Short-lived radionuclides as monitors of early crust–mantle differentiation on the terrestrial planets

Earth and Planetary Science Letters 279 (2009) 147–156

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Short-lived radionuclides as monitors of early crust–mantle differentiation on the terrestrial planets Richard W. Carlson a,⁎, Maud Boyet b a b

Department of Terrestrial Magnetism, Carnegie Institution of Washington, 5241 Broad Branch Road, NW, Washington, DC 20015, USA Laboratoire Magmas et Volcans, UMR CNRS 6524, Observatoire de Physique du Globe, Universite Blaise Pascal, 5 rue Kessler, 63038 Clermont-Ferrand Cedex, France

a r t i c l e

i n f o

Article history: Accepted 7 January 2009 Available online 12 February 2009 Editor: A.W. Hofmann Keywords: short-lived radioactivity early planetary differentiation mantle composition Moon Mars isotopes

a b s t r a c t The kinetic energy from large impacts, the gravitational energy released by core formation, and the heat provided by the decay of short-lived radioactive isotopes all drive extensive melting and chemical differentiation of silicate planets/planetesimals during and shortly after their formation. This early differentiation is best preserved on small objects such as the parent bodies of the eucrite and angrite meteorites where silicate melts were produced within 3 million years of solar system formation. The W isotopic composition of some iron meteorites testifies to core segregation on small planetesimals within as little as one million years or less of solar system formation. On larger objects, such as the Moon, Mars and Earth, the evidence for early differentiation provided by long-lived radioisotope systems has been variably overprinted by the continuing differentiation of these objects, but a clear signature of extensive early planetscale differentiation is preserved in a variety of short-lived radioisotope systems, particularly, I–Pu–Xe, Hf–W and 146Sm–142Nd. All these systems suggest that global differentiation of planetesimals and the terrestrial planets occurred during the first hundred million years of solar system history. This early processing of the Moon, Mars and Earth, may have fundamentally affected the evolution of these planets and their current internal compositional structure. © 2009 Elsevier B.V. All rights reserved.

1. Introduction With the realization that major planetary differentiation events, such as core formation and planetary-scale melting, occurred over million to hundred-million year time spans following solar system formation, deciphering the nature and timing of these early events has become a major effort. Fortunately, the solar system was salted with a number of short-lived radioisotopes (Table 1) whose decay has left indelible imprints of processes operating early in solar system history. The rapid decay of these systems allows remarkable temporal resolution on events occurring more than 4.5 billion years ago. This review will concentrate on the use of short-lived radiometric systems to track early crust–mantle differentiation on rocky planetesimals, the Moon, Mars and Earth. 2. Defining the initial conditions

Using the 26Al–26Mg system as an example, the decay equation is written as: ð26 Mg=24 MgÞΔt ¼ ð26 Mg=24 MgÞ0 þ ð26 Al=24 MgÞ0  e−λΔt

ð1Þ

where (26Mg/24Mg)0 and (26Mg/24Mg)Δt are the Mg isotopic composition at t = 0, usually the start of solar system condensation, and after a decay interval of “Δt”. (26Al/24Mg)0 is the starting abundance (at t = 0) of 26Al relative to 24Mg, and λ is the decay constant for 26Al. Because primordial 26 Al is now extinct, (26Al/24Mg)0 is rewritten as (26Al/27Al)0 × (27Al/24Mg) where 27Al/24Mg is the ratio of two stable isotopes used as a measure of the Al to Mg ratio of a sample. A line on a plot of 26Mg/24Mg vs. 27Al/24Mg thus gives a slope corresponding to (26Al/27Al)0 × e− λΔt, which is equal to the 26Al/27Al ratio at the time of closure of the Al–Mg system in the sample being studied. These ratios can be translated into relative ages with knowledge of the initial abundance of the radioactive parent (Table 1) via the equation:

2.1. Extinct radioisotope systematics Because the short-lived isotopes present in the early solar system have long-since decayed away, their presence can only be inferred through their contribution to the abundance of their decay products. ⁎ Corresponding author. E-mail address: [email protected] (R.W. Carlson). 0012-821X/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2009.01.017

Δt ¼ −ðln½ð26 Al=27 AlÞΔt =ð26 Al=27 AlÞ0 Þ=λ

ð2Þ

and the assumption that the parent element was isotopically uniform throughout the solar nebula. This relative time scale is then linked to absolute ages by using the few samples that also can be dated by a still extant decay system, usually the U–Pb system because of its ability to provide ages resolved to better than a million years.

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Table 1 Short-lived radioactive elements useful for examining early planetary differentiation Parent isotope

Daughter isotope

Half-life (Ma)

Initial abundance

26

26

0.73 3.7 6.5 15.7 103 9 82

26

Al 53 Mn 107 Pd 129 I 146 Sm 182 Hf 244 Pu

Mg 53 Cr 107 Ag 129 Xe 142 Nd 182 W Heavy Xe

Al/27Al = 4.96 ± 0.25 × 10− 5a Mn/55Mn = 6.3 ± 0.7 × 10− 6b 107 Pd/108Pd = 5.9 ± 2.2 × 10− 5c 129 127 I/ I = 1.07 ± 0.04 × 10− 4d 146 Sm/144Sm = 7.6 ± 1.7 × 10− 3e 182 Hf/180Hf = 1.07 ± 0.10 × 10− 4f 244 Pu/238U = 6.8 ± 1.0 × 10− 3g 53

a Based on the AJEF CAI from the Allende CV3 chondrite and a Pb–Pb age of 4567.6 Ma (Jacobsen et al., 2008). b From the Orgueil CI chondrite (Trinquier et al., 2008). c Derived from the Pd/Ag vs. 107Ag/109Ag correlation of whole rock carbonaceous chondrites (Schonbachler et al., 2008). d Average for 10 chondrites at the uncertain time of closure of the I–Xe system (Hohenberg et al., 1967). e Calculated for an age of 4567.6 Ma from the Sm–142Nd isochron for the angrite LEW86010 and the Pb–Pb age of 4.5578 Ga (Lugmair and Galer, 1992). f From internal isochrons for CAIs from the Allende CV3 chondrite, Kleine et al. (2005). g At the assumed 4.56 Ga closure age for Pu–Xe in the St. Severin chondrite (Hudson et al., 1989).

of the following possibilities; 1) a non-uniform distribution of at least some of the short-lived radioactive species, 2) the different pinning points used and hence inaccurately estimated solar system initial abundances of the parent isotope, 3) different closure temperatures of the various radiometric systems and protracted cooling, or later partial disturbance, of either D'Orbigny or the samples used as pinning points for the absolute time scale; 4) terrestrial Pb contamination and hence inaccuracies in either the CAI or D'Orbigny U–Pb ages or, 5) analytical errors, particularly in the measurement of the small isotopic variations observed for the short-lived systems. While the issue of the uniform distribution of short-lived isotopes will remain a topic of concern, the

The most commonly used “pinning point” to translate the shortlived radioisotope relative time-scales into absolute ages are the calcium–aluminum rich inclusions (CAI) found in some types of carbonaceous chondritic meteorites (MacPherson, 2003). The key assumption here is that CAIs sample the first solids that condensed from an originally very hot solar nebula because they contain refractory mineral phases that, with a few exceptions (e.g. Meyer and Zinner, 2006), have isotope compositions similar to average solar system materials. CAIs have the added advantage that, because they are volatile-element depleted, their high U/Pb ratios allow determination of very precise internal Pb–Pb ages, for example 4567.2 ± 0.7 Ma and 4567.4 ± 1.1 Ma for two CAIs from the Efremovka chondrite (Amelin et al., 2002) and 4567.6 ± 0.4 Ma for a CAI from the Allende chondrite (Jacobsen et al., 2008). Unfortunately, CAIs have proven difficult to date with 53Mn–53Cr and 146Sm–142Nd. Consequently, the angrite meteorites, a family of quickly-cooled planetesimal melts, also are commonly used for cross calibration of the short-lived systems with absolute U–Pb ages (Lugmair and Galer, 1992; Amelin, 2008). 2.2. An initially uniform distribution of short-lived nuclides? Whether the short-lived isotopes were uniformly distributed in the solar nebula before planet formation started remains a topic of concern, particularly for very short-lived isotopes such as 26Al and 60Fe that likely were injected into the solar nebula from a nearby stellar source just prior to the initiation of solar nebula collapse (e.g. Jacobsen, 2005). There is the additional question of whether some of the short-lived radioactive elements (10Be, 26Al, 41Ca, 53Mn) could have been generated within the solar system by irradiation close to the proto-Sun (e.g. Gounelle et al., 2001). Whether these isotopes were uniformly distributed in the early solar nebula can be checked by comparing ages for the same sample obtained by different radioisotope systems. A good example of this approach is the angrite D'Orbigny that provided a Pb–Pb age of 4564.42 ± 0.12 Ma (Amelin, 2008), and initial 53Mn/55Mn= (3.23 ± 0.04) × 10− 6 (Glavin et al., 2004), 26Al/27Al = (5.1 ± 0.3) × 10− 7 (Spivack-Birndorf et al., 2005), and 182Hf/180Hf = (7.4 ± 0.2)× 10− 5 (Markowski et al., 2007) (Fig. 1). Using the initial abundances of these isotopes at an assumed 4.5676 Ga age for the solar system (Table 1) provides the following absolute ages: Al–Mg= 4562.8 ± 0.5 Ma, Mn–Cr= 4564.0 ± 0.6 Ma, and Hf–W = 4562.8 ± 1.5 Ma. The Mn–Cr age agrees well with the Pb–Pb age while the Al–Mg and Hf–W ages agree well with one another, but are 1.6 Ma younger than the Pb–Pb age. Using the values in Table 1, the Al–Mg and Hf–W ages are pinned to data for CAIs whereas the Mn–Cr age is pinned to data from the CI chondrite Orgueil. The age difference thus could reflect any or all

Fig. 1. Examples of various radioisotope ages for the angrite D'Orbigny. The time intervals from the short-lived systems are calculated using the isochron slopes and the solar system initial parent radioisotope abundances given in Table 1. Panel A — U–Pb data from Amelin (2008). Squares and circles are the residues of different leaching procedures while the diamonds are the leaches of these splits. Panel B shows the Mn–Cr data for D'Orbigny (filled squares = olivine, open circles = glass, filled circle = total rock, diamonds = pyroxene, open square = spinel) from Glavin et al. (2004). Panel C — Hf–W data for D'Orbigny (open squares) and the similar angrite SAH 99555 (filled squares) from Markowski et al. (2007). The Hf–W isochron for CAIs from the Allende chondrite (Kleine et al., 2005) is shown by the dashed line. ε182W is the difference, in parts in 10,000, between the measured 182W/184W and that of a terrestrial standard.

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results discussed above suggest that this issue only affects the chronologies provided by these systems at the 1 to 2 million year precision level, which is similar to, or just slightly greater than the precisions dictated by analytical errors in the measurements of isotopic variations in natural samples. 2.3. On the edge of analytical precision This comparison also highlights a critical issue with the application of short-lived radioisotope systems — analytical precision. Fig. 1 shows a total range of about 17 parts in 10,000 in 182W/184W and about 8 parts in 10,000 in 53Cr/52Cr — only about 100 times greater than analytical precision. At the whole-rock scale, these variations can be even smaller. For example, the Moon shows a total variation in 142Nd/ 144 Nd of only 60 ppm and the Earth of only 35 ppm. Limited analytical accuracy postponed the discovery of a difference between terrestrial and chondritic 182W/184W for 7 yr (Lee and Halliday, 1995; Kleine et al., 2002; Schoenberg et al., 2002; Yin et al., 2002). Similarly, after the first reports of 142Nd/144Nd variation in meteorites (Lugmair and Marti, 1977), 28 yr passed until a difference in 142Nd/144Nd between meteorites and modern terrestrial rocks was confidently resolved (Boyet and Carlson, 2005). Even now, a difference of some 15 ppm exists between measurements of 142Nd/144Nd in the same lunar rocks (Nyquist et al., 1995; Rankenburg et al., 2006; Boyet and Carlson, 2007) whose explanation, though likely analytical, remains elusive. Obtaining isotope ratio precisions in the few ppm range require careful evaluation of a wide variety of instrumental parameters during mass spectrometry. More problematical is the issue of adequately correcting for the mass fractionation that occurs in nature and during mass spectrometry. At the few ppm precisions now reported, deviations in commonly used mass fractionation corrections due both to the behavior of samples in the mass spectrometer (Hart and Zindler, 1989; Upadhyay et al., 2008) and by chemically induced mass fractionation of either standards or samples (O'Neil et al., 2008) can create offsets of a few ppm that potentially can impact the interpretation of short-lived radioisotope data.

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Ranen and Jacobsen (2006) suggested that nucleosynthetic heterogeneity, and not the decay of 146Sm, could explain the difference in 142Nd/144Nd of chondritic meteorites compared to modern terrestrial rocks (Boyet and Carlson, 2005). Indeed, primitive carbonaceous chondrites do show nucleosynthetic anomalies in 142Nd/ 144 Nd that are accompanied by deficiencies in 144Sm and isotopically anomalous Ba (Andreasen and Sharma, 2006; Carlson et al., 2007). Assuming that the deficits in 142Nd and 144Sm are related to a deficit in p-process isotopes in C-chondrites (Andreasen and Sharma, 2006) allows a correction to be applied that results in an average carbonaceous chondrite 142Nd/144Nd that is between 17 and 24 ppm lower than the modern terrestrial value (Carlson et al., 2007). For an average chrondritic 147Sm/144Nd = 0.1960 (Bouvier et al., 2008b), ordinary and enstatite chondrites that show no nucleosynthetic effects in Ba, Nd or Sm outside of measurement uncertainty (Andreasen and Sharma, 2007; Carlson et al., 2007) average to a 142 Nd/144Nd ratio that is lower than the modern terrestrial value by 16 ppm (Carlson and Boyet, 2008). Thus, nucleosynthetic heterogeneity in the early solar system does appear to be responsible for the range in 142Nd/144Nd seen in carbonaceous chondrites, but not for the offset of 142Nd/144Nd between chondrites and Earth. For the latter, decay of 146Sm in an Earth reservoir with non-chondritic Sm/Nd ratio remains the more likely explanation. 3. A quick beginning to planetary differentiation The short-lived radioisotope systems show unambiguously that planetary differentiation started quickly. The most obvious indication of this is the low 182W/184W ratios found in iron meteorites (Fig. 2) (Harper and Jacobsen, 1996; Lee and Halliday, 1996; Horan et al., 1998; Kleine et al., 2002, 2005; Markowski et al., 2006b; Schersten et al., 2006). Because W is siderophile, but Hf is lithophile, iron meteorites have Hf/W ratios near zero compared to chondritic Hf/W of about 1.1. Ingrowth of

2.4. Distinguishing radiogenic contributions from other causes of isotopic variation A variety of other natural processes can produce isotope shifts of comparable magnitude to those caused by the decay of short-lived radioactive isotopes. These include nuclear reactions that occur as a result of exposure of samples to high-energy cosmic rays. This issue caused considerable confusion in the interpretation of W isotope data for meteorites (Leya et al., 2000) and lunar samples (Lee et al., 2002) and is at least one component of the discrepancy currently present between different studies of 142Nd/144Nd in lunar rocks (Nyquist et al., 1995; Rankenburg et al., 2006; Boyet and Carlson, 2007). Another cause of isotope variation is imperfect mixing of the different stellar nucleosynthetic contributions to the proto-solar nebula. Just a few years ago, the only elements that showed resolvable isotope variation at the whole-rock scale in meteorites were oxygen (Clayton et al., 1973) and noble gases (Black and Pepin, 1969). Now, this list has expanded to elements involved in short-lived radioisotope systems including Cr (Trinquier et al., 2007) and Sm–Nd (Andreasen and Sharma, 2006; Carlson et al., 2007). This isotopic variability comes as no surprise because presolar grains with isotopic compositions hugely different from the solar system have been known to exist in primitive meteorites for some time (e.g. Meyer and Zinner, 2006). Nucleosynthetic variability is a potential issue for the 53Mn–53Cr system, although presolar carriers with highly anomalous 54Cr show only minor variation in 53Cr (Qin et al., 2008a), and hence are unlikely to greatly affect the 53Mn–53Cr results except in some cases that used 54 Cr for mass fractionation correction (Lugmair and Shukolyukov, 1998).

Fig. 2. Measured W isotopic compositions for a variety of iron meteorites and metal-rich chondrites with cosmic-ray exposure ages of 300 Ma or less. The cross-hatched bar marks the initial W isotopic composition of the solar system as determined from Hf–W isochrons for Allende CAIs (Kleine et al., 2004). The grey-bar shows the present-day W isotopic composition of chondrites (Kleine et al., 2002; Schoenberg et al., 2002; Yin et al., 2002). ε182W is the difference in parts in 10,000 in sample 182W/184W compared to a terrestrial W standard. Data from Horan et al. (1998); Kleine et al. (2005); Markowski et al. (2006b); Qin et al. (2008b).

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182 W thus stops in iron metal as soon as it is segregated from the 182Hf that is retained in the silicate fraction of a differentiating planetesimal. 182 W/184W ratios also can be lowered by interaction with cosmic rays (Markowski et al., 2006a) so concentrating only on those irons with relatively short exposure ages and for those where corrections for cosmic ray effects can be applied (Markowski et al., 2006b; Qin et al., 2008b), the 182W/184W of most magmatic irons are indistinguishable from the CAI initial 182W/184W (Fig. 2). This result indicates that metal– silicate separation started on planetesimals within less than a million years of CAI formation (Kleine et al., 2005; Markowski et al., 2006b; Qin et al., 2008b). Data from the non-magmatic groups of iron meteorites (IAB, IIICD and IIE) have slightly more radiogenic W with formation ages of 5 to 14 Ma after CAIs that presumably reflect secondary processes, perhaps impacts, on their parent bodies (Kleine et al., 2005; Markowski et al., 2006b; Qin et al., 2008b). Other indicators of the rapidity of planetesimal differentiation are the very old crystallization ages of some angrites (Fig. 1) and the 4565 Ma Mn–Cr age obtained for whole-rock eucrites (Lugmair and Shukolyukov, 1998; Trinquier et al., 2008) that presumably dates an initial global differentiation event on the eucrite parent body. This very quick melting is consistent with the presence of live 26Al when some eucrites crystallized (Srinivasan et al., 1999; Nyquist et al., 2003) and when the eucrite, angrite, and mesosiderite parent bodies initially differentiated from solar Al/Mg ratios (Bizzarro et al., 2005). These very old ages suggest that 26Al provided an important heat source to fuel early planetesimal differentiation, which is likely only during the first ~1 Ma of solar system history (Bizzarro et al., 2005). A corollary is that any planetesimal that grew large enough during this time interval to retain the heat of 26Al decay experienced extensive melting and differentiation. Thus, the terrestrial planets likely accumulated from already differentiated objects rather than from primitive chondrites.

systematics of a variety lunar rocks (Tera and Wasserburg, 1974), some analyses of terrestrial Pb data (Halliday, 2008), and the limited constraints provided by the low 87Sr/86Sr of the lunar highlands (Carlson and Lugmair, 1988). The relatively young age of the Moon and its compositional similarity to Earth provides additional support for the idea that the Moon formed as the result of a giant impact into the Earth during the latter stages of planetary accumulation (Hartmann and Davis, 1975; Canup and Asphaug, 2001) some hundred million years after CAI formation. Although lunar rocks show no evidence for live 182Hf at the time of their formation, they display a sizeable range in 142Nd/144Nd that testifies to a first stage of lunar differentiation that occurred while 146Sm was still extant (Nyquist et al., 1995; Rankenburg et al., 2006; Boyet and Carlson, 2007). The lunar 142Nd/144Nd data correlate to some degree with Sm/Nd ratio defining a rough trend that corresponds to an age of some 200–250 Ma after solar system formation (Fig. 3). The exact meaning of this age, however, is not clear. The oldest precise age for the

4. The growth and differentiation of the terrestrial planets 4.1. The Moon Starting this section with the Moon may seem like an odd choice, but the returned lunar samples provided the first impetus to consider that the early evolution of planets is dominated by global melting and differentiation. The overwhelming abundance of plagioclase in the lunar highlands crust coupled with the complementary Eu depletion in basalts derived from the lunar mantle led to the idea that separation of the lunar crust and mantle was caused by crystal fractionation from a largely molten Moon. That this crust–mantle separation was accomplished early in solar system history is shown clearly by the very low 87Sr/86Sr of lunar highlands rocks (Nyquist et al., 1974), the circa 4.4 to 4.5 Ga model ages for various lunar samples in the Sm–Nd (Carlson and Lugmair, 1979), and U–Pb (Tera and Wasserburg, 1974) systems, and the 4.3 to 4.5 Ga crystallization ages of the few lunar anorthosites that have been dated (e.g. Norman et al., 2003). Early work on Hf–W systematics of lunar samples found large variation in 182W/184W that would have pushed Moon formation and differentiation back before 4.5 Ga (Lee et al., 1997), but the lunar W story changed when it was noted that 182W is produced by cosmic-ray interaction with 181Ta (Lee et al., 2002). Separation of Ta-free, but Wrich, iron metal from lunar samples now shows not only that there is no W isotope variation on the Moon caused by 182Hf decay, but that the Moon's W isotopic composition is identical to that of Earth's mantle (Touboul et al., 2007). Thus the Hf–W system constrains the timing of Moon formation and differentiation to have occurred after 182 Hf was extinct, or at least ~60 Ma after solar system formation (Touboul et al., 2007). The revised W result returns us to a discussion point of the 1970's and 1980's; the age of the Moon and the Earth. The oldest ages of lunar crustal rocks provide a minimum age of 4.46 Ga for the Moon (Norman et al., 2003). This relatively young age is consistent with the U–Pb

Fig. 3. Lunar Sm-Nd data from Nyquist et al. (1995) (grey symbols); Rankenburg et al. (2006) (small open symbols) and Boyet and Carlson (2007) (filled symbols). The large white circles along the 4.45 Ga line show the model composition for the highlands crust and KREEP, low-Ti, and high-Ti basalts under the assumption that they all formed instantaneously at the completion of magma ocean crystallization at 4.45 Ga from a high Sm/Nd ratio bulk Moon (Boyet and Carlson, 2007). The white circle plotting just below the 4.32 Ga line shows the position of the high-Ti basalt source if it did not form until 4.30 Ga. A best-fit line to the model KREEP, low-Ti and 4.30 Ga high-Ti basalt source points gives an age of 4.32 Ga, similar to that of the best-fit line through the lunar basalt data. On panel B, the long line has a slope corresponding to an age of 4.45 Ga and the short 3.7 Ga line provides a slope similar to the age of the high-Ti basalts. Neither are best-fit lines to the data.

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lunar crust is the 4.44 ± 0.02 Ga age for ferroan anorthosite 60025 (Carlson and Lugmair, 1988). If the lunar 142Nd array dates the completion of magma ocean differentiation, this would imply that the magma ocean continued to differentiate for some 70–120 Ma after lunar crust formation. The Eu depletion in some lunar basalts shows that their source regions must have formed after crust formation, which is consistent with this possibility. Assigning chronological significance to a whole-rock isochron, however, assumes that all the samples were produced in a single differentiation event and started with the same initial isotopic composition. If instead, the differentiation of the lunar interior was not completed in a single step, then the time provided by a whole rock isochron may give a time intermediate between the start and finish of early planetary differentiation. An example of this is shown by the large white circles in Fig. 3A. A combination of geodynamic and petrologic models show that crystallization of a magma ocean creates a buoyantly unstable cumulate package that will undergo convective overturn, mixing, and likely additional melting episodes (ElkinsTanton et al., 2002). For the Moon, both the presence of 142Nd/144Nd variation and the rough alignment of the lunar Sm–143Nd data along the 4.45 Ga reference lines in Fig. 3B indicate that the early differentiation event exerted the major influence over the compositional variation of lunar rocks. The alignment of whole rock Sm–Nd data for circa 3.7 Ga high-Ti basalts along a 3.7 Ga slope (Fig. 3B) shows that the partial melting event that created these magmas did fractionate Sm–Nd at 3.7 Ga, but not by enough to move them well away from the 4.45 Ga reference line. As we will see in the next sections, continued differentiation of both Earth and Mars has severely clouded the picture of early events given by long-lived radioisotope systems, but not those provided by the short-lived systems. 4.2. Mars The shergottite, nakhlite and chassignite (SNC) meteorites sample chemically distinct materials on their parent body, presumably Mars, that retain huge variations in the isotopic composition of long-lived isotope systems such as Sr, Nd, Hf and Os (Jagoutz, 1991; Borg et al., 1997; Blichert-Toft et al., 1999; Brandon et al., 2000; Borg et al., 2003; Debaille et al., 2008), and preserve clear evidence for early differentiation in variable 182W/184W and 142Nd/144Nd ratios (Harper et al., 1995; Kleine et al., 2004; Foley et al., 2005; Debaille et al., 2007; Caro et al., 2008). The most obvious sign of very rapid differentiation of the Martian interior is the range in 182W/184W observed in the Martian meteorites, from just slightly higher than terrestrial (ε182W = 0.3) for the enriched shergottites to a high of ε182W = 3.1 for the nakhlites (Kleine et al., 2004; Foley et al., 2005). The Hf/W ratio is fractionated by core formation, but also through fractionation of minerals such as clinopyroxene, garnet, and particularly ilmenite because W is more incompatible than Hf in these minerals (Shearer and Righter, 2003). Given the large uncertainty in estimates of bulk-Mars Hf/W ratio, SNC W isotope data constrain Mars core formation only to have occurred earlier than 10 Ma after solar system formation (Nimmo and Kleine, 2007). If silicate differentiation occurred at the same time, the 2.8 epsilon unit variation in 182W/184W in Martian samples requires about a factor of 2 variations in Hf/W ratio in their source regions. The variation in W isotopic composition is poorly correlated with the wide range in 142Nd/144Nd displayed by the SNC meteorites, although the 142Nd/144Nd ratios reported by (Foley et al., 2005) do not always agree with more recent, and more precise, measurements of whole rock SNC's (Debaille et al., 2007; Caro et al., 2008). In the more recent data, ε142Nd ranges from −0.27 to 0.65 in shergottites and averages 0.64 ± 0.07 (n = 4) for the nakhlites (Debaille et al., 2007; Caro et al., 2008). The 142 Nd/144Nd variation in the shergottites is well correlated with Sm/Nd ratio (Fig. 4A). The correlations correspond to ages of 4335 ± 14 Ma with

Fig. 4. Sm–Nd (A–B) and Lu–Hf (C) isochron diagrams for SNC whole rock data. Circles show data for nakhlites, black and grey squares for circa 170 Ma and 327–575 Ma shergottites, respectively. The open symbols in panel A show the source Sm/Nd ratios calculated from the initial 143Nd/144Nd at the eruption age of each sample. The open diamond in each figure shows the average value for ordinary chondrites (Boyet and Carlson, 2005; Bouvier et al., 2008b). The lines in panel A are best-fit lines to the shergottite data using measured Sm/Nd ratios (the 4.34 Ga age) excluding the two shergottites (Dhofar 19, NWA1183) that lie well off this line, and to all the shergottite data using calculated source Sm/Nd ratios (4.50 Ga). The remaining lines are reference lines through various sets of data with their measured parent/daughter ratios and isotopic compositions simply to illustrate the slopes expected for various ages. Data shown in these figures from Blichert-Toft et al. (1999); Bouvier et al. (2005); Foley et al. (2005); Debaille et al. (2007, 2008); Caro et al. (2008); Symes et al. (2008).

an initial ε142Nd= −0.13 (relative to chondritic) using measured Sm/Nd ratios, and 4505 ± 13 Ma to 4535 ± 7 using various approaches to calculate source Sm/Nd ratios (Borg et al., 2003; Debaille et al., 2007; Caro et al., 2008) with initial ε142Nd= +0.15 (relative to chondritic), which is within error of modern terrestrial Nd. These results illustrate that the value deduced for bulk-Mars Sm–Nd parameters (Caro et al., 2008) is very sensitive to the model one uses to calculate source Sm/Nd ratios for the shergottites, which, in turn, is sensitive to the assumed crystallization ages of these rocks. While this very large range in 142Nd/ 144 Nd is consistent with rapid differentiation of the Martian interior, the combination of Sr, 142Nd–143Nd, Hf and Pb isotope systematics of the SNCs provide an excellent example of the complexity of early planetaryscale differentiation events, and the difficulty of interpreting this evolution based on a small number of samples.

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Nakhlites/chassignites provide a relatively simple picture of early Mars differentiation. They have a narrow range of mineral ages (Nyquist et al., 2001) and have similar initial Nd (Caro et al., 2008) and W (Foley et al., 2005) isotopic compositions. At the circa 1.3 Ga crystallization ages of these meteorites, they have high 142Nd/144Nd, 143 Nd/144Nd (ε143Nd = 15.2), and 182W/184W, all of which suggest a source that was depleted in incompatible elements while 146Sm and 182 Hf were extant. In contrast, the rocks themselves have very low Sm/ Nd ratios that can be explained by fractionation during the partial melting and fractionation/crystal-accumulation involved in creating the nakhlites and chassignites at 1.3 Ga. Mineral Rb–Sr, Lu–Hf and Sm–Nd isochrons for shergottites divide them roughly into two age groups, one at 170–190 Ma that contains a very wide compositional spectrum in its members, and a second, compositionally restricted, group with a range of ages between 330 and 575 Ma (Nyquist et al., 2001; Borg et al., 2003; Bouvier et al., 2005). Essentially all the mineral isochrons show data scatter well beyond analytical error, leading to debates over whether the ages reflect igneous crystallization (Borg et al., 2003; Gaffney et al., 2007) or resetting of mineral radiogenic isotope systematics through element mobilization caused by shock metamorphism (Nyquist et al., 1979; Jagoutz, 1991; Blichert-Toft et al., 1999; Bouvier et al., 2008a). The debate is compounded by the fact that Pb–Pb mineral data for several shergottites provide ages of over 4 Ga (Bouvier et al., 2005; Gaffney et al., 2007; Bouvier et al., 2008a). What the SNC Pb data show clearly is that the SNC's derive from a source characterized by a very low U/Pb ratio (238U/204Pb ~ 5). What is less clear is whether the old Pb ages provide “source” ages for young magmas derived from old sources, the true crystallization age of the shergottites (Jagoutz et al., 2006; Bouvier et al., 2008a), or simply mixing lines between unradiogenic Martian Pb and modern terrestrial Pb, which unfortunately, and indisputably, also would produce slopes in Pb isochron diagrams corresponding to ages above 4 Ga (Gaffney et al., 2007). The circa 4 Ga Pb–Pb ages are supported by a similar whole rock Rb–Sr age (Jagoutz, 1991) and the wide variation in 142Nd/144Nd. In contrast, the 147 Sm–143Nd and Lu–Hf data of the shergottites show no slopes consistent with such an ancient age (Fig. 4B,C). Unlike the nakhlites, the shergottites have measured Sm/Nd ratios much higher than would be inferred for their sources based on the initial 143Nd/144Nd at the young ages given by mineral isochrons (Borg et al., 1997; Nyquist et al., 2001). This sense of Sm–Nd fractionation is not expected in a melt because Nd is more incompatible than Sm during melting. In contrast, the measured Lu/Hf ratios of the shergottites are lower than those calculated for their sources, as expected for partial melts. This unusual sense of Sm–Nd fractionation and the apparent decoupling from Lu–Hf can be explained if the depleted shergottites are not the first melts from the ancient source, but are instead produced by remelting the residues of a previous melt extraction event (Borg et al., 1997; Debaille et al., 2008). Several versions of this type of multiple-melting event scenario for shergottite genesis have been proposed, extending to the suggestion that as many as 4 or 5 discrete melting and melt-residue mixing events may be involved (Borg et al., 1997; Borg and Draper, 2003; Debaille et al., 2007, 2008). While certainly not impossible, this type of multiple melting scenario is observed rarely, if at all, on the Earth and Moon. The alternative to such complex multiple-melting petrogenetic scenarios is to assume that the composition of the shergottites was strongly modified during the events recorded by the mineral isochrons, perhaps by the loss of phosphorous-rich fluids during impact metamorphism (Blichert-Toft et al., 1999). These issues highlight two difficulties in deciphering the early evolution of Mars. First, the relatively young slopes seen in long-lived radioisotope systems (Fig. 4B,C) reflect the overprinting of early differentiation by much later events. Second, the method of “sampling” of the Martian surface provided by the SNC meteorites has resulted in substantial shock metamorphic effects that severely

complicate the interpretation of age and chemical results. The shortlived radioisotope systems show clearly that differentiation of the Martian interior started early, perhaps with a Martian magma ocean, but later events have overprinted aspects of this story making it difficult to decipher the details of the early differentiation event. 4.3. Earth The first indication from short-lived isotope systems of the importance of early events to Earth differentiation came from the discovery of 129Xe excesses in atmospheric xenon compared to primitive meteoritic Xe (Reynolds, 1960). This provided unquestionable evidence that Earth's interior had lost much of its Xe while 129I was still present. Similarly, once early errors in the determination of the W isotopic composition of chondritic meteorites were corrected, all Earth rocks so far measured have more radiogenic W than chondritic meteorites (Kleine et al., 2002; Schoenberg et al., 2002; Yin et al., 2002). Translating the Xe and W isotope data into ages of atmosphere and core formation depends on whether these reservoirs formed in a single, or more likely, multiple events. The reader is referred to recent reviews of the 129I–244Pu–Xe constraints on the timing and mechanism of atmosphere formation (Pepin, 2006) and the Hf–W constraints on core formation (Halliday and Kleine, 2005; Halliday, 2008) for more detail on the possible complexities. For the purposes of this paper, the key conclusion from these systems is that both atmosphere and core formation occurred quickly, within the first tens to a hundred million years of solar system formation. The strongest suggestion for early differentiation of the silicate earth comes from the observation that terrestrial rocks show nearly as much variability in 142Nd/144Nd ratios as do lunar rocks. Variable 142 Nd/144Nd has now been reported in circa 3.8 Ga rocks from Isua, Greenland (Harper and Jacobsen, 1992; Boyet et al., 2003; Caro et al., 2003; Boyet and Carlson, 2006; Bennett et al., 2007), and Anshan, China (Bennett et al., 2008), and in the Nuvvuagittuq greenstone belt in northern Quebec, Canada (O'Neil et al., 2008) (Fig. 5). Rocks from both Isua and Anshan range to high 142Nd/144Nd, indicative of sources with superchondritic Sm/Nd ratios. Although excesses in 142Nd/144Nd

Fig. 5. Variation in 142Nd/144Nd (ε142Nd is the difference in 142Nd/144Nd, in parts in 10,000 between the sample and a terrestrial laboratory Nd) for chondrites, basaltic eucrites and a wide variety of terrestrial rocks. The grey band shows a typical ± 0.05 analytical error about the mean of modern terrestrial Nd. Data from Nyquist et al. (1995); Caro et al. (2003); Boyet and Carlson (2005, 2006); Rankenburg et al. (2006); Bennett et al. (2007); Carlson et al. (2007); Andreasen et al. (2008); O'Neil et al. (2008).

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as high as 30 ppm were originally reported for these rocks (Harper and Jacobsen, 1992; Boyet et al., 2003), higher precision measurements have confirmed values only 15 to 20 ppm higher than modern terrestrial Nd (Caro et al., 2003; Boyet and Carlson, 2006; Bennett et al., 2007). The Nuvvuagittuq samples also have higher than chondritic 142Nd/144Nd, but range to lower values than modern terrestrial rocks. Furthermore, the Nuvvuagittuq rocks show a Sm/Nd vs. 142Nd/144Nd correlation that, if interpreted as an isochron, indicates either that these rocks are 4.28+− 0.05 0.08 Ga old, or that their source rocks formed at this time (O'Neil et al., 2008). These variations show clearly that the silicate earth was being differentiated while 146Sm was still extant, and that the differentiated reservoirs produced over the first few hundred million years of Earth history still existed by 3.8 Ga to be sampled by these various Eoarchean rocks. Between 3.8 and 3.5 Ga, however, the 142Nd/144Nd variability of mantle-derived rocks became smaller, so that all rocks younger than 3.5 Ga have the same 142Nd/144Nd (Fig. 5) (Boyet et al., 2005; Boyet and Carlson, 2006; Caro et al., 2006; Andreasen et al., 2008). A likely explanation for this is that mixing, caused by convective stirring of the mantle, was effectively mixing away the evidence for early differentiation (Caro et al., 2006; Bennett et al., 2007; Carlson and Boyet, 2008). The peculiar aspect of the data, however, is that post-3.5 Ga terrestrial rocks have 142Nd/144Nd 16 ± 3 ppm higher (Boyet and Carlson, 2005; Carlson and Boyet, 2008) than the average for ordinary chondrites when the chondrite data are adjusted to an average chondrite 147Sm/144Nd= 0.1960 (Bouvier et al., 2008b). If the superchondritic 142Nd/144Nd of all modern terrestrial rocks reflects decay of 146Sm in a reservoir with superchondritic Sm/Nd ratio, the minimum necessary 147Sm/144Nd of this reservoir would be between 0.206 and 0.207, given uncertainties in the initial abundance of 146Sm. The requisite Sm/Nd ratio thus is roughly 6% higher than the average value measured for chondrites. Today, this source would have an ε143Nd between +5.9 and +6.5. This subtly higher than chondritic Sm/Nd ratio inferred for the source of ALL modern terrestrial rocks, including both continental crust and mantle-derived rocks, has several important implications for our understanding of the compositional structure of the Earth (Boyet and Carlson, 2006). 1) The incompatible-element-depleted mantle that serves as source of mid-ocean ridge basalts (MORB) must occupy the major fraction of the mantle. Conventionally, the incompatible-element-depletion characterizing the MORB source is assumed to be complementary to the incompatible-element-enrichment of the continental crust. This simple mass balance can be expressed through the equation: ½XDMM  MDMM + ½XCC  MCC = ½XBSE  ðMDMM + MCC Þ

ð3Þ

which can be rewritten as:

MDMM = MCC  ½XCC −½XBSE = ½XBSE −½XDMM

ð4Þ

where [X] is the concentration of element “X” in a given mass (M) of depleted MORB source mantle (DMM), continental crust (CC), and the assumed bulk-silicate-earth (BSE) starting composition that is assumed to have chondritic relative abundances of the refractory lithophile elements (e.g. McDonough and Sun, 1995). Casting Eq. (4) in terms of Sm/Nd ratio gives:


MDMM = MCC  ðSm=NdÞBSE −ðSm=NdÞCC = ðSm=NdÞDMM −ðSm=NdÞBSE  ½NdCC =½NdDMM

ð5Þ If one replaces (Sm/Nd)BSE in Eq. (5) with a starting reservoir that has high enough Sm/Nd ratio to evolve the superchondritic 142Nd/ 144 Nd observed for both continental and oceanic rocks, then the mass of the depleted mantle must increase compared to a calculation that assumes a chondritic starting Sm/Nd ratio. These mass balance approaches suggest that the MORB source mantle

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occupies at least 74 to 100% of the mantle (Boyet and Carlson, 2006; Carlson and Boyet, 2008). 2) The association of mantle-derived rocks with “primitive” rare gas isotopic compositions (e.g. low 4He/3He) yet Sr, Nd and Hf isotopic compositions suggesting depleted incompatible element abundances (Class and Goldstein, 2005) is expected if the “primitive” mantle supplying surface volcanism currently has an ε143Nd of +6.5 (with 87Sr/86Sr = 0.7032 and ε176Hf = +12) instead of ε143Nd = 0 as assumed for a primitive mantle with chondritic Sm/Nd ratio (Boyet and Carlson, 2006). 3) The mantle source of MORB is slightly less depleted in incompatible elements compared to models that form the MORB source by continent extraction from a mantle that initially had chondritic relative refractory element abundances (Workman and Hart, 2005). Because the incompatible element patterns for the MORB source mantle are calculated using the Sr, Nd and Hf isotope composition of MORB (Workman and Hart, 2005), if one assumes that the starting mantle has, for example, chondritic Sm/Nd (and, hence, initially ε143Nd = 0), the ε143Nd ~ +10 of modern MORB would imply a higher present-day source Sm/Nd ratio, and hence be more incompatible-element-depleted, than would a model that assumes the starting mantle was depleted early in Earth history and thus started with a positive ε143Nd before continent formation began (Boyet and Carlson, 2006). Because the MORB source mantle would occupy most, if not all of the mantle, however, the bulk heat-generating capability of the MORB mantle (e.g. the sum of U, Th and K concentrations) is only about one-quarter that of a “chondritic” mantle (Boyet and Carlson, 2006). Perhaps the most important consequence of the superchondritic Nd/144Nd of modern terrestrial rocks depends on whether the implied superchondritic Sm/Nd ratio is a bulk-planet property or reflects only the characteristics of an incompatible-element-depleted reservoir formed during an early differentiation event on Earth. If indeed formed in an early differentiation, coupling 142Nd and 143Nd systematics requires separation of complementary reservoirs enriched and depleted in incompatible-elements no later than 30 Ma after solar system formation (Boyet and Carlson, 2005). The prime argument in favor of a bulk-earth with chondritic Sm/ Nd ratio is that Sm and Nd are both refractory lithophile elements. They thus should not be fractionated from one another by either volatilization or metal-silicate fractionation; the two primary processes responsible for causing variable composition in the chondritic meteorites that presumably represent planetary building blocks. This assumption is supported by the fact that many types of chondritic meteorites show only a very small range in Sm/Nd ratio with average 147 Sm/144Nd = 0.1960 ± 0.0024 (2σ of the population of 36 analyzed chondrites (Bouvier et al., 2008b). Not a single chondrite in the summary presented by Bouvier et al. (2008b) has a Sm/Nd ratio as high as needed to explain the elevated 142Nd/144Nd of modern terrestrial rocks. Chondrules, the igneous droplets that give chondritic meteorites their name, can have high Sm/Nd ratios (Amelin and Rotenburg, 2004), but it is not yet clear if the high Sm/Nd ratio of these chondrules results from nebular processes or equilibration with the low-Sm/Nd ratio phosphate minerals in the chondrite during parent body metamorphism. In either case, the physical mechanism that would preferentially accumulate high Sm/Nd chondrules to form the terrestrial planets is not clear. Large scale fractionation processes such as the explosive loss to space of incompatible-element-rich partial melts (Warren, 2008), or collisional erosion of incompatible-elementrich crusts on small planetesimals (Agnor and Asphaug, 2004; O'Neill and Palme, 2008) have been suggested as mechanisms to preferentially remove low Sm/Nd ratio material from the building blocks of the planets. Whether such mechanisms are physically plausible remains to be demonstrated. 142

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Fig. 6. Composition of various terrestrial geochemical reservoirs. Continental crust composition from Rudnick and Gao (2003). Composition of the other reservoirs are from the mass balance modeling of Boyet and Carlson (2005, 2006) and Carlson and Boyet (2008). Numbers after each reservoir show the percentage of the volume of BSE occupied by that reservoir.

assumptions (Boyet and Carlson, 2005; Carlson and Boyet, 2008), but its composition depends strongly on its size, for which there are few constraints. One important characteristic of the hidden reservoir, however, is that it must be enriched in incompatible trace elements compared to even chondritic estimates of the primitive mantle (Fig. 6). In particular, its abundances of U, Th and K would produce something like 9 TW of present day heat generation (Boyet and Carlson, 2005). If this layer were small, for example the D″ layer at the base of the mantle, then this amount of heat generation likely would cause extensive melting. The D″-layer shows areas with exceptionally low seismic velocities that have been interpreted as indicating the presence of as much as 2 to 15% melt (Lay et al., 2004). For these liquids to remain in the D″ for any length of time, they must be either neutrally buoyant or denser than surrounding solids. Recent investigations support the high-density of melts at pressures corresponding to the D″ layer (Mosenfelder et al., 2007). If Earth were largely molten during its early history, the downward fractionation of dense incompatible-element-rich melts and their sequestration in the deepest parts of the mantle could well explain why the portion of the mantle sampled by melting near Earth's surface has a higher Sm/ Nd ratio than the bulk-planet (Labrosse et al., 2007). A dense, inviscous, melt at the base of the mantle also would explain why this material has not been re-entrained into mantle convention (Namiki, 2003; Zhong and Hager, 2003), eventually to be sampled in surface volcanism. 5. Conclusions

Arguments in favor of a bulk-earth with non-chondritic Sm/Nd ratio include the observation that the Moon (Boyet and Carlson, 2007) and Mars (Caro et al., 2008) also have superchondritic Sm/Nd, and that the complementary low 142Nd/144Nd reservoir has never been observed. The first of these depends strongly on the interpretation of 142Nd/144Nd from Martian and lunar rocks. For the Moon, the data of Rankenburg et al. (2006) are consistent with a chondritic bulk Moon, whereas the lunar data of Nyquist et al. (1995) and Boyet and Carlson (2007) suggest a bulk Moon with 142Nd/144Nd ratio as high, or higher, than the modern terrestrial mantle. Unfortunately, the cause of this data discrepancy, though likely analytical, has not yet been resolved. In the model presented by Boyet and Carlson (2007), the Moon formed from an already differentiated Earth, sampling a terrestrial exterior that was as incompatible-element-depleted as the source of the high 142 Nd/144Nd Eoarchean rocks from Greenland and China. A problem with this interpretation, as it is for all models that show a close compositional similarity between Earth and Moon, is that giant impact models suggest that the Moon would form primarily from the impactor, not the Earth (Canup, 2008). Consequently, some postimpact equilibration process (e.g. Pahlevan and Stevenson, 2007) must be proposed to explain the compositional similarity of Earth and Moon. As discussed in the previous section, the Martian data suggest a terrestrial Sm/Nd ratio for Mars only in the simple model of a single Martian magma ocean that was followed by extreme trace element fractionation during the young igneous events that produced the shergottites (Caro et al., 2008). Whether a handful of related rocks that are as strongly, and multiply, differentiated as the shergottites can provide the bulk composition of their parent planet clearly depends on the correct interpretation of their petrogenetic history, which, as outlined in Section 4.2 is not simple. The lack of expression of the complementary early-formed incompatible-element-enriched reservoir in terrestrial rocks is a significant concern if the bulk-earth indeed is chondritic. “Hidden” reservoirs, however, are a common construct to explain geochemical data; for example the core is routinely used to explain the deficiency of siderophile elements in the mantle. The characteristics of the hidden reservoir required to balance the superchondritic 142Nd/144Nd of the outer Earth can be calculated through a variety of mass balance

Improvements in the ability to determine isotope ratios precisely over the last decade have opened up new avenues to investigate the first few hundred million years of the evolution of the terrestrial planets. The results for the Moon, Mars and Earth reveal that this early period of planetary differentiation may have had profound implications for compositional variation in planetary interiors with consequent implications for the continuing differentiation of the planet. Given the short-time for which the evidence for early planetary differentiation has been available, and the limited information inhand for the physical properties of silicate minerals and melts at great depth in a planet, many questions concerning global-scale differentiation through extensive melting during planet formation remain to be solved. The ability of short-lived radioisotope systems to peer through 4.5 billion years of planet differentiation is opening an era of investigation that hopefully will transform our understanding of the processes of early planetary differentiation from broad-brush cartoons to more detailed models of the physical and chemical processes of early differentiation and the role that they played in influencing the current characteristics of the terrestrial planets. Acknowledgements Detailed reviews by Bernard Bourdon, Klaus Mezger and an anonymous reviewer are much appreciated as are the editorial comments and handling by Al Hofmann. RWC acknowledges support from the NASA Cosmochemistry Program for some aspects of the work described in this paper. References Agnor, C.B., Asphaug, E., 2004. Accretion efficiency during planetary collisions. Astrophys. J. 613, L157–L160. Amelin, Y., 2008. U–Pb ages of angrites. Geochim. Cosmochim. Acta 72, 221–232. Amelin, Y., Rotenburg, E., 2004. Sm–Nd systematics of chondrites. Earth Planet. Sci. Lett. 223, 267–282. Amelin, Y., Krot, A.N., Hutcheon, I.D., Ulyanov, A.A., 2002. Lead isotopic ages of chondrules and calcium–aluminum-rich inclusions. Science 297, 1678–1683. Andreasen, R., Sharma, M., 2006. Solar nebula heterogeneity in p-process samarium and neodymium isotopes. Science 314, 806–809.

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Richard Carlson obtained a PhD in earth science at the Scripps Institution of Oceanography in 1980 and then joined Carnegie Institution of Washington's Department of Terrestrial Magnetism as a staff scientist. His specialty is the field of trace element and isotope geochemistry and geochronology with research interests in: timescales and mechanisms of crust formation and mantle differentiation on the terrestrial planets; origin of large-volume volcanism; characteristics of sub-continental mantle and its role in continent formation and preservation; formation of the differentiated meteorites; radiometric dating techniques; and techniques for high precision chemical and isotope analysis.

Maud Boyet joined Carnegie Institution of Washington's Department of Terrestrial Magnetism in 2003 as Carnegie Fellow following a PhD in earth and planetary sciences undertaken at the Ecole Normale Superieure de Lyon (France). She is now working in the “Laboratoire Magmas et Volcans” in Clermont-Ferrand (France). Her research focuses on understanding early silicate history of planetary bodies using geochemistry and more specifically short-lived radionuclides.