Does the Moon Have a Metallic Core?

Icarus 158, 1–13 (2002) doi:10.1006/icar.2002.6859

Does the Moon Have a Metallic Core? Constraints from Giant Impact Modeling and Siderophile Elements Kevin Righter Lunar and Planetary Laboratory, University of Arizona, Tucson, Arizona 85721 E-mail: [email protected] Received July 20, 2001; revised January 11, 2002

may represent a chemical “fingerprint” of core formation, having been differentially extracted into a metallic core according to the magnitude of metal/silicate partition coefficients. In the past 30 years, estimates of siderophile elements in the Moon’s mantle (Figs. 1 and 2) have been used to argue for the presence of a small (0.1–5.5 mass%; Table I) metallic core, based mainly on experimental partition coefficients determined at low temperature and pressure (T = 1200–1300◦ C and P = 1 bar; e.g., Newsom 1984, Ringwood 1979, O’Neill 1991, Drake 1987). Several developments have made a reexamination of this issue prudent. First, metal–silicate partitioning experiments have been carried out at higher pressures and temperatures, closer to the conditions prevailing in a lunar magma ocean. Although it was demonstrated by Righter and Drake (1996) that the Moon could have separated a small core at moderate pressure (50 kbar) and temperature (1900◦ C) conditions, this conclusion was based on fewer experimental partition coefficient data; more recent compilations (Righter and Drake 1999) include close to 300 data for Ni and 200 for Co and cover a wide P, T, oxygen fugacity (fO2 ), and composition range. Second, many previous models of core formation in the Moon have assumed a lunar bulk composition similar to Earth’s primitive upper mantle (e.g., Delano and Ringwood 1978; Ringwood and Seifert 1986, Righter and Drake 1996). However, in the most successful dynamical models for making the Moon (e.g., Cameron 2000, Canup and Asphaug 2001), the Moon is made from material derived from the mantle of the impactor. Third, estimates of the size of the lunar core have improved through acquisition of new data in the past several years from the Lunar Prospector and Clementine missions (Hood and Zuber 2000). Here the issue of whether the Moon has a small metallic core will be reevaluated in light of these new constraints.

The issue of whether the Moon has a small metallic core is reexamined in light of new information: improved dynamical modeling, new constraints on core size, and high temperature and pressure metal–silicate partition coefficients. Addressed specifically is the question of whether the Moon’s siderophile element budget can be explained by derivation of the Moon from a differentiated impactor or proto-Earth (stage 1), followed by formation of a small metallic core within the Moon (stage 2). If the Moon is made of mantle material from either a “hot” impactor or a “warm” impactor or proto-Earth, a small metallic core (0.7 to 2 mass%) is predicted. If the Moon is made from mantle material from a “hot” protoEarth, the lunar mantle would be more depleted in W or Re than is observed. Scenarios in which the Moon is made from impactor or proto-Earth mantle material that has equilibrated with metal at low pressures and temperatures (“cold” scenarios) would yield a much larger metallic core than observed. Finally, the greater depletions of Ni, Mo, and Re in the Moon (relative to the Earth) can be explained by low PT and reduced metal–silicate equilibrium in an impactor without later core formation in the Moon (i.e., no stage 2), but depletions of Co, Ga, and W cannot. Altogether, geochemically unlikely or geophysically inadequate non-metallic core alternatives, substantial geophysical evidence for a metallic core, and the successful models presented here for siderophile element depletions all favor the presence of a small lunar metallic core. Previous geochemical objections to an impactor origin of the Moon are eliminated because siderophile element concentrations in the lunar mantle are consistent with separation of a small core from a bulk Moon derived from impactor mantle material. c 2002 Elsevier Science (USA) Key Words: geochemistry; cosmochemistry; Moon; Earth; interiors.

1. INTRODUCTION

Siderophile (iron-loving) elements are sensitive indicators of metallic core formation in planets and comprise nearly 30 elements. Chemical analysis of lunar meteorites and returned samples permits the estimation of the abundances of siderophile elements in the lunar mantle (e.g., W¨anke et al. 1974, Newsom 1984). The depletion of siderophile elements in the lunar mantle

2. BACKGROUND

2.1. Giant Impact Scenarios The initial giant impact models involved a giant, late impact of a Mars-sized impactor with a nearly fully formed proto-Earth 1 0019-1035/02 $35.00 c 2002 Elsevier Science (USA)  All rights reserved.

2

KEVIN RIGHTER

of various impact conditions, such as total mass, impactor– proto-Earth mass ratios, and total angular momentum (e.g., see Cameron 1997, Canup et al. 2001), to determine under what conditions enough material could be placed into Earth’s orbit and outside the Roche limit. There are currently two successful classes of Moon-forming impact models, satisfying both angular momentum and orbiting mass constraints: those involving an early impact with a half-formed proto-Earth and an impactor– proto-Earth mass ratio of 3 : 7 (e.g., Cameron 2000), and those involving a late impact with an impactor–proto-Earth mass ratio of 1 : 9 (Canup and Asphaug 2001). In either scenario, the Moon is made of mantle material from the impactor and was subjected to very high temperatures. Under such high temperature conditions, a metallic core would separate instantaneously, but postimpact conditions could range from “hot” (∼3500◦ C) to “warm” (∼2000–2500◦ C) to “cold” (1200–1400◦ C), depending upon the specific cooling history as will be discussed further below. Based on Hf/W isotopic measurements (Lee et al. 1997), and Hf and W partitioning constraints (Shearer and Righter 2000) the Moon formed approximately 50 Ma after T (0) and the lunar magma ocean system cooled rapidly (<10 Ma). In a previous attempt to combine impact modeling with siderophile element partitioning between core and mantle in the Moon (Newsom and Taylor 1989), two scenarios were identified that could potentially explain the siderophile element composition of the Moon: the impactor mantle was similar to Earth’s mantle and required segregation of a small lunar core, and the impactor mantle was similar to the Moon and required no lunar core. This contribution focuses on these two general possibilities and utilizes the expanded siderophile element partition coefficient database that is now available (see next section).

FIG. 1. Depletion diagrams for the compatible siderophile elements Ni and Co, using Ni–MgO, Co-(MgO + FeO) correlations. Lunar and terrestrial Ni and Co data are from Delano (1986) and references therein. Crosses show the range of Moon compositions (MgO and FeO + MgO/MgO), which allow estimates of Ni and Co by extension of the basalt trend. See text for further explanation.

2.2. Metal–Silicate Experimental Studies Almost all previous siderophile element modeling efforts have used a set of partition coefficients with fixed values, determined at low temperatures (1200–1400◦ C) and pressure (1 bar) (e.g., Newsom 1984, Drake 1987). It is generally recognized, however, that the environment of the early Moon was hot enough to melt the mantle (e.g., Shearer and Papike 1999). Metal–silicate

(e.g., Cameron and Benz 1991, Newsom and Taylor 1989). Because it was realized later that much of the material ejected into orbit after such an impact would be tidally disrupted and accreted back to the Earth (within the Roche limit; Canup and Esposito 1996, Ida et al. 1997), there began an exploration

TABLE I Summary of Results of Previous Lunar Core Siderophile Elements

Core mass Core Ni (%) x Bulk Moon

1

2

3

4

5

6

2–5.5% 25–12 0.02–0.09 CI

∼1 35–55 0 PUM/CI/H

0.4 40 0 PUM

1 43 1.0 PUM/CI/H

5 8.3 1.0 PUM/CI/H

0.7–2.0 20.0–31.3 1.0 (1) protoE (2) impactor (3) PUM

Note. (1) Newsom (1984). (2) O’Neill (1991). (3) Ringwood and Seifert (1986). (4, 5) Righter and Drake (1996). (6) This study. Bulk Moon abbreviations: PUM, primitive upper mantle; CI, CI chondrite; H, H chondrite; protoE, proto-Earth.

DOES THE MOON HAVE A CORE?

3

FIG. 2. Depletion diagrams for the incompatible siderophile elements, Mo, W, Re, P, and Ga, using Mo-Pr, W-Ba, P-La, Re-Yb, and Ga-Al correlations. (A) explains how depletions can be defined for several hypothetical planets, and data for the five incompatible elements considered in this work are presented in (B)–(F). Data sources: Lunar Mo and Pr: Taylor et al. 1971; Newsom and Palme 1984; terrestrial Mo and Pr: Newsom et al. 1986; Lunar P, La, Ga, Al, W, and Ba: BVSP 1981; Warren et al. 1986, W¨anke et al. 1972, 1974, 1977; Palme et al. 1978; terrestrial W and Ba: Newsom et al. 1996; terrestrial P and La: BVSP 1981, Newsom and Drake 1983; and references therein; terrestrial Ga and Al: Dickey et al. 1977, BVSP 1981; Norman and Garcia 1999; Frey et al. 1985. Terrestrial Re and Yb data from Meisel et al. 1996, Walker et al. 1988, 1991; lunar Re and Yb data from compilation of Righter et al. 2000.

4

KEVIN RIGHTER

equilibrium in the early Moon most likely took place at conditions of high pressure and temperature. Extensive experimental studies in the past decade have led to the ability to predict siderophile element partition coefficients as a function of pressure, temperature, and composition. As a result, estimates may be made of impactor and proto-Earth mantle siderophile element compositions, and modeling of lunar siderophile element concentrations can be made at conditions most relevant to the early Moon. Partitioning of siderophile elements between metal and silicate liquid is measured by means of a partition coefficient D, defined as (wt% element M in metal)/(wt% element M in silicate melt), and is known to be a function of thermodynamic intensive variables: temperature, pressure, the fugacities of oxygen and sulfur, and silicate melt structure and composition. Advances in high pressure technology in the 1990s have permitted the experimental determination of metal/silicate partition coefficients at high pressures and temperatures (Walker et al. 1993, Hillgren et al. 1994, 1996, Ito et al. 1998). In particular, the effects of T, P, fO2 , silicate melt composition and S- and C-content of the metallic liquid on metal/silicate partition coefficients have been parameterized (Righter et al. 1997, Righter and Drake 1999). Expressions of the form, ln D(M) = a ln fO2 + b/T + cP/T + d(nbo/t) + e ln(1 − X S ) + f ln(1 − X C ) + g,

(1)

have been derived for Fe, Ni, Co, Mo, W, P, Ga, Sn, and Cu (Righter and Drake 1997a, 1999, 2000), where nbo/t refers to the degree of melt polymerization (where a value of 0 is polymerized and 4 is depolymerized; after Mysen 1991), and X S and X C refer to the mole fractions of sulfur and carbon, respectively, in the metallic liquid. Using these expressions it is possible, for instance, to calculate metal/silicate Ds along an adiabatic gradient in a deep magma ocean system (Fig. 3). It is readily apparent that the values of partition coefficients will be quite different at shallow vs deep conditions. 2.3. Geophysical Evidence for a Lunar Core A variety of geophysical techniques have been used to constrain the size of a lunar core, including gravity, seismic, electromagnetic sounding, laser ranging, and paleomagnetic data (see the recent review by Hood and Zuber 2000). Previous assessments of the lunar geophysical data had limited the size of a metallic core to <5 mass% (e.g., Hood 1986). Gravity and electromagnetic sounding data from the Lunar Prospector mission have narrowed the potential core size range to 1–3 mass% (Hood and Zuber 2000), or a core radius between 350 (pure Fe) and 430 km (pure FeS). 2.4. Source of Metal for a Lunar Core Even though a lunar metallic core is surmised to be small, the source of the metal must be considered. In impact modeling,

FIG. 3. Effect of pressure and temperature on metal–silicate partition coefficients for Ni, Co, and W using predictive expressions presented by Righter and Drake (1999). Conditions are an adiabatic temperature gradient with the oxygen fugacity fixed at 0.5 log fO2 unit below the IW buffer, peridotite magma (nbo/t = 2.8), and no light elements the metallic liquid. Also shown as vertical shaded regions are the CMB pressures of the proto-Earth and impactor before the Moon-forming collision (Cases 1 to 4 in Figs. 5 and 6). The depth of this interface may have controlled the PT conditions of metal–silicate equilibration in these bodies, and thus also the bulk siderophile element content of the Moon (if the Moon is made from the proto-Earth or impactor mantle).

there is often (but not always) a small amount of impactor core that ends up in Earth’s orbit (e.g., typically <4% in Canup and Asphaug 2001). Such material could provide metal for a small lunar core. It is also possible that high pressure and temperature mantle material ejected into Earth’s orbit could reequilibrate and be reduced at the relatively low pressures, but high temperatures in a circumterrestrial disk. For example, if the oxygen fugacity within the impactor mantle was buffered by C–CO–CO2 equilibria at the high pressures within the proto-planet, metal would not be stable. But if the same material reequilibrated at the relatively lower fO2 defined by C–CO–CO2 at the lower pressures of a circumterrestrial disk, it would be in the metal stability field and some reduction of iron may occur (e.g., Arculus et al. 1990). Whether metal came from the impactor or later reduction of FeO at relatively low PT conditions, such a small core would cause depletions of siderophile elements in the lunar mantle. 3. SIDEROPHILE ELEMENT DEPLETIONS IN THE IMPACTOR (STAGE 1)

3.1. Modeling Concentrations of Siderophile Elements in a Mantle after Core Formation Under the conditions of planetary differentiation or a Moonforming impact, both metal and silicate would be molten and thus equilibrium partitioning would be between metallic and

5

DOES THE MOON HAVE A CORE?

TABLE II Siderophile Element Compositions of Proto-Earth and Impactor Mantles for Scenarios and Previous Models Discussed in Text Case 1a hot pE Interfacec Ni (ppm) Co (ppm) Mo (ppb) W (ppb) P (ppm) Re (ppt) Ga (ppm) Pressure Temp. (◦ C)  IWd XSd XCd Core mass Bulk X

6900 120 390 1 126 3.2 13.0

Case 2 hot I

Case 3a warm I/pE

Case 3b warm I/pE

Case 3c warm I/pE

Case 4 warm I

Case 5 warm I

CMB CMB MO MO MO CMB CMB (3480, 10600) 2900 (1235, 5700) 1600 (630, 3730) 1960 (770, 4500) 2340 (920, 5400) 1085 (420, 2600) 560 (214, 1420) (54, 240) 120 (56, 230) 103 (46, 210) 134 (60, 275) 166 (74, 340) 92 (37, 200) 43 (17, 104) (130, 810) 114 (32, 336) 85 (23, 280) 193 (53, 630) 156 (43, 508) 91 (26, 275) 5.3 (1.4, 20) (0.09, 11) 5 (0.4, 39) 10 (0.9, 63) 19 (1.8, 130) 16 (1.5, 110) 22 (2.6, 81) 3.7 (0.3, 31) (10.7, 820) 420 (47, 1100) 615 (75, 1400) 560 (68, 1300) 615 (75, 1400) 840 (155, 1290) 132 (11, 740) (0.53, 19) 12 (2.0, 73) 6.9 (1.1, 42) 12 (2.0, 74) 13 (2.2, 80) 27 (4.4, 162) 4.4 (0.7, 27) (10.6, 13.9) 10.8 (8.9, 11.4) 13.7 (11.8, 14.4) 9.4 (6.3, 11.0) 8.2 (5.5, 9.6) 11.4 (9.3, 12.2) 6.5 (3.0, 9.0) 610 2925 −0.5 0.12 0.08 30% CI

350 2325 −0.75 0.1 0.1 15% CI

290 1925 −0.3 0.1 0.1 32% CI

290 1925 −0.3 0.1 0.1 32% CV

290 1925 −0.3 0.1 0.1 32% H

225 2025 −0.5 0.1 0.1 20% CI

PUMb

1890 105 65 16 95 0.28 4.0

200 2025 −1.7 0.0 0.05 10% CI

a Bolded numbers are the calculated concentrations for a given siderophile element; numbers in parentheses are the low and high estimates based on the 2σ error on the calculated (bolded) values. b PUM, primitive upper mantle of Earth; values taken from McDonough and Sun (1995). c Interface refers to the location of metal–silicate equilibration in the proto-Earth (pE) or Impactor (I); CMB, core-mantle boundary; MO, magma ocean. d Abbreviations: IW, oxygen fugacity in logfO units relative to the iron–w¨ ustite buffer at the specified temperature; X S and X C , the mole fraction of sulfur 2 and carbon, respectively, in the metallic liquid.

silicate liquids. The expression determining the concentration of a siderophile element, M, in a molten planet mantle is (Hillgren 1991) M CLS =

M Cbulk  M  , x + (1 − x) DLM/LS

(2)

M M where x is the silicate fraction of the planet, CLS and Cbulk are concentrations of M in the magma ocean and bulk planet, reM spectively, and DLM/LS is calculated using Eq. (1). Using this expression, one can calculate the concentrations of siderophile elements in the mantles of the impactor or the proto-Earth. BeM fore we do this it is useful to discuss our constraints on Cbulk , x, M and DLM/LS . M 3.1.1. Bulk composition. The bulk composition (Cbulk ) of proto-planets and terrestrial planets is uncertain. For instance, Earth’s major, minor, and isotopic composition is unlike any known chondrite group or mixture of chondrite groups (Drake and Righter 2002). The bulk composition of Mars is unknown, but it too is unlike any known chondrite group, and there is even debate about fundamental properties such as the MgO/(FeO + MgO) composition of its mantle (Bertka and Fei 1998). Recent dynamic modeling has shown that the material making up inner rocky planets such as Mercury, Venus, Earth, and Mars is mixed substantially over several astronomical units (Chambers 2001). The Moon is known to be volatile element depleted, FeO rich, and dry (compared to the Earth; Newsom and Taylor 1989), and thus clearly formed from a different inner Solar System reservoir

than Earth. With this perspective, bulk compositions for both the impactor and proto-Earth for most of the calculations were set to that of CI chondrites. Although in detail these chondrites may not represent the most accurate starting material for a terrestrial planet or impactor, the concentrations of the siderophile elements of interest to this work vary by a factor of only ∼1.5 within chondrite groups. For instance, calculations using a CV and H chondrite bulk composition are also listed in Table II, illustrating that the modeling is not particularly sensitive to the specific chondrite group used. 3.1.2. Silicate fraction (or core mass) of the impactor. Rocky planets in our Solar System have metallic cores that become less massive (relative to total planet mass) with distance from the Sun (Fig. 4). This may be related to a greater degree of oxidation with distance outward from the Sun, because planet FeO contents increase in the same direction (e.g., Mercury = <5 wt%, Earth and Venus = ∼8 wt%, Mars and Vesta = ∼18 wt%; BVSP 1981, Righter and Drake 1997b, Longhi et al. 1992). Because protoplanets and impactors may also have had core masses across this range, the mass of a core will be an important factor to fix in determining the magnitude of the siderophile element depletion. To illustrate this, one can use Eqs. (1) and (2), fixed bulk composition, P, T, fO2 , and magma and metal composition (to calculate D), to calculate the concentration of siderophile elements in a mantle that has equilibrated with cores of different masses (i.e., the variable x). Concentrations of Ni, Co, and Mo all decrease with increasing core mass, as expected. The concentration of W remains relatively stable with increasing core size, because its metal/silicate partition coefficient is close to unity.

6

KEVIN RIGHTER

Concentration in mantle (ppm)

Mercury

Earth Venus

Mars Vesta

Moon

100

10

1

Ni Co Mo W 0.3

0.4

0.5

0.6

0.7

0.8

0.9

1.0

mass fraction silicate (x) FIG. 4. Effect of core mass (or silicate mass fraction, x) on mantle concentrations of Ni, Co, Mo, and W using Eqs. (1) and (2). Partition coefficients were calculated at fixed conditions of 50 kbar, 2000◦ C, IW = −1, X S = 0.1, and nbo/t = 2.8 (peridotite); bulk composition is that of CI chondrites. Core masses of the terrestrial planets (top of diagram) are from BVSP (1981)—Venus, Mercury; Hood and Zuber (2000)—Moon; Righter and Drake (1997b)—Vesta; Longhi et al. (1992)—Mars.

One clear distinction between the composition of the Moon and the terrestrial mantle is the higher FeO content of the Moon. It is possible, given the correlations noted above between FeO content of mantle and core size, that the impactor had a small core. For this reason, several of the modeled impactors have smaller core masses of 15 and 20%. A core mass of 20% is used for the late Mars-sized impactor because this is similar to the mass of Mars’ core. In addition, core masses of 30–32% are used for the early and late proto-Earth and early impactor, since those masses are close to the median values. 3.1.3. Magnitude of D LMM / L S and the thermal state of the impactor. The concentrations of siderophile elements in an impactor or proto-Earth mantle will depend upon the depth at which metal and silicate may have equilibrated. It is likely that differentiated planets and planetesimals in the 20–50 Ma interval of terrestrial planet accretion retained enough volatiles to maintain thick atmospheres and thus surface temperatures hot enough to stabilize magma oceans (Abe et al. 2000). If the surface temperature is as high as 1700–2000◦ C, then an adiabatic temperature gradient through the mantle would keep the mantle molten at least to a depth of the stability field of Mg-perovskite (or even deeper), for which there is a significant change in slope of the melting curve (Fig. 5). Because a deep Mg-perovskite layer can stop metal from raining out of a magma ocean, and thus only allowing metal–silicate equilibrium to the depth of metal droplet stability in the magma ocean (Rubie et al. 2001), it would be useful to know at what point Mg-perovskite became stable in the mantle of the growing proto-Earth and also the impactor. Four different thermal cases will be examined which illustrate “hot” (Cases 1 and 2) or “warm” (Cases 3 and 4) conditions in the

impactor and/or proto-Earth. One successful model for a Moonforming impact is for a system with 0.65 m E-M (mass of the Earth–Moon system), which is 0.65 × 6 × 1024 or 3.9 × 1024 kg (e.g., Cameron 2000). The mass ratio of proto-Earth to impactor would be 7 : 3 or 2.7 × 1024 : 1.2 × 1024 , respectively. Assuming a mean density of 5530 kg/m3 for the bulk proto-Earth and impactor, a mantle density of 3300 kg/m3 , and that they separate cores of the same size as the Earth’s (32 mass%), the pressure at the core–mantle boundary in each can be calculated as 610 and 350 kbar (these assumptions yield values of g of 7.58 and 5.71 m2 /s for the proto-Earth and impactor, respectively; Figs. 5 and 6). In a very hot thermal state, where the mantle is molten nearly to the core–mantle boundary (CMB), the siderophile element concentrations in their mantles were moderated by high pressure and temperature metal–silicate equilibrium (Cases 1 and 2 for proto-Earth and impactor, respectively; Figs. 5 and 6). In a relatively warm thermal state, where the mantle is molten only to the depth of Mg-perovskite stability, the siderophile element concentrations in the proto-Earth and impactor will be moderated by high pressure and temperature metal–silicate equilibrium at this depth (Case 3 for both the proto-Earth and impactor; Figs. 5 and 6). A second successful model for a Moon-forming impact involves a late impactor and a roughly Mars-sized impactor (Canup and Asphaug 2001). In this case, the system has 1.0 m E-M , which is 6.0 × 1024 kg (e.g., Cameron 2000). The mass ratio of proto-Earth to impactor would be 9 : 1 or 5.4 × 1024 : 6.0 × 1023 kg, respectively. The depth to the CMB in such a Mars-sized impactor would be ∼23 GPa, slightly shallower than the appearance of Mg-perovskite on a

peridotite melting curves

Temperature (K)

Effect of core size on siderophile element depletion 1000

4000

proto-Earth and impactor adiabats CMB-Case 1

3000 CMB-Case 2 CMB-Case 3 CMB-Case 4

2000

present day mantle adiabat

Mg-perovskite stability

0

20

40

60

80

Pressure (GPa) FIG. 5. P-T phase diagram for peridotite (Zerr et al. 1998) along with four different adiabatic gradients (Cases 1–4; Fig. 6 and Table II) for a hot surface imposed by a thick blanketing atmosphere, and the present day mantle gradient for comparison. Inflections in the adiabats reflect a change in slope due to the latent heat of crystallization between the liquidus and solidus. Shaded region labeled “Mg-perovskite stability” represents the depth at which MgSiO3 perovskite becomes stable and the arrow indicates stability to greater depths.

7

DOES THE MOON HAVE A CORE?

peridotite liquidus. Siderophile element concentrations in the mantle of such an impactor, would again be moderated by metal– silicate equilibrium at the CMB (Case 4; Figs. 5 and 6). These four scenarios will be considered here, as well as a discussion of a shallow magma ocean system (“cold”) on the proto-Earth and impactor.

nearly complete melting total mass = 0.65 M(E) M(pE):M(I) =7:3

Case 1

"hot " impactor CMB = 350 kb

"hot" proto-Earth CMB = 610 kb

mass = 0.65 M(E) Case 3 { total M(pE):M(I) =7:3

Case 4 {

3.2. Calculated Siderophile Element Concentrations in the Impactor Mantle

Case 2

total mass = 1.0 M(E) M(pE):M(I) = 9:1

"cold" proto-Earth or impactor base of ocean is shallow

"warm" proto-Earth or impactor base of ocean defined by Mg-perovskite stability

FIG. 6. Schematic diagram of four different thermal scenarios for impactor and proto-Earth: hot, warm, and cool (Fig. 5 and Table II). Abbreviations are CMB, core-mantle boundary; M(pE), mass of proto-Earth; M(I), mass of impactor, and m(EM), mass of total Earth–Moon system. See text for further explanation.

The calculated compositions (using Eq. (1)) are listed in Table II, where they are also compared to Earth’s primitive upper mantle (PUM) and the lunar mantle. Because of the high pressures and temperatures of Cases 1 and 2, siderophile element partition coefficients are lowered substantially, and mantles from both of these bodies will have excesses of all seven siderophile elements over those estimated for the lunar mantle (Table III). For the intermediate (“warm”) scenarios of Cases 3 and 4, in which the depth of metal–silicate equilibrium is moderated by the depth to the Mg-perovskite stability field and the CMB, respectively (Fig. 5), mantles from both of these bodies will have excesses of all seven siderophile elements over those estimated for the lunar mantle (Table III). These four cases of the Moon being made from hot impactor or proto-Earth, or warm impactor or proto-Earth will be considered quantitatively below. A fifth case is that in which the impactor undergoes core formation and metal–silicate equilibrium, but there is no later core formation in the Moon. Attempts to match the depletions of Ni, Mo, and Re in the Moon in this way are discussed in detail in Section 5.2.

TABLE III Siderophile Elements in the Lunar Mantle Resulting from Segregation of a Metallic Core for Bulk Compositions in Table II

Ni (ppm) Co (ppm) Mo (ppb) W (ppb) P (ppm) Re (ppt) Ga (ppm) Pressure Temp. (◦ C)  IWc XSc XCc Core mass Core radius (km) Core Ni% x

Case 1a pEc

Case 2 Ic

Case 3a pE/I

Case 3b pE/I

Case 3c pE/I

Case 4 pE/I

Moon mantleb

640 (250, 1500) 64 (38, 89) 5.3 (1.4, 20) 0.6 (0.1, 1.0) 8.3 (6.1, 8.6) 0.015 (0.0024, 0.087) 12.7 (11.5, 13.1)

500 (205, 1040) 78 (52, 99) 11 (3.2, 34) 4.5 (2.2, 5.0) 28 (25, 29) 0.31 (0.052, 1.7) 10.6 (10.2, 10.8)

490 (215, 900) 80 (59, 94) 13 (3.7, 31) 15 (10, 16) 41 (39, 42) 0.34 (0.059,1.7) 13.7 (13.5, 13.8)

432 (184, 853) 103 (76, 119) 6.7 (1.8, 23.4) 15 (4.7, 19) 37 (33, 38) 0.15 (0.025, 0.87) 9.4 (9.1, 9.5)

426 (177, 880) 107 (70, 136) 15 (4.1, 44) 15 (8.8, 16) 41 (33, 4.2) 0.15 (0.025, 0.86) 8.1 (7.8, 8.2)

315 (142, 572) 74 (55, 84) 9 (2.9, 28) 21 (15, 22) 56 (52, 57) 0.19 (0.034, 8.6) 11.3 (11.1, 11.4)

470 90 2.2 18 20 1.0 0.4

50 1625 −1.0 0.0 0.2 2.0% 373 31.3 1.0

50 1725 −0.90 0.05 0.15 1.0% 296 23.0 1.0

50 1700 −0.75 0.05 0.15 0.5% 243 20.0 1.0

50 1600 −1.0 0.0 0.2 0.7% 263 21.9 1.0

50 1625 −1.0 0.0 0.1 0.8% 268 24.0 1.0

50 1625 −1.5 0.08 0.12 0.3% 198 25.7 1.0

a Bolded numbers are the calculated concentrations for a given siderophile element; numbers in parentheses are the low and high estimates based on the 2σ error on the calculated (bolded) values. b Moon mantle values are taken from Walter et al. (2000). c Abbreviations: pE, proto-Earth; I, impactor; IW, oxygen fugacity in logfO units relative to the iron–w¨ ustite buffer at the specified temperature; X S and X C , 2 the mole fraction of sulfur and carbon, respectively, in the metallic liquid. d Core radii are calculated from mass balance constraints and densities of Fe–S–C liquid alloys that are estimated assuming additivity of ternary end member component densities (Iida and Guthrie 1988), where ρ(Fe) = 7870, ρ(S) = 2070, and ρ(C) = 2290 Kg/m3 .

8

KEVIN RIGHTER

Finally, for the “cold” end member in which the depth to metal–silicate equilibrium in either the proto-Earth or impactor is shallow, partition coefficients would be much higher, and the resulting silicate mantle would have much lower siderophile element contents than estimated for the lunar mantle. However, because there may have been only a small amount of partially molten mantle present, metal will not be mobile and thus will become trapped in the silicate mantle (e.g., Rushmer et al. 2000). If this metal-rich cold mantle were then ejected in an impact to form the Moon, the metal would form a core much more massive than the Moon’s. For these reasons, “cold” pre-impact, protoEarth, and impactor scenarios will not be pursued quantitatively. 4. SIDEROPHILE ELEMENT DEPLETION IN THE MOON (STAGE 2)

The bulk composition of the Moon has been the focus of intense study (e.g., BVSP 1981), yet there is little consensus. The relative contribution of the impactor and proto-Earth to the bulk Moon is uncertain (e.g., Warren 1992). In this contribution, end members (100% impactor or 100% proto-Earth) will be utilized in the modeling. Because there are successful models for both kinds of bulk Moon compositions, it is reasonable to surmise that mixtures of impactor and proto-Earth may also work. 4.1. Previous Estimates of Siderophile Element Depletions in the Lunar Mantle A summary of siderophile element concentrations and depletions are presented here to highlight differences between the Earth and Moon. The importance of these differences will become clear in later discussions. Because the origin of the Moon is likely linked directly to the Earth, the lunar siderophile element abundances will be compared to those in the Earth. We have pieces of the terrestrial mantle in our collections, and thus estimates of siderophile elements in Earth’s mantle are more straightforward. Because this is an enormous field of literature, it will not be reviewed here, but the reader is directed to several key articles by Jagoutz et al. (1979), Palme and Nickel (1986), and recent reviews by Walter et al. (2000) and Righter et al. (2000). Because we have no direct samples of the lunar mantle, siderophile element concentrations in the Moon’s mantle must be estimated from the values measured in lunar basalt. During basalt genesis (melting of the mantle), the distribution of an element between solid and liquid can be monitored with the partition coefficient, D, which is the weight ratio of an element between a solid phase and silicate melt. Elements exhibit either compatible (D > 1) or incompatible (D < 1) behavior, so that compatible elements have lower concentrations in melts and incompatible elements have higher concentrations in melts, compared to the mantle. Correlations between a siderophile element and refractory lithophile element of equal compatibility or incompatibility can be used to estimate concentrations of siderophile elements in the original mantle. The siderophile element abundances discussed below have all been estimated

in this manner (see Newsom 1986, for a summary of this approach), using the data in Figs. 1 and 2 (Table II). There is some scatter in many of the diagrams, and the correlations are not perfect; this can be due to a number of factors, such as heterogeneity arising from small sample analytical mass (e.g., Ryder and Schuraytz 2001), diverse materials (breccias and coarse- and fine-grained materials), and contamination of volcanic materials by chondritic material at the surface of the Moon. 4.1.1. Compatible elements—Ni and Co. Since both of these elements exhibit compatible behavior, due to olivine and chromite (both phases have D > 1 for these elements) fractionation, their concentrations in planetary mantles can be estimated by correlations with MgO and FeO. This approach was pioneered by W¨anke and Dreibus (1986), who showed that the Co and Ni concentrations in the lunar mantle could be estimated from such correlations (Figs. 1A and 1B) by using an independent estimate of MgO or (MgO + FeO) in the mantle. For example, estimates of MgO range from 32.0 to 35.5 wt%, allowing estimation of Ni contents of 470 ± 50 ppm (Delano 1986). Nickel is approximately five times more depleted in the lunar mantle relative to the terrestrial mantle, while the lunar Co depletion is very similar. 4.1.2. Incompatible elements—Mo, W, Re. Many siderophile elements are incompatible during melting of a metal-free mantle, and when coupled with a refractory lithophile element of nearly equal incompatibility, their correlation can be used to estimate the original mantle concentration (Fig. 2A). The correlation line is well below the chondritic siderophile element values (at a given refractory lithophile element abundance), and this “depletion” is due to metal–silicate equilibrium (core formation) in that particular body. Hypothetical planets (A, B, C, etc.) would then have slightly different depletions of siderophile elements (Fig. 2A). An exception to this occurs when either metal or sulfide is present in the mantle source, in which case the concentration of the siderophile element will be fixed or buffered at variable refractory lithophile element concentrations (Fig. 2A). Neither metal nor sulfide affect the elements considered here, however, because none of the correlations discussed below are flat, and the Moon is known to be S poor such that there would be no residual sulfide in the mantle (e.g., Danckwerth et al. 1979, Gibson et al. 1977). In many basalt suites, including lunar samples, moderately siderophile elements such as Mo, W, and Re are incompatible and thus positively correlated with other incompatible, refractory lithophile elements (Pr, Ba, and Yb, respectively; Figs. 2B– 2D). Correlations of Mo with Pr, W with Ba, and Re with Yb, have been determined previously and are discussed in more detail by Newsom and Palme (1984), Newsom et al. (1996) and Righter et al. (1998). It is important to note, however, that because many terrestrial basalts are sulfide saturated (e.g., midocean ridge and oceanic island basalt), and sulfide/silicate partition coefficients for Re are as high as 1200, Re is mildly compatible. As a result, terrestrial basalts are not plotted in Fig. 2D.

9

DOES THE MOON HAVE A CORE?

4.1.3. Incompatible, volatile elements—Ga and P. Many siderophile elements are also volatile, and Ga, Sn, Cu, and P are good examples. As a result, their depletion can be due to either core formation or volatility-controlled processes. Depletions of Ga are best defined by correlations with Ti, due to their similar behavior during mantle melting (Righter and Drake 2000). However, the bimodal Ti-poor and Ti-rich nature of lunar basalts make this element pair more difficult to apply. Instead, Ga-Al correlations are used to define the Ga depletion in lunar samples (Fig. 2F). Depletions of P in the Moon are defined here with correlations with the refractory lithophile element La, as they were also previously defined by Newsom and Drake (1983) and Drake (1987). Although some argued previously that the lunar and terrestrial depletions of P were very similar (Ringwood 1989), it is clear that two distinct trends exist in Fig. 2E. Although there is metal/silicate partitioning data for Sn and Cu (see summary in Righter and Drake 2000), the wide range of Sn and Cu concentrations measured in a small number of lunar basalts is difficult to interpret. These two elements, while perhaps depleted to similar levels as Ga and P, await future assessment. To estimate the depletion of Ga and P due to volatility, these two elements must also be compared or normalized to lithophile elements of nearly equal volatility. Good candidates for such a comparison are the alkali elements—Na, K, Rb, and Cs. It has been recognized for some time that these elements are depleted in the lunar mantle relative to the Earth (e.g., Kreutzberger et al. 1986), and thus Ga and P must be understood in the context of this general volatile element depletion. It has been found that Ga is no more depleted than Na in lunar highland feldspars (Norman et al. 1995). In fact, neither Ga nor P are depleted relative to the lithophile, volatile elements (Li, Na, K, Rb, Cs; Fig. 7), suggesting that neither has been depleted substantially by lunar core formation. Overall the siderophile elements show increasing depletion with increasing siderophility. Ni, Mo, and Re exhibit a further depletion over that in the Earth, and Ga, Co, P, and W depletions are similar to those in the Earth. Both of these observations argue qualitatively for the presence of a metallic lunar core. It is interesting to note that the depletion of W in the eucrite parent body (Vesta) is also very similar for the Earth and Moon (e.g., Righter and Drake 1996), indicating that W is not a very sensitive indicator of the pressure and temperature of core formation in differentiated planets. 4.2. Modeling Results for the Moon Given the bulk Moon compositions calculated in Section 3.2 and knowledge of the partition coefficients for Ni, Co, Mo, W, P,

1

depletion (relative to BSE)

Instead, only sulfide undersaturated komatiites, in which Re behaves incompatibly at high degrees of melting, are used to estimate terrestrial Re depletions. The data for the Earth and Moon reveal that the lunar mantle is substantially more depleted in Mo and Re than the terrestrial mantle, and that the W depletions are very similar (Figs. 2B–2D).

0.1

Alkali depletion of Moon relative to Earth

0.01

Li

Na

P

K

Ga

Rb

Cs

FIG. 7. Depletion of Ga and P relative to the alkali elements, Li, Na, K, Rb, and Cs, and normalized to terrestrial mantle concentrations (BSE, taken from Jones and Palme 2000). No depletion of Ga and P relative to the alkalis suggests that these elements may have been depleted during core formation but only due to volatility.

Ga, and Re, the question of what size core could have separated from these compositions can now be addressed. 4.2.1. P, T, fO2 of Moon for calculating partition coefficients. Knowing the pressure, temperature, and oxygen fugacity conditions within the Moon is critical to calculating relevant metal– silicate partition coefficients. The central pressure of the Moon is approximately 50 kbar; because a lunar core would be small, this value has been fixed at 50 kbar for all calculations that follow. If the Moon was derived from material ejected into Earth’s orbit following an impact, and then underwent a magma ocean stage, equilibration of metal and silicate would have occurred at temperatures close to or above the liquidus of the Moon’s mantle. As a result, temperatures were held to greater than 1600◦ C, consistent with either complete or very extensive degrees of melting. Finally, the fO2 of lunar materials has been constrained by intrinsic oxygen fugacity measurements and thermodynamic calculations and ranges from 0.75 to 2 log fO2 units below the iron–w¨ustite (IW) oxygen buffer. These low values are consistent with the presence of Fe and FeNi alloys in lunar basalts (BVSP 1981). 4.2.2. Bulk Moon compositions (Section 3.2) and D LMM / L S . Again using Eq. (2), but in this case for the bulk Moon composition of the four different mantles (Cases 1–4), one can calculate the resulting lunar mantle compositions by segregation of a small core. Metal–silicate partition coefficients are calculated using Eq. (1). For each of the cases below, partition coefficients were calculated with the aim of fitting as many of these seven elements as possible. Case 1 yielded the worst results (fits to four elements), Cases 2 and 3 slightly better (fits to five or six elements), and Case 4 allowed a fit to all seven elements.

10

KEVIN RIGHTER

Case 4:“warm” impactor or proto-Earth (late impact): Using the mantle composition of a “warm” impactor or proto-Earth (with a 20 mass% core) as a bulk Moon composition, segregation of a 0.3 % Fe-rich metallic core is consistent with all seven elements, Ga, Co, W, Ni, Mo, P, and Re concentrations estimated for the lunar mantle (Table III; Fig. 8). This core (again S- and C-bearing) would equilibrate with the molten peridotite mantle at 1625◦ C, 50 kbar, and 1.5 log fO2 units below the IW buffer. 5. DISCUSSION

5.1. Geochemical Plausibility of an Impactor-Derived Moon

FIG. 8. Depletion of siderophile elements in the Moon, normalized to CI chondrites and Si (solid circles; from Walter et al. 2000). Cases 1–5 are calculated depletions for different lunar bulk compositions as described in the text. Lunar P and Ga concentrations are corrected for volatility (factor of 5 for both elements) according to Fig. 7; terrestrial P and Ga concentrations are corrected for volatility (factor of 3 and 5, respectively) according to the volatility diagram of McDonough and Sun (1995). Terrestrial depletions (horizontal shaded bands) are also normalized to CI chondrites, Si, and volatility corrected.

Case 1: “hot” proto-Earth: Using the mantle composition of a hot proto-Earth (with a 30 mass% core) as a bulk Moon composition, segregation of a 2 mass% Fe-rich metallic core is consistent with Ga, Co, Ni, P, and Mo concentrations, but would leave far less W and Re than is estimated (Table III; Fig. 8). In such a case, a C-bearing core would equilibrate with a molten peridotite lunar mantle at 1625◦ C, 50 kbar, and 1.0 log fO2 units below the IW buffer (Table III). Case 2:“hot” impactor: Using the mantle composition of a hot impactor (15 mass% core) as a bulk Moon composition, segregation of a 1 mass% Fe-rich metallic core is consistent with the Ga, Co, P, Ni, Mo, and Re concentrations estimated for the lunar mantle (Table III; Fig. 8). Calculated concentrations of W, however, are slightly lower than estimates. This small, Sand C-bearing core would equilibrate with the molten peridotite mantle at 1725◦ C, 50 kbar, and 0.9 log fO2 units below the IW buffer. Case 3:“warm” impactor or proto-Earth (early impact): Using the mantle composition of a “warm” impactor or proto-Earth (with 32 mass% core) as a bulk Moon composition, segregation of a 0.5 mass% Fe-rich metallic core is consistent with the Ga, Co, W, Ni, Mo, P, and Re concentrations estimated for the lunar mantle (Table III; Fig. 8). This core (S- and C-bearing) would equilibrate with the molten peridotite mantle at 1700◦ C, 50 kbar, and 0.75 log fO2 units below the IW buffer (Case 3a). There are minor differences between scenarios 3a, 3b, and 3c, where the original bulk composition is CI, CV, or H chondrite, and the core sizes are 0.7 mass% and 0.8 mass% for 3b and 3c. In Case 3c, for example, the calculated Mo contents are too high (Fig. 8).

Many geochemical similarities between the Earth’s mantle and the Moon have led geochemists to conclude that the Moon may be made of material from the Earth. This is counter to all dynamic modeling conclusions that indicate the Moon must be made primarily from material from the impactor. There are a few notable exceptions to this view—for example, McFarlane (1989) concluded that the mantle of the impactor is more likely to have the geochemical characteristics required for the bulk Moon. And Jones and Hood (1990) concluded that despite the similarities between the Earth and Moon, there are important differences such as a higher FeO content in the lunar mantle. A significant finding of the current work is that the siderophile element concentrations in the lunar mantle are consistent with an origin from an impactor, given a few reasonable assumptions of the prior thermal state, composition, and degree of differentiation of an impactor. As a result any previous geochemical-based objections to an impactor origin of the Moon are removed. 5.2. No Metallic Core? Some have suggested the possibility that the Moon has no metallic core (Khan and Mosegaard 2001, Newsom and Taylor 1989). In such a case, the mantle of the impactor would have to be stripped of Ni, Mo, and Re relative to Earth, and the density concentration near the center of the Moon would be attributed to non-metallic material. Producing an impactor mantle with larger depletions of Ni, Mo, and Re relative to the Earth could have been accomplished in several ways. One way to do this would be to have metal– silicate equilibrium at lower pressures and temperatures, because the metal–silicate partition coefficients would be higher. A second factor that might contribute to a greater depletion of these elements relative to Earth would be for the impactor to have undergone metal segregation at a lower fO2 , also promoting large metal–silicate partition coefficients and larger siderophile element depletions. A limitation with this scenario is that such a reduced impactor mantle might have a low FeO content, whereas the Moon has a relatively high FeO content (e.g., Jones and Palme 2000). On the other hand, the mantle of Vesta may have as much as 18 wt% FeO and yet is as reduced as the Moon (Righter and Drake 1997b).

11

DOES THE MOON HAVE A CORE?

Using the approach taken in Section 3, P, T, fO2 space was explored to see if a set of conditions exist that would produce the lunar siderophile element concentrations without the need for later segregation of a metallic core. A match to the observed Ni, Mo, and Re depletions was found using a CI chondrite bulk composition, a modest core (10 mass%), and metal–silicate partition coefficients calculated at 200 kbar, 2300 K, IW = −1.7, for a peridotite magma ocean (nbo/t = 2.8), and a sulfur-free metallic core that has a small fraction of C (Table II). This match is possible by a combination of the effects described above— lower PT conditions and a lower fO2 , but it is also clear that the calculated Co, Ga, and W concentrations do not match those observed. Nonetheless, if the Moon has no metallic core, then there must be a non-metallic material responsible for the density anomaly at its center. Candidates for dense non-metallic lunar cores include oxide or silicate melt, again, both of which are unlikely. Although dense, TiO2 -rich silicate melt could potentially sink to the center of the Moon relative to the less dense lunar mantle, such melts would only have a density of ∼3500 or 3600 kg/m3 (Circone and Agee 1997) at 50 kbar, too low to explain the density anomaly of 4700 kg/m3 (Khan and Mosegaard 2001). Ilmenite cumulates from a magma ocean could also form a dense core, but even a 0.5 mass% core of ilmenite would yield a bulk Moon composition that is unusually Ti-rich, and thus unlikely on geochemical grounds (i.e., it would contain more Ti than most chondrites and peridotites). Furthermore, neither ilmenite nor dense silicate magma can generate a magnetic field, one of the compelling features of a metallic core. When the siderophile element arguments are combined with the evidence from paleointensity data, center of gravity–center of figure offset, and paleogravity surfaces (e.g., Runcorn 1996, Hood and Zuber 2000) it seems an unavoidable conclusion that the Moon possesses a small metallic core. Although some have argued that the Moon’s core formed 500 Ma after the Moon accreted cold (Runcorn 1996), a scenario of extensive melting in a post-impact, rapid, accretionary event also seems unavoidable. The Moon may have started out hot and cooled over time to reach its current thermal state in which a metallic core would be solid (e.g., Pritchard and Stevenson 2000). Perhaps the 500 Ma delay in core formation proposed by Runcorn (1996) is simply the time at which a solid inner core nucleated to help drive the magnetic field, as has been proposed for the Earth’s dynamo (Glatzmaier and Roberts 1996). 6. CONCLUSIONS

Given new lunar formation models, the ability to predict metal–silicate partition coefficients at elevated temperatures and pressures, and new lunar geophysical data, the question of whether the Moon has a metallic core has been reexamined. Estimates of siderophile element concentrations for Ni, Co, Mo, W, P, Ga, and Re in the lunar mantle are consistent with scenarios in which the Moon is made from material ejected from

the mantle of either the “hot” impactor or a “warm” impactor or proto-Earth, and then separates a small (2 mass% or less) metallic core. Segregation of a small metallic core from a bulk Moon made of ejected material from a “hot” proto-Earth mantle would leave the lunar mantle too depleted in the elements W and Re. The dynamical modeling conclusion that the Moon is made of material from the impactor is consistent with the distribution of siderophile elements in the Moon. The lunar mantle concentrations of Ni, Mo, and Re can also be explained by previous core formation within an impactor, not requiring a lunar metallic core, but this result is at odds with Ga, Co, and W concentrations and geophysical data favoring a metallic core. ACKNOWLEDGMENTS Funding was from NASA Grant NAG5-9435. This work was stimulated by discussions with colleagues at the LPI-hosted “New Views of the Moon, III” meeting, October 2000, held in Houston. It also benefited from discussions and comments of R. Canup, M. J. Drake, M. Wieczorek, M. Norman, E. McFarlane, and two anonymous reviews.

REFERENCES Abe, Y., E. Ohtani, T. Okuchi, K. Righter, and M. J. Drake 2000. Water in the early Earth. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 413–434. Univ. of Arizona Press, Tucson. Arculus, R. J., R. D. Holmes, R. Powell, and K. Righter 1990. Metal–silicate equilibrium and core formation. In The Origin of the Earth (H. Newsom and J. H. Jones, Eds.), pp. 251–271. Oxford Univ. Press, London. Basaltic Volcanism Study Project 1981. Basaltic Volcanism on the Terrestrial Planets. Pergamon, New York. Bertka, C. M., and Y. Fei 1998. Implications of Mars pathfinder data for the accretion history of the terrestrial planets. Science 281, 1838–1840. Cameron, A. G. W. 1997. The Origin of the Moon and the single impact hypothesis V. Icarus 126, 126–137. Cameron, A. G. W. 2000. Higher resolution simulations of the giant impact. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 133–144. Univ. of Arizona Press, Tucson. Cameron, A. G. W., and W. Benz 1991. The Origin of the Moon and the single impact hypothesis IV. Icarus 92, 204–216. Canup, R. M., and E. Asphaug 2001. The Moon forming giant impact near the end of Earth’s formation. Nature 412, 708–712. Canup, R. M., and L. W. Esposito 1996. Accretion of the Moon from an impactgenerated disk. Icarus 119, 427–446. Canup, R. M., W. R. Ward, and A. G. W. Cameron 2001. A scaling relationship for satellite-forming impacts. Icarus 150, 288–296. Chambers, J. 2001. Making more terrestrial planets. Icarus 152, 205–224. Circone, S., and C. B. Agee 1997. Compressibility of molten high-Ti mare glass: Evidence for crystal–liquid density inversions in the lunar mantle. Geochim. Cosmochim. Acta 60, 2709–2720. Danckwerth, P. A., P. C. Hess, and M. J. Rutherford 1979. The solubility of sulfur in high TiO2 mare basalts. Proc. Lunar Planet. Sci. Conf. 10th, 517–530. Delano, J. W. 1986. Abundances of cobalt, nickel, and volatiles in the silicate portion of the Moon. In Origin of the Moon (W. K. Hartmann, R. J. Phillips, and G. J. Taylor, Eds.), pp. 231–248. Lunar and Planetary Institute, Houston. Delano, J. W., and A. E. Ringwood 1978. Siderophile elements in the lunar highlands: Nature of the indigenous component and implications for the origin of the Moon. Proc. Lunar Planet. Sci. Conf. 9th, 111–159.

12

KEVIN RIGHTER

Dickey, J. S., F. A. Frey, S. R. Hart, E. B. Watson, and G. Thompson 1977. Geochemistry and petrology of dredged basalts from the Bouvet triple junction, South Atlantic. Geochim. Cosmochim. Acta 41, 1105–1118. Drake, M. J. 1987. Siderophile elements in planetary mantles and the origin of the Moon. Proc. Lunar Planet. Sci. Conf. 17th, J. Geophys. Res. 92, E377– E386. Drake, M. J., and K. Righter 2002. Determining the composition of the Earth. Nature 416, 39–44. Frey, F. A., C. J. Suen, and H. W. Stockman 1985. The Ronda high temperature peridotite: Geochemistry and petrogenesis. Geochim. Cosmochim. Acta 49, 2469–2491. Gibson, E. K., R. Brett, and F. Andrawes 1977. Sulfur in lunar mare basalts as a function of bulk composition. Proc. Lunar Sci. Conf. 8th, 1417–1428. Glatzmaier, G. A., and P. H. Roberts 1996. Rotation and magnetism of Earth’s inner core. Science 274, 1887–1890. Herzberg, C., and J. Zhang 1996. Melting experiments on anhydrous peridotite KLB-1: Compositions of magmas in the upper mantle and transition zone. J. Geophys. Res. 101, 8271–8295. Hillgren, V. J. 1991. Partitioning behavior of Ni, Co, Mo, and W between basaltic liquid and Ni-rich metal—Implications for the origin of the moon and lunar core formation. Geophys. Res. Lett. 18, 2077–2080. Hillgren, V. J., M. J. Drake, and D. C. Rubie 1994. High pressure and temperature experiments on core-mantle segregation in the accreting Earth. Science 264, 1442–1445. Hillgren, V., M. J. Drake, and D. C. Rubie 1996. High pressure and high temperature metal–silicate partitioning of siderophile elements: The importance of silicate liquid composition. Geochim. Cosmochim. Acta 60, 2257–2263. Hood, L. L. 1986. Geophysical constraints on the lunar interior. In Origin of the Moon (W. K. Hartmann, R. J. Phillips, and G. J. Taylor, Eds.), pp. 361–409. Lunar and Planetary Institute, Houston. Hood, L. L., and M. T. Zuber 2000. Recent refinements in Geophysical constraints on lunar origin and evolution. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 397–410. Univ. of Arizona Press, Tucson. Ida, S., R. M. Canup, and G. R. Stewart 1997. Lunar accretion from an impactgenerated disk. Nature 389, 353–357. Iida, T., and R. I. L. Guthrie 1988. The Physical Properties of Liquid Metals. Oxford Univ., New York. Ito, E., T. Katsura, and T. Suzuki 1998. Metal/silicate partitioning of Mn, Co, and Ni at high pressures and high temperatures and implications for core formation in a deep magma ocean. In Properties of Earth and Planetary Materials at High Pressure and Temperature, Geophysical Monograph 101 (M. H. Manghnani, Ed.), pp. 215–225. Amer. Geophys. Union, Washington, DC. Jagoutz, E., H. Palme, H. Baddenhausen, K. Blum, M. Cendales, G. Dreibus, B. Spettel, V. Lorenz, and H. W¨anke 1979. The abundances of major, minor and trace elements in the Earth’s mantle as derived from primitive ultramafic nodules. Proc. Lunar Planet. Sci. Conf. 10th, 2031–2050. Jones, J. H., and L. L. Hood 1990. Does the Moon have the same chemical composition as the Earth’s upper mantle? In Origin of the Earth (H. E. Newsom and J. H. Jones, Eds.), pp. 85–98. Oxford Univ. Press, New York. Jones, J. H., and H. Palme 2000. Geochemical constraints on the origin of the Earth and Moon. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 197–216. Univ. of Arizona Press, Tucson. Khan, A., and K. Mosegaard 2001. New information on the deep lunar interior from an inversion of lunar free oscillations periods. Geophys. Res. Lett. 28, 1791–1794. Kreutzberger, M. E., M. J. Drake, and J. H. Jones 1986. Origin of Earth’s Moon: Constraints from alkali volatile trace elements. Geochim. Cosmochim. Acta 50, 91–98. Lee, D.-C., A. N. Halliday, G. A. Snyder, and L. A. Taylor 1997. The age and origin of the Moon. Science 278, 1098–1103.

Longhi, J., E. Knittle, J. R. Holloway, and H. W¨anke 1992. The bulk composition, mineralogy and internal structure of Mars. In Mars (H. H. Kieffer, B. M. Jakosky, C. W. Snyder, and M. S. Mathews, Eds.), pp. 184–208. University of Arizona Press, Tucson. McDonough, W. F., and S. S. Sun 1995. The composition of the Earth. Chem. Geol. 120, 223–253. McFarlane, E. A. 1989. Formation of the Moon in a giant impact: Composition of the Impactor. Proc. Lunar Planet. Sci. Conf. 19th, 593–605. Meisel, T., R. J. Walker, and J. W. Morgan 1996. The osmium isotopic composition of the Earth’s primitive upper mantle. Nature 383, 517–519. Mysen, B. O. 1991. Relations between structure, redox equilibria of iron, and properties of magnetic liquids. In Physical Chemistry of Magmas (L. L. Perchuk and I. Kushiro, Eds.), pp. 41–98. Springer-Verlag, New York. Newsom, H. 1984. The lunar core and the origin of the Moon. EOS 65, 369–370. Newsom, H. 1986. Constraints on the origin of the Moon from the abundance of molybdenum and other siderophile elements. In Origin of the Moon (W. K. Hartmann, R. J. Phillips, and G. J. Taylor, Eds.), pp. 203–229. Lunar and Planetary Institute, Houston. Newsom, H., and M. J. Drake 1983. Experimental investigations of the partitioning of phosphorus between metal and silicate phases: Implications for the Earth, Moon, and eucrite parent body. Geochim. Cosmochim. Acta 47, 93–100. Newsom, H., and S. R. Taylor 1989. Geochemical implications of the formation of the Moon by a single giant impact. Nature 338, 29–34. Newsom, H. E., and H. Palme 1984. The depletion of siderophile elements in the Earth’s mantle: New evidence from molybdenum and tungsten. Earth Planet. Sci. Lett. 69, 354–364. Newsom, H. E., W. M. White, K. P. Jochum, and A. W. Hofmann 1986. Siderophile and chalcophile element abundances in oceanic basalts, Pb isotope evolution and growth of the Earth’s core. Earth Planet. Sci. Lett. 80, 299–313. Newsom, H. E., K. W. W. Sims, P. D. Noll, Jr., W. L. Jaeger, S. A. Maehr, and T. B. Beserra 1996. The depletion of tungsten in the bulk silicate Earth: Constraints on core formation. Geochim. Cosmochim. Acta 60, 1155–1169. Norman, M. D., and M. O. Garcia 1999. Primitive magmas and source characteristics of the Hawaiian plume: Petrology and geochemistry of shield picrites. Earth Planet. Sci. Lett. 168, 27–44. Norman, M. D., K. Keil, W. L. Griffin, and C. G. Ryan 1995. Fragments of ancient lunar crust: Petrology and geochemistry of ferroan noritic anorthosites from the Descartes region of the Moon. Geochim. Cosmochim. Acta 59, 831– 847. O’Neill, H. St. C. 1991. The origin of the Moon and the early history of the Earth—A chemical model. Part 1: The Moon. Geochim. Cosmochim. Acta 55, 1143–1158. Palme, H., and K. G. Nickel 1986. Ca/Al ratio and composition of the Earth’s primitive upper mantle. Geochim. Cosmochim. Acta 49, 2123–2132. Palme, H., H. Baddenhausen, K. Blum, M. Cendales, G. Dreibus, H. Hofmeister, H. Kruse, C. Palme, B. Spettel, E. Vilczek, H. W¨anke, and G. Kurat 1978. New data on lunar samples and achondrites and a comparison of the least fractionated samples from the Earth, Moon and the eucrite parent body. Proc. Lunar Planet. Sci. Conf. 9th, 25–57. Pritchard, M., and D. J. Stevenson 2000. Thermal aspects of a lunar origin by giant impact. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 179–196. Univ. of Arizona Press, Tucson. Righter, K., and M. J. Drake 1996. Core formation in Earth’s Moon, Mars and Vesta. Icarus 124, 513–529. Righter, K., and M. J. Drake 1997a. Metal–silicate equilibrium in a homogeneously accreting Earth: New results for Re. Earth Planet. Sci. Lett. 146, 541–553. Righter, K., and M. J. Drake 1997b. A magma ocean on Vesta: Core formation and petrogenesis of eucrites and diogenites, Met. Planet. Sci. 32, 929–944.

DOES THE MOON HAVE A CORE? Righter, K., and M. J. Drake 1999. Effect of water on metal–silicate partitioning of siderophile elements: A high pressure and temperature terrestrial magma ocean and core formation. Earth Planet. Sci. Lett. 171, 383–399. Righter, K., and M. J. Drake 2000. Metal–silicate equilibrium in the early Earth: New constraints from volatile moderately siderophile elements Ga, Sn, Cu, and P. Geochim. Cosmochim. Acta 64, 3581–3597. Righter, K., M. J. Drake, and G. Yaxley 1997. Prediction of siderophile element metal–silicate partition coefficients to 120 kb and 2800◦ C: The effect of pressure, temperature, fO2 and silicate and metallic melt composition. Phys. Earth Planet. Interiors 100, 115–134. Righter, K., R. L. Hervig, and D. Kring 1998. Accretion and core formation in Mars: Molybdenum contents of melt inclusion glasses from three SNC meteorites. Geochim. Cosmochim. Acta 62, 2167–2177. Righter, K., R. J. Walker, and P. W. Warren 2000. The origin and significance of highly siderophile elements in the lunar and terrestrial mantles. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 291–322. Univ. of Arizona Press, Tucson. Ringwood, A. E. 1979. Origin of the Earth and Moon. Springer-Verlag, New York. Ringwood, A. E. 1989. The Earth–Moon connection. Z. Naturforsch. 44a, 891– 923. Ringwood, A. E., and S. Seifert 1986. Nickel–cobalt abundance systematics and their bearing on lunar origin. In Origin of the Moon (W. K. Hartmann, R. J. Phillips, and G. J. Taylor, Eds.), pp. 249–277. Lunar and Planetary Institute, Houston. Rubie, D. C., H. J. Melosh, J. Reid, and K. Righter 2001. Constraints on core formation mechanisms from metal–silicate equilibration kinetics. EOS 82, F29. Runcorn, S. K. 1996. The formation of the lunar core. Geochim. Cosmochim. Acta 60, 1205–1208. Rushmer, T., W. G. Minarik, and G. J. Taylor 2000. Physical processes of core formation. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 227–243. Univ. of Arizona Press, Tucson. Ryder, G., and B. Schuraytz 2001. Chemical variation of the large Apollo 15 olivine-normative mare basalt rock samples. J. Geophys. Res. 106, 1435– 1451. Shearer, C. K., and J. J. Papike 1999. Magmatic evolution of the Moon. Am. Mineral. 84, 1469–1494. Shearer, C. K., and K. Righter 2000. Tungsten partitioning in silicates: A key to understanding the early evolution of the Moon. Lunar Planet. Sci. XXXI, 1411 (abstract). Stevenson, D. J. 1990. Fluid dynamics of core formation. In The Origin of the

13

Earth (H. Newsom and J. H. Jones, Eds.), pp. 231–249. Oxford Univ. Press, London. Taylor, S. R., R. Rudowski, P. Muir, A. Graham, and M. Kaye 1971. Trace element chemistry of lunar samples from the Ocean of Storms. Proc. Lunar Science Conf. 2nd, 1083–1099. Walker, D., L. Norby, and J. H. Jones 1993. Superheating effects on metal/silicate partitioning of siderophile elements. Science 262, 1858–1861. Walker, R. J., S. B. Shirey, and O. Stecher, 1988. Comparative Re–Os, Sm–Nd and Rb–Sr isotope and trace element systematics for Archean komatiite flows from Munro Township, Abitibi Belt, Ontario. Earth Planet. Sci. Lett. 87, 1–12. Walker, R. J., L. M. Echeverria, S. B. Shirey, and M. F. Horan 1991. Re–Os isotopic constraints on the origin of volcanic rocks, Gorgona Island, Colombia: Os isotopic evidence for ancient heterogeneities in the mantle. Contrib. Mineral. Petrol. 107, 15–162. Walter, M. J., H. Newsom, W. Ertel, and A. Holzheid 2000. Siderophile elements in the Earth and Moon: Metal/silicate partitioning and implications for core formation. In Origin of the Earth and Moon (R. M. Canup and K. Righter, Eds.), pp. 265–290. Univ. of Arizona Press, Tucson. W¨anke, H., and G. Dreibus 1986. Geochemical evidence for the formation of the Moon by impact-induced fission of the proto-Earth. In Origin of the Moon (W. K. Hartmann, R. J. Phillips, and G. J. Taylor, Eds.), pp. 649–672. Lunar and Planetary Institute, Houston. W¨anke, H., H. Baddenhausen, A. Balacescu, F. Teschke, B. Spettel, G. Dreibus, H. Palme, M. Quijano-Rico, H. Kruse, F. Wlotzka, and F. Begemann 1972. Multi-element analyses of lunar samples and some implications of the results. Proc. Lunar Sci. Conf. 3rd, 1251–1268. W¨anke, H., H. Palme, H. Baddenhausen, G. Dreibus, E. Jagoutz, H. Kruse, B. Spettel, F. Teschke, and R. Thacker 1974. Chemistry of Apollo 16 and 17 samples: Bulk composition, late stage accumulation and early differentiation of the Moon. Proc. Lunar Sci. Conf. 5th, 1307–1335. W¨anke, H., H. Baddenhausen, K. Blum, M. Cendales, G. Dreibus, H. Hofmeister, H. Kruse, E. Jagoutz, C. Palme, B. Spettel, R. Thacker, and E. Vilczek 1977. On the chemistry of lunar samples and achondrites. Proc. Lunar Planet. Sci. Conf. 8th, 2191–2213. Warren, P. H. 1992. Inheritance of silicate differentiation during lunar origin by giant impact. Earth Planet. Sci. Lett. 112, 101–116. Warren, P. H., D. N. Shirley, and G. W. Kallemeyn 1986. A potpourri of pristine lunar rocks, including a VHK mare basalt and a unique, augite-rich Apollo 17 anorthosite. Proc. Lunar Planet. Sci. Conf. 16th; J. Geophys. Res. 91, D319–D330 (supplement). Zerr, A., A. Diegeler, and R. Boehler 1998. Solidus of Earth’s deep mantle. Science 281, 243–246.