Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Evaporative fractionation
of volatile stable isotopes
and their bearing on the origin
of the Moon
rsta.royalsocietypublishing.org
James M. D. Day1 and Frederic Moynier2
1 Scripps Isotope Geochemistry Laboratory, Geosciences Research
Review
Cite this article: Day JMD, Moynier F. 2014
Evaporative fractionation of volatile stable
isotopes and their bearing on the origin of the
Moon. Phil. Trans. R. Soc. A 372: 20130259.
http://dx.doi.org/10.1098/rsta.2013.0259
One contribution of 19 to a Discussion Meeting
Issue ‘Origin of the Moon’.
Subject Areas:
geochemistry, geology, petrology, Solar
System, space exploration, astrochemistry
Keywords:
volatile elements, Moon, the Earth, Mars,
isotopic fractionation, zinc
Author for correspondence:
James M. D. Day
e-mail: [email protected]
Division, Scripps Institution of Oceanography, La Jolla,
CA 92093-0244, USA
2 Institut de Physique du Globe de Paris, Université Paris Diderot,
Sorbonne Paris Cité, 1 rue Jussieu, 75005 Paris, France
The Moon is depleted in volatile elements relative
to the Earth and Mars. Low abundances of volatile
elements, fractionated stable isotope ratios of S,
Cl, K and Zn, high µ (238 U/204 Pb) and long-term
Rb/Sr depletion are distinguishing features of the
Moon, relative to the Earth. These geochemical
characteristics indicate both inheritance of volatiledepleted materials that formed the Moon and
planets and subsequent evaporative loss of volatile
elements that occurred during lunar formation
and differentiation. Models of volatile loss through
localized eruptive degassing are not consistent with
the available S, Cl, Zn and K isotopes and abundance
data for the Moon. The most probable cause of
volatile depletion is global-scale evaporation resulting
from a giant impact or a magma ocean phase where
inefficient volatile loss during magmatic convection
led to the present distribution of volatile elements
within mantle and crustal reservoirs. Problems
exist for models of planetary volatile depletion
following giant impact. Most critically, in this model,
the volatile loss requires preferential delivery and
retention of late-accreted volatiles to the Earth
compared with the Moon. Different proportions of
late-accreted mass are computed to explain presentday distributions of volatile and moderately volatile
elements (e.g. Pb, Zn; 5 to >10%) relative to highly
siderophile elements (approx. 0.5%) for the Earth.
Models of early magma ocean phases may be more
effective in explaining the volatile loss. Basaltic
materials (e.g. eucrites and angrites) from highly
differentiated airless asteroids are volatile-depleted,
2014 The Author(s) Published by the Royal Society. All rights reserved.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
The origin of the Moon and the chemical and physical evolution of the Earth–Moon system
remain unresolved, fundamental science problems. Dynamical models of planet formation and
giant impact provide a broad range of possible formation scenarios, some of which appear to
overcome previous angular momentum constraints [1–3]. Chronological evidence for the timing
of formation of the Moon remains controversial [4–6], with age uncertainties resulting from lack
of resolution of parent/daughter ratios, limitations in precision of measured isotopic ratios and
restrictions in accurate dating of the oldest available crustal samples that have ubiquitously
suffered later impact processes at the lunar surface. Some of the strongest available evidence for a
co-genetic origin for the Moon and the Earth comes from the identical stable isotope compositions
of oxygen [7–9], titanium [10] and silicon [11,12]. These results implicate the formation of the
Moon from a silicate reservoir that was isotopically identical to the bulk silicate Earth for these
elements. Whether this isotopic homogeneity reflects mixing during a giant impact [1,2], isotopic
homogenization between the proto-Earth and a silicate melt–vapour disc [13], or if the terrestrial
planet formation region was already isotopically well mixed prior to the formation of the Earth,
the Moon and Venus from previously differentiated planetary embryos, remains unresolved.
The Moon is depleted with respect to the bulk silicate Earth in elements that are
cosmochemically volatile or moderately volatile (figure 1a; e.g. [19–21]). Volatile and moderately
volatile elements are defined as those elements for which 50% condensation temperatures at
cosmochemical gas pressures (less than 10−4 atm) are less than 665 and 1135 K, respectively [14].
These properties extend to volatile species such as H2 O (or OH). Hydroxyl has been detected in
some lunar pyroclastic glasses [22], as well as apatites formed during the final stages of fractional
crystallization within mare basalts [23], but water appears to be more depleted in the Moon than
in the Earth. For moderately volatile elements, such as S, K, Cl, Zn, Rb, Cd and Pb, the degree of
depletion in the Moon is similar to other volatile-depleted and differentiated planetesimal bodies
in the Solar System, such as the parent bodies of eucrite or angrite meteorites (figure 1b).
Volatile and moderately volatile element depletions in the Moon are not recent effects, nor are
their depletions solely related to removal in early, hot nebular gas [24]. The currently known
chronology of volatile depletion for the Moon suggests both inheritance of volatile-depleted
materials to form the Moon (and the Earth), as well as later volatile depletion. The ubiquitously
low measured 87 Sr/86 Sr ratios (less than 0.7; initial 87 Sr/86 Sr = 0.69906; e.g. [5,25]) and extremely
high µ (238 U/204 Pb) values for lunar mare basalts (approx. 300) and lunar pyroclastic glass beads
(19–55; for comparison, the lowest terrestrial 87 Sr/86 Sr ratio comes from 3450 Ma barite, 0.70053,
e.g. [26]; compared with terrestrial 238 U/204 Pb ratios estimated at 6–8 [27]), indicate that Rb and
Pb depletion occurred early, either prior to or during formation of the Moon. The timing of loss
of moderately volatile elements is not well constrained, but is within current estimates of the
timing of the Moon-forming event [4–6]. Equally, marginally more radiogenic measured 87 Sr/86 Sr
ratios in lunar samples, compared with angrites and eucrites, indicate early loss of Rb on some
differentiated asteroidal bodies (within the first 2–3 Ma after Solar System formation for angrites
and eucrites; e.g. [28–30]).
Given the evidence for depletion of abundances of volatile and moderately volatile elements
for the Moon relative to the Earth, an increasing number of studies are examining evaporative
fractionation effects on stable isotopes of moderately volatile elements [31–34]. Their aim is to
constrain the nature of volatile element depletion and so potentially to reveal how and when such
depletion occurred and whether it relates to nebula (i.e. prior to the formation of planets and
.........................................................
1. Introduction
2
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
like the Moon, whereas the Earth and Mars have proportionally greater volatile contents.
Parent-body size and the existence of early atmospheres are therefore likely to represent
fundamental controls on planetary volatile retention or loss.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
atmophile
0.1
volatile
moderately
volatile
refractory
0.01
0.001
highly siderophile
elements
0.0001
0
500
1000
1500
2000
50% condensation temperature (K) at 10–4 bar
3
EL H, L, LL
1
CM
CV
EH
CO
Mars
0.1
Earth
Moon
0.01
HED
0.001
angrites
0.0001
1
0.001 0.01 0.1
K/U (normalized to CI chondrite)
Figure 1. (a) Estimated abundances of the elements classified according to their 50% condensation temperatures at 10−4 bar
(from [14]) in the mantles of the Earth (unfilled circles), Moon (filled circles) and the HEDPB (grey diamonds). (b) Carbonaceous
Ivuna (CI) chondrite normalized Rb/Sr versus K/U ratios in bulk meteorites and planets. Data in (a) from Dreibus & Wanke [15],
O’Neill [16], McDonough & Sun [17], with highly siderophile element abundance data for the Moon from Day et al. [18]. The
dashed line in (a) denotes the region in which the highly siderophile elements are located for the Earth and the Moon. Panel (b)
is modified from the data compilation of Davis [19].
1.0
0.6
0.2
6
4
carbonaceous
chondrites
a = 0.999
–0.2
ordinary
chondrites
d 66Zn (‰)
100
2
0
a = 0.993
a=1
–2
–4
–6
maximum vapour compositions
using fractionation factors (a)
from 0.9995 to 0.993
–8
0.1
1
10
Zn (ppm)
eucrites
low-Ti mare basalts
high-Ti mare basalts
martian meteorites
terrestrial basalts
100
1000
Figure 2. Zinc isotope compositions and abundances of planetary materials versus ordinary (plus symbols) and carbonaceous
chondrites (cross symbols). Rayleigh distillation curves (dashed lines) for fractionation factors (α) of 1, 0.999 and 0.993
(upper limit of ideal fractionation behaviour for ZnCl) are shown for melts with a carbonaceous chondrite initial composition
(δ 66 Zn = 0.33; 300 ppm Zn). The corresponding field for maximum calculated vapour compositions for limited (0.9995) to
ideal fractionation behaviour for ZnCl is shown for comparison. Inset shows expanded view of chondrite–Earth–Mars data. Data
are from [34–37].
planetesimals), large-scale global (e.g. giant impact, lunar magma ocean) or localized (e.g. lava
flow or fire fountain degassing) evaporation. Here, we review the underlying theory, seeking
evidence for evaporative fractionation of stable isotopes of moderately volatile elements in
planets and, using the currently available S, Cl, K and Zn data, place new constraints on how
and when such processes occurred in the Moon. Our prime motivation is to explain the evidence
for lower Zn abundances and high 66 Zn/64 Zn ratios measured in lunar mare basalts relative to
martian or terrestrial basalts or ordinary and carbonaceous chondrites (figure 2). These differences
have previously been interpreted to reflect a global-scale evaporation event associated with lunar
.........................................................
1
Rb/Sr (normalized to CI chondrite)
(b)
10
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
abundances normalized to CI and Si
(a)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Element volatility is a fundamental chemical property occurring in response to temperature
variations experienced by a chemical system, the associated condensation temperatures of the
elements, and the two processes of evaporation and condensation that define volatility. Under
solar nebula conditions (i.e. gas pressures of less than 10−4 atm; [14]) all meteorite classes, and
terrestrial and lunar rocks are depleted in the volatile elements compared with carbonaceous
Ivuna-type (CI) chondrites (figure 1). CI chondrites are generally regarded as being representative
of the bulk Solar System elemental composition (with the exception of abundances of highly
volatile elements (H, C, N, O, noble gases) and elements that are reactive in stellar environments
(Li, Be, B)), because of the similarity of their elemental distribution with those determined
spectroscopically in the solar photosphere [14,38–40]. Volatile depletions for various types of
undifferentiated planetary materials, relative to CI chondrites, differ, but for the most part have
the common characteristic of being reasonably smooth functions of volatility (expressed as 50%
condensation temperatures). The volatile element patterns of differentiated planetary rocks are
typically less smooth owing to the effects of igneous processes. Some incompatible elements,
however, are not significantly decoupled by igneous differentiation, and ratios of volatile to
refractory incompatible elements can be used to assess the level of volatile depletion in planets
and parent bodies [19,39,41].
A logical system for defining cosmochemical volatility has been adopted that divides the
elements into different groups based on their physical and chemical properties [42]. The primary
cosmochemical classification of the elements is based on volatility, with the elements divided
into those that are refractory (having equilibrium condensation temperatures higher than those
of the most abundant rock-forming elements, magnesium, silicon and iron), moderately volatile
(having condensation temperatures lower than those of the refractory elements, but higher than
that of FeS), and highly volatile (having condensation temperatures below that of FeS). For a gas
of solar composition at a total pressure of 10−4 atm, refractory elements have 50% condensation
temperatures above 1335 K, moderately volatile elements between 1335 and 665 K, and volatile
elements below 665 K [14]. A fourth group, the atmophile elements (hydrogen, carbon, nitrogen,
oxygen, and the noble gases), have condensation temperatures of 180 K or lower and are depleted
in CI chondrites, relative to the Sun. The other important classification of elements is based on
which mineral phase, or phases, elements concentrate into: siderophile elements concentrate
into iron–nickel metal, chalcophile elements into sulfides, and lithophile elements into oxides
and silicates. An example of how this concentration behaviour affects elemental distributions is
oxygen; although it is an atmophile element, it behaves as somewhat refractory owing to its early
incorporation into oxides and silicates. Clearly, the timing of volatile depletion, coupled with
the incorporation of elements into phases based on their concentration behaviour, is profoundly
important in the current distribution of elements in the Solar System.
Volatile element depletion can be simplified into two processes: partial condensation from a
cooling gas; and partial evaporation from condensed dust or melt. For elements with more than
one isotope, a physical kinetic or fractional equilibrium stable isotope effect during evaporation
then provides a measure of the extent of partial evaporation. The kinetic isotope effect can be
best understood using the moderately volatile elements such as S, Cl, K and Zn, for which
there is no evidence for large fractionations of their isotopes during planetary differentiation and
which have well-determined abundance depletions. In classical theory, natural mass-dependent
stable isotopic fractionation of an element occurs during any exchange reaction and is due to
the difference in vibrational energy of the bonds formed by the different isotopes at equilibrium
or under kinetic conditions [43,44]. The general rule behind isotopic fractionation is that heavy
.........................................................
2. Isotopic fractionation of moderately volatile elements in cosmochemistry
4
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
formation [34]. In assessing evidence from fractionations observed in stable isotopes of S, Cl,
K and Zn, we find the data to be most consistent with global-scale evaporative fractionation
of moderately volatile and volatile elements immediately in the aftermath of the Earth–Moon
forming event.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
In the following section, we review available stable isotopic and associated elemental abundance
data for cosmochemically moderately volatile elements (S, Cl, K, Zn) for which there are
comparable data for lunar, terrestrial and martian basalts, as well as data for eucrite meteorite
samples. We note that there are numerous moderately volatile or volatile elements that have
two or more isotopes and can potentially provide information on evaporation and condensation
processes operating during the evolution of the Moon, applying the caveats described in §2
(e.g. 11 B/10 B, 908 K, moderately volatile; 65 Cu/63 Cu, 1037 K, moderately volatile siderophile
(MVS); 71 Ga/69 Ga, 968 K, MVS; 74,73,72,70 Ge, 883 K, MVS; 78,77,76 Se, 697 K, highly volatile
chalcophile (HVC); 81 Br/79 Br, 546 K, highly volatile; 87 Rb/85 Rb, 800 K, moderately volatile;
Cd isotopes, 652 K, HVC; Sn isotopes, 704 K, HVC; Hg isotopes, 252 K, HVC). For the rest
of this paper, isotopic ratios are reported as delta (parts per one thousand or per mil ())
notation, where
δX =
(X/Y)Sample
(X/Y)Standard
− 1 × 1000,
(3.1)
where X is the heavier mass isotope (34 S, 37 Cl, 41 K, 66 Zn) and Y is the lighter mass isotope (32 S,
35 Cl, 39 K, 64 Zn).
.........................................................
3. The S, Cl, K and Zn compositions of lunar, terrestrial and martian rocks
5
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
isotopes of an element will have a tendency to be enriched in the compounds where that element
forms the stiffest (usually the shortest and strongest) bonds [45]. A natural implication of this
effect is that, during evaporation/condensation processes, the lightest isotopes will have the
tendency to be enriched in the vapour phase while the heaviest isotopes will be enriched in the
condensed phase. The amplitude of these effects is proportional to the mass difference between
the isotopes and therefore it is most notable for light elements (e.g. H, Li, O, S). The development
of a new generation of multi-collector inductively coupled plasma mass spectrometers
(MC-ICP-MS) in the late 1990s has allowed high-precision measurements of the isotopic ratios of
heavier elements (e.g. Fe, Zn, Cu) and small but resolvable isotopic variations are now measurable
[46]. For geochemical, cosmochemical and lunar applications, much of the theoretical form of
work has been reviewed previously [19,45,47] and the reader is referred to these articles for
further details.
In order to choose stable isotopes that can effectively demonstrate isotopic fractionation
during moderately volatile element evaporation, some of the same rules apply as for ‘standard’
stable isotope systems, such as oxygen. Classic stable isotopic theory states that elements
that show stable isotopic variations in nature generally have: (i) low atomic mass; (ii) large
relative mass differences between stable isotopes; (iii) a tendency to form highly covalent bonds;
(iv) multiple oxidation states or other chemical variability; and (v) two isotopes or more with
sufficient abundance to make measurements feasible (cf. [45,48]). While these rules still apply,
development of clean chemistry and precise and accurate MC-ICP-MS methods has expanded
the ability to measure higher atomic masses at smaller relative mass differences (e.g. Zn, U).
Additionally, stable isotope systems such as Li, Mg, Si, Ca and Zn have a single predominant
oxidation state that can be advantageous for simplifying interpretation of stable isotope data.
For some volatile and moderately volatile elements, the determining factor in stable isotope
ratio measurement is absolute abundance in natural materials. For example, low abundances
of Hg, Cd and Sn in lunar rocks makes measurement of their isotopic ratios difficult given the
available sample masses of lunar and asteroidal materials for study from the Apollo missions and
from meteorite falls and finds. Zinc, on the other hand, while depleted in absolute abundances
in lunar rocks, occurs in sufficient quantity (more than 1 µg g−1 ) that its isotopic ratios can be
measured precisely.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
(a)
(b)
6
regolith/soil
regolith/soil
–2.0 –1.5 –1.0 –0.5
–4 –2
0
2
4
6
8
0
0.5
Moon
Mars
Earth
1.0
10 12 14 16
d 34S (‰)
(c)
–4 –2 0
2
4
6
8 10 12 14 16 18
d 37Cl (‰)
(d)
regolith/soil
regolith/soil
mare
basalts
mare
basalts
PGB
condensation
–6 –4 –2 0
0
2
4
6
d 41K
8
(‰)
10
12
14
0
2
4
d 66Zn
2
4
6
6
8
8
(‰)
Figure 3. Comparisons of (a) S, (b) Cl, (c) K and (d) Zn isotope ratios (reported as per mil values) measured in lunar mare basalts
compared with lunar regolith/soil, terrestrial lavas and martian shergottites. All data are for bulk rock samples. Inset panel in
(d) shows pyroclastic glass beads and mare basalts with isotopically light Zn inherited from post-crystallization condensation.
Data for (a) from [53,54,56], for (b) from [32,57–59], for (c) from [31] and for (d) from [34,37].
(a) Sulfur
Sulfur has four stable isotopes, 32 S (95.02%), 33 S (0.75%), 34 S (4.21%) and 36 S (0.02%), with the
typical mass-dependent stable isotope value reported as 34 S/32 S (δ 34 S) relative to Canyon Diablo
Troilite (CDT; 34 S/32 S = 0.045). Sulfur has the lowest 50% condensation temperature (664 K) of the
elements discussed here and, as well as being highly volatile under cosmochemical conditions,
is also chalcophile [14], making its behaviour distinct from those of Cl, K or Zn. Early studies of
lunar samples revealed a relatively large spread in δ 34 S values for mare basalts (−1.6 to +0.9;
[49]), and significantly 34 S-enriched components in lunar regolith and soil materials (+6.1 to
+13.5 and 320–790 ppm; [50]), interpreted to reflect the loss of 20–30% of the abundance of
S (and K) from the lunar regolith by ion sputtering [51]. A recent study has revealed a restricted
range in δ 34 S for mare basalts of +0.57 ± 0.09 [52], while work to identify mass-independent S
fractionation in shergottites indicates an average δ 34 S for martian basalts of −0.05 ± 0.3 [53].
High-precision 34 S/32 S data have been obtained for terrestrial mid-ocean ridge basalts
(MORB), giving an average δ 34 S of −0.98 ± 0.51 [54,55]. Correlations with Sr and Nd isotope
ratios in MORB suggest mixing with a low-δ 34 S mantle and a high-δ 34 S crustal component,
in turn indicating that the depleted mantle has low δ 34 S (−1.28 ± 0.33) when compared to
lunar rocks (+0.57 ± 0.09), potentially indicating a core fractionation effect [55] and/or a
loss of the lighter S isotopes from the Moon. The range of S concentrations in lunar basalts
(730 to >1500 ppm; [52]), martian shergottites (300 to >1500 ppm; [53]) and terrestrial MORB
.........................................................
mare
basalts
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
mare
basalts
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Chlorine has two stable isotopes, 35 Cl (75.78%) and 37 Cl (24.22%), and the delta notation
(δ 37 Cl) is typically expressed relative to standard mean ocean chloride (SMOC). Chlorine has
a 50% condensation temperature estimated as 948 K and is highly volatile [14]. It is also
volatile at magmatic temperatures, incompatible during silicate melting and water-soluble, and
it has consequently been identified as a powerful potential tracer of interactions between the
terrestrial hydrosphere, crust and mantle [57,59]. Lack of significant Cl isotope fractionation in
the terrestrial mantle, crust and carbonaceous chondrites led Sharp et al. [57] to conclude lack
of nebular fractionation of Cl isotopic compositions, or of late accretion Cl additions to the
Earth. Compilations of terrestrial MORB data give an average Cl content of 205 ppm (range = 42–
1400 ppm), with δ 37 Cl of −0.7 ± 0.6 [57,59,62]. The large range in Cl contents and elevated δ 37 Cl
relative to estimated mantle composition (−1.6) indicates ubiquitous alteration of terrestrial
volcanic rocks by seawater, brines or hydrothermally altered surface rocks. Since terrestrial
magmas and lavas show a limited range of δ 37 Cl, it has been argued that there is always an
excess of H relative to Cl, and that the dominant volatile chloride species is therefore HCl gas
with a limited isotopic fraction between basaltic melts [32]. However, more recent work by
Sharp et al. [33] provided evidence for significant isotopic fractionation of Cl isotopes in HCl.
Limited data for martian shergottites also indicate a limited range in Cl isotope compositions
(δ 37 Cl = −0.4 ± 0.4) and a range in Cl contents from 24 to 115 ppm [58].
Chlorine isotopes have been strongly fractionated in some lunar materials (−1 to +24),
leading Sharp et al. [32] to suggest a mechanism of volatilization of Cl as metal halides (e.g.
NaCl, ZnCl2 , FeCl2 ), loss of Cl during basalt eruption on the lunar surface and, consequently,
an essentially anhydrous lunar interior. Volatilization of Cl on the Moon is possible because of the
low calculated H/Cl ratio for lunar basalts and from the presence of metal halides on the surface
coatings of lunar pyroclastic glass beads [63]. Sharp et al. [32] presented Cl isotope and abundance
data for bulk rock analyses of lunar mare basalts, regolith materials and lunar pyroclastic glass
beads by gas-source mass spectrometry, as well as Cl isotope analyses of apatites performed by
secondary ionization mass spectrometry. Lunar regolith components (+5.6 to +16 and 8–117
ppm Cl) are enriched in 37 Cl relative to mare basalts (0 to +4.2 and 5–25 ppm Cl) and apatites
within mare basalts and KREEP-rich (potassium–rare earth elements–phosphorus) rocks (lunar
sample no. 72275) are also enriched in 37 Cl (+13.1 to +24.5).
Sharp et al. [32] reported apatite in sample 12040 with δ 37 Cl of +17.2 and 0.83 wt% Cl and
reported a bulk composition for the same mare basalt of δ 37 Cl = 0 and 11 ppm Cl. Assuming
0.1–0.2 modal per cent phosphate in 12040—a reasonable approximation for a primitive low-Ti
mare basalt [64–66]—and using the Cl concentration measured by Sharp et al. [32] indicates that
approximately 9 ppm (out of the 11 ppm measured [32]; i.e. approx. 80% of the total Cl) is hosted
within apatite minerals in mare basalts and implies either that the isotopic composition of the
bulk should be much heavier than that reported by Sharp et al. [32], or that the approximately
20% of Cl not hosted within apatites should be extremely (and improbably) isotopically light
(approx. −70). Variable δ 37 Cl values occur for apatite in individual lunar samples, and δ 37 Cl
in excess of +25 has been measured in apatites with more than 5000 ppm chlorine and with
relatively high OH contents [67]. Therefore, values provided by individual apatites can be variable
within a single rock and may reflect late-stage, secondary volatile-loss effects during extreme
fractional crystallization. Given uncertainties in how to interpret the apatite Cl isotope and OH
contents, the possibility for heterogeneous distribution of δ 37 Cl within individual apatites and
.........................................................
(b) Chlorine
7
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
(728–1221 ppm; [54]) are similar, but lunar mare basalts have isotopically heavier δ 34 S values
(figure 3a). The similarities in the range of S concentrations for lunar, terrestrial and martian
rocks may at first appear an impediment for volatile loss. However, S is typically exhausted after
more than 20% partial melting, with most mantle-derived partial melts from the Earth or the
Moon being S-undersaturated [60,61], indicating that enrichment of S is expected in mantle melt
products.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Pioneering work in the 1990s [31,68] revealed limited isotopic fractionation of the two stable
isotopes of potassium, 39 K (93.3%) and 41 K (6.7%; the long-lived radioisotope 40 K makes up
0.012% of the natural abundance), in planetary materials. The work of Humayun and others
reports 41 K/39 K ratios as δ 41 K relative to Suprapur KNO3 [68,69]. Limited work on K has
been done in the past 20 years, although a study of K in terrestrial tektites revealed no
discernible isotopic fractionation [69]. Given that potassium is moderately volatile and has a 50%
condensation temperature of 1006 K, a lack of isotopic fraction led Humayun & Clayton [68] to
conclude that K and other moderately volatile alkalis (e.g. Na, Rb, Cs) must have been depleted
prior to chondrule, chondrite or planet formation, probably from a hot dust stage in the early
solar nebula.
Despite lack of 41 K/39 K isotopic fractionation between terrestrial basalts (δ 41 K = +0.28 ±
0.21), martian shergottites (+0.5), eucrites (+0.9 to +2.7) and mare basalts (+0.1 to +1.5),
Humayun & Clayton [68] found isotopically light δ 41 K in a ferroan anorthosite (60015, −3.9),
and isotopically heavy δ 41 K in lunar soils (+5 to +12.7) [31] (figure 3c). Additionally, these
authors confirmed previous studies [51] that had observed higher K contents in terrestrial basalts
and martian shergottites (typically greater than 1000 ppm), relative to mare basalts (300–750 ppm)
or eucrites (160–300 ppm). The lack of isotopic fractionation of 41 K/39 K in basalts formed by
high-temperature igneous processes, or in tektites formed by dry melting during hypervelocity
impacts, indicates lack of K fractionation, either during melting, or by partial evaporation [69].
Recasting the K isotopic data suggests variable and slight heavy isotopic enrichment in mare
basalts relative to terrestrial or martian igneous rocks (figure 3c).
(d) Zinc
Numerous studies have been performed to characterize the zinc isotopic composition of planetary
materials [34–37,70,71]. Zinc is composed of five stable isotopes, 64 Zn (48.6%), 66 Zn (27.9%),
67 Zn (4.1%), 68 Zn (18.8%) and 70 Zn (0.6%), but, because of its low natural abundance, 70 Zn is
rarely reported in high-precision isotopic measurement studies. The only recent study reporting
high-precision 70 Zn/64 Zn ratio showed that this ratio is mass-dependently fractionated within
chondritic material [72]. Along with Cl, Zn is classified as highly volatile, with a 50% condensation
temperature of 726 K [14]. Zinc isotope ratios are reported as 66 Zn/64 Zn (δ 66 Zn), 67 Zn/64 Zn
(δ 67 Zn) or 68 Zn/66 Zn (δ 68 Zn), typically relative to the ‘Lyon’ Zn standard JMC 3–0749 L [46].
All the Zn isotopic data measured in Solar System materials to date fall on a slope 2 line on
a δ 68 Zn versus δ 66 Zn plot, conforming to the expected mass-dependent mass-fractionation law.
The same is true for 67 Zn/64 Zn and 70 Zn/64 Zn [72]. Terrestrial, martian and lunar rocks all lie on
the same mass-dependent mass-fractionation line, along with all classes of chondritic meteorites
[34,35,73], indicating that Solar System Zn evolved from a single, isotopically homogeneous
reservoir. All reported isotopic variations are therefore due to mass-dependent fractionations.
Mass-dependent fractionation of Zn isotopes also rules out secondary neutron capture effects,
which would lead to non-mass-dependent behaviour.
Zinc is more volatile than both K and Cl under solar nebula conditions [14] and exhibits limited
isotopic fractionation during igneous processes [37]. On the Earth, and in primitive meteorites
such as chondrites, variations of δ 66 Zn larger than 1 have only been observed associated with
evaporation events [70,73,74]. Any Zn isotopic fractionation, combined with differences in Zn
concentrations between planetary igneous rocks from differentiated bodies such as the Earth,
.........................................................
(c) Potassium
8
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
the late-stage fractionation of apatite within lunar mare basalts, we only report bulk rock gassource mass spectrometry analyses of Cl for mare basalts in figure 3b. Bulk rock data indicate
that martian and terrestrial basalts have similar δ 37 Cl, whereas mare basalts are characterized by
37 Cl-enriched compositions and Cl contents that are typically 2–40 times lower than basalts from
the Earth or Mars.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
An important aspect of the Zn isotope data for the Moon is that the isotopic fractionation does
not conform to ideal isotopic fractionation by distillation of ZnCl2 or ZnS, which are both possible
(although not the only) species involved in lunar degassing (figure 2). For ideal evaporation of a
liquid into a vacuum, the Rayleigh distillation equation is
δfinal = δinitial + [(1000 + δinitial )(F(α−1) − 1)],
where F is the fraction of
vapour and melt:
64 Zn
(4.1)
remaining in the melt and α is the fractionation factor between
α = (1000 + δvapour )(1000 + δmelt ).
(4.2)
The fractionation factor (α) can be estimated as an upper limit for ideal behaviour using the
square-root of the light isotopologue over the heavy isotopologue of the vaporizing species in
the case of kinetic fractionation [76],
my
(4.3)
α=
mx
where my and mx are the masses of the light and heavy isotopologue, respectively. For ZnCl2 the
α value is 0.993. Extreme fractionations of Zn isotopes in the melt and associated vapour phase
are computed for reducing Zn concentrations in the melt fraction. Fractionations greater than
those measured can be calculated, suggesting either incomplete evaporation, or the competing
effect of diffusion rates being insufficient to maintain isotopic homogeneity in lunar melts (effect
exemplified as lower α value of 0.999, for reference; figure 2). Fractionations less than theoretical
calculations are observed experimentally [56,77] and would be expected if diffusion rates are
insufficient to maintain isotopic homogeneity of the melt.
Potential analogues to explain inefficient and partial evaporation of Zn and, by extension,
stable isotope fractionations of other volatile elements, are tektites. Tektites are terrestrial natural
glasses that are produced during hypervelocity impacts or airbursts of a projectile into terrestrial
target rocks or soils [78,79]. Some tektites are extremely water-poor (less than 0.02 wt% H2 O)
.........................................................
4. Evidence for a non-ideal Rayleigh regime during isotopic fractionation of Zn
9
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
the Moon and Mars, can reveal important differences in volatile depletion and replenishment
events for these bodies. Extensive studies of terrestrial basalts have revealed a limited variation
in Zn isotopic compositions and a precise bulk silicate Earth δ 66 Zn value of +0.28 ± 0.05 (2σ ),
with terrestrial basalts typically having between 60 and more than 160 ppm Zn, the concentration
of which broadly correlates with indices of differentiation [37]. Martian meteorites span a
limited range in δ 66 Zn (+0.13 to +0.35), with an average δ 66 Zn of +0.25 ± 0.03 [34]. Zinc
concentrations in shergottite meteorites are similar to those in terrestrial basalts at between 41
and 130 ppm, and Zn concentrations in nakhlite meteorites lie within the range of shergottite
meteorites at between 61.5 and 89 ppm [34].
Lunar samples exhibit the largest range of δ 66 Zn from any differentiated planetary body in
the Solar System. However, systematic variations are observed between: (i) lunar pyroclastic
glass beads, which have negative δ 66 Zn values (−2.9 to −4.2) and high Zn concentrations
(129–231 ppm); (ii) pristine low-Ti mare basalts, which have an average δ 66 Zn value of +1.31 ±
0.13 and are isotopically identical to high-Ti mare basalts (δ 66 Zn = +1.39 ± 0.31), and
have Zn concentrations (0.6–12 ppm) at least a factor of 10 lower than in lunar pyroclastic
glass beads; and (iii) lunar regolith, soils and anorthositic regolith breccia meteorites (MAC
88105), which have isotopically heavy δ 66 Zn (+2.6 to +5.6) (figure 3d) [34,70,71]. The high
δ 66 Zn of the regolith and soil samples reflect both sputtering effects and impact gardening
of the lunar regolith [70,71] and do not represent primary magmatic compositions. There is
a class of mare basalts with isotopically light δ 66 Zn (figure 3d, inset) that result in a similar
range in lunar δ 66 Zn compared with eucrite meteorites (−7.8 to +6.2) [36] and have been
interpreted as the consequence of condensation of isotopically light zinc after mare basalt
crystallization [34,75].
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
(a) Nebula versus planetary volatile loss
Extreme volatile depletion of the Moon, as implied from the low concentration and isotopically
heavy Zn isotope enrichment seen in lunar mare basalts, contrasts with more elevated Zn
abundance and broadly chondritic Zn isotopic compositions for the Earth and Mars [34].
Comparison of volatile and moderately volatile element abundance and isotopic compositions
indicate that mare basalts are systematically volatile-depleted and are enriched in the heavy
isotopes of moderately volatile elements compared with terrestrial or martian basalts (e.g.
figures 1 and 3). Taken as a whole, the data indicate that moderately volatile element depletion
is ubiquitous among mare basalts, through a large- rather than a small-scale process, or by
hydrodynamic escape [83]. Given the differences in moderately volatile element abundances
between the Moon and the Earth, it is not likely that this variation solely reflects depletion prior
to planet formation [24,31]. A depletion process prior to the formation of chondrules, chondrites
and planet formation in the hot, dusty solar nebula can explain the systematic depletion of volatile
elements in chondrites [19], which are proxy materials for the building blocks of planets, but other
depletion and/or replenishment events are also required to explain the differing volatile budgets
among differentiated planetary bodies.
At least two phases of volatile and moderately volatile depletion must have occurred during
the formation of the Solar System. The first is at the nebula scale, prior to or during the formation
of the first solids (cf. calcium–aluminium inclusions, chondrules). The second volatile depletion
event, or series of events, is associated with planetesimal and planet formation and appears to
relate with escape velocity (figure 5) and whether the body could retain a long-term atmosphere
(e.g. the Earth, Mars, Venus) or not (e.g. Moon, Mercury, angrite or howardite–eucrite–diogenite
parent bodies (HEDPB)), both during and after their differentiation. We refer to the first event as
nebula volatile depletion and to the second as planetary volatile depletion. It is planetary volatile
depletion that is considered in detail here, and that is applied to understanding the Moon’s origin.
We suggest two process-driven mechanisms for planetary volatile depletion for the Moon and
the Earth after the initial phase of nebula volatile depletion.
First, evaporation of volatile and moderately volatile elements could have occurred during the
high-energy Earth–Moon forming event [34,84], which was then followed by later replenishment
.........................................................
5. Mechanisms for generating moderately volatile element depletion
in the Moon
10
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
compared with precursor rocks (sediments with more than 1 wt% H2 O), making them some
of the most anhydrous naturally occurring terrestrial materials. Although volatilization at high
temperature is a probable cause of this effect, K and Cd isotope studies to seek evidence for
evaporative stable isotope fractionations have yielded equivocal results [69,80,81]. In contrast,
some tektites are enriched in the heavy isotopes of Zn compared to upper continental crust
[74], and Zn abundance is negatively correlated with δ 66 Zn, indicating that isotopic fractionation
occurred during evaporation under the high-temperature conditions of tektite formation (above
2273 K; figure 4). A similar isotopic effect has been observed for Cu isotopes [82]. Tektites do
not follow the calculated Rayleigh distillation curves for Zn isotope fractionation, which has
been interpreted as evaporative isotopic fractionation in a diffusion-limited regime in which the
molten tektite is advecting by viscous coupling with air [74]. The tektite is not re-homogenized
instantly after evaporation, as it is in ideal Rayleigh distillation, and the magnitude of the isotopic
fractionation decreases. Therefore, both natural [74] and experimental [77] evidence points to
inefficient evaporative fractionation behaviour of moderately volatile elements. By analogy, the
restricted and consistent Zn isotope fractionations observed for lunar mare basalts (approx.
+1.3), compared with ideal Rayleigh behaviour for the Zn concentrations observed (δ 66 Zn at
least more than 5; figure 2) provides an important constraint on the type of process responsible
for the approximately 1 difference in δ 66 Zn between the Earth and the Moon.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
(a)
5
11
EL
a = 0.999
d 66 Zn (‰)
3
Shallowater
2
a = 0.9998
1
CI
a=1
0
EH5
0
(b) 2.0
100
EL3
200
Zn (ppm)
EH
300
400
Moon
( present Moon and early Earth)
1.5
Zn evaporation associated
with giant impact
d 66 Zn (‰)
1.0
0.5
range of
chondrites
(protoplanet
composition)
remixing of
‘chondritic’ Zn
during late
accretion
0 Earth
–0.5
0
50
100
150 200
Zn (ppm)
250
300
350
Figure 4. (a) Zinc isotope and abundance systematics in chondrites [35,73] and terrestrial tektites [74] versus estimates of lunar,
terrestrial and martian mantle composition (from [34]). Enstatite chondrites (EH, EL3, EH5, EL and Shallowater) are indicated on
the figure. Rayleigh distillation curves of α = 1, 0.9998, 0.999 and 0.993 are shown for comparison. (b) A possible scenario of
volatile element depletion associated with the giant impact, from a chondritic precursor, followed by stochastic late accretion
in higher proportions to the Earth than to the Moon. See text for details.
of volatile elements to the Earth and Mars by proportionally larger amounts of late accretion than
to the Moon [85]. This process would be unique to the Earth–Moon system and irrelevant to Mars
or small planetesimals, such as the angrite or eucrite parent bodies, which did not obviously suffer
such a high-energy and late-stage (probably more than 100 Ma after Solar System formation; [4–6])
event.
Second, planetary differentiation processes may have led to volatile and moderately volatile
elements being lost on smaller (cf. lower mass, lower escape velocity) airless bodies, such
as the Moon, more than on the Earth or Mars. These latter planets are considered to have
formed nascent atmospheres early in their formation histories which may subsequently have
been blown off, resulting in complex volatile retention trends [86]. An inevitable consequence
of differentiation and incompatible volatile element loss on smaller bodies would also have been
the formation of nascent atmospheres during outgassing. The likelihood of loss of atmospheres,
either through smaller parent-body size or from impact-induced atmospheric blow-off, may have
been a critical ingredient in the loss of volatile and moderately volatile elements in planets.
There are two possible locations where volatile and moderately volatile elements could have
been lost during planetary differentiation: (i) either during a large-scale magmatic event on the
Moon, such as a magma ocean; or (ii) from local eruption and evaporation processes acting
.........................................................
4
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
Earth
Moon
Mars
C chondrites
E chondrites
O chondrites
tektites
a = 0.993
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
K/U
Rb/Sr
d 66 Zn
100
chondrit
ic d 66 Zn
ratio
1
= 0.25 ±
0.1
0.05‰
6
r = 0.42
tic Rb/S
chondri
0.0001
0
2
Mars
Moon
eucrite
PB
0.001
Earth
0.01
8
4
6
escape velocity (km s–1)
10
12
Figure 5. Escape velocity versus abundance ratios of K/U and Rb/Sr, and 66 Zn/64 Zn ratios (as δ 66 Zn in per mil) for the Earth, Mars,
Moon, the eucrite parent body (eucrite PB) and the angrite parent body (on the y-axis; data reported in table 1). Exponential
regressions are shown for δ 66 Zn (R2 = 0.66; solid line), Rb/Sr (R2 = 0.55; dashed line) and K/U (R2 = 0.54; dashed line). Ratios
of Rb/Sr and K/U decrease, and δ 66 Zn shows heavy-isotope enrichment, with lower escape speeds. Chondritic values for Rb/Sr,
K/U and 66 Zn/64 Zn are reported along the exponential regression lines.
during basaltic eruption. Below, we review the evidence for these potential processes, in turn. The
evidence suggests that giant impact and late accretion—or localized eruption processes—require
special circumstances in order to explain the moderately volatile element compositions of
lunar mare basalts. Volatile outgassing during a magma ocean event offers much promise for
understanding planetary volatile depletion, if mechanisms for evaporative vapour loss can be
effectively identified.
(b) Volatile loss during the Earth–Moon forming giant impact?
Since the mid-1970s, the concept of a giant impact has become the favoured theory for the
formation of the Earth and the Moon [87], and remains consistent with the physical, geochemical
and age constraints [1,5,7,8]. Such a large-scale and high-energy event can theoretically explain
planetary volatile depletion in the Earth–Moon system owing to the potential for large thermal
and chemical gradients in the associated transient protoplanetary disc (figure 6). Owing to
the differences in moderately volatile elements retained between the Earth and the Moon, this
scenario would require either: (i) more effective retention of volatiles in the Earth, perhaps due
to shielding effects from a proto-atmosphere, a stronger gravity field than on the Moon, or
differences in mantle oxidation state; or (ii) replenishment of volatiles and moderately volatile
elements to the Earth after formation of the Moon (§5c).
The current lack of resolution in models for planetary oxidation events restricts quantitative
constraints [88], but oxidation state could effectively control the stability of mineral species that
retain moderately volatile elements. For example, a difference in oxidation state between the
proto-Earth and silicate liquid disc could explain the higher FeO contents in the present-day
Moon relative to the terrestrial mantle [16]. For Zn, there is evidence from chondrite meteorites
for the control of volatility through oxidation state. Enstatite chondrites, which are more reduced
than ordinary or carbonaceous chondrites, show a trend of increasing δ 66 Zn with decreasing Zn
abundance, and are consistent with evaporative fractionation [73]. With more reducing conditions
from carbonaceous to enstatite chondrites, Zn can be preferentially partitioned into sulfides (or
chlorides) relative to silicates [35,73]. On reduced bodies, such as the Moon, desulfidation and
dechlorination will also lead to pronounced Zn loss and Zn isotopic fractionation. Depending
.........................................................
c K/U = 74
chondriti
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
10
12
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Example locations for planetary
volatile loss
liquid/vapour
exchange
proto-Earth
redox control?
silicate liquid disc
magma
ocean
silicate vapour atmosphere
(b) magma ocean event
( post-giant impact lunar magma ocean; e.g. Wood et al. [115])
convecting
magma
ocean
transient
atmosphere
and loss of
volatiles to
space
vacuum
(c) surface eruptions and lava flow outgassing
(mare basalt lavas on the Moon; Sharpe et al. [32])
vacuum
condensation and
redeposition of
volatile elements
onto surface
eruption of magma
onto surface and
evaporation of
volatile elements
(d) pyroclastic glasses and impact gardening
(lunar volcanic glasses and regolith processes; Herzog et al. [71])
magmatic
fire fountain
vacuum
quench
micrometeorites
glass spheres
vapour/
condensation
reactions
vent
ion sputtering,
and condensation
gas expansion
and acceleration
Figure 6. Illustrations of some of the possible conditions under which evaporation/condensation of moderately volatile
elements can occur during planet formation and differentiation processes, with examples from the Moon: (a) a giant impact
or high-temperature Earth–Moon forming event; (b) a lunar magma ocean event and prior to extensive anorthositic crust
formation; (c) localized degassing of mare basalts on the lunar surface more than 1 Ga after initial lunar differentiation;
(d) during evaporation in a magmatic fire fountain (cf. pyroclastic glass beads) and from impact gardening processes subsequent
to lunar differentiation. Condensation of isotopically light vapours is expected as a consequence of events in (c) and (d), and
possibly (b).
.........................................................
(Earth–Moon forming giant impact; Pahlevan & Stevenson, [13])
hydrodynamic escape
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
(a) planet–moon forming disc
13
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
The elevated volatile contents of the Earth relative to the Moon can potentially be explained
by replenishment of these elements from late accretion after the giant impact, a model that has
been used to explain the inventory of volatiles in the present-day Earth [16,90–92]. A schematic
illustration of how this would work for Zn is shown in figure 4. Depletion of volatile elements,
such as Zn, occurs through evaporation during the giant impact and, subsequently, volatileelement-enriched ‘chondritic’ materials are accreted preferentially to the Earth, relative to the
Moon, by post-core formation late accretion. This scenario is consistent with observations of
chondrite-relative highly siderophile element (HSE: Os, Ir, Rh, Rh, Pt, Pd, Re, Au) abundances
in the mantles and mantle products of the Earth, the Moon and Mars [61,93]. In this scenario,
the approximately 20–40 times lower HSE abundances of the lunar mantle [18], compared with
the Earth [94], can be explained by stochastic delivery of proportionally massive impactors to the
Earth and the Moon [85].
Using estimated lunar (δ 66 Zn = +1.4; 2 ppm Zn) and terrestrial (δ 66 Zn = +0.27; 37 ppm
Zn) bulk Zn compositions (table 1), approximately 12% late accretion addition of Zn is required
to explain the differences observed between the Earth and the Moon by late accretion. These
calculations, using concentrations estimated partly from basaltic melt products from the Earth
and the Moon, assume identical partitioning of Zn from the mantle into the melt. Employing
higher Zn contents in the bulk Moon proportionally reduces the amount of late accretion mass
required (approx. 5.5% at 20 ppm Zn in the Moon), but cannot replicate the terrestrial Zn isotope
composition. The estimated amount of mass of impactor material from Zn (approx. 12%) contrasts
with that required for Pb (approx. 5%) [92] and HSE (0.5% or more) [93] abundances in the Earth’s
mantle and at first sight suggests that late accretion cannot be responsible for volatile distributions
between the Earth and the Moon.
Albarède et al. [92] have investigated a physical mechanism to explain the different apparent
proportions of elements such as Zn, Pb and the HSE added to the Earth from late accretion.
Models show that massive impactors involved in late accretion would have experienced partial
vaporization [97], leading to a physical process of separation of elements according to phases that
.........................................................
(c) A role for stochastic late accretion?
14
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
on speciation, this mechanism may also explain why some moderately volatile elements, like
K, do not appear to show strong isotopic fractionation between terrestrial and lunar materials.
Oxidation of the Earth’s mantle could have been triggered by sequestration of the impactor’s core
into the terrestrial core, or through the sudden increase in size of the Earth from accretion of the
Moon-forming impactor, increasing generation of Mg-rich perovskite at the base of a terrestrial
magma ocean, forcing disproportionation of ferrous iron into ferric iron plus metal [88,89]. On the
other hand, new high-pressure elemental partitioning experiments have shown that the Earth’s
core could have been formed under rather oxidizing conditions [62].
There are some problems with a model of the Earth retaining volatile elements preferentially
to the Moon during a giant impact event. Mars and differentiated asteroidal parent bodies
(howardite–eucrite–diogenites (HED) or angrite meteorites) do not appear to have suffered giant
impact events, yet the latter are more volatile-depleted than the Earth, while the former is not.
It is also unclear how effective retention or loss of volatiles would be between a proto-Earth and
silicate liquid disc. Pahlevan & Stevenson [13], in their analysis of this type of system, estimated
equilibration times for oxygen on time scales of hundreds to thousands of years; moderately
volatile element equilibration would presumably have been even more rapid [84]. It therefore
seems likely that moderately volatile elements could not be effectively fractionated between
different portions of the system following a giant impact, and that preferential replenishment
of volatile and moderately volatile elements would be required in order to have a volatile-poor
Moon relative to the Earth. However, it is also important to note that no studies have considered
the subsequent volatile evolution of a planetary disc, so volatile depletion in giant impact disc
environments remains an open question.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Table 1. Escape velocities, K/U, Rb/Sr and Zn isotope and mantle abundance estimates for some planets and asteroids.
Rb/Sr
0.30
δ 66 Zn()
0.28
±2σ
0.05
Venus
10.36
?
?
?
—
?
Mars
5.03
16.4
0.070
0.25
0.03
66
Mercury
4.25
?
?
?
—
?
Moon
2.38
3.0
0.009
1.4
0.1
2
[16,34,75]
EPB
0.36
3.7
0.003
6.2
—
—
[36]
APB
<0.1
0.6
0.001
—
—
—
estimated from [96]
CI chondrite
—
74.3
0.426
0.25
0.05
310
[17,35]
sources
[17,40]
..........................................................................................................................................................................................................
..........................................................................................................................................................................................................
[34]
..........................................................................................................................................................................................................
..........................................................................................................................................................................................................
..........................................................................................................................................................................................................
..........................................................................................................................................................................................................
..........................................................................................................................................................................................................
..........................................................................................................................................................................................................
they concentrate into. For example, partial vaporization of a large impactor would lead to all of the
lithophile and volatile elements being retained in the silicate mantle or resulting vapour cloud.
In contrast, the HSE and Fe would mostly pool and sequester into the Earth’s core, with some
fraction of the HSE also being retained in the vapour cloud that would subsequently rain down
on the Earth’s surface, imparting a ‘late accretion’ signature (figure 7). Albarède et al. [92] argued
that, during post-core formation accretion, partial vaporization could lead to the proportionally
larger amounts of late accretionary mass required for volatiles, moderately volatile elements and
Pb relative to the HSE. The current inventory of the HSE (approx. 0.5% late accretion addition)
relative to Pb (approx. 5%) or Zn (approx. 12%) constrains the hypothetical vaporization ratios
at 1 : 10 to 1 : 25. Furthermore, Albarède et al. [92] argued that this form of model would not
violate W isotope constraints, because much of the impactor W would be segregated into the
terrestrial core, along with the HSE. Models of addition of CI chondrite material to a presentday Moon composition for Zn show that simple mixing does not yield a present-day terrestrial
composition (figure 7b). However, a 1 : 10 mixing model, incorporating accretion and partial
vaporization of massive impactors, models the present-day composition of the terrestrial mantle
quite well.
A model of volatile element loss during giant impact, followed by stochastic late accretion of
volatiles to the Earth and the Moon, has significant deficiencies. It is not clear how reasonable 5
to 10% or more late accretion for Pb and volatile elements is after the giant impact, given that the
Earth–Moon forming event is widely considered the last major accretion event [2]. Nor is it known
if the giant impact was a global ‘clearing house’ event for the HSE in the Earth. Understanding
if both the Earth and the Moon were systematically depleted in volatile and highly siderophile
elements during their formation is a crucial constraint for assessing whether a late accretion model
for volatile enrichment can work. Scenarios that use a high ice/rock ratio (e.g. comets) could help
to reduce the mass balance problem, but cometary material does not seem consistent with volatile
element isotopic data [91]. The similar estimated HSE abundances of the Earth and Mars [61] also
require explanation, given the large differences in timing of differentiation of the two bodies [99].
Recent work to constrain the timing of late accretion to the HED parent body(ies) has shown rapid
addition, within 2–3 Ma, and that late accretion is simply the result of cessation of core formation
on differentiated planetary bodies [98]. Thus, why Mars, whose core probably formed and ceased
growth more rapidly than for the Earth [5,99], would have similar estimated HSE abundances in
its mantle to the Earth, remains an outstanding science question. Given the discussion above, a
giant impact and/or late accretion scenario to explain moderately volatile element distributions
in the Earth and the Moon remains unconstrained, and alternative explanations need
to be sought.
.........................................................
(K/U)/1000
11.8
Zn estimated
mantle
conc (ppm)
37
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
planet/asteroid
Earth
escape
velocity
(km s−1 ) [95]
11.19
15
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
metal
core
(b)
2.5
HED
d 66 Zn (‰)
2.0
1.5
1.0
Moon
1%
0.5
1%
0
5%
20%
50%
X CI
100%
–0.5
0.01
1
0.0001 0.001
0.1
est. mantle HSE abundance/CI chondrite
10
Figure 7. Impact vaporization as a process to explain differential proportions of volatile and siderophile late accretion.
(a) Schematic diagram of an impactor striking a planet, partially vaporizing and forming a globally shrouding cloud enriched
in all elements with concomitant differentiation of the melted fraction of the impactor into a metal fraction (removing the
HSE) and a silicate melt fraction (containing lithophile elements). A similar model has been proposed to explain the different
proportions of volatiles and siderophiles apparently added to the Earth after differentiation and metal–silicate equilibration
[92]. (b) Zinc isotope compositions versus total HSE abundances for the Moon, the Earth, Mars, CI chondrites (data from
[34,61]) and HED meteorites (data from [36,98] and J.M.D.D, 2010, unpublished data). Dashed mixing lines are for a eucrite
composition (δ 66 Zn = +6; HSE abundance = 0.0001 × CI chondrite) and a lunar composition (δ 66 Zn = +1.4; HSE
abundance = 0.0001 × CI chondrite) with CI chondrite. The solid grey line is for mixing with a lunar composition with CI
chondrite, but assuming partitioning of HSE from volatiles like Zn using an impact vaporization scenario as depicted in (a) and
assuming a 1 : 10 ratio of vaporized (step 3) to segregated (step 2) HSE.
(d) Local eruption processes?
The large range in δ 37 Cl ratios measured for lunar rocks possibly indicates local eruption
processes causing evaporation and volatile loss [32] (figure 6c,d). Independent evidence also exists
that local eruption processes can lead to evaporation of moderately volatile elements and so
volatile depletion in the Moon could reflect late-stage magmatic processes. For example, elevated
volatile and siderophile abundances are surface-correlated in some lunar pyroclastic glass beads
[63,100–106]. These observations can be reconciled by deposition of volatiles (e.g. Zn, S, Cl, F,
Pb) occurring upon quenching of the beads within the associated vapour jet that expelled them,
reduction of FeO to Fe-rich metal phases at the surface of the glass beads [107] and fractionation
of Zn isotopes [71].
Localized eruption processes and impact gardening are important for partial condensation
of isotopically light S, Cl, K and Zn on the lunar surface. For example, some mare basalts
have δ 66 Zn < 0.5 (e.g. 10017 and 12018) [34,75], and have higher Zn abundances (average =
4.2 ppm) than the majority (80%) of mare basalts (average = 1.7 ppm). Previously, light Zn
.........................................................
silicate
mantle
16
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
(1) high-velocity impact
(2) melting, vaporization
and segregation of some
metal directly into planetary
core
(3) vaporized
matter, including
fraction of iron
and HSE, form
transient cloud;
precipitation
onto surface leads
to late accretion
(a)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
The distinct geochemical traits of mare basalts and other lunar rocks underpin the concept of a
lunar magma ocean to account for the plagioclase (Eu)-enriched crust overlying a plagioclase
(Eu)-depleted mare basalt source region [114,115]. This form of ‘planet-scale’ melting and
differentiation is probably an inevitable consequence of planetary differentiation, even on small
bodies such as angrite or eucrite meteorites [116]. The chronology of events is poorly constrained
[6], but the composition of mare basalts is consistent with derivation from mantle sources with
different proportions of ilmenite [113]. The lunar magma ocean hypothesis predicts mineralogical
variations in lunar cumulate reservoirs and in the source regions of low- and high-Ti mare
basalts. Such features can also explain the significant stable O–Mg–Fe isotope variations measured
for lunar basalts [8,117,118]. These studies have shown that, from existing knowledge of hightemperature mineral–melt partition coefficients, magmas generated from sources with variable
modal mineralogies and/or mineral-source compositions would have fractionated stable isotope
compositions.
Magma ocean differentiation offers an alternative explanation for the volatile element
distribution and isotopic variations observed for the Moon relative to the Earth, Mars and other
differentiated bodies (figure 6b). Outgassing of a magma ocean would lead to evaporation of
volatile elements and the systematic and progressive heavy-isotope enrichment in the condensate
relative to the vapour. An important aspect of outgassing of a magma ocean is that perfect
volatility behaviour would not be expected, for three principal reasons. First, the duration of a
lunar magma ocean phase is important [116], prior to generation of a permanent crust, which
would presumably impede volatile element loss. It is possible that the shorter duration of a
.........................................................
(e) Lunar magma ocean differentiation
17
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
isotope enrichment in pyroclastic glass beads has been attributed to Zn evaporation during
fire fountaining, with condensation of isotopically light Zn on the surfaces of the ejected
beads [70,75], as also proposed for 32 S enrichments [104], and consistent with elevated Zn
concentrations on their exteriors [63,71]. Vapour loss should increase the δ 66 Zn value of the
residual magma, and this process can be modelled assuming exhalative evaporation of Zn as
ZnCl2 on the Moon [63], versus for the Earth, where Zn is evaporated as Zn2+ [108], with
essentially no fractionation between vapour and melt (α = 1). Paniello et al. [34] showed that
Rayleigh fractionation calculations can explain isotopically light vapour Zn isotope compositions
conforming to those measured in 12018, 10017 and pyroclastic glass beads through evaporation
and re-deposition of ZnCl2 from lunar melts (e.g. figure 2). A similar process would take place
during impact gardening, with sputtering and re-deposition of isotopically light Zn on the lunar
surface. We therefore exclude mare basalts with isotopically light Zn values because they are
likely to reflect post-crystallization contamination at the lunar surface.
Both high- and low-Ti mare basalts with low Zn concentrations have δ 66 Zn of 1.4 ± 0.3
[34,36,75]. The identical Zn isotopic compositions of low- and high-Ti mare basalts indicate
identical processes acting on both basalt types, despite conspicuously differing crystallization
and cooling histories and lunar mantle sources. Cooling and crystallization rates for mare basalts
vary dramatically [109–111], with evidence for quenched and coarse-grained basalts at individual
Apollo sites (see fig. 2 in [65], for an example at the Apollo 15 site). Additionally, low-Ti mare
basalts are the predominant form of basalt exposed on the lunar surface (approx. 90%; [112]),
and their prevalence over high-Ti mare basalts is consistent with magma ocean crystallization
models that indicate that more than 78% of the lunar mantle would be dominated by olivine
and orthopyroxene, whereas ilmenite-rich cumulates formed after 95% crystallization [113]. If
localized eruption processes were the cause of Zn isotopic variations in lunar mare basalts,
correlations might be expected between crystallization history or mantle source composition, yet
the δ 66 Zn values seem invariant to these parameters. Furthermore, our reanalysis of published S,
Cl, K and Zn isotope data (cf. figure 3) suggests that a large-scale global differentiation process,
rather than local eruption processes, is a far more likely cause of the systematic depletion of these
elements in the sources of mare basalts.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
(c) cessation of evaporative fractionation due to
formation of lunar anorthosite-rich crust;
formation of ilmenite and KREEP
crust
(d ) 4.5–3 Ga magmatism following
cumulate overturn of ilmenite
low-Ti
(d 66Zn = +1.4‰)
high-Ti
(d 66Zn = +1.4‰)
ilmenite and
KREEP layers
LLMOC
LLMOC
ULMOC
(d 66Zn > +1.4‰?)
LPG
(d 66Zn < 0.5‰)
LLMOC
HMS
(d 66Zn > 1.5‰?)
Figure 8. Conceptual model of volatile outgassing during a lunar magma ocean, and the evolution of Zn isotopes.
(a) Immediately after lunar accretion, an interior forms that does not witness interaction with the lunar surface, leading to
a volatile-rich (Earth-like?) primitive composition. Convection and inefficient outgassing of the overlying magma ocean lead
to formation of less (b) and more (c) volatile-depleted lower and upper lunar magma ocean cumulates (LLMOC, ULMOC),
respectively. Outgassing and volatile depletion are eventually terminated by formation of the plagioclase-rich flotation crust.
Continued fractional crystallization of the lunar magma ocean leads to formation of ilmenite and a KREEP-rich layers with the
most volatile-depleted signatures. (d) Cumulate overturn [119] leads to formation of high-Ti mare basalt and high-Ti lunar
pyroclastic glass (LPG) source regions. Partial melting of volatile-depleted (LLMOC and ULMOC) cumulates to form mare basalts
and of primitive lunar interior to form LPG can explain the volatile abundance differences estimated for their source regions
[22,34]. If high-Mg suite rocks form from LLMOC, a testable hypothesis is that they should have isotopically heavier δ 66 Zn.
Conversely, ferroan anorthosites should have δ 66 Zn ranging from positive to negative values in response to volatile loss in their
source region and later condensation (cf. figure 6d).
lunar magma ocean prior to crust formation, relative to more prolonged crustal growth above
a magma ocean on smaller planetesimals, can partly explain the greater enrichment of δ 66 Zn in
some unbrecciated eucrites [36]. Second, for loss of volatiles, atoms would be required to exceed
the escape velocity, or would need to be lost during a process such as atmospheric blow-off, and
this process may inhibit ideal Rayleigh distillation of moderately volatile stable isotopes. On small
bodies (less than 500 km diameter), such as differentiated planetesimals, this would be a trivial
issue, but the escape velocity of the present-day Moon is non-trivial (table 1) and an additional
.........................................................
LLMOC
(d 66Zn = +1.4‰)
18
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
(a) accretion of Moon, formation of lunar interior (b) continued lunar magma ocean overturn and
and magma ocean convection; diffusion-limited
formation of lower lunar magma ocean cumulates
evaporation at the surface
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Available volatile element stable isotope variations indicate that a global-scale evaporation
event was the most likely cause of volatile depletion in the Moon, and by extension in other
fully differentiated and airless bodies. Despite this evidence, a number of fundamental science
problems remain to be resolved in order to improve understanding of volatile element depletion:
(1) Additional high-quality S, Cl, K and Zn isotope data are needed for the Moon, the Earth
and other planetary bodies, while elements with different volatilities, atomic masses
or behaviours (e.g. Cu, Rb, Ga, Cd) would place additional constraints on some of the
proposed models for planetary volatile element depletion.
(2) Analysis of both bulk rock samples and mineral phases will be important for
understanding volatile element depletion and the siting of these elements within mantlederived magmatic rocks from differentiated planetary bodies. Comparison of volatile
and moderately volatile isotopic data from mineral phases with bulk rocks needs to be
approached with caution, particularly if the mineral phase is a late-stage crystallization
product (e.g. Cl in apatite).
(3) Further study of surface-correlated effects in the pyroclastic glass beads and careful
assessment of surface coatings on the exteriors of lunar rocks are required in order to
more fully understand localized eruptive processes. Apollo samples were all collected
from the lunar regolith, so isotopically light condensation of moderately volatile elements
from eruption and impact gardening could hamper resolution of magmatic and mantle
signatures.
(4) The relationship between planetary escape velocity and volatile elements may be
important in assessing degassing and volatile loss in a magma ocean. Current evidence
suggests that the lunar pyroclastic glass eruptions led to condensation of isotopically
light Zn [71], implying that lunar escape velocities were not effectively exceeded by this
process. Similarly, if escape velocities were not exceeded during the earliest magma ocean
phases, then an ephemeral atmosphere in a low-gravity environment followed by erosion
of this atmosphere would be required. Rapid erosion of an atmosphere rich in isotopically
light vapour may offer an explanation for the loss of volatiles from a lunar magma ocean.
.........................................................
6. Conclusion and outlook
19
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
mechanism (e.g. vapour partial pressure above the magma ocean and formation and subsequent
erosion of a transient atmosphere) would become critical. Third, evaporation of volatile elements
would be expected to follow non-ideal Rayleigh fractionation. This is because only the upper
portions of a magma ocean would be exposed to vacuum, and convection in the magma ocean
would be a rate-limiting step in the efficiency of volatile loss in a lunar magma ocean.
There are a number of consequences of non-ideal Rayleigh fractionation and volatile loss in a
lunar magma ocean that appear to fit most closely with the available volatile element data. First,
ideal Rayleigh distillation behaviour would not be predicted, and this is certainly not observed
(figure 2). Second, early-formed cumulates would be expected to record less volatile loss than
later-formed cumulates, at least prior to the formation of the lunar anorthositic crust. Some deep,
early-formed cumulates could retain a higher volatile inventory and more chondritic volatile
stable isotope ratios and such sources could potentially fit with the deep mode of origin of lunar
pyroclastic glass beads [119], their lower μ-values compared with mare basalts [27] and their high
apparent water contents [22]. A conceptual model of magma ocean outgassing and volatile loss
can explain some of the observed values for moderately volatile element abundances and stable
isotope ratios (figure 8). Further work is required, however, to test the model (e.g. Zn isotope
analyses of high-Mg suite rocks, which should be isotopically heavy) and to assess mechanisms
that would explain the mass balance of light isotope depletion of moderately volatile elements
in the Moon (e.g. transient atmospheric erosion or blow-off; storage of light isotopes of volatile
elements in the lunar crust or regolith).
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
Acknowledgements. The authors thank the organizers of the Royal Society Origin of the Moon conference,
D. Stevenson and A. Halliday, for the invitation to participate. We thank R. W. Carlson and T. Elliott for
constructive review comments, A. Halliday for editorial handling, and F. Albarède, J. Dhaliwal, B. Fegley, T.
Grove, S. Mukhopadhyay, Z. Sharp and J. Wasson for helpful suggestions.
Funding statement. This work was supported by NASA LASER (NNX11AG34G) and Cosmochemistry
(NNX12AH75G) grants to J.M.D.D. and by a Chair of Excellence from the Sorbonne Paris Cité and NASA
Cosmochemistry and Exobiology grants to F.M.
References
1. Canup RM. 2012 Forming a Moon with an Earth-like composition via a giant impact. Science
338, 1052–1055. (doi:10.1126/science.1226073)
2. Cuk M, Stewart ST. 2012 Making the Moon from a fast-spinning Earth: a giant impact
followed by resonant despinning. Science 338, 1047–1052. (doi:10.1126/science.1225542)
3. Reufer A, Meier MMM, Benz W, Wieler R. 2012 A hit-and-run giant impact scenario. Icarus
221, 296–299. (doi:10.1016/j.icarus.2012.07.021)
4. Touboul M, Kleine T, Bourdon B, Palme H, Wieler R. 2007 Late formation and prolonged
differentiation of the Moon inferred from W isotopes in lunar metals. Nature 450, 1206–1209.
(doi:10.1038/nature06428)
5. Halliday AN. 2008 A young Moon-forming giant impact at 70–110 million years
accompanied by late-stage mixing, core formation and degassing of the Earth. Phil. Trans.
R. Soc. A 366, 4163–4181. (doi:10.1098/rsta.2008.0209)
6. Borg LE, Connelly JN, Boyet M, Carlson RW. 2011 Chronological evidence that the Moon
is either young or did not have a global magma ocean. Nature 477, 70–72. (doi:10.1038/
nature10328)
7. Wiechert U, Halliday AN, Lee DC, Snyder GA, Taylor LA, Rumble D. 2001 Oxygen
isotopes and the Moon-forming giant impact. Science 294, 345–348. (doi:10.1126/
science.1063037)
8. Spicuzza MJ, Day JMD, Taylor LA, Valley JW. 2007 Oxygen isotope constraints on the
origin and differentiation of the Moon. Earth Planet. Sci. Lett. 253, 254–265. (doi:10.1016/j.
epsl.2006.10.030)
.........................................................
Volatile depletion in planets remains a major challenge for understanding planet formation
processes and the cosmochemistry of the earliest Solar System. Mare basalts from the Moon
come from mantle source regions that were volatile-depleted relative to the source mantle of
terrestrial or martian basalts and require a planetary volatile depletion separate from nebular
volatile depletions observed in chondrites. It remains to be resolved what the exact process(es)
of volatile depletion was (were) for the Moon, but the challenges associated with addressing this
problem remain some of the most important in planetary science.
20
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
(5) Refined experimental data will aid in understanding diffusion-limited regimes relevant
to volatile loss [77]. Controversy surrounds the expected behaviour of volatiles into
the hard vacuum at the lunar surface [120], but current evidence suggests inefficient
isotopic fractionation of volatile stable isotopes. Further experimental and theoretical
work is also required to understand the behaviour of moderately volatile elements with
varying speciation. This form of approach is crucial for understanding the role of putative
vapour–liquid Moon-forming discs after a giant impact scenario [84].
(6) More direct comparison of moderately volatile and highly volatile elements and
associated species, such as H2 O and OH, are desirable, given the competing evidence for
a relatively ‘wet’ [22,23] or ‘dry’ Moon [32,34]. Did portions of the Moon retain a higher
volatile element inventory in a magma ocean (figure 8), or were the source materials
that formed the Moon ubiquitously volatile-depleted during a giant impact? If so, can
heterogeneous addition of late accretion materials account for ‘wet’ regions of the Moon
as sampled by lunar pyroclastic glasses?
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
21
.........................................................
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
9. Hallis LJ, Anand M, Greenwood RC, Miller MF, Franchi IA, Russell SS. 2010 The oxygen
isotope composition, petrology and geochemistry of mare basalts: evidence for largescale compositional variation in the lunar mantle. Geochim. Cosmochim. Acta 74, 6885–6899.
(doi:10.1016/j.gca.2010.09.023)
10. Zhang J, Dauphas N, Davis AM, Leya I, Fedkin A. 2012 The proto-Earth as a significant
source of lunar material. Nat. Geosci. 5, 251–255. (doi:10.1038/ngeo1429)
11. Armytage RMG, Georg RB, Williams HM, Halliday AN. 2012 Silicon isotopes in lunar rocks:
implications for the Moon’s formation and the early history of the Earth. Geochim. Cosmochim.
Acta 77, 504–514. (doi:10.1016/j.gca.2011.10.032)
12. Fitoussi C, Bourdon B. 2012 Silicon isotope evidence against an enstatite chondrite Earth.
Science 335, 1477–1480. (doi:10.1126/science.1219509)
13. Pahlevan K, Stevenson DJ. 2007 Equilibration in the aftermath of the lunar-forming giant
impact. Earth Planet. Sci. Lett. 262, 438–449. (doi:10.1016/j.epsl.2007.07.055)
14. Lodders K. 2003 Solar System abundances and condensation temperatures of the elements.
Astrophys. J. 591, 1220–1247. (doi:10.1086/375492)
15. Dreibus G, Wanke H. 1980 The bulk composition of the eucrite parent asteroid and its
bearing on planetary evolution. Z. Naturforsch. 35a, 204–216. See http://zfn.mpdl.mpg.de/
data/Reihe_A/35/ZNA-1980-35a-0204.pdf.
16. O’Neill HStC. 1991 The origin of the Moon and the early history of the Earth—a
chemical model. Part 1: The Moon. Geochim. Cosmochim. Acta 55, 1135–1157. (doi:10.1016/
0016-7037(91)90168-5)
17. McDonough WF, Sun S-S. 1995 The composition of the Earth. Chem. Geol. 120, 223–254.
(doi:10.1016/0009-2541(94)00140-4)
18. Day JMD, Pearson DG, Taylor LA. 2007 Highly siderophile element constraints on
accretion and differentiation of the Earth–Moon system. Science 315, 217–219. (doi:10.1126/
science.1133355)
19. Davis A. 2006 Volatile evolution and loss. In Meteorites and the early solar system II (eds DS
Lauretta, HY McSween), pp. 295–307. Tucson, AZ: University of Arizona Press.
20. Wolf R, Anders E. 1980 Moon and Earth: compositional differences inferred from
siderophiles, volatiles, and alkalis in basalts. Geochim. Cosmochim. Acta 44, 2111–2124.
(doi:10.1016/0016-7037(80)90208-2)
21. Jones JH, Palme H. 2000 Geochemical constraints on the origin of the Earth and Moon. In
Origin of the Earth and Moon (eds RM Canup, K Righter), pp. 197–216. Tucson, AZ: University
of Arizona Press.
22. Hauri EH, Weinreich T, Saal A, Rutherford M, Van Orman JA. 2011 High pre-eruptive
water contents preserved in lunar melt inclusions. Science 333, 213–215. (doi:10.1126/science.
1204626)
23. McCubbin FM, Steele A, Hauri EH, Nekvasil H, Yamashita S, Hemley RJ. 2010
Nominally hydrous magmatism on the Moon. Proc. Natl Acad. Sci. USA 107, 11 223–11 228.
(doi:10.1073/pnas.1006677107)
24. Halliday AN, Porcelli D. 2001 In search of lost planets—the paleocosmochemistry of the
inner Solar System. Earth Planet. Sci. Lett. 192, 545–559. (doi:10.1016/S0012-821X(01)00479-4)
25. Carlson RW, Lugmair GW. 1988 The age of ferroan anorthosite 60025: oldest crust on a young
Moon. Earth Planet. Sci. Lett. 90, 119–130. (doi:10.1016/0012-821X(88)90095-7)
26. McCulloch MT. 1994 Primitive 87 Sr/86 Sr from an Archean barite and conjecture on the
Earth’s age and origin. Earth Planet. Sci. Lett. 126, 1–13. (doi:10.1016/0012-821X(94)90238-0)
27. Tatsumoto M, Premo WR, Unruh DM. 1987 Origin of lead from green glass of Apollo
15426: a search from primitive lunar lead. J. Geophys. Res. 92, E361–E371. (doi:10.1029/
JB092iB04p0E361)
28. Papanastassiou DA, Wasserburg GJ. 1978 Strontium isotopic anomalies in the Allende
meteorite. Geophys. Res. Lett. 5, 595–598. (doi:10.1029/GL005i007p00595)
29. Moynier F, Day JMD, Okui W, Yokoyama T, Bouvier A, Walker RJ, Podosek FA. 2012
Planetary-scale strontium isotopic heterogeneity and the age of volatile depletion of early
Solar System materials. Astrophys. J. 758, 45. (doi:10.1088/0004-637X/758/1/45)
30. Hans U, Kleine T, Bourdon B. 2013 Rb–Sr chronology of volatile depletion in differentiated
protoplanets: BABI, ADOR and ALL revisited. Earth Planet. Sci. Lett. 374, 204–214. (doi:10.
1016/j.epsl.2013.05.029)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
22
.........................................................
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
31. Humayun M, Clayton RN. 1995 Precise determination of the isotopic composition of
potassium: application to terrestrial rocks and lunar soils. Geochim. Cosmochim. Acta 59,
2115–2130. (doi:10.1016/0016-7037(95)00131-X)
32. Sharp ZD, Shearer CK, McKeegan KD, Barnes JD, Wang YQ. 2010 The chlorine isotope
composition of the Moon and implications for an anhydrous mantle. Science 329, 1050–1053.
(doi:10.1126/science.1192606)
33. Sharp ZD, Barnes JD, Fischer TP, Halick M. 2010 An experimental determination of
chlorine isotope fractionation in acid systems and application to volcanic fumeroles. Geochim.
Cosmochim. Acta 74, 264–273. (doi:10.1016/j.gca.2009.09.032)
34. Paniello RC, Day JMD, Moynier F. 2012 Zinc isotopic evidence for the origin of the Moon.
Nature 490, 376–379. (doi:10.1038/nature11507)
35. Luck J-M, Ben Othman D, Albarède F. 2005 Zn and Cu isotopic variations in chondrites
and iron meteorites: early solar nebula reservoirs and parent-body processes. Geochim.
Cosmochim. Acta 69, 5351–5363. (doi:10.1016/j.gca.2005.06.018)
36. Paniello RC, Moynier F, Beck P, Barrat J-A, Podosek F, Pichat S. 2012 Zinc isotopes in HEDs:
clues to the formation of 4-Vesta and the unique composition of Pecora Escarpment 82502.
Geochim. Cosmochim. Acta 86, 76–87. (doi:10.1016/j.gca.2012.01.045)
37. Chen H, Savage PS, Teng F-Z, Helz RT, Moynier F. 2013 Zinc isotopic fractionation during
magmatic differentiation and the isotopic composition of the bulk Earth. Earth Planet. Sci.
Lett. 369–370, 34–42. (doi:10.1016/j.epsl.2013.02.037)
38. Anders E, Grevesse N. 1989 Abundances of the elements: meteoritic and solar. Geochim.
Cosmochim. Acta 53, 197–214. (doi:10.1016/0016-7037(89)90286-X)
39. Palme H, Jones A. 2003 Solar system abundances of the elements. In Treatise on geochemistry
(eds HD Holland, KK Turekian), vol. 1, Meteorites, comets, and planets (ed. AM Davis), pp.
41–61. Amsterdam, The Netherlands: Elsevier. (doi:10.1016/B0-08-043751-6/01060-4)
40. Barrat JA, Zanda B, Moynier F, Bollinger C, Liorzou C, Bayon G. 2012 Geochemistry of CI
chondrites: major and trace elements, and Cu and Zn isotopes. Geochim. Cosmochim. Acta 85,
79–92. (doi:10.1016/j.gca.2011.12.011)
41. Palme H, Larimer JW, Lipschutz ME. 1988 Moderately volatile elements. In Meteorites and the
early Solar System (eds JF Kerridge, MS Matthews), pp. 436–460. Tucson, AZ: University of
Arizona Press.
42. Larimer JW. 1988 The cosmochemical classification of the elements. In Meteorites and the early
Solar System (eds JF Kerridge, MS Matthews), pp. 375–389. Tucson, AZ: University of Arizona
Press.
43. Urey HC. 1947 The thermodynamic properties of isotopic substances. J. Chem. Soc. (London),
1947, 562–581. (doi:10.1039/JR9470000562)
44. Bigeleisen J, Mayer MG. 1947 Calculation of equilibrium constants for isotopic exchange
reactions. J. Chem. Phys. 15, 261–267. (doi:10.1063/1.1746492)
45. Schauble EA. 2004 Applying stable isotope fractionation theory to new systems. Rev. Mineral.
Geochem. 55, 65–111. (doi:10.2138/gsrmg.55.1.65)
46. Maréchal C, Telouk P, Albarède F. 1999 Precise analysis of copper and zinc isotopic
compositions by plasma-source mass spectrometry. Chem. Geol. 156, 251–273. (doi:10.1016/
S0009-2541(98)00191-0)
47. Humayun M, Cassen P. 2000 Processes determining the volatile abundances of the meteorites
and terrestrial planets. In Origin of the Earth and Moon (eds RM Canup, K Righter), pp. 3–23.
Tucson, AZ: University of Arizona Press.
48. O’Neil JR. 1986 Theoretical and experimental aspects of isotopic fractionation. Rev. Mineral.
16, 1–40.
49. Gibson EK, Chang S, Lennon K, Moore GW, Pearce GW. 1975 Sulfur abundances and
distributions in mare basalts and their source magmas. In Proc. 6th Lunar Science Conf.,
Houston, TX, 17–21 March, pp. 1287–1301. New York, NY: Pergamon.
50. Kerridge JF, Kaplan IR, Petrowski C. 1975 Evidence for meteoritic sulfur in the lunar regolith.
In Proc. 6th Lunar Science Conf., Houston, TX, 17–21 March, pp. 2151–2162. New York, NY:
Pergamon.
51. Clayton RN, Mayeda TK, Hurd JM. 1974 Loss of oxygen, silicon, sulfur, and potassium from
the lunar regolith. In Proc. 5th Lunar Science Conf., Houston, TX, 18–22 March, pp. 1801–1809.
New York, NY: Pergamon.
52. Wing B, Farquhar J. In press. Sulfur isotope homogeneity of lunar mare basalts. Geochim.
Cosmochim. Acta.
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
23
.........................................................
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
53. Franz HB et al. 2014 Isotopic links between atmospheric chemistry and the deep sulphur
cycle on Mars. Nature 508, 364–368. (doi:10.1038/nature13175)
54. Labidi J, Cartigny P, Birck J-L, Assayag N, Bourrand JJ. 2012 Determination of multiple
sulfur isotopes in glasses: a reappraisal of the MORB δ 34 S. Chem. Geol. 334, 189–198.
(doi:10.1016/j.chemgeo.2012.10.028)
55. Labidi J, Cartigny P, Moriera M. 2013 Non-chondritic sulphur isotope composition of the
terrestrial mantle. Nature 501, 208–211. (doi:10.1038/nature12490)
56. Richter FM, Janney PE, Mendybaev RA, Davis AM, Wadhwa M. 2007 Elemental and isotopic
fractionation of type B CAI-like liquids by evaporation. Geochim. Cosmochim. Acta 71, 5544–
5564. (doi:10.1016/j.gca.2007.09.005)
57. Sharp ZD, Barnes JD, Brearley AJ, Chaussidon M, Fisher TP, Kamenetsky VS. 2007 Chlorine
isotope homogeneity of the mantle, crust and carbonaceous chondrites. Nature 446, 1062–
1065. (doi:10.1038/nature05748)
58. Sharp ZD, Shearer CK, Agee CB, McKeegan KD. 2011 The chlorine isotope composition of
Mars. In Proc. 42nd Lunar and Planetary Science Conf., The Woodlands, TX, 7–11 March, Abstract
2534. See http:www.lpi.usra.edu/meetings/lpsc2011/pdf/2534.pdf.
59. Bonifacie M, Jendrzejewski N, Agrinier P, Humler E, Coleman M, Javoy M. 2008 The
chlorine isotope composition of Earth’s mantle. Science 319, 1518–1520. (doi:10.1126/
science.1150988)
60. Bombardieri DJ, Norman MD, Kamenetsky VS, Danyushevsky LV. 2005 Major element and
primary sulfur concentrations in Apollo 12 mare basalts: the view from melt inclusions.
Meteorit. Planet. Sci. 40, 679–693. (doi:10.1111/j.1945-5100.2005.tb00973.x)
61. Day JMD. 2013 Hotspot volcanism and highly siderophile elements. Chem. Geol. 341, 50–74.
(doi:10.1016/j.chemgeo.2012.12.010)
62. Siebert J, Badro J, Antonangeli D, Ryerson FJ. 2013 Terrestrial accretion under oxidizing
conditions. Science 339, 1194–1197. (doi:10.1126/science.1227923)
63. Chou C-L, Boynton WV, Sundberg LL, Wasson JT. 1975 Volatiles on the surface of Apollo 15
green glass and trace-element distributions among Apollo 15 soils. In Proc. 6th Lunar Science
Conf., Houston, TX, 17–21 March, pp. 1701–1727. New York, NY: Pergamon.
64. Day JMD, Taylor LA, Floss C, Patchen AD, Schnare DW, Pearson DG. 2006
Comparative petrology, geochemistry and petrogenesis of evolved, low-Ti lunar mare basalt
meteorites from the La Paz icefield, Antarctica. Geochim. Cosmochim. Acta 70, 1581–1600.
(doi:10.1016/j.gca.2005.11.015)
65. Schnare DW, Day JMD, Norman MD, Liu Y, Taylor LA. 2008 Laser-ablation ICP-MS study
of Apollo 15 low-titanium olivine-normative and quartz-normative mare basalts. Geochim.
Cosmochim. Acta 72, 2556–2572. (doi:10.1016/j.gca.2008.02.021)
66. Liu Y, Floss C, Day JMD, Hill E, Taylor LA. 2009 Petrogenesis of lunar mare basalt meteorite
MIL 05035. Meteorit. Planet. Sci. 44, 261–284. (doi:10.1111/j.1945-5100.2009.tb00733.x)
67. Boyce JW, Guan Y, Treiman AH, Greenwood JP, Eiler JM, Ma C. 2013 Volatile components
in the Moon: abundances and isotope ratios of Cl and H in lunar apatites. In Proc. 44th
Lunar and Planetary Science Conf., The Woodlands, TX, 18–22 March, Abstract 2851. See
http://www.lpi.usra.edu/meetings/lpsc2013/pdf/2851.pdf.
68. Humayun M, Clayton RN. 1995 Potassium isotope geochemistry: genetic implications
of volatile element depletion. Geochim. Cosmochim. Acta 59, 2115–2130. (doi:10.1016/
0016-7037(95)00131-X)
69. Humayun M, Koeberl C. 2004 Potassium isotopic composition of Australasian tektites.
Meteorit. Planet. Sci. 39, 1509–1516. (doi:10.1111/j.1945-5100.2004.tb00125.x)
70. Moynier F, Albarède F, Herzog G. 2006 Isotopic composition of zinc, copper, and iron in lunar
samples. Geochim. Cosmochim. Acta 70, 6103–6117. (doi:10.1016/j.gca.2006.02.030)
71. Herzog GF, Moynier F, Albarède F, Berezhnoy AA. 2009 Isotopic and elemental abundances
of copper and zinc in lunar samples, Zagami, Pele’s hairs, and a terrestrial basalt. Geochim.
Cosmochim. Acta 73, 5884–5904. (doi:10.1016/j.gca.2009.05.067)
72. Moynier F, Dauphas N, Podosek FA. 2009 Search for 70 Zn anomalies in meteorites. Astrophys.
J. Lett. 700, L92–L95. (doi:10.1088/0004-637X/700/2/L92)
73. Moynier F, Paniello RC, Gounelle M, Albarède F, Beck P, Podosek F, Zanda B. 2011 Nature
of volatile depletion and genetic relationships in enstatite chondrites and aubrites inferred
from Zn isotopes. Geochim. Cosmochim. Acta 75, 297–307. (doi:10.1016/j.gca.2010.09.022)
74. Moynier F, Beck P, Jourdan F, Yin Q-Z, Reimold U, Koeberl C. 2009 Isotopic fractionation of
zinc in tektites. Earth Planet. Sci. Lett. 277, 482–489. (doi:10.1016/j.epsl.2008.11.020)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
24
.........................................................
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
75. Kato C, Valdes M, Dhaliwal J, Day JMD, Moynier F. 2013 Origin of isotopically light Zn in
lunar samples through vaporization and the Zn isotope composition of the lunar mantle. In
American Geophysical Union Fall Meeting 2013, Abstract V33D-2788.
76. Esat TM. 1988 Physicochemical isotope anomalies. Geochim. Cosmochim. Acta 52, 1409–1424.
(doi:10.1016/0016-7037(88)90211-6)
77. Xue S, Yu Y, Hewins RH, Hall GS, Herzog GF. 1996 Zinc losses from and zinc isotopic
abundances in residues formed by heating of Zn, ZnO, ZeS, and Zn-doped silicate glass.
In Proc. 27th Lunar and Planetary Science Conf., Houston, TX, 18–22 March, pp. 1461–1462.
New York, NY: Pergamon.
78. Koeberl C. 1994 Tektite origin by hypervelocity asteroidal or cometary impact: target rocks,
source craters, and mechanisms. In Large meteorite impacts and planetary evolution (eds BO
Dressler, RAF Grieve, VL Sharpton), Geological Society of America Special Paper, 293,
pp. 133–152.
79. Wasson JT. 2003 Large aerial bursts: an important class of terrestrial accretionary events.
Astrobiology 3, 163–179. (doi:10.1089/153110703321632499)
80. Herzog GF, Alexander CMO’D, Glass BP, Berger EL, Delaney JS. 2005 Potassium isotope
fractionation in Australasian microtektites: evidence for evaporation and recondensation in
a vapor plume. In Proc. 36th Lunar and Planetary Science Conf., League City, TX, 14–18 March,
Abstract 1167. See http://www.lpi.usra.edu/meetings/lpsc2005/pdf/1167.pdf.
81. Wombacher F, Rehkamper M, Mezger K, Munker C. 2003 Stable isotope compositions of
cadmium in geological materials and meteorites determined by multiple-collector ICPMS.
Geochim. Cosmochim. Acta 67, 4639–4654. (doi:10.1016/S0016-7037(03)00389-2)
82. Moynier F, Koeberl C, Beck P, Jourdan F, Telouk P. 2010 Isotopic fractionation of Cu in tektites.
Geochim. Cosmochim. Acta 74, 799–807. (doi:10.1016/j.gca.2009.10.012)
83. Albarède F. 2004 The stable isotope geochemistry of copper and zinc. Rev. Mineral. Geochem.
55, 409–427. (doi:10.2138/gsrmg.55.1.409)
84. Visscher C, Fegley B. 2013 Chemistry of impact-generated silicate melt–vapor debris disks.
Astrophys. J. 767, L12. (doi:10.1088/2041-8205/767/1/L12)
85. Bottke WF, Walker RJ, Day JMD, Nesvorny D, Elkins-Tanton L. 2010 Stochastic late accretion
to Earth, the Moon, and Mars. Science 330, 1527–1530. (doi:10.1126/science.1196874)
86. Genda H, Abe Y. 2005 Enhanced atmospheric loss on protoplanets at the giant impact phase
in the presence of oceans. Nature 433, 842–844. (doi:10.1038/nature03360)
87. Hartmann W, Davis D. 1975 Satellite-sized planetesimals and lunar origin. Icarus 24, 504–514.
(doi:10.1016/0019-1035(75)90070-6)
88. Wade J, Wood BJ. 2005 Core formation and the oxidation state of the Earth. Earth Planet. Sci.
Lett. 236, 78–95. (doi:10.1016/j.epsl.2005.05.017)
89. Frost DJ, Liebske C, Langenhorst F, McCammon CA, Trønnes RG, Rubie DC. 2014
Experimental evidence for the existence of iron-rich metal in the Earth’s lower mantle. Nature
428, 409–412. (doi:10.1038/nature02413)
90. Albarède F. 2009 Volatile accretion history of the terrestrial planets and dynamic
implications. Nature 461, 1227–1233. (doi:10.1038/nature08477)
91. Marty B. 2012 The origins and concentrations of water, carbon, nitrogen and noble gases on
Earth. Earth Planet. Sci. Lett. 313–314, 56–66. (doi:10.1016/j.epsl.2011.10.040)
92. Albarède F, Ballhaus C, Blichert-Toft J, Lee C-T, Marty B, Moynier F, Yin Q-Z. 2013
Asteroidal impacts and the origin of terrestrial and lunar volatiles. Icarus 222, 44–52.
(doi:10.1016/j.icarus.2012.10.026)
93. Walker RJ. 2009 Highly siderophile elements in the Earth, Moon and Mars: update
and implications for planetary accretion and differentiation. Chem. Erde 69, 101–125.
(doi:10.1016/j.chemer.2008.10.001)
94. Becker H, Horan MF, Walker RJ, Gao S, Lorand J-P, Rudnick RL. 2006 Highly
siderophile element composition of the Earth’s primitive upper mantle: constraints from
new data on peridotite massifs and xenoliths. Geochim. Cosmochim. Acta 70, 4528–4550.
(doi:10.1016/j.gca.2006.06.004)
95. Lodders K, Fegley B. 1998 The planetary scientists companion. Oxford, UK: Oxford University
Press.
96. Riches AJV, Day JMD, Walker RJ, Simonetti A, Liu Y, Neal CR, Taylor LA. 2012.
Rhenium–osmium isotope and highly-siderophile-element abundance systematics of angrite
meteorites. Earth Planet. Sci. Lett. 353/354, 208–218. (doi:10.1016/j.epsl.2012.08.006)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
25
.........................................................
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
97. O’Keefe JD, Ahrens TJ. 1977 Impact-induced energy partitioning, melting, and vapourization
on terrestrial planets. In Proc. 8th Lunar Science Conf., Houston, TX, 14–18 March, pp. 3357–
3374. New York, NY: Pergamon.
98. Day JMD, Walker RJ, Qin L, Rumble D. 2012 Late accretion as a natural consequence of
planetary growth. Nat. Geosci. 5, 614–617. (doi:10.1038/ngeo1527)
99. Dauphas N, Pourmand A. 2011 Hf–W–Th evidence for rapid growth of Mars and its status
as a planetary embryo. Nature 473, 489–492. (doi:10.1038/nature10077)
100. Chou C-L, Baedecker PA, Bild RW, Wasson JT. 1974 Volatile-element systematics and green
glass in Apollo 15 lunar soils. In Proc. 6th Lunar Science Conf., Houston, TX, 18–22 March, pp.
1645–1657. New York, NY: Pergamon.
101. Meyer C, McKay DS, Anderson DH, Butler P. 1975 The sources of subliminates on the Apollo
15 green and Apollo 17 orange glass samples. In Proc. 6th Lunar Science Conf., Houston, TX,
17–21 March, pp. 1673–1699. New York, NY: Pergamon.
102. Cirlin EH, Housley RM. 1979 Scanning Auger microprobe and atomic absorption studies of
lunar volcanic volatiles. In Proc. 10th Lunar and Planetary Science Conf., Houston, TX, 19–23
March, pp. 341–354. New York, NY: Pergamon.
103. Krähenbühl U. 1980 Distribution of volatile and non-volatile elements in grain-size fractions
of Apollo 17 drive tube 74001/2. In Proc. 11th Lunar and Planetary Science Conf., Houston, TX,
17–21 March, pp. 1551–1564. New York, NY: Pergamon.
104. Ding TP, Thode HG, Rees CE. 1983 Sulphur content and sulphur isotope composition of
orange and black glasses in Apollo 17 drive tube 74002/1. Geochim. Cosmochim. Acta 47, 491–
496. (doi:10.1016/0016-7037(83)90271-5)
105. Delano JW, McGuire J. 1992 Abundances of sodium, sulfur, and potassium in lunar volcanic
glasses. LPI Technical Report, 92–09, Part 1, pp. 7–9. Lunar and Planetary Institute.
106. Delano JW, Hanson BZ, Watson EB. 1994 Abundance and diffusivity of sulfur in lunar picritic
magmas. In Proc. 25th Lunar and Planetary Science Conf., Houston, TX, 14–18 March, Abstract
1163, pp. 325–326. See http://www.lpi.usra.edu/meetings/lpsc1994/pdf/1163.pdf.
107. Weitz CM, Rutherford MJ, Head JW, McKay DS. 1999 Ascent and eruption of a lunar hightitanium magma as inferred from the petrology of the 74001/2 drill core. Meteorit. Planet. Sci.
34, 527–540. (doi:10.1111/j.1945-5100.1999.tb01361.x)
108. Symonds RB, Rose WI, Reed MH, Lichte FE, Finnegan DL. 1987 Volatilization, transport
and sublimation of metallic and non-metallic elements in high temperature gases at
Merapi Volcano, Indonesia. Geochim. Cosmochim. Acta 51, 2083–2101. (doi:10.1016/00167037(87)90258-4)
109. Lofgren G, Donaldson CH, Williams RJ, Mullins O, Usselman TM. 1974 Experimentally
reproduced textures and mineral chemistry of Apollo 15 quartz normative basalts. In
Proc. 5th Lunar Science Conf., Houston, TX, 18–22 March, pp. 549–567. New York, NY:
Pergamon.
110. Usselman TM, Lofgren GE, Donaldson CH, Williams RJ. 1975 Experimentally reproduced
textures and mineral chemistries of high-titanium mare basalts. In Proc. 6th Lunar Science
Conf., Houston, TX, 17–21 March, pp. 997–1020. New York, NY: Pergamon.
111. Day JMD, Taylor LA. 2007 On the structure of mare basalt lava flows from textural analysis
of the LaPaz icefield and Northwest Africa 032 lunar meteorites. Meteorit. Planet. Sci. 42, 3–18.
(doi:10.1111/j.1945-5100.2007.tb00213.x)
112. Giguere TA, Taylor GJ, Hawke BR, Lucey PG. 2000 The titanium contents of lunar mare
basalts. Meteorit. Planet. Sci. 35, 192–200. (doi:10.1111/j.1945-5100.2000.tb01985.x)
113. Snyder GA, Taylor LA, Neal CR. 1992 A chemical model for generating the sources of
mare basalts: combined equilibrium and fractional crystallization of the lunar magmasphere.
Geochim. Cosmochim. Acta 56, 3809–3824. (doi:10.1016/0016-7037(92)90172-F)
114. Smith JA, Anderson AT, Newton RC, Olsen EJ, Wyllie PJ, Crewe AV, Isaacson MS, Johnson D.
1970 Petrologic history of the Moon inferred from petrography, mineralogy, and petrogenesis
of Apollo 11 rocks. In Proc. Apollo 11 Lunar Science Conf., pp. 1149–1162. New York, NY:
Pergamon.
115. Wood JA, Dickey JS, Marvin UB, Powell BN. 1970 Lunar anorthosites and a geophysical
model of the Moon. In Proc. Apollo 11 Lunar Science Conf., pp. 965–988. New York, NY:
Pergamon.
116. Elkins-Tanton LT. 2012 Magma oceans in the inner Solar System. Annu. Rev. Earth Planet. Sci.
40, 113–139. (doi:10.1146/annurev-earth-042711-105503)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 17, 2017
26
.........................................................
rsta.royalsocietypublishing.org Phil. Trans. R. Soc. A 372: 20130259
117. Liu Y, Spicuzza MJ, Craddock PD, Day JMD, Valley JW, Dauphas N, Taylor LA.
2010 Oxygen and iron isotope constraints on near-surface fractionation effects and the
composition of lunar mare basalt source regions. Geochim. Cosmochim. Acta 74, 6249–6262.
(doi:10.1016/j.gca.2010.08.008)
118. Sedaghatpour F, Teng FZ, Liu Y, Sears DWG, Taylor LA. 2013 Magnesium isotopic
composition of the Moon. Geochim. Cosmochim. Acta 120, 1–16. (doi:10.1016/j.gca.
2013.06.026)
119. Elkins-Tanton LT, Van Orman JA, Hager BH, Grove TL. 2002 Re-examination of the lunar
magma ocean cumulate overturn hypothesis: melting of mixing is required. Earth Planet. Sci.
Lett. 196, 239–249. (doi:10.1016/S0012-821X(01)00613-6)
120. O’Hara MJ. 2000 Flood basalts, basalt floods or topless bushvelds? Lunar petrogenesis
revisited. J. Petrol. 41, 1545–1651. (doi:10.1093/petrology/41.11.1545)