origin and evolution of water in the moon's interior · 2014. 6. 4. · ea45ch04-hauri ari 14...

EA45CH04-Hauri ARI 14 August 2017 12:38

Annual Review of Earth and Planetary Sciences

Origin and Evolution of Waterin the Moon’s InteriorErik H. Hauri,1 Alberto E. Saal,2 Miki Nakajima,1

Mahesh Anand,3 Malcolm J. Rutherford,2

James A. Van Orman,4 and Marion Le Voyer5

1Department of Terrestrial Magnetism, Carnegie Institution of Washington, Washington,DC 20015; email: [email protected] of Geological Sciences, Brown University, Providence, Rhode Island 029123Department of Physical Sciences, The Open University, Milton Keynes MK7 6AA,United Kingdom4Department of Earth, Environmental, and Planetary Sciences, Case Western ReserveUniversity, Cleveland, Ohio 441065Department of Geology, University of Maryland, College Park, Maryland 20742

Annu. Rev. Earth Planet. Sci. 2017. 45:89–111

First published as a Review in Advance on May 24,2017

The Annual Review of Earth and Planetary Sciences isonline at earth.annualreviews.org

https://doi.org/10.1146/annurev-earth-063016-020239

Copyright c© 2017 by Annual Reviews.All rights reserved

Keywords

Moon, water, volcanism, pyroclastic, hydrogen, deuterium

Abstract

Nearly forty years after the return of the first lunar samples to Earth, im-provements in laboratory detection limits made possible the first definitivediscovery of magmatic water in lunar volcanic samples. The interveningdecade has seen an exponential increase in the amount of data on the abun-dance of magmatic water, and its hydrogen isotope composition, in variousrock types recovered from the Moon. Here we review these data and de-scribe how the abundance of water in the lunar interior places importantconstraints on models for the high-temperature origin and evolution of theMoon.

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ANNUAL REVIEWS Further

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http://www.annualreviews.org/doi/full/10.1146/annurev-earth-063016-020239


1. INTRODUCTION

The discovery of magmatic water in volcanic samples returned from the Moon demonstratedthat the Moon’s origin and evolution must involve processes that allow for the accretion andretention of the most volatile of Solar System compounds (Saal et al. 2008, 2013; Hauri et al.2011; Furi et al. 2014; Chen et al. 2015). Parallel studies of water in lunar apatite, measured atabundances analogous to terrestrial apatite, provided further evidence that the Moon’s interiorcontains significant amounts of magmatic water (Boyce et al. 2010, 2014; McCubbin et al. 2010;Greenwood et al. 2011; Barnes et al. 2013, 2014; Tartese & Anand 2013; Tartese et al. 2013,2014a,b; Anand et al. 2014). The presence of water in the lunar interior places a severe constrainton high-temperature models of the formation and evolution of the Moon that begin with a giantimpact creating a protolunar disk of vaporized material and end with a fully coalesced satellitecovered in a magma ocean (Cameron & Benz 1991, Canup 2004). In this article, we review thedata on H2O abundances in lunar rocks and their minerals and discuss how the abundance of waterin lunar samples constrains existing models for the formation and evolution of the Moon.

2. VOLCANIC GLASSES, MARE BASALTS, AND PLUTONIC ROCKS

2.1. Overview

The oldest rocks on the Moon, derived from the lunar highlands, cluster at ages of 4.40–4.30 billion years (e.g., Carlson et al. 2014); these rocks are thought to be the first samples oflunar crust formed as Ca-rich feldspar floated to the surface of the lunar magma ocean (LMO)after it had cooled and crystallized most of its mass (e.g., Elkins-Tanton et al. 2011). After the for-mation of the lunar crust, volcanic eruptions took place on the Moon over the time period of 3.9–3.2 billion years ago, or 500–1,000 million years after the solidification of the lunar interior (Span-gler & Delano 1984, Spangler et al. 1984, Nyquist & Shih 1992). These eruptions are evidence fora protracted period of internal heating sufficient to generate partial melts of the lunar interior todepths of 200–500 km and deliver them to the lunar crust (Elkins-Tanton et al. 2011). This heatis likely the result of a combination of primordial heat trapped from the coalescence of the Moonfrom the protolunar disk and a concentration of radiogenic heat from elevated concentrations ofK, U, and Th in the mantle of the lunar nearside (Wieczorek et al. 2013).

The water content of lunar samples is typically expressed as the oxide equivalent (H2O) ofhydrogen, whereas the species typically measured in lunar glasses and minerals is hydroxyl (OH−).The water content of lunar samples places constraints on a number of processes that affect thewhole Moon as well as individual lunar samples. The isotopic composition of hydrogen, given bythe ratio of deuterium (2H, or D) to hydrogen (1H), is typically expressed as a permil deviationfrom the isotopic composition of Vienna Standard Mean Ocean Water (VSMOW):

δD (�) = [(D/H)sample/(D/H)VSMOW − 1

] × 1,000,

where (D/H)VSMOW is the hydrogen isotope composition of VSMOW (Gonfiantini et al. 1995).The hydrogen isotope composition of lunar samples is an effective tracer for the origin of theMoon’s water, given the large range of δD in Solar System objects (Figure 1), though hydrogenisotopes are also sensitive to processes (such as magmatic degassing, hydrodynamic escape, andcosmic ray spallation) that have the potential to alter the δD of specific lunar samples or even theentire Moon (Saal et al. 2013).

The water content of the lunar interior, and the isotopic composition of hydrogen, has been in-ferred from measurements on three sample types: primitive pyroclastic volcanic glasses (Saal et al.

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–1,000 0 1,000 2,000 3,000 4,000 5,000

δD (‰)

Sun: solar windSaturn

Jupiter

Uranus

Neptune

Titan

Earth

Bulk carbonaceous chondrites

Mars mantle Mars atmosphere

Interplanetary dust particles:chondrite organics

to +20,000‰

Comets

+100,000‰Venus atmosphere

Lunar glasses

Lunar mantle

Figure 1The range of hydrogen isotope compositions in Solar System objects (Saal et al. 2013 and references therein).

2008, 2013; Hauri et al. 2011; Furi et al. 2014; Chen et al. 2015), mare basalt lava flows (Boyce et al.2010; McCubbin et al. 2010; Greenwood et al. 2011; Barnes et al. 2013; Tartese & Anand 2013;Tartese et al. 2013, 2014a,b; Chen et al. 2015), and various plutonic rocks (McCubbin et al. 2010,Greenwood et al. 2011, Hui et al. 2013, Barnes et al. 2014). Secondary ion mass spectrometry(SIMS) has been the most widely used method in these studies, achieving low detection limits(∼0.1 ppm H) and precise data for hydrogen isotopes (Hauri et al. 2002, 2006; Greenwood et al.2011; Barnes et al. 2013) as well as simultaneous measurements of C, F, S, and Cl; Fourier trans-mission infrared (FTIR) spectroscopy has also been used for water abundance measurements (Huiet al. 2013). Analysis of major elements in lunar volcanic glasses and melt inclusions has pro-vided a direct measure of specific magmatic compositions—there are roughly 25 distinct chemicalcompositions that represent the most primitive magmas on the Moon (Delano 1986). Analysis ofH2O in minerals requires the use of mineral-melt partition coefficients in order to estimate theH2O content of a coexisting melt; fortunately, there are several SIMS studies that have measuredvarious mineral-melt partition coefficients for H2O in laboratory melting experiments (e.g., Hauriet al. 2006; McCubbin et al. 2015a,b).

2.2. Water and Other Volatiles in Lunar Apatite

Apatite has been the main target mineral for in situ studies [e.g., SIMS, Electron Probe Mi-croprobe Analysis (EPMA)] investigating the abundance and isotopic composition of indigenouslunar volatiles such as H, F, and Cl. Apatite, with an idealized formula of Ca5(PO4)3(F,Cl,OH),can contain water bound into the crystal structure as OH as well as other volatiles such as F and Cl.Apatite occurrence has been recorded in minor or trace quantities from almost all lunar lithologies,

www.annualreviews.org • Water in the Moon 91

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100 µm

60 µm60 µm

ol

ap

plag

ilm

a

b

c

~10 µm

Figure 2(a) Backscattered electron image of apatite (ap) surrounded by olivine (ol), ilmenite (ilm), and plagioclase feldspar (plag) in Apollo 17lunar breccia 79215. (b) Reflected light image of lunar volcanic glass beads from Apollo 15 pyroclastic glass sample 15426.(c) Backscattered electron image of an olivine-hosted melt inclusion from Apollo 17 pyroclastic glass sample 74220.

ranging from mare basalts to magnesian-suite and alkali-suite rocks. Apatite grains are typicallysubhedral to anhedral, ranging in size from ∼2 μm to >200 μm in their longest dimension, andare in textural association with minerals that form during the end stages of melt crystallization(Figure 2).

Typically, SIMS studies of lunar apatite include measurements for the abundances of F, Cl,and OH and the isotopic composition of H and Cl (e.g., Anand et al. 2014). McCubbin et al.(2007, 2008) were the first to estimate OH contents in lunar apatites on the basis of stoichiometriccalculations of EPMA data. Subsequently, a number of studies performed direct analyses for water(measured as either H or OH but reported in terms of equivalent H2O) by SIMS and confirmedthat lunar apatite indeed contains variable and sometimes appreciable amounts of water, rangingfrom a few tens of parts per million up to weight percent levels (Boyce et al. 2010; McCubbin et al.

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Mareapatite

NonmareapatiteδD

(‰ v

ersu

s V

SMO

W)

H2O (ppm)

a b1,500

1,000

Mare basalts

Nonmaresamples

500

50

–500

–1,000100 101 102 103 104

H2O (ppm)101100 102 103 104

KREEP basalts

Lunar glasses

Figure 3(a) H2O content (in parts per million by weight) versus hydrogen isotope composition of apatite from mare basalts (gray circles),KREEP basalts (red circles), and other nonmare samples (blue circles). (b) H2O content (in parts per million by weight) versus hydrogenisotope composition of glasses and melt inclusions from lunar pyroclastic samples. Data are from Anand et al. (2014), Barnes et al.(2013, 2014, 2016), Greenwood et al. (2011), Hauri et al. (2011), Robinson & Taylor (2014), Robinson et al. (2016), Saal et al. (2008),Tartese & Anand (2013), and Tartese et al. (2013, 2014a,b). Abbreviations: KREEP, potassium/rare earth element/phosphorus;VSMOW, Vienna Standard Mean Ocean Water.

2010; Greenwood et al. 2011; Barnes et al. 2013, 2014; Tartese et al. 2013, 2014a,b). Similarly,Cl and F contents display a large range; the F end-member is more dominant in mare basalts,whereas Cl-rich apatite appears to be more common in nonmare samples that also show increasedinvolvement of KREEP (K, potassium; REE, rare earth element; P, phosphorus) components(McCubbin et al. 2011).

Hydrogen isotopes in apatite, as well as hydrogen abundances, can be affected not only byprocesses that formed the Moon but also by processes that have affected specific lunar samples, in-cluding degassing, hydrogen diffusion, and production of deuterium from cosmic ray spallation—though most apatites have sufficiently high hydrogen abundances that the latter process is typicallynot important within analytical uncertainties. H2O contents and D/H ratios of lunar apatite areshown in Figure 3. Greenwood et al. (2011) carried out the first H isotopic measurements in lunarapatite, mainly from mare basalts, and reported elevated δD values ranging from approximately+400� to +1,000�. Subsequent studies have expanded this H isotope data set on apatite froma diverse set of lunar lithologies and have reported δD values ranging from −750� to +1,100�(Barnes et al. 2013, 2014; Tartese et al. 2013, 2014a,b; Robinson & Taylor 2014; Robinson et al.2016). Several processes and sources have been invoked to reconcile this large range in measuredδD values. In the case of mare basalts, degassing of H2 is thought to have played a major role ingiving rise to elevated δD values in the residual melts from which apatite subsequently crystallized(Tartese & Anand 2013, Tartese et al. 2013). In cases in which degassing either did not play asignificant role or did not occur, the δD values are comparable to those of the terrestrial mantle,strongly suggesting a common origin for at least some of the indigenous water in the Earth-Moonsystem (Barnes et al. 2014, Tartese et al. 2014a,b), a conclusion also reached by a separate H iso-tope study on lunar pyroclastic glasses and their melt inclusions (Saal et al. 2013). Interestingly,apatite in some highly evolved lunar rocks displays very low δD values, almost approaching solarvalues (Robinson & Taylor 2014). There is no obvious signature of solar wind involvement in thepetrogenesis of these samples; therefore, taken at face value, these depleted H isotopic signatures


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have been interpreted as evidence for the presence of a distinct low-δD reservoir in the lunarinterior (Robinson & Taylor 2014). Additional work focusing on similar lithologies from otherApollo sites will be necessary to further evaluate this hypothesis.

Not all recent laboratory investigations of lunar samples support a hydrous lunar interior. Forexample, Sharp et al. (2010) measured chlorine isotopic compositions of a range of Apollo samplesand concluded that the lunar interior had to be anhydrous. The δ37Cl values they measured in asuite of lunar samples (both bulk samples and in situ apatite) ranged from −1� to +24�. Theδ37Cl value is expressed as a permil deviation from the isotopic composition of Standard MeanOcean Chloride (SMOC):

δ37Cl(�) =[(37Cl/35Cl

)sample/

(37Cl/35Cl)

SMOC − 1]

× 1,000,

where (37Cl/35Cl)SMOC is the Cl isotope composition of SMOC. In comparison, the Cl isotopecomposition of terrestrial and a vast majority of nonlunar extraterrestrial materials is clusteredaround ± 2� (Sharp et al. 2007, 2013). Later studies (Tartese et al. 2014a, Boyce et al. 2015,Barnes et al. 2016) further confirmed the extreme enrichment in 37Cl of some lunar apatite (δ37Clvalues of up to +36�). Sharp et al. (2010) interpreted the enrichment in 37Cl as a result ofthe different behavior of the two Cl isotopes during near-surface magma degassing—the lighterisotope is preferentially partitioned into the vapor phase, whereas the heavier isotope, bondedwith metals as metal chloride, is either left behind in the melt or redeposited from the vaporphase on volcanic glasses, leading to a large fractionation of the two isotopes. The presence of anywater in the system would have significantly influenced the behavior of chlorine isotopes such thatthere would not have been any significant isotopic fractionation, as observed in terrestrial systems.Sharp & Draper (2013) expanded their petrogenetic model and argued that the requirement foran anhydrous melt during Cl degassing did not preclude the possibility that significant H-bearingspecies existed in the melt before Cl loss and were lost more quickly than Cl, as also observed inexperiments performed by Ustunisik et al. (2011, 2015). Furthermore, a recent study reporting themost comprehensive Cl isotope data set on lunar apatite to date demonstrated that the elevatedCl isotopic signature of lunar apatite is well correlated with the geochemical signature of KREEPin specific samples (Barnes et al. 2016). It is hypothesized that the lunar interior acquired a heavyCl isotopic composition during the earliest stages of LMO evolution (Boyce et al. 2015, Barneset al. 2016) such that the KREEP layer acquired a heavy Cl isotope signature, which is reflected inlunar samples that assimilated variable KREEP components during later stages of lunar magmatichistory (Barnes et al. 2016).

Despite apatite being the common target for H and Cl isotopic investigations, as describedabove, it is not straightforward to interpret the range of isotopic fractionation recorded in apatitefor these two elements in terms of a single process or event. There is no obvious correlationbetween the H and Cl isotopic composition of lunar apatite, suggesting involvement of multipleprocesses, possibly at different stages of the lunar differentiation history. Under highly reducingoxygen fugacity conditions in the lunar interior, the speciation of H (and Cl) in the melt and thevapor phase would have certainly influenced the extent of isotopic fractionation during degassing.However, it remains unclear whether successive episodes of volatile degassing from a single magmacould explain both elevated δD and elevated δ37Cl values measured in lunar apatite. As Barneset al. (2016) pointed out, the requirement of very low H/Cl ratios in the melt to account forthe elevated δ37Cl values measured in mare apatite, as proposed by Sharp et al. (2010), seemsincompatible with the occurrence of OH-rich apatite relative to Cl in the mare basalts (Tarteseet al. 2013; McCubbin et al. 2015a,b) given the relative preference of Cl over H into crystallizingapatite (Boyce et al. 2014; McCubbin et al. 2015a,b), which should result in very OH-poor apatite

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(Ustunisik et al. 2011). Therefore, it is likely that variations of apatite δD and δ37Cl values resultedfrom the interplay between different magmatic processes such as degassing or mixing, potentiallyoverprinting any geochemical signatures inherited from previously formed isotopic reservoirs.

2.3. Water in Primitive Lunar Magmas

The H2O contents of lunar volcanic glasses are shown in Figure 3. The H2O contents of pyro-clastic glass beads from various Apollo missions, representing different eruptive events, are almostall below 100 ppm and in many cases approach SIMS and FTIR detection limits. These low con-centrations were determined to be the result of volcanic degassing via diffusive loss of H2O andother volatiles through melt droplets; Saal et al. (2008) estimated that lunar magmas from theApollo 15 and 17 sites had lost over 90% of their pre-eruptive budget of water via degassing. Thiswas confirmed by Hauri et al. (2011), who measured up to 1,200 ppm H2O in melt inclusionscontained within olivine crystals from the Apollo 17 orange glass and demonstrated that this lu-nar magma contained F, S, and Cl in abundances similar to those in terrestrial mid-ocean ridgebasalts. Importantly, the range of concentrations of H2O and other volatiles in the melt inclusionsrevealed a degassing trend connecting the melt inclusion volatile concentrations to those of themore degassed pyroclastic glasses from the same eruption. It is thus thought that lunar magmaslikely had much more water (and C, F, S, and Cl) prior to eruption than we currently measure inthe degassed melt droplets that have quenched to glass.

As with apatite, the isotopic composition of hydrogen in lunar glasses and melt inclusionscan be affected by whole-Moon processes as well as sample-specific processes such as degassing,diffusion, and D production by cosmic ray spallation. These data are shown in Figure 3; at thevery lowest H2O contents, the δD values of lunar samples are dominated by production of D bycosmic ray spallation (Merlivat et al. 1976, Reedy 1981, Saal et al. 2013, Furi & Deloule 2016),but this effect is reduced as a function of increasing H2O concentration as the budget of magmaticwater dominates the spallogenic component at H2O abundances >20 ppm (Figure 3). The dataare systematic in that the lowest δD values are associated with the highest H2O abundances, in theApollo 17 orange melt inclusions. Correcting the δD values of low-H2O samples is complicatedby uncertainties in the spallogenic production of D by cosmic rays; when the production ratedetermined by Merlivat et al. (1976) is used, the corrected δD values suggest an increase withdecreasing H2O that can be explained by degassing of H2 (Saal et al. 2013). However, the higherproduction ratio reported by Furi & Deloule (2016), from analyses of lunar olivine in samples witha range of cosmic ray exposure ages, results in a likely overcorrection that produces very negativeδD values with decreasing H2O.

2.4. Degassing of Lunar Magmas and Pre-Eruptive Volatile Estimates

Saal et al. (2008), Hauri et al. (2011), and Wetzel et al. (2015) demonstrated that H2O, C, F, S,and Cl in lunar volcanic glass beads have been affected by magmatic degassing. This degassinginvolved release of volatiles as a separate phase during magma ascent and eruption, followed bydiffusion-limited kinetic degassing of volatiles out of melt droplets after magma fragmentation(Rutherford et al. 2015, Wetzel et al. 2015). There exist two primitive magma compositions withestimates of pre-eruptive volatile contents; according to the classification of Delano (1986), thesemagmas are Apollo 15 green very low Ti (VLT) glass (Saal et al. 2008) and Apollo 17 orangeglass melt inclusions from lunar sample 74220 (Hauri et al. 2011). Diffusion modeling of thedegassing profiles in Apollo 15 green VLT glass indicated a best-fit prediffusion H2O content of


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745 ppm, and a minimum of 265 ppm at the 95% confidence limit (Saal et al. 2008). This modelingindicated that during the diffusive stage of degassing following magma fragmentation, Apollo 15green VLT magma lost ∼98% of its initial H2O and ∼57% of its initial Cl, but only ∼45% ofits initial F and ∼19% of its initial S content. Comparison of Apollo 17 orange glasses and meltinclusions (Hauri et al. 2011) indicates ∼99% degassing of H2O, 85–90% degassing of F and Cl,and 61% degassing of S, all of which indicate a greater extent of degassing of Apollo 17 orange74220 magma compared with the Apollo 15 green VLT magma.

Analogously to volatile studies on terrestrial volcanic glasses, the systematics of volatile elementscan be compared with a nonvolatile element that would otherwise behave similarly during partialmelting and magmatic crystallization. Ratios of volatile to nonvolatile elements, such as H2O/Ceand C/Nb, have seen wide use for inferring the extent of degassing in terrestrial submarine volcanicsystems (Dixon & Clague 2001, Dixon et al. 1995, Cartigny et al. 2008) and identifying regionsof Earth’s interior where volatile abundances may be heterogeneous (Michael 1995, Cooper et al.2012). Trace element data for Apollo 15 green VLT glass (Hauri et al. 2015) result in H2O/Ce =169, F/Zr = 0.4, F/Nd = 3.67, S/Dy = 129, Cl/Nb = 0.23, and Cl/Ba = 0.021. The H2O/Ceratio is within the terrestrial mantle range, whereas the S/Dy ratio is low by a factor of two andthe F/Zr, F/Nd, Cl/Nb, and Cl/Ba ratios are low by factors of 5–20 compared with Earth’s uppermantle. The Apollo 17 orange melt inclusions are somewhat more depleted in highly volatileelements, with H2O/Ce = 67–77, F/Zr = 0.44–0.46, F/Nd = 4.23–4.46, S/Dy = 90–94,Cl/Nb = 0.19–0.20, and Cl/Ba = 0.035–0.041. The F/Zr, F/Nd, S/Dy, Cl/Nb, and Cl/Ba ratiosare similar to the estimated ratios in the Apollo 15 green VLT glass, whereas the pre-eruptiveH2O/Ce ratio for the 74220 magma is 40–48% lower than that estimated for the Apollo 15 greenVLT magma. These data suggest either that the 74220 magma was more degassed or that itsmantle source was lower in volatile elements compared with the Apollo 15 green VLT magmasource.

3. THE ORIGIN OF THE MOON’S INTERIOR WATER

The composition of the bulk silicate Moon (BSM) was estimated by Hauri et al. (2015) on the basisof the moderately and highly volatile element concentrations observed in lunar pyroclastic glasses.This BSM composition differs from all others in that it is less depleted in water and other volatileelements, and this conclusion is based on two assumptions: (a) that prior estimates, based on marebasalt compositions, were biased to low volatile abundances due to higher extents of degassing inmare basalts compared with lunar glasses and (b) that the estimated mantle source compositionsinferred from lunar volcanic glasses apply to the whole Moon. On this second point, Chen et al.(2015) examined melt inclusions from a variety of mare basalts and demonstrated that severaldifferent chemical types of lunar basalt had concentrations of F, Cl, and S similar to those of themelt inclusions derived from the pyroclastic high-Ti sample 74220; this argues strongly that thehigh volatile abundances of the 74220 melt inclusions are not anomalous but are representativeof several different types of lunar magma compositions.

Apart from highly and moderately volatile elements, the elemental composition of the BSMderived by Hauri et al. (2015) is otherwise very similar to estimates of the composition of thebulk silicate Earth (cf. McDonough & Sun 1995). For nonvolatile elements, this view is supportedby data showing that Earth and the Moon have isotopic compositions that are virtually identicalwithin laboratory analytical errors (cf. Pahlevan 2014, Pahlevan et al. 2016). The list of matchingisotopic compositions includes oxygen (Wiechert et al. 2001, Spicuzza et al. 2007, Hallis et al.2010, Young et al. 2016), chromium (Lugmair & Shukolyukov 1998), silicon (Chakrabarti &Jacobsen 2010a, Armytage et al. 2012, Fitoussi & Bourdon 2012), calcium (Valdes et al. 2014),

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titanium (Zhang et al. 2012), zirconium (Schonbachler et al. 2003), magnesium (Chakrabarti &Jacobsen 2010b), iron (Beard & Johnson 1999, Liu et al. 2010), and stable isotopes of strontium(Charlier et al. 2012). These isotopic similarities are not shared by all planetary objects in theSolar System; in fact, the Moon and Earth are the only planetary pair to share such similarity forsuch a large number of elements. The only isotopic exceptions to Earth-Moon similarity are smallmeasured differences in the isotopes of several volatile elements: copper and zinc (Herzog et al.2009, Paniello et al. 2012), chlorine (Sharp et al. 2010, Boyce et al. 2015), potassium (Wang &Jacobsen 2016), and high-precision measurements of hydrogen isotopes (Saal et al. 2013). Despitethe temptation to ascribe these isotopic shifts to losses from the high-temperature proto-lunar diskfollowing the giant impact, all of the isotopic fractionations observed in these volatile elementscan be explained as a consequence of either volcanic degassing or degassing of the LMO (Boyceet al. 2015, Hauri et al. 2015).

4. VOLATILE ELEMENT CONSTRAINTS ON MOON FORMATION

4.1. Atmospheric Erosion

The amount of volatiles lost from Earth by giant impact is a long-standing open question. Althoughvolatiles lost from Earth are part of the Earth-Moon budget, it should be noted that this isa completely different problem from volatile loss from the Moon-forming disk. The Moon isextremely unusual in the sense that it formed from a disk and volatiles could have been lost duringthe prolonged (100–1,000 years or longer) accretion process (Thompson & Stevenson 1988,Salmon & Canup 2012, Charnoz & Michaut 2015). In contrast, a planet can lose its volatiles dueto impact-induced ground motion over a very short timescale (<1 day), but conversely, it can alsogain volatiles from the impactor. Therefore, whether the planet gains or loses volatiles dependson the competition of these two processes, and the fact that the Moon is depleted in volatiles doesnot indicate that Earth also became depleted. This can be inferred from the observation that thelunar ratio of the moderately volatile element K to the nonvolatile element Th is much lower thanthat of Earth (Lodders & Fegley 1998, Prettyman et al. 2006), whereas Earth’s K/Th ratio is fairlyclose to those of Mars (Taylor et al. 2006a,b), Venus (Lodders & Fegley 1998), and Mercury,which may have or may not have experienced traumatic giant impacts (Peplowski et al. 2011).

A number of studies have estimated the extent of atmospheric loss by giant impacts. Ahrens(1990, 1993) suggested that a Mars-sized impact on Earth would induce ground motion (1.6–2.7 km/s), and this could completely remove Earth’s atmosphere. In contrast, Genda & Abe (2003b)performed one-dimensional numerical simulations of the atmosphere using ground motion of theplanet as a parameter and showed that the criterion suggested by Ahrens (1990, 1993) is thecriterion for onset of atmospheric loss rather than for complete loss. More recently, Genda & Abe(2005) showed that the presence of a water ocean enhances atmospheric loss due to evaporationof the ocean and its lower shock impedance compared with that of the ground. Stewart et al.(2014) later extended this work to three-dimensional calculations and confirmed that a Mars-sized impactor would not remove Earth’s atmosphere, whereas the more energetic recent impactmodels could efficiently remove it. Recently, Schlichting et al. (2015) used analytic self-similarsolutions and numerical simulations and found that rather small impactors tend to more efficiently(per mass) remove the planetary atmosphere. Therefore, a number of relatively small impactors(2 km in size) could play an important role in removing early planetary atmospheres. Furthermore,the impact angle, which has not been considered carefully, may play an important role as well(Shuvalov 2009). Further studies are needed to understand atmospheric erosion and replenishingprocesses.


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4.2. Origin of the Moon by Giant Impact

It is widely accepted that Earth’s Moon formed by a collision between the proto-Earth and animpactor approximately 4.5 billion years ago (Hartmann & Davis 1975, Cameron & Ward 1976,Ida et al. 1997). This impact created a partially vaporized disk around the planet, from which theMoon accreted. According to the standard version of this impact model, a Mars-sized impactorhit Earth with a relatively large impact angle (∼45–60◦) and an impact velocity that was similar tothe escape velocity (Canup & Asphaug 2001). This model can explain several observed signaturesof the Earth-Moon system, but it cannot easily explain the fact that Earth and the Moon havenearly identical isotopic ratios [e.g., oxygen, silicon, tungsten, chromium, and titanium (Lugmair& Shukolyukov 1998, Wiechert et al. 2001, Armytage et al. 2012, Fitoussi & Bourdon 2012, Zhanget al. 2012, Herwartz et al. 2014, Kruijer et al. 2015, Young et al. 2016). According to smoothedparticle hydrodynamic (SPH) simulations, which describe a fluid as a collection of spherical par-ticles, disk materials mainly originate from the impactor (Canup & Asphaug 2001). Because it islikely that the impactor had different isotopic ratios from those of Earth, the model would predictthat the isotopic ratios between Earth and the Moon would likely be different. It is possible thatthe impactor happened to have similar isotopic ratios of a specific element, such as oxygen (Kaib &Cowan 2015, Mastrobuono-Battisti et al. 2015), but it is unlikely that Earth and the impactor hadidentical isotopic ratios in all of the elements measured to date. Reufer et al. (2012) performed anumber of SPH simulations and showed that an impact with a steep angle could form a disk mainlyoriginating from Earth’s mantle, but the contribution from the impactor is still too large to explainthe isotopic similarity. Pahlevan & Stevenson (2007) and, more recently, Lock et al. (2015) sug-gested that mixing between Earth’s mantle and the disk may have homogenized the isotopic ratiosof the two reservoirs. The potential issue is that these models require Earth to have undergonewhole mantle convection for complete mixing. This may be problematic because Earth’s surfacewould have been more shock heated than its interior; therefore, Earth’s mantle was thermally sta-ble and would not have immediately begun convection (Nakajima & Stevenson 2015). Eventually,with further cooling, Earth’s mantle would have cooled by radiation and started convection, butthe timescale for this process would have been much longer than that for the Moon formation.

To overcome these problems, new impact models have been suggested. Cuk & Stewart (2012)and Lock et al. (2015) suggested that a small impactor hit a rapidly rotating Earth, whereasCanup (2012) suggested that two half-Earth-sized objects collided. In these models, the disk’scomposition is similar to Earth’s, and the isotopic similarities can therefore be naturally explained.These new models are promising alternatives, but they may predict that Earth’s mantle becomesmixed (Nakajima & Stevenson 2015), which may contradict geochemical observations that Earth’smantle was never fully homogenized (e.g., Willbold et al. 2011, Mukhopadhyay 2012, Touboulet al. 2012); the standard model, in contrast, would be consistent with the preservation of mantleheterogeneity. This difference occurs primarily because these recent models are more energeticthan the standard model. Another potential issue is that the angular momentum of the system afterthe impact is approximately two to three times as large as that of the current Earth-Moon system.Cuk & Stewart (2012) suggested that this excess of angular momentum could have been transferredto the Sun-Earth system due to the evection resonance, which occurs when the precession periodof the Moon is the same as the orbital period of Earth. Whether it is possible to remove the excessangular momentum efficiently is under active discussion (Wisdom & Tian 2015), but this workcertainly has opened up examination of new details of the giant impact process. To summarize, eachmodel has its own problems; the standard model has difficulty explaining the isotopic similaritybetween Earth and the Moon, whereas the new models face challenges to account for the excessof angular momentum and preservation of mantle heterogeneity.

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Figure 4Two-dimensional maps of total vapor pressure in circumterrestrial protolunar disks produced by giant impacts on the early Earth(Nakajima & Stevenson 2014a,b). Radial distance (r/R⊕) and vertical distance (z/R⊕) are normalized to the radius of Earth (R⊕), whichis centered at (0,0); vertical exaggeration is ∼2×. The disk midplane is along the horizontal axis at (z/R⊕) = 0. Individual temperaturecontours are shown for 3,500 K (stability field of silicate vapor) (thick white curve), 2,500 K, and 700 K (near the limit of the stabilityfield of water vapor). Four impact scenarios are shown: (a) the standard impact model of a Mars-size impactor impacting Earth (Canup2004); (b) a fast-spinning Earth model (Cuk & Stewart 2012); (c) a sub-Earth impact model (Canup 2012); and (d) an intermediate-entropy giant impact model (Nakajima & Stevenson 2014a,b). The low vapor pressures of the disk will lead to low water partialpressure and limited solubility of water in the molten material that eventually coalesces into the Moon.

4.3. Dynamics of the Moon-Forming Disk

Each impact model produces a Moon-forming disk around Earth within ∼24 h after the impact(Figure 4). Because the recent impact models are more energetic, they produce disks that are hotter(up to 6,000–7,000 K) and more vaporized (80–90 wt%) than that produced in the standard model[up to 4,000–5,000 K and 20–30 wt% (Nakajima & Stevenson 2014b)]. Initially, disk materials aredistributed across the Roche radius, defined as the radius within which material cannot form self-gravitating clumps due to a strong tide from the planet. The outer region of the disk (r > RRoche,where r is the distance from Earth’s spin axis and RRoche is the Roche radius) cools more quicklythan the inner region, partly because of its relatively large surface area and partly because of a lackof viscous heating (Thompson & Stevenson 1988, Salmon & Canup 2012). Subsequently, the outer


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part of the disk fragments and forms moonlets. The inner part of the disk (r < RRoche) viscouslyspreads out and slowly accretes to these moonlets while some material accretes to Earth (Salmon& Canup 2012, 2014). A higher initial disk temperature leads to a longer disk lifetime, becausethe disk lifetime is ultimately regulated by radiative cooling [∼100 to 1,000 years (Thompson &Stevenson 1988) or longer (Charnoz & Michaut 2015)]. A set of nonstandard SPH simulationsprovided additional insights into this problem (e.g., Wada et al. 2006, Canup et al. 2013, Hosonoet al. 2016).

4.4. Volatile Loss from the Moon by Impact

The Moon-forming impact has been thought to be at least partly responsible for the fact that theMoon is depleted in volatiles with respect to Earth [e.g., K, Rb, Na, and other volatile elements(Krahenbuhl et al. 1973a,b; Tera & Wasserburg 1976; Ringwood & Kesson 1977; Taylor 1979;Wolf & Anders 1980)], because the giant impact was so energetic that some volatiles may haveescaped during the process. This idea has also been supported by the observation that lunar rockshave low K/Th and K/U ratios compared with Earth (Tera et al. 1974, Prettyman et al. 2006).

It has been suggested that these volatiles were lost from the Moon-forming disk by hydro-dynamic escape (Abe et al. 2000, Genda & Abe 2003a, Desch & Taylor 2013), which is causedby light elements escaping from the high-temperature disk, dragging with them heavier atomsand molecules. The disk would initially have contained some amount of water, which would havedissociated to hydrogen and oxygen under high temperature and low pressure in the disk. Desch& Taylor (2013) suggested that hydrodynamic escape could be triggered by hydrogen in the disk,and this escape blows off volatiles when the condition λ ≡ GM⊕m/2kT < 2 is met. Here, G isthe gravitational constant, M⊕ is the mass of Earth, m is the mean molecular weight of the vaporphase, k is the Boltzmann constant, and T is the disk temperature. According to their calcula-tions, at T = 2,000 K, water is dissociated to oxygen and hydrogen (m = 6 g/mol), and thiscondition is achieved. However, this model assumes that the disk is isothermal, which is not thecase for the Moon-forming disk (Nakajima & Stevenson 2014a). Another issue is that the escapecriterion is developed for the solar wind, which primarily consists of hydrogen (Parker 1963), andthe same model may not be applicable if the disk vapor is enriched in heavier elements such asSi and O.

In fact, previous studies have shown that the presence of heavy elements alters the escaperegime, at least for planetary atmospheres. When a planetary atmosphere is dominated by heavyatoms or molecules, these elements are gravitationally bound to the planet. For light elements toescape, they must first diffuse out from the heavy element–rich atmosphere. This regime is calleddiffusion-limited escape and is much slower than atmospheric blowoff (Hunten 1973, Huntenet al. 1987, Zahnle et al. 1988). The escape flux in the diffusion-limited regime is expressed as(bf/H)homopause, where b is the binary collision parameter, f is the fraction of a light element, andH is the scale height. Here, homopause is the location where the kinetic diffusion equals themolecular diffusion. Nakajima & Stevenson (2014a) and Nakajima (2016) investigated the diskstructure assuming the disk is in thermal equilibrium and showed that the disk is dominated bysilicate vapor and that hydrogen is a minor species when the midplane (liquid-vapor interface)temperature Tmid is high (for example, f < 10−3 at >4,000 K). At lower disk temperatures (Tmid <

2,500 K), which would be achieved after the disk cools, the partial pressure of silicate vaporbecomes small and water becomes the dominant species. Under this condition, the homopausetemperature is ∼1,500 K, and water exists in its molecular form (H2O). Therefore, in any diskdominated by silicate vapor, hydrogen would not have been a dominant species ( f � 1). As aresult, the escape flux of hydrogen is small under such conditions, and a weak hydrodynamic

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wind would not efficiently drag other heavy elements (such as potassium). This analysis indicatesthat diffusion-limited hydrodynamic escape is not an efficient mechanism to remove hydrogen orother volatiles from the protolunar disk (for a detailed chemical disk model, see Visscher & Fegley2013).

If hydrodynamic escape was not significant, what caused the volatile depletion of the Moon?Hauri et al. (2015) and Canup et al. (2015) suggested on the basis of thermochemical argumentsand dynamical simulations that incomplete accretion could be responsible for the depletion. Asdescribed in Section 4.3, the Moon first accretes from the outer part of the disk, which Canupet al. (2015) estimated would be as volatile rich as Earth. Liquid from the inner part of the diskviscously spreads out and eventually crosses the Roche radius. The important aspect here is thatthis liquid is depleted in volatiles; vapor in the inner disk does not move with the liquid because itsviscosity is much less than that of the liquid. Before all the volatiles condense, inner-disk materialsbegin to accrete to Earth more efficiently than to the Moon due to gravitational scattering bythe growing Moon. As a result, Canup et al. (2015) suggested that the interior of the Moon ismade of volatile-rich materials that were initially located in the outer disk, whereas the exteriorof the Moon is made of volatile-depleted materials that were initially located in the inner disk.The key assumption here is that the Moon’s interior has not been vigorously mixed. This couldbe the case, given that the Moon may not have been completely melted or mixed down to itscore-mantle boundary, but it is still an open question whether the lunar interior is heterogeneous(e.g., MacDonald 1960, Solomon 1986, Zhang et al. 2013, Elkins-Tanton & Bercovici 2014).This incomplete accretion model can be further tested by investigating whether the lunar isotopicratios can be reproduced [such as potassium (Humayun & Clayton 1995)].

Hauri et al. (2015) described a somewhat different mechanism for nonaccretion of volatilesto the Moon; they postulated that the total vapor pressure at the midplane of the disk decreasesas a function of radius, while at the same time, the three-dimensional extent of a steam atmo-sphere surrounding the disk could occupy a larger volume than the condensed portion of the diskthroughout much of its thermal evolution (Figure 4). These two effects conspire to produce alow partial pressure of water in the region in which the Moon begins to coalesce, so that by thetime the Moon reaches its final mass, the steam atmosphere either remains largely extended orhas collapsed onto Earth. In both of these nonaccretion scenarios, volatiles are not lost from theEarth-Moon system; instead, nonaccretion appeals to a very low partial pressure of H2O in theprotolunar disk to produce a Moon depleted in water and other volatiles.

4.5. The Lunar Magma Ocean

The final stage of lunar evolution prior to solidification is the evolution of the LMO from amolten state to sufficient lithospheric thickness that the Moon becomes a one-plate body. Byfar the most uncertain aspect of lunar evolution is the initial thermal state of the Moon andits subsequent thermal evolution; the temperature of the LMO at the time of coalescence fromdisk-generated moonlets could have been anywhere between the liquidus temperature (∼1,850 K)and disk midplane temperatures (3,000–7,000 K). At the beginning of LMO crystallization, theMoon loses heat by radiation at its surface, and this mode of heat loss is not expected to transitionto conductive heat loss until the formation of a stable surface crust. The LMO could cool from3,000 K to the liquidus temperature on a timescale of <1,000 years, after which the thermalevolution would depend entirely on the details of surface heat loss. Elkins-Tanton et al. (2011)assumed that no stable surface crust forms until plagioclase begins to float, roughly 1,000 yearsafter the first crystals begin to form from the LMO. In the presence of a stable surface crust,heat loss from the LMO slows quickly as the heat flux is limited by conduction through the crust.


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Elkins-Tanton et al. (2011) predicted that the LMO, of depth 500–1,000 km, becomes fullycrystallized in ∼10 million years; however, if a stable crust exists throughout LMO crystallization,the loss of heat from the LMO is greatly reduced and the timescale for solidification can be aslong as 100–200 million years (Solomon & Longhi 1977, Meyer et al. 2010). This difference is afactor of ten, demonstrating that the timescale for full Moon solidification is highly sensitive tothe details of heat loss through the surface of the LMO.

Because the Moon is too small in mass to retain an atmosphere, convection of the LMOmagma to the Moon’s surface will release volatiles into the vacuum of space; however, the extentof degassing will be mitigated by the presence of a surface crust (Elkins-Tanton & Grove 2011).Yet this is the only window of time during which the Moon can gain water, because it is the mostprotracted period between the giant impact and full solidification of the Moon. During this time,it is also possible for impactors to be mixed by convection into the LMO interior. Assuming asuitable input of water from impacting hydrous meteorites and/or comets, the only time for theMoon to acquire additional water and transport it to the depths of picritic magma production[250–600 km (Elkins-Tanton et al. 2011)] is either during convection of the LMO or just prior tothe overturn of the lunar cumulate pile that is predicted by most models of LMO crystallization(Hess & Parmentier 1995). The sole factor that determines transport of water into the deep lunarinterior is the thickness of the lunar lithosphere; if there is no lithosphere, as is the prevailingcase during the initial phase of rapid convection of the liquid LMO, then impacting meteoritesand comets can be assimilated directly into the liquid LMO. But if the lithosphere becomessufficiently thick, impactors cannot penetrate into the lunar mantle, so they remain stalled in thecrust and thus cannot be mixed into the lunar mantle to participate in the generation of primitivemagmas.

4.6. The Moon’s Late Veneer

If the BSM contains 100–300 ppm H2O (Hauri et al. 2015), then the entire water inventory ofthe Moon amounts to 7.4 × 1018 to 2.2 × 1019 kg of H2O. If the Moon formed from a highlyenergetic giant impact that resulted in a molten protolunar disk, then the preceding discussiondemonstrates that it is unlikely that the Moon acquired this entire amount of H2O directly fromEarth. If carbonaceous chondrites are responsible for a late veneer, as appears required by theisotopic compositions of sulfur, chlorine, and hydrogen in lunar volcanic glasses (Thode & Rees1976, Ding et al. 1983, Sharp et al. 2010, Saal et al. 2013, Tartese et al. 2013, Furi et al. 2014),then constraints are placed on the implied flux of impactors. Carbonaceous chondrites contain∼7.8% H2O on average (Robert & Epstein 1982; Kerridge 1985; Alexander et al. 2012, 2013), soa minimum of 0.95–2.8 × 1020 kg of carbonaceous chondrites must be assimilated into the lunarinterior; a smaller amount could be accommodated if carbonaceous chondrites originally containedwater ice. If the lunar lithosphere develops in 10 million years (Elkins-Tanton et al. 2011), theresulting meteorite flux is 0.95–2.8 × 1013 kg/year; in a 200-million-year time frame (Solomon &Longhi 1977), the flux is 0.47–1.4 × 1012 kg/year. These fluxes amount to 0.001 to 0.004 lunarmasses, an order of magnitude higher than suggested by Albarede et al. (2015) but well within thetotal mass fluxes inferred for the tail end of planetary accretion (O’Brien et al. 2014). The fluxof carbonaceous chondrites calculated here is comparable to that estimated for Earth (Morbidelliet al. 2000, Becker et al. 2006, Halliday 2013) after accounting for the Moon-Earth difference inaccretion cross section, which is 50–200 times lower for the Moon than for Earth (Bottke et al.2007, Nesvorny et al. 2010, Schlichting et al. 2012). The required highly siderophile elementbudget estimated for the lunar mantle is <0.001 lunar masses (Day et al. 2007); the water budgetcan be made consistent with our calculations if carbonaceous chondrites contained water ice prior

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to their delivery to Earth or if some of the highly siderophile element budget were partitionedinto the lunar core.

Barnes et al. (2016) provided a more detailed analysis of the assembly of the Moon’s late veneerby working from two different starting scenarios: one in which the Moon is perfectly anhydrous(inheriting no water from Earth) and one in which variable amounts (up to 25%) of the BSMbudget of water is derived from Earth. Considering both H2O abundance and hydrogen isotopes,Barnes et al. (2016) determined that the bulk of the material assimilated into the Moon duringthe end of accretion would have to consist of CM, CV, CO, or CI carbonaceous chondrites (ora mixture of these classes), with CR carbonaceous chondrites, ordinary chondrites, and cometscontributing <20%. These results are consistent with the nitrogen isotope composition of theMoon (Furi et al. 2015, Mortimer et al. 2015).

5. FUTURE MODELS FOR THE FORMATION OF THE MOON

The isotopic compositions of refractory elements place constraints on the origin of the materialsthat compose the Moon (and Earth), while the budget of water and other volatiles not only revealsthe origin of volatile elements but also illuminates the high-temperature processes that lead to theformation of rocky planets. The Moon’s average volatile depletion factor of ∼0.25 (compared withthe bulk silicate Earth) is consistent with the ratios of volatile to refractory elements estimated byHauri et al. (2015), Delano & Livi (1981), and Ringwood (1992). The central question remains:How did the Moon acquire this composition? The answer to this question will also shed light onhow Earth acquired its volatile inventory.

There are several ways to estimate the composition of the Moon, all of them based on phys-ical processes that we cannot observe and thus need to simulate. Nearly all efforts to model theformation of the Moon have begun by assuming a single giant impact resulting in the formationof a high-temperature melt-vapor disk from which the Moon eventually forms. This so-calledhot start to lunar formation seems the most likely given our understanding of the energetics ofimpacts. However, if our intuition on the energetics is incorrect, then lower-temperature mod-els enter the realm of feasibility. At the nonmolten end of the spectrum, a cold start to lunarorigin would posit that the entirety of the Moon’s water was inherited directly from Earth’s man-tle in solid form, without the need to appeal to a melt-vapor disk, and that 75% of this waterwas outgassed by the LMO; alternatively, a warm start to lunar origin would posit that 25% ofthe Moon’s water was inherited from Earth in solid form, while the remaining 75% consists ofhighly degassed material from a protolunar disk that was merely partially molten (Hauri et al.2015). The warm start scenario is one that, in terms of water budget, is most closely matchedby Canup et al.’s (2015) recent model, though that model is a hot start model that appeals todegassing of the interior of a molten protolunar disk. Ultimately, many variants remain to beexplored, but the current frontier in understanding the formation of the Moon clearly lies in un-derstanding the dynamics and chemistry of the protolunar disk that emerges from a single giantimpact.

Additionally, N-body planetary accretion models uniformly predict several giant impacts in theformation region of the terrestrial planets (Chambers 2004, Jacobson et al. 2014, O’Brien et al.2014), and more than half of these impacts are predicted to generate circumplanetary disks (Gendaet al. 2012) that could either form satellites on their own or add material to an earlier-formedprotomoon. Canup et al. (1999) described a scenario with the occasional formation of two satellitesin orbit around Earth that eventually merge. Jutzi & Asphaug (2012) described the formation ofthe Moon in multiple accretion stages, and such a history was implicit in shallow magma oceanmodeling four decades ago (Solomon & Longhi 1977). Elkins-Tanton et al. (2011) demonstrated


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that the thickness of the anorthositic lunar crust requires a magma ocean no deeper than ∼800 km,and thus the Moon need not have formed in a fully molten state. These observations admit thepossibility of a multistage heterogeneous lunar accretion. Future modeling efforts may illuminatethese possibilities; however, a multistage history will depend on the likelihood of retention versusscattering of protomoons during multiple impact events.

The past decade of lunar research points clearly to some considerations that will be importantfor better understanding how the Moon, and indeed all of the rocky bodies in the Solar System,acquired its inventory of volatile elements. Melt inclusions from samples representing the fullspectrum of lunar magma compositions need to be identified and analyzed for the abundancesand isotopic compositions of hydrogen, sulfur, and chlorine. There exist mare basalt sampleswith similarities to nearly all of the 25 chemical categories of lunar volcanic glass identified byDelano (1986), and these need to be examined for the presence of melt inclusions. Laboratoryexperiments on the degassing and evaporation of highly and moderately volatile elements need tobe conducted under lunar surface conditions in order to reveal the volatile depletion pattern ofthe Moon and its relationship to volcanic degassing. An important research frontier is improvedthree-dimensional hydrodynamic simulations of the giant impact and protolunar disk, coupledwith improved equations of state and volatile solubility models, which will enable the formulationof hypotheses on volatile element evolution that can be tested with data from lunar samples.Finally, the next generation of lunar sample return should target pyroclastic deposits identifiedon both the nearside and the farside, as well as deep impact basins that may expose material fromthe lunar mantle that could be analyzed directly for water. These and other future considerationswill hasten progress on understanding volatiles in planetary objects as well as the utilization ofhydrogen for future mission resources.

SUMMARY POINTS

1. Previous measurements of water in lunar samples were dominated by terrestrial contam-ination and solar wind; only with improved microanalytical detection limits has it beenpossible to resolve the presence of magmatic water in lunar rocks.

2. The abundances of water and other volatiles (fluorine, sulfur, and chlorine) in primitivelunar magmas are similar to those in mid-ocean ridge basalts, and the abundances inlunar mantle are similar to those in the terrestrial upper mantle.

3. The hydrogen isotope composition of the lunar mantle overlaps with that of carbonaceouschondrites but is higher than that of Earth, possibly due to degassing of erupting lunarmagmas.

4. The isotopic compositions of hydrogen, sulfur, and chlorine are consistent withthe Moon’s volatile inventory being dominated by material similar to carbonaceouschondrites.

5. The protolunar disk produced by a giant impact is characterized by a low partial pressureof hydrogen and water; coalescence of the Moon from disk material is thought to lead toa very dry Moon covered in a global magma ocean.

6. The inventory of water and other volatiles in the lunar interior—and in Earth—was likelyderived from assimilation of late-accreting carbonaceous chondrites, the so-called lateveneer.

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FUTURE ISSUES

1. Melt inclusions from samples representing the full spectrum of lunar magma composi-tions need to be identified and analyzed for the abundances and isotopic compositionsof hydrogen, sulfur, and chlorine in order to provide further constraints on the origin ofvolatiles in the Moon.

2. Laboratory experiments on the evaporation of highly and moderately volatile elementsneed to be conducted under lunar surface conditions in order to reveal the volatile de-pletion pattern of the Moon.

3. Improved three-dimensional hydrodynamic simulations of the giant impact and proto-lunar disk, coupled with improved equations of state and volatile solubility models, arerequired for the formulation of hypotheses on volatile element evolution that can betested with data from lunar samples.

4. The next generation of lunar sample return should target pyroclastic deposits identifiedon both the nearside and the farside, as well as deep impact basins that may exposematerial from the lunar mantle that could be analyzed directly for water.

DISCLOSURE STATEMENT

The authors are not aware of any affiliations, memberships, funding, or financial holdings thatmight be perceived as affecting the objectivity of this review.

ACKNOWLEDGMENTS

The study of volatiles in the Moon stands on the shoulders of John Delano, and we thank himfor his continued encouragement. We also acknowledge the attention given to the Carnegie ionmicroprobes by Jianhua Wang and Ben Pandit. We thank NASA Curation and Analysis PlanningTeam for Extraterrestrial Materials (CAPTEM) for important sample loans and Francis Albarede,Rick Carlson, Tim Grove, Linda Elkins-Tanton, David Stevenson, Steve Desch, Robin Canup,Denton Ebel, Evelyn Furi, Marc Hirschmann, Kaveh Pahlevan, and Sarah Stewart for discussionson some of the ideas presented here. M.A. thanks Jessica Barnes, Romain Tartese, and Ian Franchifor numerous discussions about the apatite H isotope data and acknowledges funding from the UKScience and Technology Facilities Council (grant numbers ST/I001298/1 and ST/L000776/1–Project N to M.A.). This study was supported by NASA’s Cosmochemistry and LASER Programs,by the Lunar Science Institute [now the Solar System Exploration Research Virtual Institute(SSERVI)], and by the Carnegie Institution of Washington.

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EA45_FrontMatter ARI 11 August 2017 17:21

Annual Reviewof Earth andPlanetary Sciences

Volume 45, 2017Contents

Researching the Earth—and a Few of Its NeighborsSusan Werner Kieffer � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 1

The Fascinating and Complex Dynamics of Geyser EruptionsShaul Hurwitz and Michael Manga � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � �31

Plant Evolution and Climate Over Geological TimescalesC. Kevin Boyce and Jung-Eun Lee � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � �61

Origin and Evolution of Water in the Moon’s InteriorErik H. Hauri, Alberto E. Saal, Miki Nakajima, Mahesh Anand,

Malcolm J. Rutherford, James A. Van Orman, and Marion Le Voyer � � � � � � � � � � � � � � � �89

Major Questions in the Study of Primate OriginsMary T. Silcox and Sergi Lopez-Torres � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 113

Seismic and Electrical Signatures of the Lithosphere–AsthenosphereSystem of the Normal Oceanic MantleHitoshi Kawakatsu and Hisashi Utada � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 139

Earth’s Continental Lithosphere Through TimeChris J. Hawkesworth, Peter A. Cawood, Bruno Dhuime, and Tony I.S. Kemp � � � � � � � 169

Aerosol Effects on Climate via Mixed-Phase and Ice CloudsT. Storelvmo � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 199

Hydrogeomorphic Ecosystem Responses to Natural andAnthropogenic Changes in the Loess Plateau of ChinaBojie Fu, Shuai Wang, Yu Liu, Jianbo Liu, Wei Liang, and Chiyuan Miao � � � � � � � � � � � 223

Interface Kinetics, Grain-Scale Deformation, and PolymorphismS.J.S. Morris � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 245

Back-Projection Imaging of EarthquakesEric Kiser and Miaki Ishii � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 271

Photochemistry of Sulfur Dioxide and the Origin of Mass-IndependentIsotope Fractionation in Earth’s AtmosphereShuhei Ono � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 301

Southeast Asia: New Views of the Geology of the Malay ArchipelagoRobert Hall � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 331

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EA45_FrontMatter ARI 11 August 2017 17:21

Forming Planets via Pebble AccretionAnders Johansen and Michiel Lambrechts � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 359

Tungsten Isotopes in PlanetsThorsten Kleine and Richard J. Walker � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 389

Shape, Internal Structure, Zonal Winds, and Gravitational Field ofRapidly Rotating Jupiter-Like PlanetsKeke Zhang, Dali Kong, and Gerald Schubert � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 419

Effects of Partial Melting on Seismic Velocity and Attenuation: A NewInsight from ExperimentsYasuko Takei � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 447

Origin and Evolution of Regional Biotas: A Deep-Time PerspectiveMark E. Patzkowsky � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 471

Statistics of Earthquake Activity: Models and Methods for EarthquakePredictability StudiesYosihiko Ogata � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 497

Tectonic Evolution of the Central Andean Plateau and Implicationsfor the Growth of PlateausCarmala N. Garzione, Nadine McQuarrie, Nicholas D. Perez, Todd A. Ehlers,

Susan L. Beck, Nandini Kar, Nathan Eichelberger, Alan D. Chapman,Kevin M. Ward, Mihai N. Ducea, Richard O. Lease, Christopher J. Poulsen,Lara S. Wagner, Joel E. Saylor, George Zandt, and Brian K. Horton � � � � � � � � � � � � � � 529

Climate and the Pace of Erosional Landscape EvolutionJ. Taylor Perron � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 561

The Rise of Animals in a Changing Environment: Global EcologicalInnovation in the Late EdiacaranMary L. Droser, Lidya G. Tarhan, and James G. Gehling � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 593

The Late Heavy BombardmentWilliam F. Bottke and Marc D. Norman � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 619

Reconstructing Climate from GlaciersAndrew N. Mackintosh, Brian M. Anderson, and Raymond T. Pierrehumbert � � � � � � � � 649

Autogenic Sedimentation in Clastic StratigraphyElizabeth A. Hajek and Kyle M. Straub � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � � 681

Errata

An online log of corrections to Annual Review of Earth and Planetary Sciences articlesmay be found at http://www.annualreviews.org/errata/earth

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origin and evolution of water in the moon's interior · 2014. 6. 4. · ea45ch04-hauri ari 14...

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