed of mm-sized shocked orthopyroxene that forms a cumulate, metamorphosed (note triple junctions of opx) texture with minor spinel and plagioclase in the interstices. Close inspection of WDS maps (Fig. 4r) reveals that none of the phases in the sample appear to be zoned. It is likely that any primary zoning in the sample was homogenized during an annealing (metamorphic) event which led to the formation of triple junctions. Several workers agree this rock went through at least one shock event and one alteration event (by a CO2rich fluid?) after crystallization (Mittlefehldt, 1994; Treiman, 1995; Gleason et al., 1997; Scott et al., 1998). ALH 84001 was initially identified as a diogenite until studies by Berkley and Boynton (1992) and Mittlefehldt (1994) found considerable ferric iron in chromites within the rock. This Fe3+ implied a more oxidized origin than those found for the diogenites, and a martian origin was later confirmed by oxygen isotopic analysis. Subsequent fO2 studies show this meteorite may actually be reduced, equilibrated at conditions of QFM 2.7 (Righter et al., 2008). 4.7. Dunites The last martian igneous lithology is the Dunites, which are classified as olivine–chromite cumulate rocks. The sample we study here is Chassigny, which is the type specimen for the group that only consists of two meteorites, the other being the recently discovered NWA 2737. Chassigny fell in 1815 in the Haute-Marne region of France, and one stone weighing 4 kg was recovered. Its modal mineralogy is 92% olivine, 5% pyroxene (high-Ca), 2% feldspar (Na, Krich), 1% chromite, and trace amounts of other minerals (Prinz et al., 1974). Magmatic melt inclusions found within olivine grains can be very large (190 lm) and common (Floran et al., 1978). Fig. 4s shows the dominance of cataclastic (i.e., planar deformation), chemically homogenous olivine along with minor anhedral to euhedral chromite and rare poikilitic high-Ca pyroxene. Several workers concur that Chassigny is a shocked and somewhat recrystallized

Fe-rich Dunite that crystallized from a fractionated melt (e.g., Prinz et al., 1974; Johnson et al., 1991). The age of Chassigny is 1.3 Ga, as has been determined by a number of systems including Rb–Sr, Ar–Ar, K–Ar, and Sm–Nd (Lancet and Lancet, 1971; Bogard and Nyquist, 1979; Nakamura et al., 1982; Jagoutz, 1996). Floran et al. (1978) noted significant Fe3+ in chromite grains which implies relatively oxidizing conditions (QFM) during crystallization. 5. PLAGIOCLASE FELDSPAR Plagioclase feldspar is a common mineral in all of the martian rock types and has been found in nearly all of the martian meteorites. Although plagioclase generally forms relatively late in most basaltic crystallization sequences, its structure incorporates several elements whose ionic radii are too large to fit into earlier-crystallizing ferromagnesian minerals such as olivine and pyroxene (see Papike and Cameron, 1976). Therefore, the chemistry of plagioclase carries a record of the chemical and physical processes of basaltic volcanism, but utilizing different elements from those that partition into olivine and pyroxene. Table 2 lists the average plagioclase chemical analyses of 17 of the martian meteorites in our reference suite. Note there is no plagioclase in Y 980459 or MIL 03346. Also note that all analyses listed are stoichiometric plagioclase, even though in most martian rocks crystalline plagioclase has been shocked to form a diaplectic glass (maskelynite), though the glass retains a plagioclase stoichiometry. Fig. 5a–e compares the major element chemistry of plagioclase in the five different suites by plotting anorthite content [(Ca/Ca + Na)  100] vs. K (both in atoms per formula unit, afu). Plagioclase from Chassigny is not plotted here as it contains much more K than plagioclase from the other suites. Olivine-phyric, Pyroxene-phyric basalts and Lherzolites show similar plagioclase compositional trends starting at An65 and becoming progressively more K and Na-rich with crystallization. Plagioclase grains from the Clinopyroxenites and Orthopyroxenites are much more K and Na-rich than in the other suites, which suggests these crystallized from more fractionated melts or intercumulus liquids. The plots also illustrate that in general, plagioclase compositions can be related to a rock’s depleted or enriched nature (see Fig. 3). For example, enriched samples NWA 1110, RBT 04262, NWA 3171, Los Angeles and Shergotty have plagioclase grains that are relatively high in K and Na. Likewise, depleted rocks DaG 476, QUE 94201, and SAU 005 have plagioclase compositions that are relatively low in Na and K. However, this correlation does not hold true for all meteorites unless the bulk Al2O3 is considered, which is discussed below. Fig. 6a and b further explores the relationship between plagioclase compositions and a rock’s enriched or depleted nature by multiplying the K2O/Na2O ratio of the plagioclase by the bulk rock Al2O3 content. This parameter is valid because the K2O/Na2O ratio of the plagioclase is a good proxy for LREE abundance (i.e., enriched, intermediate, depleted) in the rock, while multiplying this ratio by the bulk rock Al2O3 normalizes the data and allows for different rock types to be compared. Fig. 6a presents the range and average of the above parameter for several martian samples and

Table 2 Average plagioclase (maskelynite) composition for each meteorite. SAU 005 NWA 1110 DaG 476 NWA 1195 NWA 2046 NWA 2626 ALH 77005 LEW 88516 RBT 04262 Los Angeles NWA 3171 QUE 94201 Shergotty Nakhla Gov. Val. ALH 84001 Chassigny N = 42 N = 40 N = 46 N = 49 N = 42 N = 39 N = 27 N = 47 N = 49 N = 60 N = 47 N = 31 N = 46 N = 28 N = 14 N = 20 N = 13 52.8

56.4

52.9

53.6

53.4

52.4

55.8

56.0

55.3

55.4

55.9

53.6

55.2

60.5

60.7

60.1

64.2

29.13 0.4 12.3 4.54 0.13 0.16 99.48

27.19 1.0 10.0 5.73 0.55 0.11 100.94

29.67 0.5 12.9 4.35 0.09 0.18 100.64

29.71 0.5 12.7 4.51 0.10 0.15 101.26

29.11 0.5 12.3 4.64 0.11 0.15 100.25

29.41 0.6 12.9 4.14 0.10 0.15 99.82

28.11 0.6 10.28 5.74 0.37 0.23 101.08

28.08 0.4 10.6 5.59 0.26 0.14 101.06

27.72 0.5 10.2 5.60 0.46 0.1 99.92

27.75 0.6 10.8 5.42 0.19 0.1 100.29

27.43 0.7 10.2 5.66 0.37 0.1 100.33

29.57 0.4 11.9 4.40 0.04 0.13 100.09

27.90 0.6 10.2 5.28 0.30 0.05 99.57

24.17 1.1 6.4 7.62 0.84 0.0 100.66

24.67 1.1 6.6 7.59 0.70 0.0 101.39

25.34 0.2 6.8 7.75 0.68 0.03 100.96

22.36 0.4 3.2 8.54 2.11 0.02 100.83

Cation formula based on 8 oxygens Si 2.406 2.523 Al 1.565 1.434 3+ Fe 0.014 0.034 Total 3.986 3.991 tet. Ca 0.603 0.479 Na 0.401 0.497 K 0.007 0.031 Mg 0.011 0.008 Total 5.008 5.007 cations K2O/Na2O X 100 2.76 Or 0.7 Ab 39.7 An 59.6 a

9.58 3.1 49.3 47.5

2.389 1.578 0.018 3.985

2.401 1.569 0.018 3.988

2.417 1.552 0.018 3.987

2.387 1.578 0.022 3.986

2.491 1.481 0.020 3.992

2.498 1.478 0.015 3.991

2.499 1.476 0.017 3.993

2.494 1.473 0.022 3.989

2.515 1.454 0.022 3.992

2.420 1.573 0.014 4.007

2.499 1.488 0.021 4.008

2.691 1.268 0.037 3.996

2.680 1.284 0.036 4.000

2.662 1.323 0.008 3.993

2.830 1.162 0.014 4.006

0.623 0.381 0.005 0.012 5.006

0.609 0.392 0.006 0.010 5.004

0.594 0.407 0.006 0.010 5.004

0.631 0.365 0.006 0.010 4.999

0.492 0.497 0.021 0.0 5.018

0.506 0.484 0.015 0.009 5.005

0.496 0.490 0.026 0.007 5.012

0.521 0.473 0.011 0.007 5.000

0.494 0.493 0.021 0.004 5.004

0.578 0.385 0.002 0.009 4.980

0.494 0.463 0.017 0.004 4.986

0.305 0.657 0.048 0.003 5.009

0.314 0.650 0.039 0.001 5.004

0.325 0.666 0.038 0.002 5.024

0.151 0.730 0.119 0.001 5.006

2.02 0.5 37.8 61.7

2.22 0.6 38.9 60.5

2.37 0.6 40.4 59.0

2.48 0.6 36.4 63.0

6.53 2.1 49.2 48.7

4.74 1.5 48.2 50.3

8.20 2.6 48.4 49.0

3.58 1.1 47.1 51.8

6.55 2.1 48.9 49.0

0.88 0.2 39.9 59.9

5.59 1.7 47.5 50.7

11.06 4.7 65.1 30.2

9.24 3.9 64.8 31.3

8.78 3.7 64.6 31.7

Silicate mineralogy of martian meteorites

SiO2 (wt%)a AI2O3 Fe2O3 CaO Na20 K20 MgO Total

24.74 11.9 73.0 15.1

Electron microprobe analyses have typical errors of 3% for major elements and up to 5% for minor elements.

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0.06

(a)

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(b)

SAU 005

NWA 2626 NWA 2046 NWA 1195 DaG 476

Olivine-phyric basalts

0.05

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K (afu)

NWA 1110

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0.00 20

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50

60

70

80

90

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80

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Pyroxene-phyric basalts

(d)

ALH 77005

Shergotty QUE 94201 NWA 3171 Los Angeles

LEW 88516

0.05

0.05

RBT 04262

0.04

K (afu)

0.04

K (afu)

100

0.06

0.06

0.03

0.03

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0.02

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0.00 20

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An % 0.06

90

An %

An %

(e)

40

50

60

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80

90

100

An %

Clinopyroxenites and Orthopyroxenite

0.05

Nakhla Gov. Valadares ALH 84001

K (afu)

0.04

0.03

0.02

0.01

0.00 20

30

40

50

60

70

80

90

100

An %

Fig. 5. (a–e) K (afu) vs. anorthite % for plagioclase analyses from the martian meteorite suites.

suggests that Shergotty, Los Angeles, and NWA 1110 are relatively enriched, SAU 005, QUE 94201 and DAG 476 are relatively depleted, and RBT 04262, ALH 77005, and LEW 88516 are intermediate. Fig. 6b further illustrates the correlation by plotting the above parameter against the La/Yb ratio of the bulk meteorite. Again the three groups are separated quite nicely and there is a strong positive correlation. The upshot of this observation is that relatively simple EMP analyses (compared with trace element bulk rock analysis) of plagioclase combined with the Al content of a bulk meteorite can give a quick estimate as to whether the rock came from a depleted or enriched source.

6. OLIVINE 6.1. Introduction Olivine occurs in all of the different martian rock types (except for the Pyroxene-phyric basalts), although in the Clinopyroxenite and Orthopyroxenite groups it is mostly late-stage, Fe-rich and minor to trace in abundance. In Lherzolites and Olivine-phyric basalts, olivine is often the first phase to crystallize, and as such its chemistry records important information about the bulk rock’s igneous history. For example, in a recent study, Shearer

Silicate mineralogy of martian meteorites

7459

Martian plagioclase

(a) Pyroxene-phyric basalts Shergotty

QUE 94201 Los Angeles

Lherzolites RBT 04262 LEW 88516 ALH 77005

Olivine-phyric basalts DaG 476 NWA 1110 SAU 005

0.00

0.10

0.20

0.30

0.40

0.50

0.60

0.70

K2O/Na2O (plag) X Al2O3 wt.% (bulk rock)

La/Yb bulk rock (chondrite normalized)

1.20 Los Angeles

(b)

NWA 1110

1.00 Shergotty RBT 04262

0.80

0.60

0.40

r ic En

e hm

d en Tr t n

LEW 88516 ALH 77005

0.20 DaG 476

0.00 0.00

SAU 005

QUE 94201

0.10

0.20

0.30

0.40

0.50

0.60

0.70

K2O/Na2O (plag) X Al 2O3 (bulk rock)

Fig. 6. (a) Range of K2O/Na2O (wt% in plagioclase) multiplied by Al2O3 (wt% bulk rock) for representative meteorites from the Pyroxenephyric, Olivine-phyric and Lherzolite groups. Average values are noted by an open square. Bulk rock Al2O3 values from same as listed in Fig. 3. (b) La/Yb (bulk rock normalized to chondrite) vs. K2O/Na2O (wt% in plagioclase) multiplied by Al2O3 (wt% bulk rock) for select martian meteorites. Bulk rock values from same as listed in Fig. 3.

et al. (2008) used the chemistry of olivines (in olivine-bearing martian rocks) to address the question as to whether the olivine crystals were phenocrysts, megacrysts or xenocysts. They concluded that olivines were true phenocrysts in only one rock (Y 980459), cumulate crystals (megacrysts) from a closely related basaltic melt in most of the Olivine-phyric basalts, and xenocrysts (picked up by basalts in transit from significantly different melt systems) in rocks such as NWA 1110 and RBT 04262. The following section builds upon the Shearer et al. (2008) data and observations and also explores other chemical systematics in olivine. Table 3 lists average olivine analyses for the different meteorites, while Fig. 7a presents the range of Fo content, i.e., [(Mg/Mg + Fe)  100] (afu), from the four martian suites that contain olivine. The plot also shows (top x-axis) the calculation of the Mg# = MgO/(MgO + FeO) that the melt would have to have in order to be in equilibrium with

the corresponding olivine, using a Mg–Fe2+ exchange KD of 0.33 (Roeder and Emslie, 1970). Fig. 7b plots the average olivine core Fo value (or Mg#) from the different rocks against the measured Mg# of its bulk rock. The calculated curve reveals that only olivine in Y 980459 appears to be in equilibrium with the melt (bulk rock) from which it crystallized, and thus Y 980459 is a primitive melt from the martian mantle. This assertion was first postulated by McKay et al. (2004) and Koizumi et al. (2004), and recently confirmed by Usui et al. (2008a). The figure also illustrates that all of the other olivine-bearing martian meteorites are in varying degrees of disequilibrium with their bulk rock compositions. The reason for this is most likely fractional crystallization, whereby Olivine-phyric basalts such as DaG 476 and NWA 1110 have had some olivine accumulation; Lherzolites ALH 77005, LEW 88516, and RBT 04262 have had more olivine accumulation; and the Clinopyroxenites MIL 03346, Governador Valadares, and Nakhla probably

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Table 3 Average olivine composition for each meteorite. Y 980459 SAU 005 NWA 1110 DaG 476 NWA 1195 NWA 2046 NWA 2626 ALH 77005 LEW 88516 RBT 04262 Nakhla N = 45 N = 79 N = 35 N = 45 N = 42 N = 31 N = 46 N = 47 N = 37 N = 88 N = 49 38.6 0.05 <0.03 0.20 40.5 19.0 0.43 0.22 99.1

36.6 0.06 <0.03 0.04 32.7 28.7 0.58 0.30 99.1

Cation formula based on 4 oxygens Si 1.000 0.997 Al 0.002 0.002 Ti 0.000 0.000 Cr 0.004 0.001 Mg 1.563 1.327 0.413 0.653 Fe2+ Mn 0.009 0.013 Ca 0.006 0.009 Total cations 2.997 3.002 XFe 0.21 0.33

36.3 0.10 <0.03 0.03 29.2 32.8 0.63 0.27 99.3 1.004 0.003 0.000 0.001 1.201 0.762 0.015 0.008 2.994 0.39

Trace element abundances in ppm (wt) Ni 1252 955 1122 Co 149 211 240 Ni/Co 8.39 4.53 4.68 a

36.9 0.16 <0.03 0.07 34.1 27.6 0.53 0.29 99.7 0.992 0.005 0.000 0.001 1.366 0.620 0.012 0.008 3.005 0.31 693 211 3.28

36.0 0.02 <0.03 0.04 30.4 33.2 0.61 0.24 100.4 0.986 0.001 0.000 0.001 1.241 0.762 0.014 0.007 3.013 0.38 753 273 2.76

38.0 0.10 <0.03 0.08 38.3 23.3 0.46 0.25 100.6 0.990 0.003 0.000 0.002 1.484 0.511 0.010 0.007 3.007 0.26 908 207 4.38

37.7 0.13 <0.03 0.05 34.6 27.1 0.50 0.22 100.3 1.002 0.004 0.000 0.001 1.365 0.604 0.011 0.006 2.995 0.31 937 186 5.03

37.6 0.03 <0.03 0.03 36.6 25.8 0.55 0.23 100.9 0.989 0.001 0.000 0.001 1.434 0.567 0.012 0.007 3.010 0.28 658 183 3.60

Electron microprobe analyses have typical errors of 3% for major elements and up to 5% for minor elements.

36.8 0.04 <0.03 0.03 33.2 29.6 0.61 0.27 100.5 0.990 0.001 0.000 0.001 1.330 0.665 0.014 0.008 3.009 0.33 734 209 3.51

36.7 0.07 <0.03 <0.03 29.6 33.7 0.66 0.24 100.9 1.001 0.002 0.000 0.000 1.203 0.768 0.015 0.007 2.997 0.39 1419 248 5.72

32.6 <0.02 0.03 <0.03 11.4 55.6 1.10 0.24 101.1 1.004 0.000 0.001 0.000 0.524 1.430 0.029 0.008 2.995 0.73 297 273 1.09

33.2 0.03 0.05 <0.03 12.4 53.5 1.15 0.48 100.9 1.010 0.001 0.001 0.000 0.560 1.370 0.030 0.016 2.988 0.71 544 340 1.60

33.6 <0.02 <0.03 <0.03 14.8 51.2 1.02 0.37 101.1 1.008 0.000 0.000 0.000 0.661 1.284 0.026 0.012 2.992 0.66 713 344 2.07

37.7 0.02 <0.03 <0.03 34.7 27.9 0.54 0.16 101.0 0.999 0.001 0.000 0.000 1.367 0.616 0.012 0.004 3.000 0.31 500 123 4.07

J.J. Papike et al. / Geochimica et Cosmochimica Acta 73 (2009) 7443–7485

SiO2 (wt%)a Al2O3 TiO2 Cr2O3 MgO FeO MnO CaO Total

MIL 03346 Gov. Val. Chassigny N = 19 N = 45 N = 36

Silicate mineralogy of martian meteorites

(a) 73

7461

Mg# of melt 54

41

32

23

16

11

6

Chassigny Gov. Val. MIL 03346 Nakhla RBT 04262 LEW 88516 ALH 77005 NWA 2626 NWA 2046 NWA 1195 DaG 476 NWA 1110 SAU 005 Y 980459

90

80

70

60

50

40

30

20

10

Olivine composition (Fo%) 80

(b) ALH 77005 LEW 88516

70

DaG 476

Mg# Bulk rock

Y 980459

Chassigny RBT 04262 SAU 005

60 NWA 1110

50

Gov. Val.

Nakhla

MIL 03346

40 Calculated Mg# of melt in equilibrium with olivine core 30 90

80

70

60

50

40

30

20

Mg# Olivine core

Fig. 7. (a) Range of olivine compositions (Fo% – bottom x-axis) for the martian meteorites in this study. Also shown (top x-axis) is the calculated Mg# of a melt (hypothetical) that the olivine grains would be in equilibrium with. (b) Core Mg# of olivine grains from the martian meteorites vs. Mg# of their bulk rock compositions. The curved line denotes equilibrium between Mg# bulk rock and Mg# of the olivine cores.

have had significant olivine removal and some pyroxene accumulation. 6.2. Co–Ni systematics Fig. 8 is a plot of Ni vs. Co for olivine from select martian meteorites. This diagram is very similar to Fig. 9 in Shearer et al. (2008), except that here we have limited the number of plotted points, combined some groups that exhibit similar behavior, and added new Clinopyroxenite data. Some of the Shearer et al. (2008) observations are re-stated here: Y 980459 could be the “type” of parental melt that evolved, by fractional crystallization, to the meteorites designated as Others (several Olivine-phyric basalts, Lherzolites) in the diagram. However, RBT 04262 and NWA 1110 plot at relatively high Co values and are outliers on

this diagram and other diagrams that follow. The Clinopyroxenites fall into a separate group on this diagram, at higher Co and lower Ni. Fig. 9a–c shows the variation of Ni vs. XFe (Fe/ Fe + Mg–afu), Co vs. XFe and Co/Ni vs. XFe for olivine from select martian rocks. Fig. 9a shows a decrease in Ni in olivine with fractionation as would be expected for a compatible element. Y 980459 again plots as a potential parental melt for the Others group of meteorites. Again NWA 1110 and RBT are outliers (i.e., higher Ni) and the Clinopyroxenites fall along the trend but at the high XFe end. Fig. 9b shows Co increasing in olivine with fractionation. This trend is counter-intuitive because Co is a compatible element in olivine and should therefore decrease with increasing crystallization. Similar anomalous Co behavior was first noted by Papike et al. (1999) in a study

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cause it often crystallizes early in a cooling magma and most importantly, it can incorporate most of the major and minor rock-forming elements into its structure (but not K). Several previous studies have shown that pyroxene chemistry not only records crystallization conditions but also planetary parentage as well (see Papike, 1980, 1981, 1996, 1998; Karner et al., 2006). The following section continues that assertion by focusing on the pyroxene chemistry of the martian meteorites.

1600

Ni (ppm)

1200

800

7.2. Quadrilateral components (QUAD) = Ca, Mg, Fe2+, Mn, Si Y 980459 NWA 1110 RBT 04262 Clinopyroxenites Others

400

0 0

100

200

300

400

Co (ppm)

Fig. 8. Ni (ppm) vs. Co (ppm) systematics in olivine from select martian meteorites. “Others” points are from Olivine-phyric and Pyroxene-phyric basalts, and Lherzolites.

of Ni and Co in lunar olivine grains. They found that lunar olivine cores showed DNi (D = Ni in olivine/Ni in melt) was 9.9 and DCo was 4, so both elements were compatible in olivine, only Ni more so. Ni concentration across olivine grains decreased with fractionation as expected, but Co concentration did not vary with fractionation. They concluded that the increase of DCo with fractionation just counterbalanced the depletion of Co in the melt, leading to the flat concentration trends. The same reasoning could explain anomalous Co behavior in Fig. 9b, namely that DCo is increasing at a faster rate than the melt is depleted in Co (remember that DCo is lower than the DNi). 6.3. Mn vs. XFe systematics Fig. 10a–e shows the positive correlation of Mn with Fe for olivine in the different martian groups. Note that olivine from the Olivine-phyric basalts, Lherzolites, and Dunite share very similar trends in Mn/Fe ratios with fractionation, even though these rock types have had very different crystallization histories. Olivines from the Clinopyroxenites fall slightly off the trend and are much more Fe-rich and slightly more Mn-rich (at corresponding Fe concentrations) than olivine from the other groups. The Fe-enrichment is explained by the fact that olivine in Clinopyroxenites is a late crystallizing phase, while the Mn enrichment may suggest that there is a decoupling of Mn and Fe as crystallization proceeds to very Fe-rich olivine compositions. 7. PYROXENES 7.1. Introduction Pyroxene is the most abundant mineral by volume in most of the martian rock types and is also arguably the most abundant mineral on the surface of Mars. It is a powerful recorder of the igneous history of a basaltic melt be-

Averages of the pyroxene analyses for each meteorite in the reference suite are presented in Table 4, while the systematics of the pyroxene QUAD components (Wo-En–Fs) are illustrated in Fig. 11. The plot shows that the Olivine-phyric basalts and Lherzolites have similar pyroxene trends. The trends start with the crystallization of low-Ca pyroxene (orthopyroxene or pigeonite) that zones to augite. The pyroxene in these rocks crystallizes after olivine, which in most cases is cumulus, with the exception of olivine in Y 980459. The trajectory from low-Ca pyroxene to augite is likely caused by delayed nucleation of plagioclase; so that the activity of Ca increases in the melt and thus drives the crystallizing pyroxene to higher Ca contents. Pyroxene-phyric basalts are more evolved melts than either Olivine-phyric basalts or Lherzolites and likely have experienced previous olivine and chromite fractionation. The pyroxene crystallization in these rocks shows two types of trajectories: the first trend includes pyroxenes from LA, NWA 3171, and Shergotty, and has two pyroxenes (augite and pigeonite) coming on the liquidus and subsequently zoning to higher Fe contents. The second trend includes QUE 94201 pyroxenes and begins with pigeonite crystallization, which then zones to augite, and then zones across the quadrilateral to pyroxferroite (e.g., Karner et al., 2007a). Pyroxenes from the Clinopyroxenites usually have a single pyroxene trend from Mgrich augite to more Fe-rich augite. The pyroxene in ALH 84001 (Orthopyroxenite) shows a tight cluster of points in the orthopyroxene field, which is most likely caused by long annealing time or sub-solidus equilibration. The differences in the pyroxene trends from the martian suites demonstrate that analysis of a single pyroxene grain in the martian regolith could reveal the difference between several different basalt compositions from which it crystallized. 7.3. Others components (OTHERS) = Na, Cr3+, Fe3+, VI Al, Ti

IV

Al,

This section gives an overview of the behavior of the OTHERS components followed by a more detailed discussion keyed to a number of figures below. All OTHERS components have valence states different from the cations they replace (i.e., Ca, Mg, Fe2+, Mn2+, Si4+) and thus require charge balance couples when entering the pyroxene crystal structure. For each of the OTHERS substitutions the following charge balance equation applies: M2Na1+ + IV Al3+ (i.e., charge balance deficiencies): = VICr3+ + VI Fe3+ + VIAl3+ + 2Ti4+ (i.e., charge balance excesses).

Silicate mineralogy of martian meteorites

7463

2000

(a)

Y 980459 NWA 1110 RBT 04262 Clinopyroxenites Others

1800 1600

Ni (ppm)

1400 1200 1000 800 600 400 200 0 0.00

0.20

0.40

0.60

0.80

1.00

XFe (afu) 500

(b)

450 400

Co (ppm)

350 300 250 200 150

Y 980459 NWA 1110

100

RBT 04262 Clinopyroxenites Others

50 0 0.00

0.20

0.40

0.60

0.80

1.00

XFe (afu) 1.20

(c)

1.00

Co/Ni

0.80

0.60

0.40

Y 980459 NWA 1110 RBT 04262 Clinopyroxenites Others

0.20

0.00 0.00

0.20

0.40

0.60

0.80

1.00

XFe (afu)

Fig. 9. (a–e) Ni, Co, and Co/Ni (ppm) vs. XFe (Fe/Fe + Mg) in atoms per formula unit (afu) in olivine from select martian meteorites. “Others” points are from Olivine-phyric and Pyroxene-phyric basalts, and Lherzolites.

For example, one viable charge balance couple is Na– Fe (acmite) for Ca–Mg (diopside). The most important charge balance couples for pyroxene in the 19 martian 3+

meteorites studied are VIFe3+–IVAl, VITi4+–2AlIV, and (VIR3+, VITi4+–IVAl) where R3+ = Cr3+, Fe3+, and VIAl3+. In the last couple, the charge contributions from Cr3+,

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J.J. Papike et al. / Geochimica et Cosmochimica Acta 73 (2009) 7443–7485 0.040

0.040

0.035

0.030

0.030

0.025

0.025

Mn (afu)

Mn (afu)

0.035

Olivine-phyric basalts

(a)

0.020 0.015

0.020 0.015 0.010

0.010 SAU 005 Y 980459 NWA 1110

0.005 0.000 0.000

0.200

0.400

0.600

0.800

1.000

1.200

NWA 2626 NWA 2046 NWA 1195 DaG 476

0.005 0.000 0.000

1.400

Fe (afu)

0.400

0.600

0.800

1.000

1.200

1.400

0.040

Lherzolites

(c)

0.035

0.030

0.030

0.025

0.025

Mn (afu)

Mn (afu)

0.200

Fe (afu)

0.040 0.035

Olivine-phyric basalts

(b)

0.020 0.015

(d)

Clinopyroxenites

0.020 0.015 0.010

0.010 ALH 77005 LEW 88516 RBT 04262

0.005 0.000 0.000

0.200

0.400

0.600

0.800

1.000

1.200

Nakhla MIL 03346

0.005

Gov. Valadares

1.400

0.000 0.000 0.200 0.400 0.600 0.800 1.000 1.200 1.400 1.600 1.800

Fe (afu)

Fe (afu) 0.040 0.035

Dunites

(e)

0.030

Mn (afu)

0.025 0.020 0.015 0.010 0.005 0.000 0.000

Chassigny

0.200

0.400

0.600

0.800

1.000

1.200

1.400

Fe (afu)

Fig. 10. (a–e) Mn (afu) vs. Fe (afu) for olivine grains from the martian rock groups.

Fe3+, VIAl3+, and Ti4+ are approximately equal. The percent of OTHERS substitutions = [(Al, Na, Cr, Fe3+, Ti/4)  100] in pyroxene from each rock ranges from about 1% to 4% (see Table 4). This suggests that previously determined pyroxene QUAD phase relationships, for the pure system, are applicable to martian pyroxenes. Ferric iron is generally low in concentration, ranging from approximately 0.05 to 1.57 Fe2O3 wt% or 0.01 to 0.045 afu. Thus the total ferric iron ranges from 1% to 10% (see Table 4). Fig. 12a–e shows the behavior of Cr in pyroxene from all the martian meteorite groups. In all the meteorites Cr decreases with increasing XFe, which is consistent for a compatible element. Cr is compatible in pyroxene, but its

D varies with fO2, melt composition, and pyroxene composition. Karner et al. (2007b) found for the QUE 94201 bulk composition that DCr for pigeonite varied from 3 to 5 from IW 1 to IW + 1 while augite Cr Ds varied from 7 to 9 over the same fO2 range. The increase of DCr with fO2 results from an increase in the Cr3+/Cr2+ in the melt, where Cr3+ is much more compatible in the pyroxene crystal structure than Cr2+ (Papike et al., 2005). Chromium systematics in pyroxene can also be used to evaluate whether a Y 980459 composition could be parental to QUE 94201 composition by fractional crystallization of olivine, chromite, and pyroxene. The Cr systematics suggest that the answer is no. Note that the maximum Cr value in Y

Table 4 Average pyroxene composition for each.

Si IV

Al TotaI tet. VI AI Fe3+ Ti Cr Mg Fe2+ Mn Ca Na TotaI cations Wo En Fs XFe XFe3+ % Others a

SAU 005 N = 54

NWA 1110 N = 67

DaG 476 N = 75

NWA 1195 N = 40

NWA 2046 N = 49

NWA 2626 N = 44

ALH 77005 N = 75

LEW 88516 N = 39

RBT 04262 N = 70

Los Angeles N = 48

NWA 3171 N = 94

QUE 94201 N = 65

Shergotty N = 74

NakhIa N = 44

MIL 03346 N = 52

Gov. Val. N = 44

ALH 84001 N = 38

Chassigny N = 17

54.9 0.71 0.15 0.10 0.69 27.0 13.3 0.47 2.74 0.03 100.0

53.8 1.03 0.78 0.23 0.44 23.4 15.3 0.53 4.99 0.11 100.6

52.2 1.22 1.14 0.27 0.43 19.0 17.6 0.59 7.82 0.12 100.3

53.9 1.12 0.64 0.21 0.54 23.4 14.7 0.52 5.64 0.07 100.8

53.6 1.27 0.61 0.26 0.50 22.3 16.4 0.54 5.57 0.08 101.1

54.4 0.97 0.57 0.16 0.55 24.3 15.7 0.52 4.28 0.05 101.5

54.6 0.88 0.33 0.12 0.54 26.0 14.7 0.48 2.74 0.06 100.5

52.7 1.78 1.63 0.45 0.75 19.8 8.98 0.43 13.96 0.18 100.6

53.6 1.23 1.45 0.23 0.62 22.5 11.1 0.47 9.37 0.12 100.7

53.1 1.38 0.75 0.34 0.58 19.6 13.5 0.51 10.8 0.18 100.7

50.1 0.94 0.44 0.44 0.05 10.4 26.5 0.75 10.8 0.11 100.6

51.3 1.14 1.05 0.35 0.36 13.8 19.4 0.63 12.4 0.22 100.7

50.9 1.16 0.43 0.52 0.43 15.4 22.3 0.69 8.15 0.08 100.0

51.5 0.81 0.05 0.29 0.35 15.0 22.4 0.70 8.66 0.10 99.9

51.7 0.82 1.49 0.21 0.33 13.0 13.9 0.45 18.2 0.24 100.5

51.8 1.09 1.57 0.32 0.25 12.7 13.3 0.43 19.1 0.29 100.8

51.8 0.96 1.29 0.25 0.33 12.9 14.2 0.46 18.4 0.23 100.8

54.8 0.64 1.18 0.14 0.32 26.1 16.2 0.48 1.57 0.07 101.5

53.5 1.24 1.16 0.29 0.74 17.2 8.02 0.34 18.5 0.33 101.4

1.973 0.025 1.998 0.005 0.004 0.003 0.020 1.444 0.401 0.014 0.107 0.002 3.998 5.6 73.9 20.5 0.22 0.01 1.5

1.959 0.039 1.997 0.006 0.021 0.006 0.013 1.270 0.468 0.016 0.195 0.008 4.000 10.1 65.7 24.2 0.27 0.05 2.3

1.948 0.048 1.996 0.006 0.032 0.008 0.013 1.054 0.550 0.019 0.314 0.009 4.000 16.4 54.9 28.6 0.34 0.06 2.9

1.957 0.040 1.998 0.008 0.017 0.006 0.015 1.267 0.447 0.016 0.220 0.005 3.999 11.4 65.5 23.1 0.26 0.04 2.3

1.955 0.044 1.998 0.011 0.017 0.007 0.014 1.209 0.501 0.017 0.218 0.006 3.998 11.4 62.6 25.9 0.29 0.04 2.5

1.962 0.036 1.997 0.006 0.015 0.004 0.016 1.299 0.475 0.016 0.167 0.004 3.999 8.7 66.8 24.5 0.27 0.03 2.0

1.969 0.029 1.998 0.009 0.009 0.003 0.015 1.392 0.446 0.015 0.107 0.004 3.999 5.6 71.5 22.9 0.25 0.02 1.7

1.923 0.073 1.995 0.004 0.045 0.012 0.022 1.076 0.273 0.013 0.547 0.013 4.000 29.0 56.7 14.4 0.20 0.15 4.2

1.942 0.051 1.994 0.002 0.040 0.006 0.018 1.213 0.336 0.014 0.368 0.008 4.000 19.4 63.1 17.5 0.22 0.12 3.1

1.950 0.048 1.998 0.012 0.021 0.009 0.017 1.071 0.415 0.016 0.426 0.013 3.998 22.4 56.0 21.6 0.28 0.05 3.0

1.963 0.036 1.999 0.007 0.013 0.013 0.001 0.603 0.875 0.025 0.454 0.009 3.999 23.5 31.2 45.3 0.59 0.02 2.0

1.952 0.045 1.997 0.006 0.030 0.010 0.011 0.784 0.621 0.020 0.504 0.016 4.000 26.4 41.1 32.5 0.44 0.05 3.0

1.951 0.047 1.998 0.005 0.012 0.015 0.013 0.868 0.724 0.023 0.334 0.006 3.998 17.4 45.1 37.5 0.45 0.02 2.4

1.976 0.024 2.000 0.013 0.001 0.008 0.010 0.851 0.725 0.023 0.356 0.007 3.995 18.4 44.0 37.5 0.46 0.00 1.6

1.959 0.036 1.995 0.001 0.043 0.006 0.010 0.732 0.443 0.014 0.739 0.018 4.000 38.6 38.2 23.2 0.38 0.09 2.8

1.951 0.045 1.996 0.003 0.045 0.009 0.008 0.715 0.419 0.014 0.770 0.021 4.000 40.4 37.5 22.1 0.37 0.10 3.3

1.957 0.039 1.996 0.003 0.037 0.007 0.010 0.723 0.448 0.015 0.745 0.017 4.000 38.9 37.7 23.4 0.38 0.08 2.8

1.965 0.027 1.992 0.000 0.032 0.004 0.009 1.396 0.487 0.015 0.060 0.005 4.000 3.1 71.8 25.1 0.26 0.06 1.9

1.950 0.048 1.999 0.005 0.032 0.008 0.021 0.936 0.244 0.010 0.721 0.023 4.000

Silicate mineralogy of martian meteorites

SiO2 (wt%)a Al2O3 Fe2O3 TiO2 Cr2O3 MgO FeO MnO CaO Na2O Total

Y 980459 N = 47

37.9 49.2 12.8 0.21 0.12 3.4

Electron microprobe analyses have typical errors of 3% for major elements and up to 5% for minor elements.

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Olivine-phyric 50 50

50 50

50 50

50 50

50 50

50 50 60 60

40 40

60 60

60 60

40 40

40 40 70 70

30 30

70 70

70 70

30 30

30 30 80 80

20 20

80 80

80 80

20 20

20 20 90 90

10 10

90 90

90 90

10 10

10 10

Y 980459 100

90

80

70

60

50

40

30

SAU 005

100 100

20

100

10

90

80

70

60

50

40

30

NWA 1110

100 100

20

100

10

90

80

70

60

50

40

30

20

100 100

10

Olivine-phyric 50 50 60 60 70 70 80 80 90 90

70

60

50

40

30

100

10

90

80

70

60

50

40

30

10

50 50

40 40

10 10

NWA 2626 40

30

20

1010

100

10

90

90 90

80

70

60

50

40

30

20

10 10

LEW 88516

100 100 100

10

90

40

30

20

70 70

30 30 80 80

20 20

20 20 90 90

90 90

10 10

10 10

Los Angeles

100 100

100

10

90

80

70

60

50

40

30

20

NWA 3171

100 100

100

10

90

30

20

20 20 90 90

90

90

10 10

10 10

Shergotty

100 100

100

10

90

80

70

60

50

40

30

20

Nakhla

100 100

100

10

Clinopyroxenites 50 50

30

20

60

Chassigny

80

80

20 20 90

90 90

90

10 10

10 10

Gov. Valadares

100 100

10

70

70

30 30

20 20

10 10

100

100 100

10

80 80

90 90

MIL 03346

20

60

40 40

30 30

20 20

30

40

50

70 70

80 80

40

50

50

40 40

30 30

50

60

60 60

40 40

60

70

50 50

50 50 60 60

70

80

Orthopyroxenite and Dunite

50 50

80

90

50 50

70 70

100 100

10

80 80

80

80

QUE 94201 40

20

30 30

20 20 90 90

30

70 70

70

70

30 30

10 10

40

40 40

40 40

80 80

50

60 60

60

60

60 60

40 40

20 20

60

50 50

50 50

70 70

70

50 50

50

50

50 50

30 30

80

Clinopyroxenites

50 50

90

100 100

10

60 60

Pyroxene-phyric

100

20

80 80

RBT 04262

50

30

40 40

30 30

10 10

60

40

70 70

90 90

70

50

50 50

40 40

20 20

80

60

60 60

80 80

90

70

50 50

50 50

30 30

100

80

Pyroxene-phyric

70 70

50

100 100

10

80 80

50 50

40 40

60

20

20 20

ALH 77005

100 100

60 60

70

30

70 70

90 90

50 50

80

40

30 30

80 80

50 50

90

50

60 60

Lherzolites

100

60

40 40

70 70

2020 90 90

50

70

50 50

3030 80 80

20 20

60

80

50 50

60 60

4040 70 70

30 30

70

90

50 50

5050 60 60

80

100

Lherzolites

50 50

90

NWA 2046

100 100

20

Olivine-phyric

100

90 90

10 10

NWA 1195

100 100

20

80 80

20 20 90 90

10 10

10 10

DaG 476

70 70

30 30 80 80

20 20

20 20

80

60 60 40 40

70 70

30 30

30 30

90

50 50 50 50

60 60

40 40

4040

100

50 50

50 50

5050

90

80

70

60

50

40

30

20

10

ALH 84001

100 100

100

90

80

70

60

50

40

30

20

100 100

10

Fig. 11. QUAD compositions for pyroxene from each martian meteorite. Arrows indicate crystallization trends.

980459 pyroxene (Fig. 12a) is not higher than the maximum Cr in QUE 94201 pyroxene (Fig. 12c). Chromium would have to have been drastically depleted in QUE 94201 pyroxene to allow for the model that links Y 980459-type melts to

QUE 94201-type melts. The parental melt to QUE 94201 must have had higher Cr concentrations than Y 980459. Fig. 13a–f shows the Ti vs. XFe systematics for martian pyroxenes. The systematics are clearest for the

Silicate mineralogy of martian meteorites 0.035

0.035 0.030 0.025 0.020

0.025 0.020 0.015

0.010

0.010

0.005

0.005

0.20

0.40

0.60

0.80

ALH 77005 LEW 88516 RBT 04262

0.030

0.015

0.000 0.00

Lherzolites

(b)

SAU 005 NWA 2626 NWA 2046 NWA 1195 DaG 476 Y980459 NWA 1110

Cr (afu)

Olivine-phyric basalts

(a)

Cr (afu)

7467

0.000 0.00

1.00

0.20

XFe (afu)

0.40

0.60

1.00

0.035

0.035

(c)

Pyxroxene-phyric basalts

Clinopyroxenites

(d)

Shergotty QUE 94201

0.030

Gov. Valadares

0.025

Los Angeles

Cr (afu)

0.025 0.020 0.015

0.020 0.015

0.010

0.010

0.005

0.005

0.000 0.00

0.20

0.40

0.60

0.80

Nakhla MIL 03346

0.030

NWA 3171

Cr (afu)

0.80

XFe (afu)

1.00

XFe (afu)

0.000 0.00

0.20

0.40

0.60

0.80

1.00

XFe (afu)

0.035

(e)

Orthopyroxenite and Dunite

ALH 84001 Chassigny

0.030

Cr (afu)

0.025 0.020 0.015 0.010 0.005 0.000 0.00

0.20

0.40

0.60

0.80

1.00

XFe (afu)

Fig. 12. (a–e) Cr (afu) vs. XFe (afu) for pyroxene grains from the meteorites.

Pyroxene-phyric basalts, which show an increase in Ti with increasing XFe (fractionation). In addition, Karner et al. (2007b) show that there is a decrease of Ti in pyroxene when ulvo¨spinel and/or ilmenite come onto the liquidus. This trend is evident in both the LA and QUE 94201 systematics (see Fig. 13d) where Ti in pyroxene increases to 0.6–0.7 XFe (respectively) and then decreases sharply. Fig. 14a–f shows the systematics of total Al vs. XFe for martian pyroxenes. The results, though not spectacular, do indicate the extent to which OTHERS components are entering the pyroxene. Because IVAl is the most important charge deficiency cation in martian pyroxene (discussed

below), the higher the Al content, the higher the amount of OTHERS components and thus the deviation from rigorous quadrilateral phase relationships. Fig. 15a–f shows the Na vs. XFe behavior for pyroxenes in all the meteorites. Although Na is in low abundance, there is a correlation between elevated Na in the pyroxene M2 site and Ca. This was first noted by Papike (1981) with the rationale that the similar ionic radii of Ca and Na allows easier Na access to a Ca containing M2 site compared with a Mg, Fe2+-bearing M2 site. The observation was originally made for the high-Ca vs. low-Ca trend in Shergotty pyroxenes, but here the plots show that pyroxenes (all high-Ca augites) in the Clinopyroxenites (Fig. 15e), the

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J.J. Papike et al. / Geochimica et Cosmochimica Acta 73 (2009) 7443–7485 0.030

0.030 0.025

0.025

0.020

0.020

Ti (afu)

Ti (afu)

Olivine-phyric basalts

(b)

Olivine-phyric basalts

(a)

0.015

0.015 0.010

0.010

NWA 2626 SAU 005 Y 980459 NWA 1110

0.005 0.000 0.00

0.20

0.40

0.60

NWA 2046 NWA 1195 DaG476

0.005

0.80

0.000 0.00

1.00

0.20

XFe (afu)

1.00

Pyroxene-phyric basalts

(d)

0.025

0.025

0.020

0.020

Ti (afu)

Ti (afu)

0.80

0.030

Lherzolites

(c)

0.015 0.010

0.015 0.010

ALH 77005 LEW 88516 RBT 04262

0.005 0.000 0.00

0.20

0.40

0.60

0.80

Shergotty QUE 94201 NWA 3171 Los Angeles

0.005 0.000 0.00

1.00

0.20

0.030

0.60

0.80

1.00

0.030

Clinopyroxenites

(e)

0.40

XFe (afu)

XFe (afu)

Orthopyroxenite and Dunite

(f)

0.025

0.025

0.020

0.020

Ti (afu)

Ti (afu)

0.60

XFe (afu)

0.030

0.015 0.010

0.015 0.010

Nakhla MIL 03346 Gov. Valadares

0.005 0.000 0.00

0.40

0.20

0.40

0.60

0.80

1.00

XFe (afu)

0.005

ALH 84001 Chassigny

0.000 0.00

0.20

0.40

0.60

0.80

1.00

XFe (afu)

Fig. 13. (a–f) Ti (afu) vs. XFe (afu) for pyroxene grains from the martian meteorites.

high-Ca pyroxenes in NWA 3171 (Fig. 15d), and the augite in Chassigny (Fig. 15f), all show elevated Na. Fig. 16a–f demonstrates that pyroxene Fe3+ and XFe have no clear correlation, but close inspection shows that Fe3+ prefers Ca-rich pyroxene to Ca-poor pyroxene, with all other factors being equal. There is also no clear correlation between estimated ferric iron content and previous estimates of fO2 (for the individual meteorites). For example, Shergotty is generally estimated to be the most oxidized martian meteorite (IW + 2.5) but has very low ferric iron in its pyroxenes (Fig. 16d). Conversely, QUE 94201 has been found to be one of the most reduced martian meteor-

ites (IW + 1), but has relatively high ferric iron in its pyroxenes (Fig. 16d). Several factors (in addition to fO2) affect the incorporation of ferric iron into pyroxene including the amount of Ca in the pyroxene M2 site, with DFe3+ greater for augite than pigeonite. Fig. 17a–e shows the tight, positive correlation of Mn and Fe2+ in pyroxene from the different martian groups. Virtually all the pyroxene analyses fall on the same Mn/ Fe ratio regardless of the pyroxene XFe value, as opposed to the olivine data which showed Mn and Fe decoupling in high-Fe olivines. Like the olivine data, the Mn/Fe trends in pyroxene are similar for all the rocks from Mars. Distinct

Silicate mineralogy of martian meteorites 0.160

0.160

Olivine-phyric basalts

(a) 0.140

SAU 005 Y980459 NWA 1110

0.120

Total Al (afu)

0.100 0.080 0.060

0.100 0.080 0.060

0.040

0.040

0.020

0.020

0.000 0.000

0.200

0.400

0.600

0.800

0.000 0.000

1.000

0.200

0.400

XFe

0.600

0.800

1.000

XFe 0.160

0.160

Lherzolites

(c)

ALH 77005

(d)

LEW 88516

0.140

Pyroxene-phyric basalts

Total Al (afu)

0.120 0.100 0.080 0.060

QUE 94201 NWA 3171

0.120

Los Angeles

0.100 0.080 0.060

0.040

0.040

0.020

0.020

0.000 0.000

0.200

0.400

0.600

0.800

0.000 0.000

1.000

0.200

0.400

XFe

0.600

0.800

1.000

XFe

0.160

Clinopyroxenites

(e)

Shergotty

0.140

RBT 04262

Total Al (afu)

NWA 2626 NWA 2046 NWA 1195 DaG 476

Olivine-phyric basalts

(b) 0.140

0.120

Total Al (afu)

7469

0.160

Nakhla

(f)

MIL 03346

0.140

ALH 84001 Chassigny

Orthopyroxenite and Dunite

0.140

Gov. Valadares

0.120

Total Al (afu)

Total Al (afu)

0.120 0.100 0.080 0.060

0.100 0.080 0.060

0.040

0.040

0.020

0.020

0.000 0.000

0.200

0.400

0.600

0.800

1.000

XFe

0.000 0.000

0.200

0.400

0.600

0.800

1.000

XFe

Fig. 14. (a–f) Total Al (IVAl + VIAl–afu) vs. XFe (afu) for pyroxene grains from the different martian meteorites.

Mn/Fe ratios in materials from the different planets has been known for some time (e.g., Dymek et al., 1976; Drake et al., 1989); the cause of these distinct ratios in olivine and pyroxene from the planets will be further discussed in the following section. Lastly, Fig. 18 presents the types of coupled substitutions that are required for charge balance when OTHERS cations substitute into the pyroxene crystal structure. Papike (1981) used a statistical method to assess the substitution couples, but here we use a graphical method for easier comprehension. The plot shows the contribution of each element (Deficiency or Excess) in afu to the charge

balance couple. Note both IVAl and Na contribute to charge balance deficiencies in all the meteorites, but the contribution of IVAl is greater than Na in most cases. Also highlighted for each sample is the most important (qualitative) charge couple. 8. COMPARATIVE PLANETOLOGY 8.1. Mars, Earth, Moon, 4 Vesta comparisons Before we focus on the igneous evolution of Mars, we would like to make some first order comparisons of the four

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J.J. Papike et al. / Geochimica et Cosmochimica Acta 73 (2009) 7443–7485

0.030 Olivine-phyric basalts

(a) 0.025

0.030

SAU 005 Y 980459 NWA 1110

0.020 0.015

0.015

0.010

0.010

0.005

0.005

0.000 0.000

0.200

0.400

0.600

0.800

0.000 0.000

1.000

0.200

0.400

XFe

0.800

1.000

0.030 Lherzolites

(c)

ALH 77005 LEW 88516 RBT 04262

0.025

(d)

Pyroxene-phyric basalts

0.025

0.020

Shergotty QUE 94201 NWA 3171 Los Angeles

0.020

Na (afu)

Na (afu)

0.600

XFe

0.030

0.015

0.015

0.010

0.010

0.005

0.005

0.000 0.000

0.200

0.400

0.600

0.800

0.000 0.000

1.000

0.200

0.400

XFe

0.600

0.800

1.000

XFe 0.030

0.030 Clinopyroxenites

(e)

Nakhla

0.025

ALH 84001 Chassigny

0.020

Na (afu)

0.020 0.015

0.015

0.010

0.010

0.005

0.005

0.000 0.000

Orthopyroxenite and Dunite

(f)

MIL 03346 Gov. Valadares

0.025

Na (afu)

NWA 2626 NWA 2046 NWA 1195 DaG 476

0.020

Na (afu)

Na (afu)

Olivine-phyric basalts

(b) 0.025

0.200

0.400

0.600

0.800

1.000

XFe

0.000 0.000

0.200

0.400

0.600

0.800

1.000

XFe

Fig. 15. (a–f) Na (afu) vs. XFe (afu) for pyroxene grains from the martian meteorites.

differentiated planetary bodies on which we focused our past comparative planetary mineralogy studies, Earth, Moon, Mars, and 4 Vesta. These are listed in the order of our degree of understanding of their igneous evolution. Our understanding of the igneous evolution of Mars is at a rather immature stage compared to our understanding of Earth and Moon. Table 5 presents estimates of the bulk-silicate compositions of the four planetary bodies for select oxides and oxide ratios, while Table 6 presents a crude attempt at giving an average basalt composition for each body. In these comparisons we did not include exotic basalts that are unique to one body, such as high-Ti basalts from the Moon. We tried to compare similar types of basalts as much as possible.

We first consider the bulk composition of the planets. This is a very difficult estimate to make and several authors have used different approaches, and thus the compositions in Table 5 are representative of the range of bulk compositional estimates reported in the literature. Examining Table 5, two observations stand out: (1) the high FeO concentration in Mars and 4 Vesta, which leads to high FeO/ MgO for those planets when compared to Earth and Moon, and (2) the narrow range of CaO/Al2O3 for all the planets, which is in stark contrast to the wide range of CaO/Al2O3 seen in the planetary basalt compositions (Table 6). More important for this paper, however, is the compositional characteristics of the mantle reservoirs from which the basalts were extracted. These reservoirs are affected by core

Silicate mineralogy of martian meteorites

7471

0.100

0.100

Olivine-phyric basalts

(a) 0.080

SAU 005 Y 980459 NWA 1110

Olivine-phyric basalts

(b)

NWA 2626 NWA 2046 NWA 1195

0.080

Fe3+ (afu)

Fe3+ (afu)

DaG 476

0.060

0.040

0.060

0.040

0.020

0.020

0.000 0.000

0.200

0.400

0.600

0.800

0.000 0.000

1.000

0.200

0.100

0.600

ALH 77005 LEW 88516 RBT 04262

0.060

0.040

Pyroxene-phyric basalts

(d) 0.080

Fe3+ (afu)

0.080

0.020

1.000

Shergotty QUE 94201 NWA 3171 Los Angeles

0.060

0.040

0.020

0.000 0.000

0.200

0.400

0.600

0.800

0.000 0.000

1.000

0.200

0.400

0.600

0.800

1.000

XFe (afu)

XFe (afu) 0.100

0.100

Clinopyroxenites

(e)

Nakhla

(f)

Orthopyroxenite and Dunite

MIL 03346

0.080

0.060

3+

(afu)

0.060

ALH 84001 Chassigny

0.080

Gov. Valadares

Fe

Fe3+ (afu)

0.800

0.100

Lherzolites

(c)

Fe3+ (afu)

0.400

XFe (afu)

XFe (afu)

0.040

0.020

0.000 0.000

0.040

0.020

0.200

0.400

0.600

0.800

1.000

XFe (afu)

0.000 0.000

0.200

0.400

0.600

0.800

1.000

XFe (afu)

Fig. 16. (a–f) Fe3+ (afu) vs. XFe (afu) for pyroxene grains from the martian meteorites. Ferric iron in pyroxene estimated by EMP analysis stoichiometry (Droop, 1987).

formation, homogeneous or heterogeneous accretion of the planetary body, the presence or absence of an early magma ocean, the size of the planetary body which defines the pressure regimes and thus the nature of the phases in the potential basalt source regions, the timing of the separation of distinct geochemical reservoirs, and the presence or absence of volatile phases in the distinct reservoirs. This is not a comprehensive list of factors that define the distinct geochemical reservoirs but it is illustrative. Since we cannot sample these distinct geochemical reservoirs directly (i.e., we cannot drill that deep, mantle nodules are unavailable or often compromised by alteration), we commonly use

basalts as “windows into the mantle sources”. Thus, Table 6 attempts to define average planetary basalts, and in examining these compositions several observations can be made: (1) the high Al content for terrestrial basalts and low Al content of martian basalts, which leads to a high-Ca/Al ratio (i.e., super-chondritic) for martian basalts relative to terrestrial, (2) the relatively high Na content of terrestrial basalts relative to those from other planetary bodies, and (3) the relatively high-Ca/(Mg + Fe) value for Earth relative to the other bodies. This ratio, combined with the high Al content of terrestrial basalts, leads to high modal plagioclase relative to olivine + pyroxene for Earth

7472

J.J. Papike et al. / Geochimica et Cosmochimica Acta 73 (2009) 7443–7485 0.040

0.040

0.035

0.030

0.030

0.025

0.025

Mn (afu)

Mn (afu)

0.035

Olivine-phyric basalts

(a)

0.020 0.015

Olivine-phyric basalts

(b)

0.020 0.015 0.010

0.010 SAU 005 Y 980459 NWA 1110

0.005 0.000 0.000

0.200

0.400

0.600

0.800

1.000

NWA 2626 NWA 2046 NWA 1195 DaG 476

0.005

1.200

0.000 0.000

1.400

0.200

0.400

Fe (afu)

1.000

1.200

1.400

0.040

Lherzolites

(c)

0.035

0.030

0.030

0.025

0.025

Mn (afu)

Mn (afu)

0.800

Fe (afu)

0.040 0.035

0.600

0.020 0.015

Pyroxene-phyric basalts

(d)

0.020 0.015 Shergotty QUE 94201 NWA 3171

0.010

0.010 ALH 77005 LEW 88516 RBT 04262

0.005 0.000 0.000

0.200

0.400

0.600

0.800

1.000

1.200

0.005

Los Angeles

1.400

0.000 0.000

0.200

0.400

0.600

0.800

1.000

1.200

1.400

Fe (afu)

Fe (afu) 0.040 0.035

(e)

Clinopyroxenites, Orthopyroxenite and Dunite

0.030

Mn (afu)

0.025 0.020 0.015 Nakhla MIL 03346 Gov. Valadares ALH 84001 Chassigny

0.010 0.005 0.000 0.000

0.200

0.400

0.600

0.800

1.000

1.200

1.400

Fe (afu)

Fig. 17. (a–e) Mn (afu) vs. Fe2+ (afu) for pyroxene grains from the martian meteorites.

(Fig. 19), and the prominence of high-Ca pyroxene relative to low-Ca pyroxene for terrestrial basalts relative to other planetary basalts (Papike, 1981). Fig. 19 also shows the near absence of olivine from basalts from 4 Vesta. The usual interpretation of this observation is that olivine is contained deep within an early magma ocean cumulate horizon, although there is also some olivine in the diogenites (Orthopyroxenites). A final observation from Fig. 19 is that “Others”, which includes glass, oxides, phosphates, hydrous silicates, sulfides, etc., are highest for terrestrial basalts.

8.2. Comparative planetary mineralogy Comparative planetary mineralogy studies work on the premise that the compositions of the silicate phases in planetary basalts should also reflect the differing chemical and physical conditions of the melts from which they crystallized. Several studies have shown that the premise holds true (see Papike, 1981, 1998; Karner et al., 2003, 2004, 2006 – Appendix data for these papers is also available at www.minsocam.org in electronic deposit). Here we add our new data on martian silicates to the existing comparative

Silicate mineralogy of martian meteorites

7473

Olivine-phyric basalts 0.12

0.12

Deficiency

Excess

0.10

Y 980459 (R3+, Ti4+) - IVAl

0.08

0.12

Deficiency

Excess

0.10

Deficiency

Fe3+ - IVAl

0.08

0.06

0.06

0.04

0.04

0.04

0.02

0.02

0.02

0.00

0.00 VI

Na

Al

Fe3+

Al

Fe3+ - IVAl

0.00 IV

Cr

2Ti

NWA 1110

0.08

0.06

IV

Excess

0.10

SAU 005

Na

Al

VI

Al

Fe3+

IV

Cr

2Ti

VI

Na

Al

Fe3+

Al

2Ti

Cr

Olivine-phyric basalts 0.12

Deficiency

0.12

Excess

0.10

Deficiency

DaG 476 (R3+, Ti4+) - IVAl

0.08

Deficiency

(R3+, Ti4+) - IVAl

0.08

0.06

0.06

0.04

0.04

0.02

0.02

Al

Na

Al

Fe

3+

Cr

2Ti

0.12

0.00 IV

Na

Al

Olivine-phyric basalts

VI

3+

Fe

Al

2Ti

Cr

IV

Lherzolites

Excess

Deficiency

(R3+, Ti4+) - IVAl

0.08

ALH 77005 Fe3+ - IVAl

0.08

LEW 88516

0.06

0.04

0.04

0.04

0.02

0.02

0.02

0.00

0.00

Na

VI

Al

Fe3+

2Ti

Al

Na

VI

Al

Fe3+

2Ti

0.12

Deficiency

Excess

(R3+, Ti4+) - IVAl

Excess

(R3+, Ti4+) - IVAl

0.08

0.04

0.04

0.04

0.02

0.02

0.02

0.00

0.00

Al

Fe3+

2Ti

Fe3+ - IVAl

0.00 IV

Cr

NWA 3171

Al

Na

VI

Al

Fe3+

2Ti

Cr

IV

Deficiency

Ti - 2IVAl

0.08

Excess

0.10

QUE 94201

Deficiency

Ti - 2IVAl

0.08

0.06

0.06

0.04

0.04

0.02

0.02

0.02

0.00

0.00

Na

VI

Al

Fe3+

2Ti

Na

Al

VI

Al

Fe3+

2Ti

Deficiency

Al

Fe3+ - IVAl

0.08

VI

Al

Fe3+

2Ti

Cr

0.12

Deficiency

MIL 03346

Na

Orthopyroxenite and Dunite

0.12

Excess

0.10

Excess

Fe3+ - IVAl

IV

Cr

Clinopyroxenites 0.12

Cr

0.00 IV

Cr

2Ti

Nakhla

0.08

0.04

Al

Fe3+

0.10

Shergotty

0.06

IV

Al

0.12

Deficiency

Excess

VI

Clinopyroxenites

0.12

0.12

Na

Al

Pyroxene-phyric basalts

0.10

Cr

Excess

0.08

0.06

VI

2Ti

Deficiency

0.06

Na

Fe3+

Al

0.10

Los Angeles

0.06

Al

VI

0.12

Deficiency 0.10

RBT 04262

IV

Na

Al

Pyroxene-phyric basalts

0.12

0.08

IV

Cr

Lherzolites

0.10

Fe3+ - IVAl

0.00 IV

Cr

Excess

0.08

0.06

Al

Cr

2 Ti

0.10

0.06

IV

Fe3+

Al

Deficiency

Excess

0.10

NWA 2626

VI

Na

Al

0.12

0.12

Deficiency 0.10

(R3+, Ti4+) - IVAl

0.02

0.00 VI

NWA 2046

0.08

0.04

IV

Excess

0.10

NWA 1195

0.06

0.00

Charge balance cations (afu)

0.12

Excess

0.10

Excess

Deficiency

Gov. Valadares

0.10

ALH 84001, Chassigny

Fe3+ - IVAl

0.08

Excess

0.10

0.08

Fe3+ - IVAl 0.06

0.06

0.06

0.04

0.04

0.04

0.02

0.02

0.02

0.00

0.00 IV

Al

Na

VI

Al

Fe3+

2Ti

Cr

0.00 IV

Al

Na

VI

Al

Fe3+

2Ti

Cr

IV

Al

Na

VI

Al

Fe3+

2Ti

Cr

Fig. 18. Range of charge balance excess and deficiency cations (afu) in pyroxene grains for each of the martian meteorites. Highlighted in each panel is the most important charge balance couple for each meteorite.

planetary mineralogy dataset, highlighting only those plots where distinct planetary trends are defined.

We have discussed plagioclase feldspar in a comparative planetary mineralogy context in three previous papers,

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Table 5 Bulk-silicate compositions of the planets.

SiO2 Al2O3 TiO2 MgO FeO CaO Na2O CaO/Al2O3 FeO/MgO a b

Eartha

Marsa

Moonb

Vestab

46.00 4.20 0.23 36.80 7.58 3.54 0.39 0.84 0.21

44.40 3.02 0.14 30.20 17.90 2.45 0.50 0.81 0.59

43.40 6.00 0.30 32.00 10.70 4.50 0.09 0.75 0.33

42.50 2.50 0.12 30.50 20.50 1.85 0.06 0.74 0.67

Wa¨nke and Dreibus (1988). Taylor (1982).

Table 6 Averagea planetary basalt compositions. Earth SiO2 (wt%) Al2O3 TiO2 C2O3 MgO FeO MnO CaO Na2O K2O P2O5

50.0 15.6 1.38 0.07 7.92 10.1 0.17 10.8 2.40 0.43 0.18

Total CaO/Al2O3 FeO/MgO CaO/(MgO + FeO)

99.0 0.7 1.4 0.6

Mars 48.9 6.55 0.96 0.47 13.5 18.0 0.48 9.13 0.91 0.07 0.95 100.0 1.5 1.7 0.3

Moon

Vesta

46.6 10.0 1.87 0.46 9.16 20.4 0.26 10.7 0.20 0.04 n.d.

48.0 12.3 0.72 0.31 6.53 19.1 0.56 10.2 0.51 0.04 0.11

99.8 1.1 2.3 0.4

98.4 0.8 2.9 0.4

a Average of a range of basalt types from each planet: Earth: ocean floor, Keweenawan, Island Arc, Hawaiian, Columbia Plateau, and Taos Plateau (BVSP, 1981). Mars: Shergotty, QUE 94201, Y 980459, DaG 476 (Meyer, 2008). Moon: 12009, 15058, Luna 24 (BVSP, 1981). Vesta: Pasamonte (BVSP, 1981).

Papike (1981, 1998) and Karner et al. (2004). In those studies it was noted that martian plagioclase plotted at higher Na and K values than terrestrial basaltic plagioclase and significantly higher than from the volatile depleted planetary bodies Moon and 4 Vesta. Fig. 20a shows these trends. In past versions of Fig. 20a it was suggested that the Mars data reflected the higher volatile (Na and K) content of the martian source compared with reservoirs on Earth. Here we question that conclusion. Table 6 suggests that terrestrial basalts have higher Na and K contents than martian basalts. Why then is plagioclase from Mars richer in Na and K than terrestrial basaltic plagioclase? The answer could be due to the relatively low Al content of martian basalts relative to terrestrial (Table 6). We make two observations: (1) albite and orthoclase, NaAlSi3O8, KAlSi3O8 only require one half the Al than does anorthite, CaAl2Si2O8 and (2) the low Al activity in martian basalts results in delayed nucleation of plagioclase while olivine and pyroxene are crystallizing. Sodium is quite incompatible in both olivine and pyroxene; estimates of partition coefficients at fO2

values of IW suggest DNa pigeonite/melt is 0.02 and 0.08 for augite/melt (Karner et al., 2007b, 2008). Therefore, while the mafic phases are crystallizing, the activity of Na is building up in the residual melt. In summary, we believe the two most important factors for the higher albite content of martian plagioclase compared with terrestrial are the low activity of Al in the martian melts and the resulting delayed nucleation of plagioclase. Fig. 20b illustrates the Cr trends for pyroxene from the four different planetary bodies. Pyroxenes from all the planets are similar in that Cr concentrations decrease with increasing XFe, but the planetary trends are distinct in that Cr is higher (at low XFe values) in terrestrial pyroxenes, and decreases more sharply with crystallization than in pyroxenes from the other planets. This may be a consequence of the Cr3+/Cr2+ ratio in the melts from the different planets, and the crystallization sequences of these melts. For instance, a high Cr3+/Cr2+ in basaltic melts on Earth leads to significant Cr (3+) entering the crystallizing pyroxene. However, Cr3+ is also entering co-crystallizing chromite and olivine. The competition for Cr leads to a rapid depletion in the melt and a subsequent rapid decrease of Cr in pyroxene. Conversely, a low Cr3+/Cr2+ in Vesta melts leads to crystallizing pyroxene without any co-crystallizing olivine or spinel and thus the Cr in the melt is drawn down much more slowly. The above paragraph gives a short explanation of the Cr systematics in pyroxenes, but the whole Cr story is much more complex and is a consequence of the Cr3+/Cr2+ ratio of planetary melts as well as the crystallization sequence of olivine, chromite and pyroxene. For example, Fig. 20c (modified from Karner et al. (2003)) shows that Cr is much more abundant in lunar olivines than in those from Mars and Earth. This trend is also seen in the bulk compositions of magmas from Earth and Moon (Papike and Bence, 1978). It was suggested that the observed depletion of Cr in terrestrial basalts as compared with lunar mare basalts was due to the possibility that most terrestrial basalts (with low Cr contents) have Cr-spinel retained in their source regions (Bence et al., 1981). Therefore, Cr in lunar melts is high, but the low fO2 on the Moon also leads to a significant portion of Cr2+ in lunar melts. Divalent Cr will not readily partition into chromite or pyroxene, but instead partitions into olivine; trivalent Cr partitions best into chromite or pyroxene, but still fits into the olivine structure. Because of its large size, Cr2+ likely enters the M1 site of olivine, similar to Ca in Ca-rich olivines. Substitution for Mg2+ or Fe2+ creates no charge-balance problems, but Cr3+ requires either a M2Na+1– M1Cr3+ or h–2Cr3+ substitutional couple. Despite this difference in substitutional mechanism of Cr2+ and Cr3+ into the olivine structure, the D-values for both in olivine are remarkably similar (Hanson and Jones, 1998). In fact, DCr (olivine/melt) remains fairly constant at 0.6 over a large range of fO2 (Mikouchi et al., 1994; Gaetani and Grove, 1996). The similarity in DCr3+ and DCr2+ in olivine is probably because, although Cr2+ substitution for Mg2+ or Fe2+ causes no charge-balance problems, the Cr2+ cation is almost too big for the olivine M1 site (similar to Mn2+). By contrast, the Cr3+ ion fits nicely into the olivine M1 site (see Papike

Silicate mineralogy of martian meteorites

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100% 90% 80%

Modal mineralogy

70% 60%

Others Plagioclase

50%

Pyroxene 40%

Olivine

30% 20% 10% 0% Earth

Moon

Mars

Vesta

Planetary basalts

Fig. 19. Histogram showing the modal mineralogy of basalts from four planets. The term “Others” includes such phases as fine-grained (glassy) matrix, oxides, sulfides, metals, phosphates, etc. Planetary basalt modes were derived from the following: Earth – average from ocean floor, Taos Plateau, Island Arc, and Hawaiian basalt samples (Karner, 2003); Lunar – average of samples 12021, 12075, and 15545 (Karner, 2003); Mars – average of Shergotty, Y 980459, QUE 94201, and DaG 476 (Stolper and McSween, 1979; Harvey et al., 1996; Mikouchi et al., 2004); Vesta – Pasamonte mode (Karner, 2003).

et al. (2005)). Fig. 20c also shows that Cr decreases with crystallization in olivines from each of the planets. We believe this trend is largely the result of early chromite crystallization, which depletes the evolving melt in Cr3+. In summary, the likely reason for the high Cr abundance in lunar olivine is that lunar melts have high Cr contents and significant Cr2+. The activity of Cr2+ in the melts is high (Hanson and Jones, 1998), and will partition into early crystallizing olivine. This scenario is supported by the XANES work of Sutton et al. (1993), which showed that Cr in lunar olivine was predominantly divalent. Fig. 20d–f shows Co, Ni systematics in olivine from Earth, Mars, and Moon. Fig. 20d shows that Ni is highest in olivine from terrestrial basalts, lower in martian olivine, and lowest in lunar olivine grains. Nickel decreases with fractionation as is expected for compatible element behavior. Fig. 20e shows the same “anomalous” behavior for Co as described in Section 6 of this paper, with a high positive slope for olivine from Earth, less for Mars, and the flat slope for olivine from the Moon. Fig. 20f shows the resulting Co/Ni variation in olivine with fractionation for the three planetary bodies. All three figures (20d–f) show distinct trends for olivines from each planet with the Mars trend being intermediary between Earth and Moon. These trends mimic the fO2 conditions of the planets where the Earth is most oxidized, Moon is most reduced and Mars intermediate. However, a rigorous experimental study by Herd et al. (2009), concluded that fO2 conditions had no significant effect on either Co or Ni partitioning into olivine. We believe the different behavior of Co and Ni in olivine for Earth, Mars, and Moon must be a result of differing melt composition and thus melt structure for the three

bodies (see Herd et al. (2009) for a comprehensive discussion). For example, Table 6 shows that average basalts from each planet have significant differences and follow systematics in the same way the different behavior of Co and Ni is exhibited. Note that FeO wt% is 10.1 for Earth, 18.0 for Mars, and 20.4 for Moon. We also note that the FeO/MgO ratio changes in the order of 1.4 for Earth, 1.7 for Mars, and 2.3 for Moon and that these differences lead to different melt structures. Although we may not know the exact reasons for the behavior of Co and Ni in basalts from these three planets, their zoning trends in a single zoned olivine crystal give strong clues to the planet from which they came. Fig. 20g and h shows that olivines and pyroxenes from Earth, Moon and Mars and Vesta have distinct Mn/Fe ratios in the order Vesta > Mars > Earth > Moon. Note that in both figures the slope of the planetary trends is given and Vesta is included only in the pyroxene plot (i.e., Vesta basalts contain no olivine). These data are similar but more comprehensive (for Mars) than what has been presented in previous papers, e.g., Papike (1998) and Karner et al. (2003, 2006). The planetary trends can be explained by the fact that Mn and Fe are two elements that behave similarly in igneous processes except that Mn is much more volatile than Fe (Schmitt and Laul, 1973; Drake, 2001). Therefore, it could be that planetary differences in Mn/Fe ratios are largely due to the volatility of Mn relative to Fe. Manganese enrichments in Mars compared with the Earth and Moon have led investigators to use the Mn/Fe ratio of materials as a fingerprint of planetary provenance (Dymek et al., 1976; Drake et al., 1989; Papike, 1998), where differences in Mn/Fe ratios can be attributed to initial accretional abundances in the planetary bodies (Drake,

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0.06 0.05 0.04

0.02

20

30

40

50

60

70

80

90

0.03 0.02

0.20

0.40

An % 500

Earth Moon Mars

(d)

Co (ppm)

Ni (ppm)

0.60

0.80

2400 1600 800

0.60

(e)

2.50

Planetary olivines

(f)

400

2.00

300

1.50

200

0.80

Earth Moon Mars

0 0.00

1.00

0.40

0.20

0.40

0.60

0.040

0.035

(h) Planetary pyroxenes

0.80

0.80

Earth Moon Mars

1.00

0.00 0.00

1.00

0.20

0.40

0.60

0.80

0.160

Vesta

(i)

y = 0.0268x + 0.0059

Earth

0.015

y = 0.0127x + 0.0004

Moon

0.010

y = 0.0059x + 0.0028

y = 0.0215x + 0.0069

0.025 0.020

Earth y = 0.0234x + 0.0016

0.015 0.010

Planetary pyroxenes

Mars Earth

0.120 0.080 0.040

0.005

0.005

0.000 0.00 0.20 0.40 0.60 0.80 1.00 1.20 1.40 1.60 1.80

0.000 0.00 0.20 0.40 0.60 0.80 1.00

Fe (afu)

Moon y = 0.0132x + 0.0023

Fe3+ (afu)

Mn (afu)

Mn (afu)

0.020

Mars

0.030

y = 0.0202x - 0.0001

0.025

1.00

XFe (afu)

Mars

0.030

1.00

0.50

0.040 Planetary olivines

0.60

Planetary olivines

XFe (afu)

XFe (afu)

(g)

0.20

XFe (afu)

100 0.40

0 0.00

1.00

Co/Ni

Planetary olivines

0.20

1000

XFe (afu)

4000

0 0.00

1500

500

0.00 0.00

100

Earth Moon Mars

Fe (afu)

1.20 1.40

0.000 0.00

0.01

0.02

0.03

0.04

Na (afu)

Fig. 20. (a–i) Summary graphs of planetary signatures found in plagioclase, olivine and pyroxene from the different bodies. Silicate data for Earth, Moon and Vesta taken from Karner et al. (2003, 2004, 2006). (a) K (afu) vs. An% in plagioclase grains from the four planetary bodies. (b) Cr (afu) vs. XFe (afu) of pyroxene grains from the four planets. (c) Cr (ppm) vs. XFe (afu) of olivine grains from the four planets. (d–f) Ni, Co, and Co/Ni (ppm) vs. XFe (afu) trends in olivine grains from the different planets. (g and h) Mn (afu) vs. Fe (afu) for olivine and pyroxene grains from the different planets along with the slopes of each array. (i) Fe3+ (afu) vs. Na (afu) for pyroxene grains from Mars and Earth.

J.J. Papike et al. / Geochimica et Cosmochimica Acta 73 (2009) 7443–7485

0

Planetary olivines

(c)

2000

0.01

0.01

0.035

Earth Moon Mars Vesta

0.04

0.03

3200

2500

Planetary pyroxenes

(b)

Cr (afu)

K (afu)

0.05

Earth Moon 4 Vesta Mars

Cr (ppm)

Planetary plagioclase

(a)

Silicate mineralogy of martian meteorites

2001). Previous studies have shown that basaltic silicate minerals (olivine, pyroxene) tend to preserve this Mn/Fe relationship (Papike, 1998; Papike et al., 2003) and now we update those Mn/Fe trends here. Furthermore, if Mn/Fe ratios are the result of initial accretional abundances and accretional position in the solar system, one would expect lunar and terrestrial silicates to have similar Mn/Fe ratios because they formed at approximately the same heliocentric distance in the solar system (Clayton et al., 1976). However, Fig. 20g and h shows that terrestrial olivines and pyroxenes have systematically higher Mn/Fe than lunar olivines and pyroxenes. Lower Mn/Fe ratios in lunar silicates are consistent with the overall volatile depletion of the Moon (Taylor, 1992), which is also consistent with a “giant impact hypothesis” for the origin of the Moon, wherein the Moon accreted from a hot, volatiledepleted debris ring around the early Earth, resulting from the giant impact. Silicates from the Moon likely record this event by crystallizing with lower Mn/Fe ratios than their terrestrial equivalents. Finally, Fig. 20i is a plot of Fe3+ (afu) vs. Na (afu) for basaltic pyroxene from Earth and Mars. Note that there is little or no ferric iron in the reduced basalts from Moon and Vesta. The plot shows a general correlation of increasing Fe3+ with increasing Na. Furthermore, while pyroxenes from the two planets overlap somewhat, terrestrial pyroxenes generally have higher ferric iron than most martian pyroxenes. The positive correlation suggests more Fe3+ can enter the pyroxene structure when there is more Na to be used in charge balancing substitutions. Higher ferric iron contents in terrestrial pyroxenes is probably due to the fact that most terrestrial basalts formed at higher fO2 conditions than martian basalts, thus there is more Fe3+ in terrestrial melts and thus more Fe3+ to partition into pyroxene. 9. MODELS FOR MARTIAN MAGMATISM 9.1. Introduction Geochemical and isotopic characteristics of martian meteorites provide a foundation for modeling of the martian mantle and basaltic magmatism. These characteristics suggest the occurrence of at least four silicate reservoirs that all formed during the primordial differentiation of Mars. As summarized by Wadhwa and Borg (2006), these reservoirs are: (1) Depeleted Mantle Reservoir 1 which is the source of LREE depleted Shergottites such as Y 980459, QUE 94201, DaG 476, and SaU 005. This source is reduced (fO2 = IW to IW + 1) and LREE depleted (147Sm/144Nd P 0.285). Further, it has additional isotopic characteristics that distinguish it from the other reservoirs: e182W P 0.6, e142Nd P 0.9 and 180 Hf/183W P 18. (2) Depeleted Mantle Reservoir 2 is the source of all the Clinopyroxenites (Nakhlites) presently in the collection of martian meteorites. Although this source is LREE depleted (147Sm/144Nd  0.255–0.266), unlike

7477

Depeleted Mantle Reservoir 1 it is oxidized (fO2 P IW + 3.5). Further, it has additional isotopic characteristics that distinguish it from the other reservoirs: e182W  2.95, e142Nd  0.74 and 180 183 Hf/ W  22–43. (3) Depeleted Mantle Reservoir 3 is the source of old martian Orthopyroxenite crust represented by ALH 84001. This source is not as depleted as the two other mantle reservoirs (147Sm/144Nd  0.214). The redox state of this source is a point of contention, but it could be between IW and IW + 1. Further, it has additional isotopic characteristics that either overlap or are somewhat different from the other depleted reservoirs: e182W  0.49, e142Nd  0.19 and 180 183 Hf/ W  19. (4) The Enriched Mantle Reservoir is either a crustal or mantle component that contributed to the LREE enriched signature of Shergottites such as Shergotty, Zagami, Los Angeles, and NWA 1110. This reservoir is oxidized (fO2 > IW + 2) and LREE enriched (147Sm/144Nd < 0.182). In addition, isotopic characteristics distinguish it from the other reservoirs: e182W 6 0.3, e142Nd 6 0.2 and 180Hf/183W 6 11. Another interesting attribute to all of these reservoirs is that they formed during the very early stages of martian differentiation. Harper et al. (1995), Borg et al. (1997, 2003) and Wadhwa and Borg (2006) concluded that the Shergottite sources formed by 4525 Ma. The source for the nahklites formed somewhat earlier at P4542 Ma (Wadhwa and Borg, 2006). In contrast, Debraille et al. (2007) concluded that the sources for the Shergottites formed early but within 100 my after core formation. All of the above workers attributed this early reservoir formation event to magma ocean formation and crystallization. An interesting conclusion drawn from these observations is that these sources have remained separate for almost the entire history of Mars (4000–4500 my). Only the Shergottite sources experienced mixing between the time of melting and transport to the martian surface. 9.2. Models for martian basaltic magmatism The identification of distinct mantle reservoirs (depleted and enriched) from which the Shergottites (basaltic Shergottites and Lherzolitic Shergottites) were derived has been placed within the context of two general models: (1) assimilation of crustal material by mantle-derived basaltic magmas (Fig. 21a) and (2) derivation of basaltic magmas from multiple mantle reservoirs (Fig. 21b and c). Here, we examine the nature and attributes of each of these end-member models. Assimilation models (Fig. 21a) propose that the Enriched Reservoir is located within the martian crust and that Depleted Reservoir 1 is the mantle source for all Shergottite basalts (i.e., Jones, 1986, 1989, 2003; Wadhwa, 2001; Herd et al., 2002; Shearer et al., 2008). Both of these reservoirs were produced during the initial stages of martian differentiation through the crystallization of a magma ocean. The upper martian mantle would be at a redox condition

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Reduced Shergottites (mantle characteristics)

Nakhlites (~1.3 Ga)

Oxidized Shergottites (crustal characteristics)

(a)

Crust Crustal Assimilation

(Enriched Reservoir)

Mantle (Depleted Reservoir 1)

Mantle (Depleted Reservoir 2)

Old Shergottites (~340 Ma)

Old Shergottites (>474 Ma)

QUE

Young Shergottites (~175 Ma)

Shergotty LA

DaG Lhzs

NWA1195

SAU Y98

Crust

(b)

NWA1110/1068

Enriched Mantle (contains trapped liquid) Enriched Source Intermediate Source

Pyroxene-phyric Depleted Source

Depleted Mantle

Intermediate Shergottites

Enriched Shergottites

Olivine-phyric

Depleted Shergottites

Nakhlites (~1.3 Ga)

(c)

Crust

Enriched Upper Mantle Pocket Nakhlite Residue Depleted Upper Mantle / Plume Mixing

Plume Head

Fig. 21. (a–c) Models of martian magmatism. (a) Crustal assimilation model modified from Herd et al. (2002). (b) Multiple mantle source model modified from Symes et al. (2008). (c) Multiple mantle source model modified from Blinova and Herd (2009). See text for detailed discussion of all three models.

of IW to IW + 1 (Shearer et al., 2008), whereas the martian crust would be fairly oxidized. The redox state of the martian crust would be controlled by hydrothermal alteration

or the emplacement of intrusions containing hydrous phases (Herd et al., 2002). In this model, varying degrees of crustal assimilation by mantle-derived basalts produces

Silicate mineralogy of martian meteorites

depleted cumulate horizons must also initiate melting. The figure illustrates scenarios where melts rise with little assimilation or crystallization until they enter magma chambers in the lower martian crust. Here crystallization commences and various mixtures of melts and crystals can be derived by fractional crystallization before being delivered to the martian surface. Symes et al. (2008) and Shearer et al. (2008) proposed that this fractional crystallization process links the Olivine-phyric basalts to Pyroxenephyric basalts within each distinct Shergottite composition (i.e., depleted vs. enriched). An example of a potential petrogenetic linkage between Olivine-phyric basalt Y 980459 and Pyroxene-phyric basalt QUE 94201 is illustrated in Fig. 22 (after Symes et al., 2008). This figure is very important because it describes the phase appearance and compositions of the only two documented liquids identified from Mars to date. Although the two-stage model has many attractive aspects explaining the martian basaltic lithologies it falls short in fully predicting the REE systematics of Shergottites. A much more complex, three stage petrogenetic model was proposed by Blinova and Herd (2009) and is illustrated in Fig. 21c. In this model, the source of the depleted Shergottites is a mixture of upper mantle cumulates that experienced prior melting and melt extraction (nahklite parental magmas) at 1.3 Ga and a deep mantle lithology that was located near the mantle–core boundary. These two sources evolved distinctly in Nd isotopic systematics due to their different Nd/Sm ratios. The deep mantle source was transported to the shallow mantle by a plume and experienced polybaric melting over a wide pressure regime. Melting the mixture of the plume and mantle residue reproduces the REE patterns of the depleted Shergottites. The enriched reservoir is a late-stage MMO cumulate lithology that exists in the upper martian mantle. The lower mantle plumes also provide a heat source to initiate melting for the enriched and intermediate martian basalts. The multiple mantle reservoir models rule out the difficulty imposed by thermodynamic and chemical requirements for assimilation of the martian crust by mantle-

1500 Spinel

Olivine Y980459 Liquidus

Fo86

1400

Temperature (oC)

the chemical and isotopic arrays exhibited by the Shergottites (i.e., Longhi, 1991; Borg et al., 1997, 2003; Jones, 2003; Borg and Drake, 2005; Wadhwa and Borg, 2006; Shearer et al., 2008). The assimilation model has the advantage of easily explaining the early separation and long term isolation of the two reservoirs; one in the crust and one in the mantle. Furthermore, the model provides a process by which the enriched and depleted components could mix during basalt generation and eruption. Multiple mantle reservoirs models: in these types of models (Fig. 21b and c), Depleted Reservoir 1 and the Enriched Reservoir are located in the martian mantle (i.e., Borg et al., 1997, 2003; Borg and Drake, 2005; Wadhwa and Borg, 2006; Shearer et al., 2008; Symes et al., 2008; Blinova and Herd, 2009). These two reservoirs must be isolated within the mantle for approximately 4000 my. These sources only mixed at the time of melting to produce the chemical and isotopic arrays exhibited by the Shergottites. A fundamental aspect of the multiple mantle reservoir models is that they are produced during the crystallization of a deep martian magma ocean (MMO). The complete crystallization sequence for the MMO is presented in Borg and Draper (2003). Within their MMO model, the enriched reservoir represents very late-stage MMO cumulates analogous to the KREEP component in lunar magma ocean models. The Depleted Reservoir 1 is MMO cumulates that preceded the crystallization of the enriched reservoir. If the magma ocean crystallized at 15 GPa at a depth of 1350 km (i.e., 0.4 of the radii of Mars), the liquidus phase would be majoritic garnet. Majorite [Mg3(Fe,Al,Si)2Si3O12] is a high pressure form of garnet where some Si is in octahedral coordination, which increases the density of the phase. Thus, the lowest cumulate layer would be majoritic garnet followed by a layer containing olivine plus majorite. The important point here is that the early garnet layers sequester significant Al. The later MMO cumulate horizons take on an increasingly high-Ca/Al ratio, which results in Shergottites having a super-chondritic Ca/Al. As the MMO continues to crystallize, late-stage cumulates became fairly oxidized, perhaps due to the disassociation of water. Alternatively, the difference in redox state between the depleted and enriched sources could be controlled by polybaric graphite–CO–CO2 equilibria (Righter et al., 2008). In contrast to the Borg and Draper model (2003), if these equilibria were responsible for the difference in the redox state of the two sources, the enriched source must be deeper in the cumulate pile than the depleted source. There are several permutations of the multiple mantle reservoir model that adjust the location of reservoirs and dynamics of reservoir mixing. Examples of two-stage (Symes et al., 2008) and three-stage models (Blinova and Herd, 2009) are presented in Fig. 21b and c and discussed below. An example of a two-stage model (after Symes et al., 2008) adapted to our sample suite, shows schematically the depleted, intermediate, and enriched source regions (Fig. 21b). Heat producing elements K, Th, U, etc. are concentrated in the late-stage crystallization products of the MMO and can be mixed by convection into the depleted cumulate horizons. To fulfill the observation that these reservoirs remain isolated, this “fertilization” of the

7479

1300

Orthopyroxene Fo81

En83Fs16

Pigeonite 1200

En73Fs22

En67Fs25

Augite

Plagioclase

QUE 94201 Liquidus

En46Fs21

En50Fs17

An71

En6Fs53

An57

1100

1000

Fig. 22. Results of MELTS algorithm modeling of Y 980459 parent composition. Crystallization sequence and mineral compositions produced from a liquid that underwent fractional crystallization. Taken from Symes et al. (2008).

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derived basalts (Symes et al., 2008). Although these models do not explain all the issues tied to the timing of reservoir isolation and mixing, they do impose their own requirements on these mantle processes. For example, mantle convection must occur on a relatively small scale if these reservoirs are separate for 4000 my. Therefore, boundaries between different reservoirs must be thermal boundaries that allow for the transport of heat but not the exchange of mass. Furthermore, the mixing process must initiate melting. Isotopic observations have been interpreted as indicating reservoirs remained isolated until just before the onset of crystallization. The petrogenesis for the Clinopyroxenites (nahklites) and Orthopyroxenites (ALH 84001) and their respective reservoirs can be placed within the context of the various Shergottite models. Models by Jones (1986, 1989, 2003) and Righter et al. (2008) place the source for the nahklites (Depleted Reservoir 2) deep within the MMO cumulate pile (e.g., Fig. 21a). Jones (2003) placed the reservoir for the nahklites in the deep mantle because it would help drive melting in the shallow mantle which consists of primarily depleted mantle with limited amounts of heat producing elements. The rationale of Righter et al. (2008) for placing the source deep is that the redox state of the Shergottite and nahklite reservoirs may be controlled by polybaric graphite–CO–CO2 equilibria. If these equilibria do control reservoir fO2, the Shergottite reservoirs must have melted between 1.2 and 3.0 GPa and the depleted reservoir must have melted to form magmas parental to the nahklites at >3.0 GPa. On the other hand, Blinova and Herd (2009) suggested that Depleted Reservoir 2 occurred in the shallow martian mantle. Part of their rationale was based on the MMO crystallization models of Borg and Draper (2003) that suggested that the upper MMO cumulate horizons were both fairly oxidized and more enriched in heat producing elements. Their solution to providing the heat required to melt depleted reservoirs was to transfer heat from the deep mantle via plumes. Depleted Reservoir 3 is the source of Orthopyroxenite ALH 84001 and is somewhat speculative because this meteorite has a crystallization age that approaches the age of martian differentiation. In addition, there is some debate concerning the redox conditions of ALH 84001. Wadhwa (2001) and Wadhwa and Borg (2006) suggested crystallization conditions of IW, whereas estimates of ferric iron in spinel based on stoichiometry suggest slightly more oxidizing conditions. Based on these redox estimates, Righter et al. (2008) suggested that the mantle reservoir from which the ALH 84001 parent magma was derived was shallower than any of the other reservoirs. Due to the expected differences in the thermal environment of Mars at 4.5 Ga compared with 1.3–0.2 Ga, it is difficult to speculate on the relationship between Depleted Reservoir 3 and other martian reservoirs. 9.3. Constraints placed on models by mineralogical observations Although the mineralogical data offered in this manuscript cannot support one specific model of martian mag-

matism over another, the data does support the basic reservoirs proposed and may also be used to constrain some aspects of the specific petrogenetic models. 9.3.1. Mineralogical fingerprints of magmatic reservoirs The different reservoirs involved in martian basalt generation have been identified chiefly on their geochemical and isotopic characteristics. However, the data presented here show that mineral compositions also allow for the identification of basalts produced from the different reservoirs. For example, the K2O/Na2O of the plagioclase and the aluminum content of the bulk rock is correlated to a rock’s depleted or enriched nature (Fig. 6). Also, Shearer et al. (2008) illustrated that the Ni–Co systematics of olivine and melt inclusions trapped in olivine are closely related to the source of the basalt. Finally, the behavior of V in olivine (Shearer et al., 2006), Cr and V in pyroxene and glass (Karner et al., 2006, 2007a,b, 2008), and Eu in pyroxene and plagioclase (Wadhwa, 2001) provide fingerprints of mantle reservoirs through the behavior of multivalent cations at different fO2 conditions. 9.3.2. Linkages among basalts As noted above, the geochemical and isotopic characteristics of martian rocks indicate three groups of Shergottites: depleted, enriched, and intermediate. The isotopic data clearly eliminates the possibility that a simple crystallization relationship exists among these three groups. However, Shearer et al. (2008) and Symes et al. (2008) suggested that basalts in each group may be linked by fractional crystallization. For example, Olivine-phyric basalts (i.e., Y 980459) could produce Pyroxene-phyric basalts (i.e., QUE 94201) through fractional crystallization. Fig. 22 illustrates this potential petrogenetic relationship. Although isotopic data, phase equilibria, and major element chemistry of silicates are consistent with a fractional crystallization linkage between these types of basalts, minor element systematics within the silicates show otherwise. For example, Cr contents in pyroxene from Y 980459 and QUE 94201 are not consistent with a fractional crystallization linkage between these two compositions (Fig. 12a and c). According to the proposed relationship, a Y 980459-type melt fractionates and eventually yields a QUE 94201-type basalt. The QUE 94201-type rock should contain pyroxenes with much lower Cr than the Y 980459 pyroxenes. Fig. 12a and c shows this is not the case. The parental Olivine-phyric magma to QUE 94201 would have be a Y 980459 bulk composition with higher Cr. 9.3.3. Location of the enriched reservoir and timing of its incorporation into martian basalts Most models of martian magmatism call upon the addition of the enriched reservoir component either through assimilation of a crustal reservoir or just prior to the melting of the mantle source. Mineralogical data constrains both scenarios. In NWA 1110, olivine (one of the earliest phases to crystallize in the Olivine-phyric basalts), appears to have crystallized from an enriched melt. Likewise, both Ni–Co–Y systematics from olivine core to rim (Shearer et al., 2008), and REE patterns from melt inclusions carry

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an enriched signature (Wadhwa and Crozaz, 2003; Shearer et al., 2008). These observations indicate that if crustal assimilation was the process that incorporated the enriched signature into the basalt, assimilation and homogenization of that component in the melt occurred before the crystallization of the olivine megacrysts. Comparisons of the Ni and Co in Y 980459 and NWA 1110 indicate that they crystallized from basalts with similar Ni and slightly different Co. This provides some constraints on both assimilation and multiple source models for the Shergottites. First, if assimilation of a crustal component is the dominant process by which the enriched reservoir signature is added to martian basalts, then the NWA 1110 parent melt would have had to have a substantially higher Mg# and Ni content than Y 980459 prior to assimilation. Assimilation and associated fractional crystallization would have resulted in substantial removal of olivine from the mantle-derived basalt. Second, according to most models, the enriched reservoir represents very latestage crystallization products of a MMO. Such a reservoir should have very low Ni abundances (analogous to lunar KREEP). The relatively high Ni abundances in olivine from NWA 1110 imply that this late-stage MMO cumulate could not be the sole contributor to the enriched Shergottites. More likely, it is consistent with mixing of late-stage MMO cumulates with Depleted Reservoir 1. Although it appears that the olivine in NWA 1110 crystallized from an enriched basalt throughout its crystallization history, Herd (2003) suggested that this early olivine crystallized under more reducing conditions than the latestage olivine. This suggests a decoupling between two characteristics of the enriched reservoir (i.e., redox state vs. enriched geochemical characteristics). This could be interpreted as indicating the equilibria controlling the fO2 in the mantle is no longer relevant because of a new crustal environment (lower P, higher redox state) of the mantle derived magma. ACKNOWLEDGMENTS This research was supported by a NASA/Cosmochemistry grant to J.J.P. which we greatly appreciate. We thank Mike Spilde for help in producing the beautiful WDS thin section maps. This manuscript was much improved by the comments and suggestions of two anonymous reviewers, Brad Joliff, and especially by Associate editor Hiroko Nagahara.

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