The bulk composition of Mars

Chemie der Erde 73 (2013) 401–420

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Chemie der Erde journal homepage: www.elsevier.de/chemer

Invited review

The bulk composition of Mars G. Jeffrey Taylor ∗ Hawaii Institute of Geophysics and Planetology, School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, HI 96822, United States

a r t i c l e

i n f o

Article history: Received 16 July 2013 Accepted 11 September 2013 Keywords: Mars Composition Planet formation Cosmochemistry Geochemistry Terrestrial planets

a b s t r a c t An accurate assessment of the bulk chemical composition of Mars is fundamental to understanding planetary accretion, differentiation, mantle evolution, the nature of the igneous parent rocks that were altered to produce sediments on Mars, and the initial concentrations of volatiles such as H, Cl and S, important constituents of the Martian surface. This paper reviews the three main approaches that have been used to estimate the bulk chemical composition of Mars: geochemical/cosmochemical, isotopic, and geophysical. The standard model is one developed by Wänke and Dreibus in a series of papers, which is based on compositions of Martian meteorites. Since their groundbreaking work, substantial amounts of data have become available to allow a reassessment of the composition of Mars from elemental data, including tests of the basic assumptions in the geochemical models. The results adjust some of the concentrations in the Wänke–Dreibus model, but in general confirm its accuracy. Bulk silicate Mars has roughly uniform depletion of moderately volatile elements such as K (0.6 × CI), and strong depletion of highly volatile elements (e.g., Tl). The highly volatile elements are within uncertainties uniformly depleted at about 0.06 CI abundances. The highly volatile chalcophile elements are likewise roughly uniformly depleted, but with more scatter, with normalized abundances of 0.03 CI. Bulk planetary H2 O is much higher than estimated previously: it appears to be slightly less than in Earth, but D/H is similar in Earth and Mars, indicating a common source of water-bearing material in the inner solar system. K/Th ranges from ∼3000 to ∼5000 among the terrestrial planets, a small range compared to CI chondrites (19,000). FeO varies throughout the inner solar system: ∼3 wt% in Mercury, 8 wt% in Earth and Venus, and 18 wt% in Mars. These differences can be produced by varying oxidation conditions, hence do not suggest the terrestrial planets were formed from fundamentally different materials. The broad chemical similarities among the terrestrial planets indicate substantial mixing throughout the inner solar system during planet formation, as suggested by dynamical models. © 2013 Elsevier GmbH. All rights reserved.

Contents 1. 2.

3. 4.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Approaches to estimating bulk composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. Models based on geochemistry and nebular components . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1.1. Wänke and Dreibus model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1.2. Morgan and Anders model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2. Estimates based on isotopic composition of Martian meteorites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3. Estimates based on geophysical properties of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.4. Summary of the models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Datasets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Complications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1. Is the Martian surface composition representative of the entire crust? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.1. Highlands megaregolith . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.2. Lava flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.3. Element ratios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2. Heterogeneity of the Martian mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

∗ Tel.: +1 808 956 3899; fax: +1 808 956 6322. E-mail address: [email protected] 0009-2819/$ – see front matter © 2013 Elsevier GmbH. All rights reserved. http://dx.doi.org/10.1016/j.chemer.2013.09.006

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G.J. Taylor / Chemie der Erde 73 (2013) 401–420

A reassessment of Martian bulk composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1. Estimating uncertainties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2. Refractory element abundances . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3. FeO and MnO . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.4. Phosphorous . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5. Moderately volatile elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6. Highly volatile elements, including halogens . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.7. Ni and Co . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8. Strongly siderophile elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.9. H2 O and D/H . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1. Mars is rich in FeO . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2. Depletion of volatile elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3. Water: abundance and source . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.4. Highly siderophile elements: implications for accretion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5. Halogen concentration of the crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.6. The core . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.7. Comparing planet compositions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction An accurate assessment of the bulk chemical composition of Mars is fundamental to the entire geologic history of Mars, including accretion, differentiation, aqueous alteration to produce sediments, and the initial concentrations of important volatile elements. For example, knowing the bulk composition in principle allows us to understand crystallization and cumulate overturn in the magma ocean (if there was one), partial melting to produce the crust through intrusion and extrusion of basaltic magmas, and the formation of the distinctive source regions (e.g., enriched and depleted shergottites) of Martian meteorites. These processes, of course, make it tricky to extract the bulk composition from geochemical and geophysical data. An example of the complexity in just modeling magma ocean crystallization and cumulate overturn can be found in Elkins-Tanton et al. (2003, 2005). Fortunately, we have a solid database for the composition of the surface of Mars and a good understanding of element behavior during petrologic processing. The database includes published meteorite analyses of the continuously expanding collection of Martian meteorites, and orbital and lander datasets. This paper reviews models for Martian bulk composition and makes a complete reassessment in light of the substantial amount of data obtained during the past two decades. It begins with an overview of the existing bulk composition models and their approaches, summarizes the datasets available and their utility, discusses the likelihood that data from the surface provides global information, and provides a complete reassessment of the bulk composition of Mars in light of new data. I emphasize the composition of bulk silicate Mars, but briefly discuss models for core composition, which are geochemically less well constrained.

2. Approaches to estimating bulk composition Three main approaches have been used to estimate the bulk chemical composition of Mars. One takes a cosmochemical approach by defining different components and determining their abundances in Mars. Another uses isotopic compositions of Martian meteorites and chondrite groups to define the abundances of the chondritic raw materials that accreted to Mars. A third uses broad geophysical properties to define mantle mineralogy. Note that these approaches have tended to be dominated by

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geochemistry, isotopic geochemistry, or geophysics, but there is no reason why a blended model cannot be used.

2.1. Models based on geochemistry and nebular components 2.1.1. Wänke and Dreibus model Wänke (1981, 1987), Dreibus and Wänke (1984, 1987), Wänke and Dreibus (1988, 1994), Longhi et al. (1992), and Halliday et al. (2001) estimate the bulk composition from element correlations in Martian meteorites, with the assumption that refractory elements are present in chondritic abundances. This model is directly tied to Mars through element abundances in Martian meteorites, and this direct link to samples of Mars has made the model exceedingly robust, explaining why it is generally considered to be the standard model for Martian bulk composition. Wänke and Dreibus assume all elements more refractory than Mn are in chondritic proportions. From the cosmochemical viewpoint, refractory and volatile tendencies are related to their condensation temperature in the solar nebula. The rough definitions of refractory and volatile and some gradations are shown in Table 1, using 50% condensation temperatures from Lodders (2003), calculated at a pressure of 10−4 atm. The 50% condensation temperature is simply the temperature at which half the mass of an element has condensed (the remainder is in the gas). The standard Wänke–Dreibus model for Mars bulk composition is shown in Tables 2 and 3. As an example of the Wänke–Dreibus approach, consider how they derive the FeO content of the mantle from the FeO/MnO ratio in meteorites. These oxides do not fractionate from each other significantly if the major phases are olivine and pyroxene (partition coefficients are about 1 for both), so igneous processes tend to preserve their mantle values. MnO values in Martian meteorites are in the range 0.4–0.6 wt%. The CI chondrite MnO concentration is 0.46 wt%, implying that Mars is not depleted in MnO. FeO/MnO in Martian meteorites is 39.1 (but see in Section 5), and FeO/MnO in CI chondrites is 100.6. Thus, if MnO is at CI abundance (0.46), then FeO in the mantle is 39.1 × 0.46 = 17.9 wt%, in reasonable agreement with bulk Martian meteorite and GRS orbital data (Taylor et al., 2006a), and with models derived from the moment of inertia (Bertka and Fei, 1998a,b). Other correlations include, for example, volatile elements with refractory elements (K/La, K/Th), correlations among the alkalis (K/Rb, Rb/Cs), and among the halogens (Br/Cl). Element correlations tend to be reliable for element pairs

G.J. Taylor / Chemie der Erde 73 (2013) 401–420

403

Table 1 Elements classified by geochemical behaviora and 50% condensations temperatures at a pressure of 10−4 atm (Lodders, 2003). Category

Temperature range (K)

Lithophiles

Siderophiles

Chalcophiles

Refractory

≥∼1300

Zr, Hf, Sc, Y, lanthanides, Th, U, Al, Ti, Ta, Nb, Ca, Sr, Ba, V, Mg, Si, Cr, (Fe) P, Mn, K, Ga, Na, Clb , Rb, Cs, F, Clb , Br, I H/H2 O, N, C

Re, Os, W, Ir, Mo, Ru, Pt, Rh, Ni, Co, (Fe), Pd

None

Au None None

As, Cu, Ag, Sb, Ge Bi, Pb, Zn, Te, Sn, Se, S, Cd, In, Tl, Hg None

Moderately volatile Highly volatile Ultra volatile

1230–800 750–250 <182

a The classification into lithophile, siderophile, and chalcophile elements is not clear-cut. In the absence of metallic iron or sulfides, for example, elements behave like lithophile elements. b Cl condensation temperature might be much lower than calculated by Lodders (2003), similar to those of Br and I.

Table 2 Major element concentrations in models for bulk silicate Mars. Wänke and Dreibusa

Morgan and Andersb

Sanloup et al.c

Lodders and Fegleyd

Khan and Connollye

SiO2 TiO2 Al2 O Cr2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O3

44.4 0.14 3.02 0.76 17.9 0.46 30.2 2.45 0.50 0.04 0.16

41.6 0.33 6.39 0.65 15.85 0.15 29.78 5.16 0.1 0.01 –

47.5 0.1 2.5 0.7 17.7 0.4 27.3 2.0 1.2 – –

45.39 0.14 2.89 0.68 17.21 0.37 29.71 2.36 0.98 0.11 0.18

44 – 2.5 – 17 – 33 2.2 – – –

Total

100.03

100.02

99.4

100.00

98.7

a b c d e

Wänke and Dreibus (1994). Morgan and Anders (1979). Sanloup et al. (1999) model EH45:H55. Lodders and Fegley (1997). Khan and Connolly (2008).

with similar bulk partition coefficients, such as the incompatible elements K and Th or La, or the compatible elements Mg and Cr. Inherent in the Wänke–Dreibus model is the concept that two components combined to form Mars (and the other terrestrial planets). It stems from ideas developed earlier by Ringwood (1979) and

Wänke (1981). Component A is reduced and does not contain elements more volatile than Mn (this was originally formulated by Wänke and Dreibus as more volatile than Na, but subsequent papers used the slightly more refractory element Mn as the cut-off). The more refractory elements are in CI abundances. Fe and siderophile

Table 3 Bulk composition of Martian crust + mantle (primitive Martian mantle).

Be (ppb) F (ppm) Na (%) Mg (%) Al (%) Si (%) P (ppm) Cl (ppm) K (ppm) Ca (%) Sc (ppm) Ti (ppm) Cr (ppm) Mn (%) Fe (%) Co (ppm) Ni (ppm) Cu (ppm) Zn (ppm) Ga (ppm) Br (ppb) Rb (ppm) Sr (ppm) Y (ppm) Zr (ppm) Nb (ppb) a b c

W-Da

M-Ab

L-Fc

52 32 0.37 18.2 1.60 20.6 700 38 305 1.75 11.3 840 5200 0.36 13.9 68 400 5.5 62 6.6 145 1.06 15.6 2.7 7.2 490

109 23.6 0.071 18.0 3.37 19.4 1985 0.88 76.5 3.68 23.5 1951 4469 0.116 12.3 – – – 41.9 2.43 4.73 0.26 35.1 6.41 38.0 1938

– 41 0.73 17.8 1.53 21.2 740 150 920 1.68 10.5 815 4640 0.284 13.4 67 140 2.0 83 4.4 940 3.5 13.5 2.8 8.3 –

Mo (ppb) In (ppb) I (ppb) Cs (ppb) Ba (ppm) La (ppb) Ce (ppb) Pr (ppb) Nd (ppb) Sm (ppb) Eu (ppb) Gd (ppb) Tb (ppb) Dy (ppb) Ho (ppb) Er (ppb) Tm (ppb) Yb (ppb) Lu (ppb) Hf (ppb) Ta (ppb) W (ppb) Tl (ppb) Th (ppb) U (ppb)

W-Da

M-Ab

L-Fc

118 14 32 70 4.5 480 1250 180 930 300 114 400 76 500 110 325 47 325 50 230 34 105 3.6 56 16

– 0.10 0.59 25.9 9.88 926 2457 309 1704 506 194 691 130 877 193 562 84 557 94 557 56 440 0.17 125 35

17 12 120 154 5.4 400 1120 167 850 250 99 395 69 454 98 300 50 277 44 229 29 80 10 56 16

Wänke and Dreibus (1994), slightly embellished by Taylor and McLennan (2009). Morgan and Anders (1979); siderophile and chalcophile elements not included. Lodders and Fegley (1997); the siderophile/chalcophile elements P, Co, Ni, Cu, Ga, Mo, In, are Tl were calculated from metal-silicate partition coefficients.

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elements are metallic in Component A. Thus, the central assumption in this model is that the refractory elements are in CI relative abundances. Component B is oxidized and contains all elements in CI chondritic abundances. These are useful constructs in understanding the components that accreted to the planets, although it is not certain that the planets were really constructed from known chondrites or their components. Morgan and Anders (1979) developed a more elaborate multi-component model.

2.1.2. Morgan and Anders model Ganapathy and Anders (1974) and Morgan and Anders (1979) modeled Mars as a mixture of chondritic materials that had been modified by the same limited set of processes that affected chondrites, such as variations in condensation temperature and fractionation of metal from silicate. The model focuses on chondritic components formed in the solar nebula. This may not be correct: as discussed below, Warren (2011) shows that terrestrial planets, differentiated meteorites, and non-carbonaceous chondrites are clearly distinguishable from carbonaceous chondrites (the low-temperature component in the Morgan and Anders, 1979, model). Morgan and Anders (1979) proposed that there are three primary condensates from the solar nebula: a high-temperature, refractory-rich condensate; Fe–Ni metal; and magnesian silicates. Morgan and Anders (1979) define a fourth component, FeS and FeO, which they postulate formed by reaction of Fe metal with H2 S and H2 O, respectively, with the FeO ending up in the magnesian silicate component. They also suggest a “remelted” component, by which they mean chondrules, hence depleted in volatiles (presumably lost during high-temperature chondrule formation). Their magnesian silicates and chondrule components are not dramatically different in composition. They further define a component rich in highly volatile elements, referred to in Morgan and Anders (1980) as the “unremelted” component. When implemented to estimate the planet’s bulk composition, they use four main components, the refractories, magnesian silicates, metallic Fe, the volatile component (including both moderately volatile and highly volatile elements). A central assumption in the Morgan and Anders approach is that elements of similar volatility do not fractionate during nebular processes, allowing them to use four “index” elements (U for refractory elements, Fe for metal, K for moderately volatile elements, and Tl or 36 Ar for highly volatile elements) to calculate the abundances of 83 elements in the planet. When Morgan and Anders reported their work in 1979, the idea of a group of meteorites being from Mars was just blossoming and quite controversial, so they did not use data from the meteorites. Their estimated Mars bulk composition is given in Table 2. Morgan and Anders (1979) defined the concentrations of the index elements for refractory (U, Th) and moderately volatile (K) elements in Mars, using gamma-ray data from the Soviet orbiter Mars 5 and thermal models available at the time to predict a value of 620 for K/Th in bulk Mars, much lower than measured by the Mars Odyssey gamma-ray spectrometer, GRS (5300; Taylor et al., 2006a) or in Martian meteorites. In spite of their estimate for K/Th being far too low, the basic approach is interesting and it is not significantly different from the Wänke–Dreibus method. The idea that refractory elements are present at CI chondrite relative abundances is common to both compositional models. The central difference is that Wänke and Dreibus, and my reassessment of the Mars bulk composition (Section 5), use multiple elements determined independently, rather than assuming, for example, that the highly volatile elements, though depleted, have CI (Wänke–Dreibus) or CV3 (Morgan and Anders) relative abundances.

2.2. Estimates based on isotopic composition of Martian meteorites Lodders and Fegley (1997), Sanloup et al. (1999), and Burbine and O’Brien (2004) focused on fitting the oxygen isotopic composition of Mars, known from Martian meteorites, to mixtures of chondritic meteorites; their results are shown in Table 2 (and Table 3 for Lodders and Fegley’s model). Sanloup et al. (1999) point out two important features of the isotopic approach. One is that oxygen is the most abundant element and other models do not determine it explicitly. The other is that a model based on oxygen isotopes has only one major assumption, in contrast to several when considering assorted components. An isotopic estimate can also be tested by data from Martian meteorites, orbiters, and landers. Lodders and Fegley’s (1997) assessment led to the estimate that Mars was constructed from a mixture of 85% H-chondrites, 11% CV-chondrites, and 4% CIchondrites. In turn, this led to a predicted value of 16,000 for K/Th in bulk Mars, much higher than GRS data indicate (5300; Taylor et al., 2006a). Sanloup et al. (1999) took a similar approach in estimating the composition of Mars, arriving at a best fit being a mixture of 45% EH and 55% H chondrites. They did not estimate the abundances of K and Th, but judging from their relatively high estimated Na content, Sanloup et al. (1999) composition, appears to be enriched in moderately volatile elements compared to Mars. Burbine and O’Brien (2004) ambitiously examined over 225 million combinations of oxygen isotopic and chemical compositions of 13 chondrite groups, testing reasonableness by comparison with assumed bulk FeO and Mg/Si and Al/Si for Mars (based on Martian meteorites). They extracted only limited major element data as their main goal was to test the feasibility that known meteorite types could be mixed to produce Earth and Mars. Burbine and O’Brien (2004) find that the average of reasonable fits to Mars bulk composition could involve contributions from all 13 chondrite types modeled, but are dominated by enstatite and ordinary chondrites, which make up about 60% of the contributors (carbonaceous chondrites make up 25% and R chondrites contribute 14%). One problem with this approach is that Mg/Si and Al/Si are not reliable parameters to distinguish Earth and Mars (Filiberto et al., 2006; McSween et al., 2009). Martian rocks, for example, range from distinctly lower than Earth in Al/Si in Martian meteorites to Earthlike or higher in rocks and soils at the Pathfinder and MER landing sites and in global GRS data (see Fig. 3 in McSween et al., 2009). Warren (2011) took a unique approach by using stable isotopes to evaluate mixing within the early solar system, specifically the isotopes of Ti, Cr, and O. His results support the notion that Mars (and Earth) could be mixtures of non-carbonaceous meteorites. He notes that for Cr and Ti especially there is a dichotomy between materials formed in the outer solar system (mainly carbonaceous chondrites) and those formed in the inner solar system (ordinary chondrites, differentiated meteorites, and the Earth, Moon, and Mars). This important observation is consistent with the results of Sanloup et al. (1999) model, which suggests that Mars is made exclusively of non-carbonaceous meteorites, and with the solutions discovered by Burbine and O’Brien (2004), which involve only 25% carbonaceous chondrites. Lodders and Fegley’s (1997) best estimate is also consistent as it indicates that the primary ingredients of Mars involved 85% H chondrites and 15% carbonaceous chondrites; the percentage of the carbonaceous chondrites component is within the limits Warren (2011) estimates for the carbonaceous contribution to Mars. As noted, the major problem with isotopic models for Martian bulk composition is that while isotopic compositions can be matched, the abundances of volatile elements are over estimated.

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2.3. Estimates based on geophysical properties of Mars Geophysical properties such as the mean density, moment of inertia, tidal Love number and dissipation factor, and radius provide independent information about the bulk composition of Mars, including the size and composition of the core. We do not yet have seismic data or heat flow for Mars, two important measurements that will greatly enhance our knowledge of the Martian interior. However, the InSight (Interior Exploration using Seismic Investigations, Geodesy and Heat Transport) mission, to be launched in 2016, will make both seismic and heat flow measurements. Even without these data, geophysical studies have contributed to an improved understanding of the Martian interior and its composition. One geophysical approach is to begin with a compositional model (usually the Wänke–Dreibus model) and calculate (e.g., McGetchin and Smyth, 1978; Longhi et al., 1992; Sohl and Spohn, 1997; Sanloup et al., 1999) or do experiments (e.g., Bertka and Fei, 1997) to determine how the mineralogy varies with pressure and temperature inside the planet. The results can be used to calculate bulk density and moment of inertia for comparison with Martian geophysical properties. A second geophysical approach is to invert geophysical data to construct mineralogical models of the interior, including phase changes (Khan and Connolly, 2008). This approach has the advantage of determining the bulk chemical composition for major elements from mineralogy and mineralogical variation with depth (pressure), with few assumptions. It does not assume CI chondrite abundances, a particular mix of chondrites, or that we can infer composition from geochemical correlations. It does not, of course, give us minor and trace element concentrations. Nevertheless, it is of great utility and serves as an independent monitor of geochemical calculations. Khan and Connolly (2008) use Gibbs energy minimization in the system CaO–FeO–MgO–Al2 O3 –SiO2 (which make up 98% of bulk silicate Mars) to calculate the equilibrium mineralogy as a function of depth and temperature inside Mars. The minimization technique to compute phase equilibria are described in detail by Connolly (2005). The results of the calculations, constrained by measured geophysical properties (mean density, moment of inertia, Love number, tidal dissipation factor) are shown in Table 2. The major element oxide concentrations are quite similar to the Wänke–Dreibus composition, though MgO is distinctly higher in the composition determined by Khan and Connolly (2008), giving Mg/Si higher than in CI chondrites, which have the highest Mg/Si of any chondrite group.

2.4. Summary of the models The Wänke–Dreibus geochemical model has become the standard for Mars bulk chemical composition. The approach used by Morgan and Anders (1979) is not fundamentally different. Its estimate is quite different from that given by Wänke and Dreibus, particularly for volatile elements for which Morgan and Anders (1979) values are significantly lower. As noted above, for example, Morgan and Anders (1979) estimate K/Th of 620, considerably less than the global surface value measured by the Mars Odyssey GRS, 5300 (Taylor et al., 2006a). However, if good GRS measurements and Martian meteorites had been available to Morgan and Anders (1979), most of the differences between their estimate and that of Wänke and Dreibus would be minimal. The essential point is that both geochemical models depict Mars as composed of a refractory component (condensation temperatures equal to or higher than that of Mn) present in chondritic (specifically CI) relative abundance, Fe partitioned between FeO and metallic Fe, and volatiles depleted.

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The oxygen isotopic approaches have the great virtue of having only one parameter to match among mixtures of chondrite groups, and oxygen is the most abundant element. Major element oxide compositions do not differ much from those of the geochemical models, though the fact is that chondrites, like planetary mantles, are all ultramafic rocks. The significant difference between isotopebased and element-based estimates is the strong enrichment in volatile elements in the isotope models (Tables 2 and 3). This enrichment is not seen in the GRS data: K/Th is 5300 for the Martian surface versus 16,400 in Lodders and Fegley’s (1997) model. Similar differences are seen in the estimated abundances of the other alkalis, halogens, and highly volatile elements such as Tl. Lodders and Fegley (1997) suggest that aqueous leaching in the mantle and hydrothermal alteration in the crust redistributed the volatile elements. Taylor et al. (2006a,b) argue against this concept on the basis of only modest variations in K/Th in the crust. It is possible that the materials that mixed to make Mars were like chondritic meteorites, as Lodders and Fegley (1997) and Sanloup et al. (1999) propose, but that these components had not yet acquired their full complement of volatiles. Alternatively, volatiles could have been lost during the accretion process (e.g., O’Neill and Palme, 2008). More likely, the accreting protoplanets were differentiated, hence with the characteristics of differentiated meteorites (e.g., lower volatile contents than chondrites). It is reassuring that the estimates based on geophysical properties of Mars are similar to those obtained by the other independent methods (Table 2). Although MgO is somewhat higher, it is probably within error of the other estimates, and FeO (17 wt%) falls in the narrow range of 15.8–17.9 wt%. The geophysical data almost certainly provide the most reasonable estimate for core size and composition, as discussed in Section 6.

3. Datasets Since the groundbreaking work of Wänke and Dreibus, substantial amounts of data have become available to allow a reassessment of the composition of Mars from elemental data. Instead of only ∼10 Martian meteorites that informed Wänke and Dreibus’ work, we now have over 60 distinct Martian meteorites. Chemical, petrologic, mineralogic, and isotopic data are compiled in the Martian Meteorite Compendium (Meyer, 2013), a crucially valuable resource. I have used the data reported in the Compendium to compile a database of Martian meteorite compositions. For each meteorite, the database consists of averages of multiple analyses of the same meteorite, but for certain sets of elements, such as many volatile chalcophile elements, data from only one study are available. A particularly useful paper besides the Compendium is the careful summary of compositional data available through 1997 provided by Lodders (1998). The meteorite data have the highest fidelity of any data available, but are restricted to the random collection of Martian meteorites. There is a bias toward younger ages in the collection, with all shergottites having ages < 500 Ga and the nakhlites and Chasigny have ages of only 1.3 Ga. The only older ages reported thus far are 4.1 Ga for ALH 84001 (Terada et al., 2003; Bouvier et al., 2009; Lapen et al., 2010) and 2.1 Ga for NWA 7034 (Agee et al., 2013). Additional data come from orbiting spacecraft and rovers. Of particular importance for understanding Martian bulk chemical properties are the data from the Mars Odyssey gammaray spectrometer (GRS). This instrument operated for eight years in Martian orbit, resulting in global data for K and Th, and equatorial data (from approximately 50◦ N to 50◦ S latitude) for Fe, Si, and H, with less quantitative information for S, Ca, and Al. Nominal spatial resolution is 5◦ , about 500 km. While not impressive spatial resolution by the standards of imaging spectrometers,

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it is ideal for global data analysis problems. The GRS probes to a depth of a few tens of centimeters. GRS data are summarized in Boynton et al. (2007, 2008) and available online at http://geo.pds.nasa.gov/missions/odyssey/grs.html. The highly successful Pathfinder and Mars Exploration Rover missions provide excellent data on rocks on the surface. Elemental data were obtained by the alpha particle X-ray spectrometer (APXS). The rocks abraded with the rock abrasion tool (RAT) on the MER missions are particularly useful as they provide complementary data to those provided by meteorites and the abrasion process limits that amount of soil masking the underlying rock. Soil analyses by themselves are useful, especially for assessing global average elemental abundances. Soil data are available online at http://pds-geosciences.wustl.edu/ missions/mer/geo mer datasets.htm. A good compilation and discussion of the data can be found in Brückner et al. (2008).

4. Complications 4.1. Is the Martian surface composition representative of the entire crust? A significant uncertainty about estimating the compositions of the crust and mantle source regions is the extent to which we can use data obtained primarily from surface samples (meteorites, rocks and soils at the landing sites, orbital data) to determine the composition of the entire crust and mantle. Taylor et al. (2006a) argue that we can use the surface value for certain parameters, such as the K/Th ratio, to represent the entire crust (and to a great extent the mantle). I summarize these and additional arguments briefly here. 4.1.1. Highlands megaregolith A substantial fraction of the crust (possibly about half; McLennan and Grotzinger, 2008; Taylor and McLennan, 2009) was constructed before the end of the heavy bombardment at ∼3.8 Ga. It would have been repeatedly excavated and mixed by impacts, especially by basin-forming events, resulting in a thick (10–20 km) megaregolith. This early period of regolith formation may have provided a substantial fraction of the present-day regolith in the absence of significant crustal recycling on Mars; that is, what goes into the Martian crust stays in the Martian crust. Thus, soils and orbital (GRS) data may be sampling more than just the most recent surface rocks and sediments. 4.1.2. Lava flows Lava flows visible at the surface of Mars are undoubtedly accompanied by larger volumes of intrusive rock stalled inside the crust. On Earth, the ratio of the intrusive to extrusive magma volumes is 5:1 (Crisp, 1984) in oceanic (basaltic) regions. If this holds for Mars, the abundant lava flows visible at the surface (themselves forming thick sequences of lavas) are accompanied by five times as much intrusive magmas with similar composition. Furthermore, magma compositions reflect the compositions of their mantle sources, assuming we can correct for fractionation processes as the magmas migrated to the surface or stalled in magma chambers. Of course, we see only the uppermost, youngest lava flows, and it is possible that magma compositions changed with time. In fact, on the basis of Mars Odyssey GRS data, Hahn et al. (2007) suggested that changes in surface compositions reflect changes in magmatic compositions with age. Nevertheless, the compositions of all available lava flows (meteorites, Gusev rocks, inferred from GRS) are highly informative about the compositions of the magmas from which the crust was constructed.

4.1.3. Element ratios The ratios of elements with very similar geochemical behavior in igneous systems will reflect their ratio in the mantle. For example, K and Th do not readily fractionate, as shown by their similar, and very low (1), crystal-melt partition coefficients (Beattie, 1993; Borg and Draper, 2003; Hauri et al., 1994). Both elements are incompatible and their concentrations in magmas are not greatly affected by source rock composition or crystallizing phase, even when garnet is involved. There are interesting exceptions, however. Th is highly compatible in phosphate minerals (Jones, 1995). Phosphates form late in the crystallization of a magma and are unlikely to be retained in a mantle source region, so probably do not play a role in fractioning K from Th during igneous processes. However, in principle, it could be significant if mantle regions were metasomatized by fluids that contained phosphate components. K is compatible in phlogopite (Halliday et al., 1995) and somewhat compatible in amphibole (Halliday et al., 1995), so if these phases were present in the Martian mantle, it could lead to fractionation of K from Th. In addition, a rock rich in K (named Jake M) has been analyzed in Gale Crater, and classified as a mugearite by Stolper et al. (2013). Terrestrial mugerites form by extensive fractional crystallization of alkaline basalts and are not abundant on Earth, implying that their formation, if accompanied by fractionation of K from Th, are not of global compositional importance. No Th data are available for the rock. Nevertheless, in general, K/Th in a lava flow reflects the ratio in its mantle source region. Similar arguments can be made for other element pairs (see Section 5). Studies of terrestrial basalts show that certain elements correlate strongly with one another (e.g., Jenner and O’Neill, 2012), implying that they have the same bulk partition coefficient. Following Wänke and Dreibus, I use such correlations as guides to searching for correlations in the Martian meteorite dataset.

4.2. Heterogeneity of the Martian mantle The shergottites exhibit a large range in geochemical and isotopic composition. Their rare earth element (REE) patterns range from severely light REE-depleted (CI-normalized La/Yb ∼0.1) with low abundances through to very slightly LREE-enriched (CInormalized La/Yb ∼1.2) with high abundances. The variations in REE patterns correlate with isotopic compositions such as initial 87 Sr/86 Sr and 143 Nd/144 Nd (e.g., Norman, 1999, 2002) and geochemical parameters such as ratios among incompatible elements (e.g., McLennan, 2003) and oxygen fugacity (e.g., Wadhwa, 2001; Herd et al., 2002). These variations have been interpreted to indicate that the shergottites represent a mixture of two distinctive sources. Nakhlites are enriched in incompatible elements (CI-normalized La/Yb ∼3) and isotopic data suggest a distinctive source region (Foley et al., 2005). The origin of these distinctive sources is important. A discussion of how they might have formed is beyond the scope of this paper, but just their existence is important for understanding how to extract the bulk composition of Mars from element ratios. In addition, Martian meteorites may not be a representative sample of the Martian crust or representative probes of its mantle. This is made particularly clear in Fig. 1, a plot of the K/Th versus K. Virtually all the meteorites are lower in K and in K/Th than the GRS global mean. This suggests that we have not sampled a significant mantle source region that is richer in volatile elements than the sources represented by the meteorites. The two prominent samples (Nakhla and lherzolite NWA 1950) with K/Th around 10,000 may be anomalous or affected by cumulate processes that fractionated K from Th. Their high K/Th does not appear to be due to terrestrial weathering: Nakhla is a fall and both have Th/U > 2, whereas samples altered on Earth have Th/U < 0.2. The central point is that

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reservoirs indicates that some caution is advised when extracting planetary bulk compositions from a non-representative set of samples. This problem is further complicated by the role of variable oxidation state throughout the mantle, which might explain some differences between the Martian meteorites and the Gusev rocks (Tuff et al., 2013), but also changes geochemical behavior of some trace elements, making the link from meteorite to mantle composition less clear. 5. A reassessment of Martian bulk composition

Fig. 1. Trace element characteristics of Martian meteorites and the average bulk surface (Taylor et al., 2006a) suggest diverse sources in the Martian mantle. Portions of the mantle must have higher K/Th to counter the lower values in meteorites compared to the Mars Odyssey gamma-ray spectrometer (GRS) mean surface composition. Similar characteristics are shown by other trace elements (McLennan, 2003; Taylor et al., 2008). This mantle diversity adds an element of uncertainty to determinations of bulk composition from trace element ratios.

Mars may contain a reservoir with higher K/Th than the GRS global mean, needed to balance the low K/Th of the meteorites (Brandon et al., 2012). The high K/Th ratio of the Martian surface measured by GRS might be caused by secondary aqueous alteration processes, implying that this ratio is not reflective of the bulk crust. Taylor et al. (2006b) explored this possibility and although plausible, a specific mechanism giving rise to a planetary-scale change in K/Th ratio of surficial materials was not identified. The existence of these

I take a geochemical approach to reassess the bulk composition of Mars using the methods outlined in Longhi et al. (1992), with emphasis on its bulk silicate composition. It is basically the same approach as reported in the insightful papers by Wänke and Dreibus, but with the addition of estimates of the uncertainties for each element and to show more of the correlations or lack of correlations than usually given. It also minimizes assumptions about condensation temperatures (e.g., components A and B), except for the unavoidable assumption that refractory elements are present in CI relative abundances. The new (only slightly modified) chemically-determined bulk composition is given in Tables 4 and 5. 5.1. Estimating uncertainties Estimates of the uncertainties in Martian bulk composition stem from analytical and sampling uncertainties for each sample or the global GRS data, and from variations from sample to sample. These can be quantified from the data. Uncertainties also

Table 4 Bulk composition in revised model for bulk silicate Mars.

Li (ppm) Be (ppb) F (ppm) Na (%) Mg (%) Al (%) Si (%) P (ppm) Cl (ppm) K (ppm) Ca (%) Sc (ppm) Ti (ppm) Cr (ppm) Mn (%) Fe (%) Co (ppm) Ni (ppm) Cu (ppm) Zn (ppm) Ga (ppm) As (ppb) Se (ppb) Br (ppb) Rb (ppm) Sr (ppm) Y (ppm) Zr (ppm) Nb (ppb) Ru (ppb) Pd (ppb) Ag (ppb)

Conc

2-␴

Methoda

Ratiob

3.0 47.7 21 0.40 18.5 1.64 20.5 675 32 309 1.74 11.0 832 4990 0.34 14.1 71 330 2.0 18.9 6.6 86 85 191 1.30 14.6 2.89 7.49 501 2.6 2.4 4.2

1.7 0.4 13 0.08 0.7 0.04 0.9 215 9 36 0.04 0.4 30 420 0.05 0.8 25 109 0.7 2.9 0.8 55 36 58 0.14 0.7 0.52 0.60 0.07 0.9 0.8 2.8

S-B S-B A-B S-A R R R S-A S-A A(GRS) R R R R S-A, D S-A S-A S-A S-A A-A S-A A-B S-A S-A S-S R R R R A-Ac A-Ac A-A

Li/Fe Be/Nd F/K Na/Al – – – P/Yb Cl/Th – – – – – Fe/Mn Co/Ni Ni/Mg Cu/Mg Zn/Sc Ga/Al As/Ce Se/Yb Br/Cl Rb/La – – – – Ru/Mg Pd/Mg Ag/Dy

Cd (ppb) In (ppb) Sn (ppb) I (ppb) Cs (ppb) Ba (ppm) La (ppb) Ce (ppb) Pr (ppb) Nd (ppb) Sm (ppb) Eu (ppb) Gd (ppb) Tb (ppb) Dy (ppb) Ho (ppb) Er (ppb) Tm (ppb) Yb (ppb) Lu (ppb) Hf (ppb) Ta (ppb) W (ppb) Re (ppb) Os (ppb) Ir (ppb) Pt (ppb) Tl (ppb) Bi (ppb) Th (ppb) U (ppb)

Conc

2-␴

Method

Ratio

9.6 6.9 38.5 36 95 4.37 439 1170 176 864 274 103 374 67.3 450 106 306 44.8 308 44.8 217 27.2 74 0.88 2.0 2.0 3.1 1.28 0.60 58 16

6.1 2.2 7.0 22 37 0.21 48 110 5.8 60 13 6 31 13.2 17 6 37 6.2 26 1.7 14 1 31 0.66 0.8 1.0 0.8 0.71 0.41 12 3

A-A S-A A-A A-A S-S R R R R R R R R R R R R R R R R R S-A A-Ac A-Ac A-Ac A-Ac S-A A-A R R

Cd/Dy In/Y Sn/Sm I/Cl Cs/La – – – – – – – – – – – – – – – – – W/Th Re/Mg Os/Mg Ir/Mg Pt/Mg Tl/Th Bi/Th – –

a R: refractory element from volatile-free CI chondrite composition. S-B: slope of correlation line for olivine-phyric and basaltic shergottites. S-S: slope of correlation line for shergottites, including lherzolitic shergottites. S-A: slope of correlation line for all martian meteorite types. A-B: average of ratio to abundance of another element for olivine-phyric and basaltic shergottites. A-A: average of ratio to abundance of another element for all martian meteorites. A(GRS): Average Mars Odyssey GRS K, Th analysis of surface. D: partition coefficient for Mn in basaltic melt divided by Mn in peridotite, using experimental data. b Ratio of element to refractory element used in slope or average methods. c Based on concentrations of element in samples containing ≥15 wt% MgO (Brandon et al., 2012).

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Table 5 Major element concentrations in revised model for bulk silicate Mars.

SiO2 TiO2 Al2 O Cr2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O3 Total

Concentration (wt%)

Uncertainty ±2 sigma

Methoda

Ratiob

43.7 0.14 3.04 0.73 18.1 0.44 30.5 2.43 0.53 0.04 0.15 99.8

1.0 0.01 0.10 0.04 1.0 0.06 0.05 0.01 0.10 0.002 0.047

Refrac Refrac Refrac Refrac Slope D Refrac Refrac Slope Average Slope

– – – – Fe/Mn – – – Na/Al K/Th P/Yb

a Refrac: refractory element from volatile-free CI chondrite composition. Slope: slope of correlation line. Average: average of ratio to abundance of another element. D: partition coefficient for Mn in basaltic melt divided by Mn in peridotite, using experimental data. b Ratio of element to refractory element used in slope or average methods.

arise from assumptions made in determining the bulk composition, such as the assumption that elements more refractory than Mn are present in CI relative abundances. I show below that to the extent we can test this assumption, the relative abundances of refractory elements are in chondritic proportions. Another assumption is that a good correlation between two elements is meaningful. This would seem to be strongest when the correlations hold for datasets involving meteorites from all mantle sources. I assessed the analytical uncertainties in several ways: For the refractory elements, I simply used the 2-sigma uncertainties determined by Lodders (2003) for CI chondrite analyses. For Mn I used the uncertainty in experimental determinations of the partition coefficients between basaltic melts and peridotites at a range of pressures. These were determined by fitting a line to the experimental data, using the least squares method outlined by York (1969). (The least-squares calculations used a spreadsheet generously provided by Randy Korotev, Washington University in St. Louis). The calculation fits a line and determines 2-sigma uncertainty in its slope and intercept. Similarly, for FeO, I used experimental data and fit a line of the crystal/liquid partition coefficient for FeO, and its 2-sigma uncertainty. For trace elements I used linear correlations when the square of the correlation coefficient (R2 ) was equal to or greater than 0.5, and applied the York approach to determine the uncertainty in the slope. Elements usually involved a refractory element

(whose composition is known through the assumption of uniform refractory element abundances) and a volatile one. In most cases, I was able to fit a line through the origin, at least to within the uncertainty of the intercept. That is, I forced the line through a zero intercept, allowing the slope to serve as a proxy for the average ratio of the two elements plotted. This approach assumes that the scatter about a line is caused by the combination of analytical precision, sampling errors, natural variation within the sample set, and small differences in bulk partition coefficients (Hanson, 1989). In some cases, R2 was less than 0.5, but the data did not scatter across a large compositional space (typically a range of less than a factor of two of the mean). In those clustered but not linearly correlated cases I calculated the standard errors of the mean of each element. The uncertainty in the ratio is given by Ra/b [( a /a)2 + ( b /b)2 ]1/2 , where Ra/b is the ratio of two elements with concentrations a and b, and  a and  b are the standard errors for elements a and b. In cases where a two-element comparison was simply too scattered to be meaningful, I did not determine the concentration of the unknown element. 5.2. Refractory element abundances The cosmochemical approach assumes that refractory elements are present in Mars at chondritic relative abundances. Below, I assume that all elements with higher 50% condensation temperatures than that of Cr are present at CI relative abundances. One way of testing this crucial, and common, assumption is to examine the abundances of the refractory lithophile elements in major chondrite groups, normalized to CI abundances (Fig. 2). In spite of some irregularities in the abundance patterns, the general impression is that the elements are largely unfractionated, as expected. Differences in the compositions of chondrite groups reflect differences in the histories and relative abundances of their constituents in the solar nebula and in their parent bodies (e.g., Huss et al., 2003). There is a curious increase in relative abundances from Y to the more refractory elements, with roughly similar slopes. (Note that the concentration axis in Fig. 2 is linear rather than the conventional logarithmic abundance.) There is also a slight decrease at the least refractory end, from Eu to Cr. The slight depletion in Cr, Si, and Mg is on the order of only 10%, which in principle could translate to a 10% uncertainty (or perhaps a systematic over estimate) in the abundances of these elements in Mars. I also tested element correlations among the refractory elements in Martian meteorites (Fig. 3), using element pairs with similar geochemical behavior. The incompatible trace element pairs Th/La,

Fig. 2. Abundances of lithophile elements (Lodders and Fegley, 1998) in chondrite groups normalized to CI chondrites (using the mean given by Lodders, 2003). Except for an increase for the two most refractory elements (Zr and Hf), relative abundances are essentially unfractionated. This gives some confidence that it is reasonable to assume that these elements are present in chondritic relative abundances in Mars.

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Fig. 3. Correlations among pairs of refractory elements, plotted to test the concept that refractory elements in Mars are present in chondritic proportions. Slopes of the correlations are within errors of the ratio in carbonaceous chondrites (using values reported by Lodders, 2003).

Zr/Hf, and Zr/Y have reasonably linear slopes with R2 > 0.7 (Fig. 3). The ratios are within the 2-sigma uncertainties of the CI ratios. The ratio of the compatible major element oxides Cr2 O3 /MgO forms a tight linear array (R2 of 0.91) with a statistically significant intercept. When the line is extrapolated to the bulk planetary MgO of 30.5 (Table 5), the MgO/Cr2 O3 ratio is 30.3 ± 4.9, within error of the CI value of 26.5. While not overwhelmingly convincing, these tests are consistent with the assumption that refractory elements are present at CI relative abundances in Mars.

Using the cosmochemical approach, I calculated the concentrations of all elements equal to or more refractory than Cr, and renormalized the sum of those elements to 100. It excludes Ni and Co, which concentrate in metallic iron during core formation. It also ignores oxygen at this stage. A complication arises for Fe as it is partitioned between both FeO and metallic Fe in Mars, indicating that in the calculation of refractory elements in bulk silicate Mars we should use only the Fe in silicate Mars. This is derived from FeO in the crust (see next section). All elements are then converted to oxides; oxide concentrations in bulk silicate Mars are given in Table 4. The uncertainties reflect the uncertainties in the average CI chondrite value and the uncertainty in FeO (see below), which matters because of the normalization to 100%. Compared to Wänke and Dreibus (1994), my revision is slightly lower in SiO2 . 5.3. FeO and MnO

Fig. 4. FeO versus MnO in Martian meteorites and rocks at the Gusev landing site on Mars (Ratted: abraded rock samples; Brushed: less abraded rock samples). Gusev olivine basalts and alkaline basalts are from a data analysis by McSween et al. (2006a,b).

Wänke and Dreibus derived the FeO concentration in Mars from the MnO content in Martian meteorites. They noted that MnO has partition coefficients for major minerals and melts close to 1, so its abundance in the mantle is the same as in the crust. The set of meteorites available to them had an average MnO content of 0.46 wt%. Noting that this was the same as in CI chondrites, and that FeO/MnO in CI chondrites was 100.6, they used the FeO/MnO ratio in Martian meteorites (39.5 ± 1.2) and the MnO in the meteorites to derive a bulk FeO of 17.9 ± 0.6. This approach is rigorous, but there are two problems. One is that the partition coefficient (MnO in peridotite/basaltic melt) is not exactly 1.0. Using experimental data reported by Takahashi and Kushiro (1983), Walter (1998), Hertzberg and Zhang (1996),

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Wasylenki et al. (2003), and Le Roux et al. (2011), it is clear that the crystal/liquid partition coefficient for MnO is less than 1.0. The average for measurements in the pressure range 10–25 kb (expected range for magma generation), is 0.93 ± 0.08 (2-␴). The revised average MnO in Martian meteorites is 0.48 ± 0.12 (ignoring Governador Valadares, which has an anomalously high MnO content in the one analysis available). The partition coefficient and the average MnO in the crust implies a bulk planetary MnO of 0.44 ± 0.08. Using the revised FeO/MnO for the larger set of Martian meteorites, 40.3 ± 2.3 (2-␴), giving an FeO concentration of 17.7 ± 2.3 wt%, essentially the same as determined by Wänke and Dreibus (17.9 wt%) In spite of the agreement between my new estimate using considerably more meteorites in the database, an important caution is in order. Igneous rocks analyzed at the Gusev landing site have considerably higher FeO/MnO than that recorded by the meteorites (Fig. 4), a point raised by McSween et al. (2009). Almost all analyses of basaltic rocks at Gusev, whether brushed or abraded to remove adhering dust and weathering products, have higher FeO/MnO than do the meteorites. McSween et al. (2006a,b) determined the compositions of two main suites of rocks at Gusev, olivine basalts and alkaline basalts, using all available data to extrapolate measured compositions to the compositions of the alteration-free igneous rocks. These are also plotted in Fig. 4 and are slightly (olivine basalts) to significantly (alkaline basalts) offset compared to Martian meteorites. This may reflect distinct source regions or oxidation state of the source regions (Tuff et al., 2013) for the magmas from which the Gusev rocks formed, and raise questions about the utility of using FeO/MnO to deduce bulk FeO in Mars. An alternative approach is to determine bulk FeO independent of its correlation with MnO. All data indicate without question that the Martian crust is richer in FeO than terrestrial basaltic rocks (∼10 wt%) as shown by FeO concentrations measured by GRS, and in meteorites, soils, and abraded rocks (Fig. 5). FeO peaks at close to 19 wt% with a global mean of 18.4 ± 0.2 (based on summing all spectra, with the 2-␴ uncertainty based on counting statistics (Taylor et al., 2006a), but lower values, down to 13 wt% in some Gusev rocks are significant. The values significantly below the meteorite and GRS distributions may indicate the presence of mantle regions with lower FeO than typical. Experimental data on FeO partitioning during melting of peridotite provide a way to calculate the mantle FeO from the mean surface FeO. Fig. 6 shows data compiled from several studies (see caption for references). It assumes that the solid/melt partition coefficient for FeO is closely approximated by the ratio of FeO in the original peridotite solid to the FeO in the melt. Data used are only for cases where the amount of melting is no more than 25%. Note the trend of decreasing D with increasing pressure. In Mars, magmas appear to form at pressures of 10–25 kb (Filiberto and Dasgupta, 2011). In this range, the apparent D(FeO)S/L straddle the D = 1 line. Using the linear fit to the data, at 20 kb, D(FeO) is 0.95 ± 0.06. Using the mean crustal FeO determined by the GRS, 18.4 wt%, and using D(FeO) of 0.95, I calculate a mantle FeO of 18.1 ± 2.2 wt% (2␴); this is the value used in Table 5. This agrees within uncertainties with the result obtained from FeO/MnO of the Martian meteorites. More importantly, the geochemical estimate agrees with the FeO determined by Khan and Connolly (2008) on the basis of bulk geophysical properties of the planet, 17 wt%. This agreement indicates that lower FeO sources like those giving rise to some Gusev basalts make up a small fraction of the Martian mantle. Agee and Draper (2004) made an independent assessment of the FeO concentration in the Martian interior through experiments on an L-chondrite composition at 5 GPa pressure. Their intent was to try to determine if the mantle source rocks for the shergottite group of Martian meteorites could be formed from a relatively FeO-rich composition like L chondrites. The L-chondrite

Fig. 5. Histograms of FeO (wt%) distributions from the Mars Odyssey gamma-ray spectrometer (GRS, from 5-degree grid points, Taylor et al., 2006a) and (bottom) Martian meteorites, abraded (Ratted) rocks from the Gusev landing site, and mean soils from Gusev, Meridiani, Pathfinder, and Viking-1 landing sites. Note the peaks at 18–19 wt%; even the lowest values are higher than typical Mid-Ocean Ridge basalts on Earth (∼10 wt%).

composition used was that of the Homestead chondrite, which is similar to the Wänke–Dreibus bulk Mars composition. They found that the Homestead composition produced magmas that contained too much FeO than even the FeO-rich shergottites contain, coupled with the correct CaO/Al2 O3 ratio, or had a reasonable FeO concentration coupled with the wrong CaO/Al2 O3 ratio. They concluded that the solution was a polybaric differentiation process,

Fig. 6. Apparent partition coefficient (concentration in initial peridotite divided by composition in equilibrium melt) versus pressure in kilobars. Dashed lines bracket typical pressures thought to represent source depths for Martian magmas (Filiberto and Dasgupta, 2011). In this pressure range, the partition coefficient is close to unity, implying that surface basalt FeO concentrations are reliable indicators of source region (hence mantle) concentrations. Data are from experimental studies by Falloon and Green (1988), Faloon et al. (1988), Takahashi and Kushiro (1983), Hertzberg and Zhang (1996), Walter (1998), Wasylenki et al. (2003), and Agee and Draper (2004).

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Fig. 7. P versus refractory element Yb for olivine phyric and basaltic shergottites. The correlation coefficient indicates a significant correlation between moderately volatile P and refractory element Yb. Knowing Yb from the assumption of chondritic refractory abundance allows calculation of the P abundance in bulk Mars.

but with a shergottites source containing an FeO content more like H-chondrites than L-chondrites, ∼15 wt% FeO. Concentrations (Table 5) of FeO (18.1 wt%) and MgO (30.5 wt%) result in an Mg# [100 × molar Mg/(Mg + Fe)] of 75 ± 4. This is in the middle of the range of Mg# in the most magnesian value in the cores of olivines in olivine-phyric shergottites, which range from 70 to 86 (Meyer, 2013). Thermal emission spectroscopy show that magnesian olivine is exposed in places on Mars, the most notable of which is a ring of the Argyre basin (Koeppen and Hamilton, 2008; Lane and Goodrich, 2010). Comparing to laboratory spectra of experimentally produced olivine suggests that the Argyre olivine occurrences are associated with olivine-rich basalts with olivine Mg# in the range 85 ± 5, within uncertainty of the most magnesian olivine in olivine-phyric basalts (86, Musselwhite et al., 2006). Taken together, these data show that the Martian mantle varies in Mg#, but that the bulk composition is close to an Mg# of 75. The somewhat lower FeO reported by Agee and Draper (2004) results in a higher Mg#, ∼79. The geophysically-determined FeO concentration of 17 wt% combined with the geochemically determined MgO of 30.5 produces a bulk Mg# of 76.8. Khan and Connolly’s (2008) MgO is higher than the geochemically based one (Table 2), and yields a bulk Mg# of 77.6. Even the highest of these values is lower than the terrestrial bulk Mg# of 89. 5.4. Phosphorous Phosphorous is a lithophile element with some siderophile tendencies. It correlates well with Yb (Fig. 7). The slope of the correlation line gives the P/Yb ratio in Mars, assuming these elements reflect the bulk composition. Multiplying by the Martian bulk Yb of 308 ppb (Table 4) gives a P concentration of 675 ppm (0.15 wt% P2 O5 ). This is somewhat depleted compared to CI chondrites (P/CI is 0.7), perhaps indicating that P was partitioned partially into the core during primary Martian differentiation. 5.5. Moderately volatile elements Gallium and sodium correlate well with Al (Fig. 8) with correlation coefficients (R2 ) of 0.92 (Ga–Al) and 0.75 (Na–Al). Using the slopes, which provide Ga/Al and Na/Al, and the concentration of Al in Mars (Table 4, 1.63 wt%), I calculate that bulk silicate Mars contains 6.6 ppm Ga and 0.40 wt% Na (0.53 wt% Na2 O). Potassium is best determined in the GRS data as it more likely represents a global average. It correlates well with Th as both are strongly incompatible large-ion lithophile elements, hence are

Fig. 8. Moderately volatile elements Ga and Na plotted against refractory element Al. Bulk Mars concentrations for Ga and Na can be determined from the Ga/Al and Na/Al ratios. Data represent analyses of all Martian meteorite types.

likely to reflect the bulk planetary ratio. K/Th varies across the Martian surface (Fig. 9). Taylor et al. (2006b) examined the possible reasons for this, with no definitive answers. It is clear, however, that the distribution is close to Gaussian and well defined. The mean K/Th is 5330 ± 440 (2-␴). In this case, the uncertainty is calculated from the sum of all the spectra obtained by the GRS, involving over 2 × 107 s of counting time. Thus, it reflects the counting statistics, not the variation in the data, but does reflect the accuracy with which we know the mean. Using the K/Th ratio and the Th concentration of 58 ppb (Table 4), I calculate a bulk K concentration of 309 ± 26 ppm. I obtain the concentrations of the other alkali elements by the correlations with La in shergottites only (Fig. 10A, C). Including the other Martian meteorites renders the correlations much weaker. Rb is particularly well correlated with La (Fig. 10A), revealing a Rb/La ratio of 2.91. Using a bulk La content of 439 ppb (Table 4) I estimate a bulk Rb concentration of 1.27 ± 0.13 ppm. Rb also strongly correlates with K and although the K/Rb ratio is commonly used to estimate planetary geochemical reservoirs, it is in principle better to use the ratio to a refractory element (La) rather than to an element (K) that is itself determined by a ratio. Nevertheless, using the strong correlation between K and Rb (Fig. 10B), I estimate a Rb content of 1.45 ± 0.25, within error of the estimate using Rb/La. This gives some credence to restricting the Rb calculation to the olivine phyric and basaltic shergottites only. The Cs (Fig. 10C) data form a good linear array. Using the bulk La value (Table 4) and a slope of 0.180, I estimate a Cs abundance of 79 ± 31 ppb. Other moderately volatile elements include As, Cu, Ag, and Sb, all somewhat chalcophile or siderophile. I found no significant correlation between Sb and any element, and the plots were so scattered that an average of Sb with a geochemically similar element would not be informative. It is not clear whether the lack of correlation is due to analytical issues or complicated geochemical behavior

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Fig. 9. Histogram from Mars Odyssey gamma-ray spectrometer determinations (5-degree grid points) of K/Th ratio of Martian surface (Taylor et al., 2006a,b). Data represent the ratio in upper few tens of centimeters of surface. Global mean and uncertainty are calculated on the basis of counting statistics of a global spectrum of over 2.4 × 107 s counting time. Uncertainty represents the uncertainty in our knowledge of the mean and does not reflect the natural variation in K/Th.

in Mars. As and Ag also do not form linear arrays when plotted against suitable elements, but do cluster sufficiently to allow estimating their abundance from averages, with Ce for As and Dy for Ag. Results are given in Table 4. Copper correlates well with Mg (Fig. 11), steadily declining with increasing MgO, indicating less compatible behavior for Cu than Mg. I model the bulk Cu as the value where MgO equals bulk Mars (30.5 wt%). This results in a Cu concentration of 2.0 ± 0.7 ppm.

5.6. Highly volatile elements, including halogens The highly volatile elements have 50% condensation temperatures less than 750 K (Lodders, 2003), and include Bi, Zn, Sn, Se, Cd, In, and Tl, plus the halogens. (Other elements in this group do not have good correlations with a refractory element. These include Pb, Te, and Hg.) The two best-determined elements are Tl and In (Fig. 12). Both correlate acceptably with a refractory element, from which we can determine their concentrations in bulk Mars (Tl is 1.4 ± 0.7 ppb, In is 6.9 ± 2.2 ppb). The other elements are determined from average values in Martian meteorites compared to the average of an element with similar geochemical behavior: Bi/Th, Zn/Sc, Sn/Sm, Se/Yb (correlates with R2 of 0.5), and Cd/Dy. Results are given in Table 4. The halogens are potentially valuable as they behave as incompatible elements, hence their concentrations in the mantle can be determined from their correlations with refractory incompatible trace elements. Furthermore, Br and I have low condensation temperatures, thus potentially helping us determine the abundance of all highly volatile elements. However, bulk Martian halogens are difficult to assess because, except for F, they are readily lost from lava flow surfaces in gas phases, so our database of mostly extrusive rocks is ambiguous with regards to the magmatic source regions in the mantle. Because loss is so common for halogens from lavas, Taylor et al. (2010) suggested that the mean surface Cl/K (1.27, determined by GRS), which is close to the chondritic value of 1.28, may reflect the bulk composition of silicate Mars. If the chondritic Cl/K is not a coincidence (a distinct possibility considering that Cl is highly mobile in aqueous fluids and heterogeneously distributed on both rover and GRS scales), then the ratio of these two incompatible elements suggests that the CI-normalized Cl abundance is about the same as that of K, 0.6. This is consistent with the concentration of Cl on the surface, 0.5 wt%. Such an elevated Cl concentration is also consistent with the similar condensation temperatures of Cl and K,

948 and 1006, respectively (Lodders, 2003). Furthermore, Cl and Br are well correlated at a ratio close to chondritic (Fig. 13B), implying that Br has a normalized abundance of 0.5, too. This is surprising in light of low 50% condensation temperature of 546 K for Br. The condensation temperatures might be wrong, of course, as they are dependent of what phases are assumed to contain trace elements. Or the chondritic Cl/K is just a coincidence and Cl is actually lower than in bulk Mars than in the uppermost crust. If we assume that the Martian meteorites did not lose their magmatic Cl, then a plot of Th versus Cl (Fig. 13A) allows us to estimate a bulk Cl value of 32 ± 9 ppm. This is lower than Wánke and Dreibus determined, 38 ppm, but within the uncertainties. Given the good correlation between Br and Cl (Fig. 13B), I estimate a Br concentration of 191 ± 58 ppb. The Br/Cl ratio is close to chondritic (Fig. 13B). Iodine was determined from the average I/Cl in Martian meteorites (data are not well correlated). The normalized abundances of Cl, Br, and I are all in the range 0.05–0.07. Low halogens are supported by analyses of melt inclusions in olivine phenocrysts in olivine-phyric shergottite Y 980459 (Usui et al., 2012). They report that only modest losses of F and Cl could have taken place because both correlate well with Na in both melt inclusions and in the glassy groundmass of the rock. Thus, it seems safe to conclude that the Martian meteorites in general record the halogen concentrations in the mantle and the original and Wánke and Dreibus approach is appropriate. 5.7. Ni and Co Ni and Co correlate well with Mg (Fig. 14). From the ratio of each to Mg and the Mg concentration in Mars (Table 4), we get a Ni concentration of 330 ± 109 ppm and a Co concentration of 71 ± 25 ppm. These are lower than the bulk planet has (∼2 wt% and ∼0.1 wt%, respectively) because most of inventory of these elements is in the core. 5.8. Strongly siderophile elements Brandon et al. (2012) present high quality analyses of highly siderophile elements (Os, Ir, Ru, Pt, and Re) in shergottites. For those with MgO greater than about 15 wt%, the siderophiles are in chondritic proportions. If those abundances represent the bulk composition of Mars, hence assuming that the siderophile elements in the other shergottites with lower MgO have been fractionated, we can estimate the abundances in bulk silicate Mars from Brandon et al.’s (2012) data. I averaged the concentrations of the elements in shergottites measured by Brandon et al. (2012), and assumed that

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Fig. 12. Correlations of highly volatile elements versus refractory elements, for the few data available for all Martian meteorite types. (A) Tl versus Th. (B) In versus Y. Correlations are only modestly strong, but sufficient to allow estimation of In and Tl in bulk silicate Mars.

Fig. 10. Correlation of moderately volatile alkali elements versus refractory La (A, C) and for Rb versus K (B). Correlations can be used to determine abundances of the alkali elements (see text). Data are for shergottites (including lherzolitic shergottites).

Fig. 11. Cu versus MgO concentrations correlate reasonably well, for all Martian meteorite types. A Martian bulk silicate concentration for Cu can be calculated from the linear fit for Cu versus MgO at the point where MgO has the bulk silicate Mars value of 30.5 wt% (Table 5).

Fig. 13. Cl–Th (A) and Br–Cl (B) correlations for all Martian meteorites. Note that Br/Cl is close to the chondritic ratio.

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Fig. 15. H2 O versus Cl for 5-degree grid points measured by Mars Odyssey gammaray spectrometer (GRS). GRS actually determines concentration of H, which has been converted to H2 O equivalent (near the surface, H could be bound as OH or H2 O). The correlation coefficient is substantially less than the 0.5 significance level required for other element pairs, but nevertheless informative (see text).

Fig. 14. Siderophile elements Ni and Co versus Mg. Correlations are reasonable and useful for deriving Ni and Co concentrations in bulk silicate Mars. Most of the Ni and Co inventories are probably in the core.

their relative and absolute abundances reflect their mantle concentrations. The results appear in Table 4 and are discussed below. This approach is probably sound for Pt because during silicate partial melting, Pt has a partition coefficient close to 1 (Jones et al., 2003). All these elements have almost identical abundances when normalized to CI chondrites (using the values for Orgueil, Horan et al., 2003), ranging from 0.0036 to 0.0045, except for Re (0.22). Excluding Re, the group of elements averages ∼50% of the value in the terrestrial primitive upper mantle (Becker et al., 2006). 5.9. H2 O and D/H Observations of the Martian surface from spacecraft revealed that its crust has been modified by substantial fluvial activity. Landed spacecraft and studies of weathering veins in Martian meteorites have repeatedly confirmed the presence of water in Mars. However, Dreibus and Wänke (1987) estimate that bulk silicate Mars contains only 39 ppm of H2 O, the equivalent of about 15% of a terrestrial ocean-equivalent and 7% of the lower estimate of 500 ppm in the bulk Earth (see review by Mottl et al., 2007). This seems insufficient considering the evidence for the role of water in the evolution of the Martian crust. Direct measurements of water in Martian meteorites indicate that the water contents of Martian magmas (e.g., McCubbin et al., 2010) are much higher than the venerable Dreibus and Wänke (1987) estimate. McCubbin et al. (2010) estimate from hydrous amphibole in melt inclusions in Chassigny that the mantle source region for the Chassigny magma contained between 130 and 250 ppm H2 O, depending on the amount of partial melting required to produce the observed magma. Leshin (2000) measured the concentration of H2 O in apatite in a depleted shergottite (QUE 94201) and McCubbin et al. (2012) measured the concentration of H2 O in apatite an enriched one (Shergotty). From the observed

concentrations and an estimate of when apatite crystallized in the lava, McCubbin et al. (2012) estimate that for these two samples, their parent magmas contained between 730 and 2870 ppm H2 O (ignoring loss from the lavas). If they were produced by 10% partial melting, their mantle source regions contained between 73 and 287 ppm H2 O. Hallis et al. (2012) measured H2 O in apatite crystals in the unweathered fall Nakhla, finding H2 O contents similar to those in Shergotty (0.46–0.64 wt%). This suggests a similar water content for the mantle source of the nakhlites; using the approach taken by McCubbin et al. (2010) I estimate a mantle concentration of 150–220 ppm. In contrast to these estimates, Usui et al. (2012) analyzed H2 O in melt inclusions in olivine in Y980459. They find no evidence for degassing of the inclusions, yet find an average of only 146 ppm in the melt inclusions, hence a pre-degassing magma content with the same value. They calculate that this indicates a mantle source containing 15–47 ppm H2 O, in the range estimated by Dreibus and Wänke (1987). One possible interpretation of these results is that water is heterogeneously distributed in the mantle. In addition, the water content of the mantle is likely to have decreased with time: the extensive fluvial alteration of the surface attests to significant water release early in Martian history. Thus, the young igneous rocks on which these mantle water estimates are based may reflect a degassed mantle that contained significantly more water initially. An independent estimate can be made from the mean concentrations of H2 O and Cl on the surface of Mars as determined by GRS (Boynton et al., 2008). Both species behave incompatibly during partial melting and both degas from lava flows. If, in spite of their complex behavior during weathering and other aqueous processing, they maintain an approximately constant ratio and both concentrate to the same extent in the upper crust, we can estimate their abundance from the H2 O/Cl ratio and our estimate bulk Cl content of 32 ppm. GRS data (which measures H in the upper few tens of centimeters) gives an equatorial (between ∼52.5 degrees north and south latitude) H2 O concentration (water equivalent of measured H) of 3.9 ± 1.9 wt% (2-␴ of total variation; standard error of 1500 points gives a 2-␴ of the mean of 0.05 wt%). The large portions of the crust at higher and lower latitudes have very high H2 O concentrations (e.g., Boynton et al., 2008), implying a much higher surface concentration, and higher H2 O/Cl. Taking a very conservative water concentration of 5 wt% for the entire surface gives H2 O/Cl of 10 and implies a bulk Mars concentration of 330 ± 10 ppm. As a rough check, a plot of H2 O versus Cl, when forced through zero (Fig. 15), gives a H2 O/Cl ratio of 8.1 ± 0.5, and a bulk Mars H2 O of 260 ± 16 (2-␴); however, the R2 is only 0.2. (The correlation line

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fit is essentially the same as using the average equatorial global H2 O/Cl ratio.) Except for the interesting results from Y980459, meteorite data and the blatant evidence for extensive water action on the surface support a relatively wet bulk composition for Mars. Geophysical modeling results seem to support the idea of a wet mantle. Thermal evolution models of crustal evolution (Hauck and Phillips, 2002; Guest and Smrekar, 2005) require water abundances of 100–1000 ppm to produce mantle viscosities low enough to allow for convection and ductile flow of mantle materials. Combining all the data, I estimate that the Martian mantle contained 300 ± 150 ppm of H2 O, but considering the morphological record for substantial surface water, the primitive mantle might have contained more than this amount. However, even a minimum value of ∼300 ppm is about a factor of ten higher than the Dreibus and Wänke (1987) estimate. Thus, their suggested mechanism for enriching FeO in Mars, reaction between H2 O and metallic iron during accretion, was not effective enough to account for the observed FeO concentration in bulk Mars. The deuterium/hydrogen ratio might be diagnostic of the sources of water to Mars (e.g., outer solar system objects such as comets versus inner solar system objects, adsorbed nebula water versus water in phyllosilicates in planetesimals). The atmosphere is enriched by a factor of 5 in D/H (␦D value of a few thousand), implying a large loss of water through sputtering by the solar wind and thermal escape of H preferentially to D (Owen et al., 1988). Weathering products in Martian meteorites also have elevated D/H (␦D of a few thousand). To understand the initial D/H in Martian water we need samples that have not been affected by atmosphere. Because plate tectonics did not recycle the crust and its modified D–H fractionated water, igneous rocks may contain the information about the isotopic composition of primary Martian water, if they have not been altered after emplacement in the crust or while on Earth. Hallis et al. (2012) measured H isotopes in apatite in Nakhla, a well-preserved fall. Terrestrial alteration is not detectable and Martian weathering is identifiable. Hallis et al. (2012) show that the Nakhla parent magma had water with a terrestrial-like ␦D (−78 to +188); the terrestrial mantle has ␦D of −140 to +60 (Boettcher and O’Neil, 1980; Michael, 1988; Ahrens, 1989; Deloule et al., 1991; Bell and Rossman, 1992; Thompson, 1992; Graham et al., 1994; Jambon, 1994; Wagner et al., 1996; Xia et al., 2002). Usui et al. (2012) measured D/H in melt inclusions in olivine in Y980459, finding a ␦D of +275. It appears that bulk Mars began with a D/H similar to that of Earth, though it could be slightly elevated.

6. Discussion The composition of bulk silicate Mars derived above is shown in Tables 4 and 5. Fig. 16 shows the results normalized to CI chondrites (Lodders, 2003, except for highly siderophile elements where I use data for Orgueil, Horan et al., 2003), and plotted in order of decreasing 50% condensation temperature (hence increasing volatility). Cl is plotted with Br and I, rather than its calculated condensation temperature. The pattern for lithophile elements is the familiar one with uniform refractory element abundances (assumed, but reasonably so, Figs. 2 and 3). The exceptions are Fe and P, where significant fractions were fractionated into the core during primary differentiation. Mars has roughly uniform depletion of moderately volatile elements (0.6 × CI), and strong depletion of highly volatile elements. The highly volatile elements are within uncertainties uniformly depleted at about 0.06 CI abundances. The highly volatile chalcophile elements (Fig. 16) are likewise roughly uniformly depleted, but with more scatter. They have normalized abundances of 0.03 × CI. I discuss these abundances in more detail

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Fig. 16. Abundances in bulk silicate Mars normalized to CI chondrites for lithophile (top), chalcophile (middle), and highly siderophile (bottom) elements. Variations in chalcophile elements probably reflect the extent to which the elements concentrate in sulfide phases or their behavior when conditions force them to behave as lithophile elements. Except for Re, highly siderophile elements are present in CI relative abundances (Brandon et al., 2012) and have abundances about a hundred times higher than expected if they equilibrated with a metallic phase during core formation.

below with the goal of showing the utility of knowing Martian (and other planetary) compositions. The striking feature of the revised composition is how similar it is to that derived by Wänke and Dreibus, in spite of analyses of numerous newly found Martian meteorites, orbital and lander geochemical data, and improved global geophysical data. The robustness of the model stems from their geochemical insight about element behavior, allowing us to determine volatile elements from the assumed abundances of refractory elements and to assess bulk FeO, and the certainty that we have meteorites from Mars. It seems remarkable that the conclusions reached by Wänke and Dreibus based on only ∼10 SNC meteorites stand the test of abundant new data. The only significant difference between my reassessment and the original Wänke–Dreibus model is in the concentration of H2 O. Other differences are a matter of degree but

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do not change the basic picture: moderately volatile elements are depleted by a small amount compared to CI chondrites (factor of about 0.6), while highly volatile elements are depleted by a factor approaching 50. 6.1. Mars is rich in FeO There is no doubt that the Martian mantle is much richer in oxidized iron than is the terrestrial mantle. Both geochemical and geophysical data confirm it. One might argue that melting of a hydrous mantle might have enriched magmas in FeO (Nekvasil et al., 2007), but Taylor et al. (2006a) marshal experimental data to argue that any magma generated by partial melting of a wet Martian mantle and subsequent fractionation of that hydrous magma would lead to lower FeO than in the original mantle peridotite. Orbital measurements of Mercury (e.g., Peplowski et al., 2011) confirm the suggestion (Robinson and Taylor, 2001) that it has a low FeO content (2–4 wt%). Robinson and Taylor (2001) argue on the basis of surface analyses of Venus (Surkov et al., 1987) that its bulk FeO is close to that of Earth, about 8 wt% (e.g., McDonough and Sun, 1995). It is tempting to suggest on the basis of these few data points (only four, though they represent 100% of the terrestrial planets!) that there is a gradient in the oxidation state throughout the inner solar system. Another way of looking at this is that the planetesimals accreting to form the terrestrial planets varied in their oxidation state (E chondrites to L chondrites, for example): Mercury received a bigger share of the reduced chondritic raw materials than did Earth than did Mars. This is consistent with the modeling designed to match oxygen isotopic compositions (Sanloup et al., 1999; Lodders and Fegley, 1998; Burbine and O’Brien, 2004). Alternatively, perhaps the FeO was made on Mars by oxidation of metallic iron. Wänke and Dreibus (1994) hypothesized that the higher FeO was caused by metal oxidation by H2 O during the accretion of Mars. As discussed by Bertka and Fei (1998b), an oxidation event like this would increase the sulfur content of a metallic core. Whatever the cause, the different FeO concentrations of the terrestrial planets must be taken into account when modeling planetary accretion. 6.2. Depletion of volatile elements Volatile elements fall into two distinct groups on the abundance plots (Fig. 16). One involves the moderately volatile (50% condensation temperatures between ∼800 and 1100 K) lithophile elements K, Ga, Na, Rb, Cs, and F, and the moderately volatile chalcophile elements As, Cu, and Ag. The lithophiles are depleted by small factors compared to CI chondrites (the mean depletion is 0.6 × CI). In contrast, the moderately volatile chalcophile elements are significantly more depleted, average 0.03 × CI. This suggests that their depletion is driven by core formation (probably S-rich, see below) rather than volatility. Highly volatile elements (50% condensation temperatures <750 K) are strongly depleted. Lithophile element (Cl, Br, I) abundances are 0.06 × CI and chalcophile elements (Bi, Zn, Sn, Se, Cd, In, and Tl) abundances are 0.03 × CI. The chalcophile elements have a large range in depletion factors, from 0.004 (Se) to 0.09 (In). This range likely reflects the combination of formation of a core rich in S and once sulfide was depleted in the mantle subsequent lithophile partitioning of the elements. All the volatile elements are combined in Fig. 17, along with a few refractory elements for reference, and compared to terrestrial abundances. Moderately volatile elements are somewhat more depleted in Earth than in Mars (0.4 versus 0.6 × CI), but on average overlap for the highly volatile elements. The depletion patterns for the two planets track each other, discrepant most strongly for Br and Tl. In spite of the difference for moderately volatile elements, the two bodies are quite similar. The biggest difference between

Fig. 17. CI-normalized abundances for Earth and Mars compared. Refractory elements between Cr and Sr are shown for context; the full list of concentrations for Mars is shown in Table 4. The overall pattern suggests similar compositions for Earth and Mars, though Mars contains somewhat higher concentrations of moderately volatile elements.

Mars and Earth is that Mars has more than double the FeO concentration and a sulfur-rich core (see below). This emphasizes the similarity among the terrestrial planets. They may have different mixtures of the ingredients that delivered the refractory elements, but the sources for the volatile elements may have been quite similar. The methods that use oxygen isotopes to derive the composition of Mars only fail because they add too much of the volatile component. Perhaps the chondrite groups proposed had similar refractory element abundances, but accretion occurred mostly before the volatiles condensed. 6.3. Water: abundance and source Mars appears to have somewhat less H2 O than does Earth, 300 ± 150 versus 500 ppm; the terrestrial value is a minimum (Mottl et al., 2007). In spite of appearing to have less water than bulk Earth, Mars has higher concentrations of moderately volatile elements and about the same concentrations of highly volatile elements. This is, of course, convoluted because most highly volatile elements are chalcophiles. If we consider only Cl, Br, and I, which are lithophile, then Mars appears to be enriched compared to Earth, 0.06 versus 0.02 × CI (Fig. 17). Assuming that the highly volatile elements were delivered to accreting Mars in waterbearing planetesimals resembling carbonaceous chondrites, then those objects contained less water than did those accreting to Earth. I estimate that the planetesimals contributing the 0.06 × CI contribution of halogens to Mars would have contained on average only 0.5 wt% H2 O to produce a bulk Martian water content of 300 ppm. The volatile-bearing planetesimals used in constructing the Earth would have contained about 3 wt% H2 O to produce a bulk water content of 500 ppm while adding highly volatile elements to bring the terrestrial inventory to only 0.02 × CI abundances. The somewhat damp but otherwise volatile-rich planetesimals are likely to have been formed from typical inner solar system materials. Otherwise, the D/H of Martian water would not be so similar to the terrestrial value (Section 5.9). Except for Jupiter-family comets, which have D/H similar to the Earth (Hartough et al., 2011), astronomical measurements of D/H in comets suggest that outer solar system materials have elevated D/H. Thus, it appears that the inner solar system objects were fed from the same source of H2 O, or at least from sources with the same D/H ratio. This analysis of total water in Mars assumes that the planet did not lose water during accretion or afterwards. It is possible, however, that water may have been lost preferentially by impact heating during accretion (e.g., Bond et al., 2010), might have reacted with metallic iron during accretion, accounting for the high FeO

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in Mars (Wänke and Dreibus, 1994), or lost during magma ocean crystallization (Elkins-Tanton et al., 2005).

Table 6 Comparisons for two important chemical parameters among inner solar system objects.

6.4. Highly siderophile elements: implications for accretion The significance of the concentrations of highly siderophile elements in Mars Earth, and the Moon were reviewed in detail by Walker (2009). New analyses and an updated discussion appear in Brandon et al. (2012). The highly siderophile elements are generally in chondritic relative abundances (except for Re) and are present at 0.04 × CI concentrations (Fig. 16). The terrestrial primitive upper mantle contains about twice this abundance (Becker et al., 2006). The abundances in Mars are quite depleted (0.04 × CI), but if equilibrated with metallic iron during core formation they would be another factor of 40–50 times lower because their lowpressure metal-silicate partition coefficients are around 10,000. As explained in detail by Walker (2009) and Brandon et al. (2012), three explanations have been advanced to explain the surprisingly high concentrations of highly siderophile elements: (1) inefficient core formation; (2) equilibrium at high pressures where the partition coefficients are lower; and (3) late addition of materials rich in siderophile elements. As Brandon et al. (2012) explained, all of these ideas have flaws, but noted that late addition may be the most straightforward explanation. One problem that deserves more attention is how a chondritic ratio can be preserved during magma ocean crystallization and subsequent mantle melting.

417

Mercury Venus Earth Moon Mars

K/Th

Refs

FeO (wt%)

Refs

5200 ∼3000 2900 360 5300

a b c d e

3 8 8 13 18

f,g f c h i

(a) Peplowski et al. (2011); (b) Surkov et al. (1987); (c) McDonough and Sun (1995); (d) Warren and Wasson (1979), Warren (1989); (e) Taylor et al. (2006a); (f) Robinson and Taylor (2001); (g) Nittler et al. (2011); (h) Taylor et al. (2006a); (i) Wänke and Dreibus (e.g., 1988) and this paper.

a core radius of 1680 km and a composition (derived from calculated core density) of 75–78 wt% Fe + Ni and 22–25 wt% S. Using the new estimated Martian bulk composition, the relative amounts of Fe and FeO (total Fe, as for all refractory elements, is about 1.9 times CI), assuming that essentially all the Ni and S are in the core (ignores small fractions in the silicate portion), and assuming an original S abundance similar to the moderately volatile elements (0.6 × CI), I estimate a core composition of 78.6 wt% Fe + Ni and 21.4 wt% S. Better estimates await a determination of the core size by seismic measurements to be done by the Interior Exploration using Seismic Investigations, Geodesy and Heat Transport (InSight) mission, scheduled to be launched in 2016.

6.5. Halogen concentration of the crust

6.7. Comparing planet compositions

If the bulk Cl abundance is as low as it appears (Section 5.6), the high Cl in the uppermost surface has interesting implications for crustal evolution. Half the K in Mars is in the mantle (Taylor et al., 2006a). This should apply to Cl as well because like K it is highly incompatible during igneous processing, implying that the crustal contribution to the total inventory is 16 ppm. Since the crust averages 57 km thick (Wieczorek and Zuber, 2004), making up about 4.6 wt% of bulk silicate Mars, its mean Cl content should be around 350 ppm. GRS and MER data show that the average surface Cl concentration is about 5000 ppm. This indicates that Cl is concentrated in the uppermost crust. Assuming that the upper few meters are not exceptionally enriched, mass balances indicate that the entire crustal inventory of Cl could be confined to the upper ∼4 km. If correct, this suggests efficient aqueous transport of Cl from deep in the crust to the surface, or continuous aqueous transport to the surface during construction of the crust. This analysis is complicated by the unknown extent to which Cl could have been lost in a gas phase even at depth in the crust; such loss could also fractionate Cl from K. Nevertheless, to first order, it seems likely that Cl is systematically concentrated toward the surface, illustrating the important role of water in the crustal geochemistry of Mars. It is interesting to consider the ramifications if Cl is much more abundant, as argued by Taylor et al. (2010) from the chondritic mean Cl/K ratio of the surface and calculated similar 50% condensation temperatures (Lodders, 2003). Considering that Cl/Br is chondritic and Cl/I close to it, this implies that all the halogens have CI-normalized abundances of 0.6. If the condensations temperatures of Br and I are as low as Lodders’ (2003) calculations indicate, then all the highly volatile elements would have been present in Mars at about the same CI-normalized abundance, implying that core formation reduced the abundances of the highly volatile chalcophile elements from 0.6 to the observed ∼0.03.

A central reason for determining planetary compositions is to compare them to deduce variations in compositions, conditions in the solar nebula and chemical uniformity of it, accretion processes (including the extent of mixing), and differentiation styles (e.g., floatation crust or not). Planetary scientists often emphasize differences among the terrestrial planets, but the similarities are striking (Table 6). The close similarities in the D/H ratios of Mars, Earth, carbonaceous chondrites, and Jupiter-family comets suggest a common source of water-bearing material in the inner solar system (Alexander et al., 2012). K/Th (Table 6) varies among the terrestrial planets: Mercury, 5200 ± 1800 (Peplowski et al., 2011); Venus, ∼3000 (Surkov et al., 1987); Earth, 2900 (Jagoutz et al., 1979; McDonough and Sun, 1995; Taylor and McLennan, 2009); Mars, 5300 ± 220 (Taylor et al., 2006a). However, these variations seem less significant when compared to the K/Th ratios of the carbonaceous chondrites (19,000 – McDonough and Sun, 1995). The terrestrial planets are depleted in volatile elements compared to carbonaceous chondrites, but the depletions are not correlated with distance from the Sun. Oxygen isotopes are distinctive among planets and meteorite groups (Mittlefehldt et al., 2008). The difference between Earth (17 O of 0‰ by definition), and Mars (17 O of + 0.25‰), is much smaller than the total range observed among chondrites 17 O of −4.3 to +2.5‰). Warren (2011) emphasizes the similarity among terrestral planets and differentiated meteorites compared to carbonaceous chondrites in the isotopic compositions of O, Cr, and Ti. Some chemical parameters do vary directly with heliocentric distance. Bulk silicate FeO (Table 6) increases from ∼3 wt% in Mercury (Robinson and Taylor, 2001; Nittler et al., 2011), to 8 wt% in Earth (McDonough and Sun, 1995) and Venus (Robinson and Taylor, 2001) to 18 wt% in Mars. FeO and the size of the metallic cores are inversely correlated. These trends suggest a range in oxidation conditions correlated with heliocentric distance, in contrast to the apparent weak correlation with heliocentric distance for D/H, K/Th, and oxygen isotopes. As these differences can be produced by varying oxidation conditions, they do not suggest the terrestrial planets were formed from fundamentally different

6.6. The core The most reliable estimate of core composition may be the one by Khan and Connolly (2008) based on geophysical data. They find

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