Meteorites and the Timing, Mechanisms, and Conditions of ... · 10 m.y. These estimates are based...

Halliday and Kleine: Meteorites and Planetary Accretion and Differentiation 775

775

Meteorites and the Timing, Mechanisms, and Conditions ofTerrestrial Planet Accretion and Early Differentiation

Alex N. HallidayUniversity of Oxford

Thorsten KleineEidgenössische Technische Hochschule (ETH) Zentrum

Isotopic studies of meteorites provide the fundamental data for determining how the terres-trial planets, including Earth, accreted. Over the past few years there have been major advancesin our understanding of both the timescales and processes of terrestrial planet accretion, largelyas a result of better definition of initial solar system abundances of short-lived nuclides andtheir daughters, as determined from meteorites. Cosmogenic effects, cross calibrations with otherisotopic systems, and decay constant uncertainties are also critical in many instances. Recentimprovements to 182Hf-182W chronology in all these areas have been particularly noteworthy.Uncertainty has surrounded the initial Hf- and W-isotopic composition of the solar system andthe meaning of the unradiogenic W-isotopic compositions of iron meteorite data, rendered com-plicated by cosmogenic effects. However, the timescales for the formation of certain iron me-teorite parent bodies would appear to be very fast (<1 m.y.). Similarly, the accretion of Marswould appear to have been very fast, consistent with rapid accretion via runaway growth. Pre-cise quantification is difficult, however, because martian meteorites display variable W-isotopiccompositions that relate in part to the levels of depletion in siderophile elements. This is asexpected from a planet that never achieved a well-mixed silicate reservoir characterized byuniform siderophile-element depletion, as is found on Earth. Therefore, attempts to apply W-isotopic models to martian meteorites need to be treated with caution because of this demon-strable variability in early source Hf/W presumably resulting from partial metal or core segre-gation. The current best estimates for martian reservoirs as represented by Zagami would implyformation within the first 1 m.y. of the solar system. The modeled timescales for metal segrega-tion from the source of other meteorites, Nakhla, for example, would appear to be more like10 m.y. These estimates are based on trace-element and isotopic data obtained from different andprobably unrepresentative aliquots. Further high-quality combined trace-element and isotopicstudies are needed to confirm this. Nevertheless, the chondritic 142Nd abundance for Zagami pro-vides powerful supporting evidence that the W-isotopic effects record extremely rapid (<1 m.y.)accretion and core formation on Mars. The timescales for Earth accretion are significantly moreprotracted. The last major stage of accretion is thought to be the Moon-forming giant impact,the most recent Hf-W age estimates for which are in the range 40–50 m.y. after the start of thesolar system. Applying this to accretion models for Earth provides evidence that some of theaccreted metal did not fully equilibrate with silicate reservoirs. This cannot explain the very lateapparent accretion ages deduced from other chronometers, in particular U-Pb. Either all theestimates for the Pb-isotopic composition of the bulk silicate Earth are in error or there wassome additional late-stage U/Pb fractionation that removed Pb from the silicate Earth. If the latterwas the case there was either late segregation of Pb to the core after W removal, or removal ofPb via atmospheric escape following the “giant impact.” Changes in the mechanisms and parti-tioning associated with core formation are indeed predicted from the stability in the mantle ofS-rich metal before, and sulfide after, the giant impact. However, losses from Earth also need tobe evaluated. Strontium-isotopic data provide evidence of major late (>10 m.y.) losses of mod-erately volatile elements from the material that formed the Moon and probably Earth. The Earth’snonchondritic Mg/Fe may similarly reflect silicate losses during growth of Earth itself or theprotoplanets that accreted to Earth. The budgets for plutonogenic Xe provide evidence that someerosion was extremely late (>100 m.y.), clearly postdating the giant impact and presumably re-lated to irradiation and bombardment during the Hadean.

776 Meteorites and the Early Solar System II

1. INTRODUCTION

For the past 50 years it has been clear that meteoritesprovide the crucial evidence that permits determination ofthe formation times of the terrestrial planets and tests of dy-namic models for their growth and early differentiation. Themost remarkable example of this was the classic experimentby Patterson (1956), in which the age of Earth was deter-mined by measuring the Pb-isotopic compositions of ironmeteorites and achondrites. Prior to this time, Houtermans(1946) and Holmes (1946) had both independently shownthat Pb isotopes could determine the age of Earth. How-ever, at that time they had to make their best estimates ofthe minimum age (roughly 3 Ga) on the basis of the Pb-isotopic compositions of terrestrial samples measured byNier et al. (1941). Although it was Houtermans who in thiscontext first coined the term “isochron,” it was Pattersonwho, by measuring meteorites, got the right points and slopethat defined the age of the solar system and Earth.

In this review the latest pertinent information on thevarious isotopic chronometers as obtained from meteoritesis summarized and it is explained how this yields constraintson the age and early evolution of the terrestrial planets.Some of the ways in which various isotopic systems can bejointly utilized are discussed. The chapter finishes with thebest current estimates of the timescales for the formationof the terrestrial planets and asteroid belt, closing with somespeculations on how the use of these isotopic systems maydevelop in the future.

2. CONSTRAINTS ON PLANET FORMATIONPROCESSES AND TIMESCALES

The timescales and processes associated with planetaryaccretion can be studied with three complementary ap-proaches.

1. Theoretical calculations and dynamic simulations pro-vide clear indications of what is expected in terms of thetimescales and mechanisms of accretion and core formationgiven certain assumptions about the prevailing conditions(see Chambers, 2004, for a recent review). However, thedependence on the amount of gas present, the potential forplanet migration since accretion, and the difficulty in deriv-ing a clearly workable mechanism to build objects that areof sufficient size for gravity to play a dominant role leaveconsiderable uncertainty about the degree to which thesemodels represent a real solar system. There is still no un-derstanding of how circumstellar dust sticks and eventuallybecomes kilometer-scale objects that can accrete with oneanother gravitationally (Blum, 2000; Benz, 2000), and/orwhether gravitational instability may play a role (Ward,2000).

2. Observations can now be made of circumstellar disks(McCaughrean and O’Dell, 1996) as well as extrasolarplanets (Mayor and Queloz, 1995). However, the extrasolarplanets found so far are almost exclusively limited by the

detection methods to Jupiter-sized objects orbiting with avery short period. Young stars that are the most interestingfrom the point of view of planet formation are sufficientlyenergetic and internally unstable that they are largely un-suitable for detecting the small Doppler shift effects usedso far to identify most extrasolar planets.

3. Isotopic measurements have long provided key evi-dence for the processes, rates, and timing of formation ofinner solar system objects. Most of this work has been re-stricted to studies of planetesimals. However, with the de-velopment of multiple-collector inductively coupled plasmamass spectrometry (MC-ICPMS) (Halliday et al., 1995),this field has taken off in several new directions involvingshort-lived isotopic chronometers, including the determina-tion of accretion rates for the terrestrial planets. It now ap-pears that calculated timescales differ depending on thechronometer. These differences may be providing us withimportant new information about how planets are built, asdiscussed below.

3. THE MAIN ISOTOPIC SYSTEMSRELEVANT TO PLANET FORMATION

The kinds of isotopic systems that could conceivably beused for studying how planets form can be crudely catego-rized as follows:

1. Long-lived radioactive decay systems have mainlybeen used to provide information on the absolute age (Pat-terson, 1956; Hanan and Tilton, 1987; Wasserburg et al.,1977a; Lugmair and Galer, 1992; Carlson and Lugmair,2000; Amelin et al., 2002). Such systems also can provideconstraints on the time-averaged chemical composition ofthe materials from which an object formed by defining theradioactive parent element/radiogenic daughter element ra-tio (Gray et al., 1973; Wasserburg et al., 1977b; Hanan andTilton, 1987; Halliday and Porcelli, 2001).

2. Short-lived nuclides have been extremely successfulas high-resolution chronometers of events in the early so-lar system, in particular 26Al (Lee et al., 1977), 53Mn (Birckand Allègre, 1988; Lugmair and Shukolyukov, 1998), 107Pd(Chen and Wasserburg, 1996), 129I (Reynolds, 1960), and182Hf (Harper et al., 1991a; Lee and Halliday, 1995). How-ever, they need to be “mapped” onto an absolute timescale.They too can provide time-averaged chemical compositions(Halliday et al., 1996; Halliday, 2004; Jacobsen and Har-per, 1996).

3. Nucleosynthetic isotopic heterogeneity in the diskfrom which the planets form has been proposed on the basisof O-isotopic data (Clayton, 1986, 1993; Clayton and May-eda, 1975, 1996), although the currently favored mecha-nism for production of mass-independent anomalies in Oisotopes is by photochemical means either in the solarnebula or in the precursor molecular cloud (Clayton, 2002;Yurimoto and Kuramoto, 2004). Whatever their cause, O-isotopic data provide a tool for studying the provenance ofearly solar system materials and the precision on these


measurements, and hence the ability to resolve small dif-ferences in average provenance has moved to new levels(Wiechert et al., 2001, 2004). Mass-independent isotopiceffects of plausible nucleosynthetic origin increasingly arebeing demonstrated as very high precision measurementsof a wide range of elements have become possible usingMC-ICPMS and thermal ionization mass spectrometry(TIMS) (Dauphas et al., 2002a,b, 2004; Yin et al., 2002a;Schönbächler et al., 2003; Podosek et al., 1997; Trinquieret al., 2005).

4. Isotopic heterogeneity in highly volatile elements likethe noble gases provides clues to the mixture of nebular anddust-derived volatile phases (Wieler et al., 1992) and havebeen crucial for understanding the timescales of volatileaccretion and loss (Porcelli et al., 2001; Porcelli and Pepin,2000).

5. Mass-dependent stable-isotope fractionation of lightelements such as H, C, N, O, Mg, Si, S, and K has beenthoroughly characterized in recent years. Such fraction-ations or lack thereof provide powerful constraints on theconditions and processes associated with planetary accre-tion and development (Young, 2000; Humayun and Clayton,1995). Evidence for large isotopic fractionations associatedwith volatile loss in the early solar system has been foundin the gases Ne and Xe (Hunten et al., 1987; Pepin, 1997,2000). This field is likely to expand rapidly and becomeincreasingly interesting as a wide range of elements fromLi to U become amenable to study at high precision withnew MC-ICPMS techniques (Galy et al., 2000; Poitrassonet al., 2004).

6. Cosmic-ray exposure ages do not provide much di-rect evidence for how the planets formed. They do provideevidence of transit times of material delivered to Earth (e.g.,Heck et al., 2004) and it has occasionally been proposedthat some isotopic effects in chondrites reflect primordialpreexposure on a meteorite parent-body regolith.

7. Local irradiation phenomena in the early Sun alsoprovide a mechanism for generating isotopic heterogene-ity and a local source of short-lived nuclides in the earlysolar system (Lee et al., 1998; Gounelle et al., 2001; Leyaet al., 2000, 2003). Studies of some of the isotopic varia-tions predicted should shed new light on how material thatformed close to the Sun was scattered across the disk (forexample, by X-winds).

This paper mainly focuses on the conclusions one de-rives from the first two of these: long- and short-lived (ex-tinct) nuclides.

4. CALIBRATION OF CHRONOMETERSFROM THE START OF THE SOLAR SYSTEM

4.1. Long-lived Chronometers

The definition of events in absolute time at the start ofthe solar system is accomplished using long-lived radionu-clides (Table 1). Although several of these have been uti-

lized, the precision and accuracy that can be achieved with238/235U-207/206Pb is so much greater that the results of everyother kind of chronometry, short- or long-lived, are refer-enced to the data derived from this system.

The 238/235U-207/206Pb system is very reliable and wellunderstood. The uncertainties on the decay constants areof potential importance in defining an absolute age (Schönet al., 2004) but are small and insignificant when it comesto resolving time differences at the start of the solar sys-tem. There does exist uncertainty concerning the amountof 247Cm in the early solar system (Stirling et al., 2005),and because this decays to 235U with a long characteristichalf-life of 12 m.y. it could potentially change early solarsystem 238/235U-207/206Pb chronometry. However, the 247Cmabundance is probably very small (Stirling et al., 2005) andso at present there exists no case for a correction.

Other long-lived isotopic systems have, in general terms,been more problematic. The greatest difficulties have beenwith 87Rb and 176Lu. In both cases it is found that to achieveconcordance with isotopic ages obtained by 238/235U-207/206Pbor 147Sm-143Nd in meteorites, a different decay constantfromthat obtained by counting experiments or by comparing theages of terrestrial samples is required. The reasons for thisare unclear at present and seem to be different in the twocases. One should also take note of the recommendationsof Steiger and Jäger (1977). Some of the recent measure-ments relevant to 176Lu are found in Scherer et al. (2001),Blichert-Toft et al. (2002), Bizzarro et al. (2003), and Söder-lund et al. (2004). A recent compilation and discussion ofthis very topical issue can be found in Scherer et al. (2003).

These decay constant uncertainties of a few percent ren-der as limited the usefulness of the decay schemes for de-fining a precise absolute timescale. However, such problemsare of lesser consequence when it comes to determiningsmall time differences between early objects and events. Adecay constant error of 1% produces a huge uncertainty inabsolute time at the start of the solar system at 4.5 Ga thatrenders it of little use given that the solar system was prob-ably completed in <50 m.y. However, this same percentageuncertainty also translates into a similar fractional error ontime differences, where it is of negligible importance. Thesimplest and most relevant example is for long-lived chro-nometers where the time difference Dt between t1 and t2 is<<1/λ (the inverse of the decay constant or mean life) andcan therefore be simplified to

∆t = λ × (Dt1 – Dt2

)/P

where D is the atomic ratio of a radiogenic daughter to anonradiogenic stable isotope of the same element and P isthe ratio of the radioactive parent nuclide to the same stabledenominator isotope of the daughter element. It can be seenthat the fractional error on λ translates linearly to a fractionalerror on ∆t. As such, Sr-isotopic compositions, for example,can be used to infer time differences of less than a millionyears at the start of the solar system for which a decay con-


stant error of 1% is completely negligible (Gray et al., 1973;Papanastassiou and Wasserburg, 1976; Wasserburg et al.,1977b; Lugmair and Galer, 1992; Halliday and Porcelli,2001).

4.2. Uncertainties in Short-lived Systems —Decay Constants

The relative importance of uncertainties for short-livedsystems is somewhat different. The errors on the decay con-stants are generally insignificant but with increasing preci-sion and greater opportunity for cross-calibration these be-come critical. The 182Hf decay constant is a particularlygood (or bad) case in point and has been the most prob-lematic in recent years when it comes to early solar systemtimescales. Unfortunately, the artificial production of sig-nificant 182Hf for activity measurements requires very largeneutron fluxes. As a consequence, little work has been doneuntil recently. The half-life of 9 m.y. has long been limitedby an uncertainty of more than ±20% (Wing et al., 1961).This stated uncertainty does not affect first-order conclu-sions regarding the accretion rates of the planets but has

started to become significant when resolving time differ-ences in the early solar system and comparing results be-tween chronometers (Halliday, 2003, 2004). Of greater con-cern perhaps is that some half-life determinations have beeninaccurate by far outside their stated uncertainties, as wasthe case with 60Fe until the work of Kutschera et al. (1984).The new study by Vockenhuber et al. (2004) addresses the182Hf issue and has produced a new, precise, and highly re-liable value of 8.904 ± 0.088 m.y. for the 182Hf decay con-stant. Therefore, this issue can now be considered closed.

4.3. Uncertainties in Short-lived Systems —Initial Abundances

The uncertainties in the bulk solar system initial (BSSI)isotopic composition of radioactive parent or radiogenicdaughter elements introduce further and frequently the larg-est error in precise calibration of the timescale. All the short-lived nuclide systems suffer from this problem but it hasbeen particularly apparent for 26Al-26Mg (Bizzarro et al.,2004; Young et al., 2005), 53Mn-53Cr (Lugmair and Shukol-yukov, 1998; Birck et al., 1999; Nyquist et al., 1997, 2001),

TABLE 1. Main isotopic decay systems in use in studyingthe origin and early evolution of the terrestrial planets.

Parent Daughter(s) Half-Life (yr) Status* Principal Applications and Comments

Long-lived40K (10.7%) 40Ar 1.25 × 109 ✓ Chronology, especially lunar bombardment40K (89.3%) 40Ca 1.25 × 109 ✓ Not used greatly87Rb 87Sr 4.75 × 1010 ✓ Chronology, refractory/volatile fractionations138La (66.4%) 138Ba 1.05 × 1011 ✓ Not used greatly138La (33.6%) 138Ce 1.05 × 1011 ✓ Not used greatly147Sm 143Nd 1.06 × 1011 ✓ Chronology and early silicate differentiation176Lu 176Hf 3.57 × 1010 ✓ Chronology and early silicate differentiation187Re 187Os 4.23 × 1010 ✓ Chronology of metals190Pt 186Os 6.5 × 1011 ✓ Not used greatly232Th 208Pb 1.40 × 1010 ✓ Time-averaged Th/U235U 207Pb 7.04 × 108 ✓ Chronology, refractory/volatile fractionations238U 206Pb 4.47 × 109 ✓ Chronology, refractory/volatile fractionations

Extinct7Be 7Li 1.459 × 10–1 ? CAI formation10Be 10B 1.5 × 106 ✓ CAI formation26Al 26Mg 7.3 × 105 ✓ CAIs, chondrules, early silicate melting, heat41Ca 41K 1.04 × 105 ✓ CAI formation53Mn 53Cr 3.7 × 106 ✓ Early silicate melting and provenance60Fe 60Ni 1.49 × 106 ✓ Massive star signature, early metals, heat92Nb 92Zr 3.6 × 107 ✓ p-process signature, disk heterogeneity93Zr 93Nb 1.53 × 106 ? Daughter is monoisotopic97Tc 97Mo 2.6 × 106 ? p-process signature, disk heterogeneity98Tc 98Mo 4.2 × 106 ? s-process signature, disk heterogeneity99Tc 99Ru 2.13 × 105 ? s-process signature, disk heterogeneity107Pd 107Ag 6.5 × 106 ✓ Metal and refractory/volatile chronology126Sn 126Te 2.35 × 105 ? r-process signature129I 129Xe 1.57 × 107 ✓ CAI, chondrule chronology, degassing135Cs 135Ba 2.3 × 106 ? Refractory/volatile fractionations146Sm 142Nd 1.03 × 108 ✓ Early solar system, crustal evolution182Hf 182W 8.9 × 106 ✓ r-process, metal, accretion chronology205Pb 205Tl 1.5 × 107 ✓ s-process signature244Pu 136Xe† 8 × 107 ✓ CAI, chondrule chronology, degassing247Cm 235U 1.6 × 107 ? Supernova r-process

*A check mark means that the nuclide’s presence or former presence has been established. Nuclides with a question mark arestill not convincingly demonstrated.

† Spontaneous fissionogenic product.


60Fe-60Ni (Shukolyukov and Lugmair, 1993; Mostefaoui etal., 2004), 92Nb-92Zr (Harper et al. 1991b; Hirata, 2001;Münker et al., 2000; Sanloup et al., 2000; Yin et al., 2000;Schönbächler et al., 2002, 2003), 107Pd-107Ag (Chen andWasserburg, 1996; Carlson and Hauri, 2001; Woodland etal., 2005), 182Hf-182W (Harper and Jacobsen, 1996; Jacob-sen and Harper, 1996; Lee and Halliday, 1995, 1996,2000a;Kleine et al., 2002; Quitté et al., 2000; Quitté and Birck,2004; Schönberg et al., 2002; Yin et al., 2002b), and 247Cm-235U (Arden, 1977; Blake and Schramm, 1973; Chen andWasserburg, 1980; Tatsumoto and Shimamura, 1980; Stir-ling et al., 2005). In the first three cases (26Al-26Mg, 53Mn-53Cr, and 60Fe-60Ni) the uncertainty is largely dominated bystudying different archives and issues having to do with thehandling and interpretation of the data. In the case of 92Nb-92Zr and 107Pd-107Ag there are apparent discrepancies in thedata produced from different laboratories that have not beensatisfactorily explained. In the case of 247Cm-235U there isa clear indication that the U-isotopic variations reported insome studies were probably analytical artifacts (Chen andWasserburg, 1980; Stirling et al., 2005). In the case of 182Hf-182W there are both discrepancies in the data and differencesin interpretation. However, it is now known that certainearly W-isotopic data for carbonaceous chondrites (Lee andHalliday, 1995, 1996) are definitely incorrect by about 180–200 ppm. The discrepancies over 107Pd-107Ag and 247Cm-235U are of little relevance to planetary timescales at thepresent time and will not be discussed further here. In thefollowing the uncertainties over the other systems are de-scribed briefly.

4.3.1. Aluminum-26–magnesium-26. The canonicalvalue of (26Al/27Al)BSSI is 5 × 10–5 (Lee et al., 1977). How-ever, new Mg-isotopic data for bulk CAIs and minerals haveled Young et al. (2005) to propose that this only representsthe isotopic abundance at the time of recrystallization of theCAIs. Young et al. present a large amount of new data forCAIs measured by laser ablation MC-ICPMS. From thescatter of data for minerals and bulk CAIs they extract aminimum apparent (26Al/27Al)BSSI ~ 7 × 10–5 using theupper portion of the scatter to the Al-Mg isochron. Thedifficulty with the approach is that nonsystematic and ma-trix-sensitive errors in element ratios can be quite signifi-cant using laser ablation. These are hard to quantify in fine-grained materials, particularly in the absence of detailedmatrix matching and laser artifact studies. As a result themeaning of the upper limit of such a scattered dataset be-comes ambiguous. A few bulk CAIs that were analyzedconventionally in the same study lie within the range of thelaser ablation measurements but do not extend to such high26Al/27Al. The Young et al. study appears to be in conflictwith the very precise (26Al/27Al)BSSI = (5.25 ± 0.10) × 10–5

obtained for bulk Allende CAIs by Bizzarro et al. (2004).The reasons for the apparent difference are complex. Thetwo groups have analyzed different CAIs, in some casesfrom different meteorites. However, they also treat the cor-rection for isotopic fractionation differently. The raw dataare distorted by both natural and instrument-induced proc-

esses following different laws. With only three isotopes, oneof which is radiogenic, a judgment has to be made abouthow best to determine the radiogenic difference. This be-comes important in strongly fractionated residues like bulkCAIs and has the potential to leave an apparent residual iso-topic anomaly that is significant relative to the small effectbeing expected from low Al/Mg CAIs. Davis et al. (2005)argue that the high 26Al/27Al reported by Young et al. (2005)could reflect inaccurate characterization of this natural frac-tionation. Young et al. claim that on its own, this is insuffi-cient to explain the total amount of the apparent discrep-ancy with the results reported by Bizzarro et al. (2004).However, Bizzarro et al. (2005) have now reported an errorin their treatment of their data resulting in a corrected 26Al/27Al = (5.83 ± 0.11) × 10–5. This is consistent with the con-ventional bulk data in Young et al. (2005). In fact, regres-sion of all data from Young et al. gives an identical, althoughless-precise, 26Al/27Al of 5.9 ± 0.3 × 10–5. However, thevery high “supracanonical” (26Al/27Al)BSSI value of 7 × 10–5

is determined as an upper limit to the scatter of the dataobtained by laser ablation and has not yet been ratified byany conventional measurements.

4.3.2. Manganese-53–chromium-53. The former pres-ence of live 53Mn in the early solar system is now well es-tablished (Birck and Allègre, 1988; Lugmair and Shukol-yukov, 1998). However, the study of different meteoritearchives has yielded different values for the initial abun-dance and this has a dramatic effect on the timescales de-duced for Mn/Cr fractionation in planetary bodies. The Mn-Cr-isotopic data for CAIs (Birck and Allègre, 1988) areconsistent with a solar system initial 53Mn/55Mn that is sig-nificantly higher than that implied from the Mn-Cr system-atics of angrites and eucrites (Lugmair and Shukolyukov,1998, 2001). The data for CAIs could conceivably reflectnucleosynthetic effects that are well established in chon-drites (Podosek et al., 1997) and/or disturbance of Mn-Cr-isotopic systematics (Papanastassiou et al., 2005). Manga-nese-chromium chronometry is left in a somewhat equivo-cal status until this is clarified. The ramifications for earlyplanetesimal accretion are discussed further below.

4.3.3. Iron-60–nickel-60. The demonstration of for-merly live 60Fe in the early solar system (Shukolyukov andLugmair, 1993) was accomplished by measuring the Ni-iso-topic compositions of basaltic achondrites and in particulareucrites. Recently other groups have demonstrated very high60Fe/56Fe in some sulfides from chondrites (e.g., Mostefaouiet al., 2004). The recently reported evidence for very high60Fe/56Fe (Moynier et al., 2004) in iron meteorites has notbeen confirmed by more recent measurements (Cook et al.,2005; Quitté et al., 2006). Exactly how high is the value of(60Fe/56Fe)BSSI is now unclear but it could be as high as10–6

in which case early solar system events recorded in eucritesare more protracted than previously thought or the Ni-iso-topic systems have been partially reset at some late stage.

4.3.4. Niobium-92–zirconium-92. In the case of 92Nbthere exist basically two values. The first is (92Nb/92Nb)BSSI ~10–3 (Münker et al., 2000; Sanloup et al., 2000; Yin et al.,


2000). The second is (92Nb/92Nb)BSSI ~ 10–5 (Harper et al.,1991b; Hirata, 2001; Schönbächler et al., 2002, 2003,2005). These two very distinct estimates result in radicallydifferent constraints on the early differentiation history ofthe terrestrial planets. If the former is correct the Zr-isoto-pic compositions for the silicate Earth imply late formationof Earth’s core and continental crust (Münker et al., 2000;Jacobsen and Yin, 2001). The reasons for the differencesare unclear. Schönbächler et al. (2003) were unable to re-produce the Zr-isotopic variations reported by Münker etal. (2000) and Sanloup et al. (2000) when similar materi-als (the same bulk chondrites and similar refractory low Nb/Zr Allende CAIs) were analyzed. Similarly, the results foriron meteorite rutiles reported by Yin et al. (2000) differfrom those reported by Harper et al. (1991b) and the re-sults for an early (meteorite) zircon with Nb/Zr = 0 reportedby Yin et al. (2000) are different from the results for similarphases studied by Hirata (2001). These apparent discrep-ancies warrant further studies to settle the issue definitively.

4.3.5. Hafnium-182–tungsten-182. In the case of thevalue of (182Hf/180Hf)BSSI the initial estimate of Harper andJacobsen (1996) and Jacobsen and Harper (1996) of ~10–5

was low because of a paucity of data (Ireland, 1991) andthe modeling available from comparisons with other isoto-pic systems (Wasserburg et al., 1994). Lee and Halliday(1995, 1996) presented the first W-isotopic data for carbo-naceous chondrites and inferred a (182Hf/180Hf)BSSI > 10–4.They then appeared to confirm this with a value of 182Hf/180Hf ~ 2 × 10–4 for internal Hf-W isochrons of rather vari-able quality for the ordinary chondrites Forest Vale and Ste.Marguerite (Lee and Halliday, 2000b).

Three groups have independently shown that some ofthese University of Michigan data are incorrect (Kleine etal., 2002; Schönberg et al., 2002; Yin et al., 2002b). Theyhave instead demonstrated a (182Hf/180Hf)BSSI ~ 1 × 10–4.The most clear-cut error is in the compositions of the car-bonaceous chondrites Allende and Murchison. These, andall carbonaceous chondrites measured thus far (Kleine etal., 2002, 2004a), have a 182W/184W that is 180–200 ppmlower than that originally determined by Lee and Halliday(1995, 1996) (Fig. 1).

The reason for the error in the Michigan measurementsis unclear. The MC-ICPMS instrument used was the firstbuilt and had relatively poor sensitivity compared with morerecent instruments. This should not matter directly unlessit resulted in a background problem.

The original internal isochrons for Forest Vale and Ste.Marguerite (Lee and Halliday, 2000b) also differ in partfrom the results presented by Kleine et al. (2002). The ex-planation could again be analytical error. There may alsohave been differences between the particular phases ana-lyzed. Clearly, the age or time of closure of a particularphase is critical to the 182Hf/180Hf determined. However,the data from CAIs and other ordinary chondrite isochrons(Yin et al., 2002b) and the data obtained on bulk meteor-ites all appear to confirm a (182Hf/180Hf)BSSI of about 1 ×10–4, rather than 2 × 10–4.

The W-isotopic composition of iron meteorites may,however, provide evidence that the initial abundance liesbetween these two values. Iron meteorites sample early solarsystem W without Hf and therefore provide an indicationof the W-isotopic composition at the start of the solar sys-tem in an analogous fashion to the way in which the initialPb-isotopic composition at the start of the solar system hasbeen defined by Canyon Diablo troilite. The most recentestimate of (182Hf/180Hf)BSSI is (1.7 ± 0.3) × 10–4 based onthe least-radiogenic W-isotopic composition of solar sys-tem materials yet determined, as measured in the iron me-teorite Tlacotopec (Quitté and Birck, 2004). This particularmeteorite has a very long cosmic-ray exposure age and theeffects on the W-isotopic composition (Leya et al., 2000,2003) could be significant. This has been confirmed bymore detailed high-precision W-isotopic measurements ofiron meteorites (Markowski et al., 2006). Although the dif-ferences over the BSSI abundance of Hf and W isotopeswould appear to be minor compared with the uncertaintiesand disputes over other isotopic systems, in particular 53Mn-53Cr, 60Fe-60Ni, and 92Nb-92Zr, they do have a nonnegligibleeffect on calculated early solar system timescales and theirinterpretation, and are particularly important because of thewider utility of 182Hf-182W.

Fig. 1. The W-isotopic compositions of carbonaceous chondrites,as determined by Kleine et al. (2002, 2004a), shown here, but alsoby Yin et al. (2002b) and Schoenberg et al. (2002), demonstratethat the earlier measurements of Allende and Murchison (Lee andHalliday, 1995, 1996) are inaccurate by roughly 180 to 200 ppm.The new values agree within error with the values determined forenstatite chondrites (Lee and Halliday, 2000a) shown here but pre-viously considered anomalous and hard to explain. The averagefor carbonaceous chondrites (vertical line) is ε182W = –1.9 ± 0.1.


4.4. Uncertainties in Short-lived Systems —Conversion to Absolute Time

To convert the results from short-lived decay systems likeHf-W into an absolute timescale one needs cross-calibrationpoints. In principle, cross-calibrations with precise 235/238U-207/206Pb ages should allow one to determine the initial abun-dances and the decay constants of the short-lived nuclidesindependently. There now exist at least four of these, all ofwhich can be tied in to absolute timescales with the veryprecise 235/238U-207/206Pb ages shown in Table 2.

1. The 26Al-26Mg chronometer yields time differences(Russell et al., 1996) that are fully consistent with differ-ences in absolute time deduced from 235/238U-207/206Pb datafor CAIs and chondrules (Amelin et al., 2002).

2. The very precise 235/238U-207/206Pb data for angrites(Lugmair and Galer, 1992) provides a powerful means ofmapping 53Mn-53Cr onto an absolute timescale (Lugmairand Shukolyukov, 1998, 2001).

3. Similarly, the 235/238U-207/206Pb data for ordinary chon-drites, in particular Ste. Marguerite (Göpel et al., 1994),allow the 182Hf-182W chronometer to be cross-calibratedonto an absolute timescale (Kleine et al., 2002, 2005b) ifthey date the same event. In fact, the U-Pb closure tempera-ture in phosphates is 400°–500°C whereas Hf-W closurebetween metal and silicates takes place at temperatures inexcess of 600°C (Kleine et al., 2005b). Using the 53Mn-53Crchronometry (Polnau and Lugmair, 2001), which itself ismapped onto an absolute timescale by cross calibration with235/238U-207/206Pb (Lugmair and Shukolyukov, 1998, 2001),

TABLE 2. Isotopic ages of early objects and best estimates of early solar system timescales.

Type of Event Object or Model Isotopic System Reference Age (Ga) Time (m.y.)

Start of solar system Efremovka CAIs 235/238U-207/206Pb Amelin et al. (2002) 4.5672 ± 0.0006 0.0 ± 0.6Start of solar system Allende CAIs 235/238U-207/206Pb Göpel et al. (1991) 4.566 ± 0.002 1 ± 2Start of solar system Allende CAIs 26Al-26Mg Bizzarro et al. (2004) 4.567 0.00 ± 0.03Chondrule formation Acfer chondrules 235/238U-207/206Pb Amelin et al. (2002) 4.5647 ± 0.0006 2.5 ± 1.2Chondrule formation UOC chondrules 26Al-26Mg Russell et al. (1996) <4.566 to 4.565 >1 to 2Chondrule formation Allende chondrule 26Al-26Mg Galy et al. (2000) <4.5658 ± 0.0007 >1.4 ± 0.7Chondrule formation Allende chondrules 26Al-26Mg Bizzarro et al. (2004) 4.567 to <4.565 0 to ≥1.4H chondrite parent body Ste. Marguerite phosphate 235/238U-207/206Pb Göpel et al. (1994) 4.5627 ± 0.0006 4.5 ± 1.2 metamorphismAsteroidal core formation Magmatic irons 182Hf-182W Markowski et al. (2006) >4.566 <1Vesta differentiation Silicate-metal 182Hf-182W Kleine et al. (2002) 4.563 ± 0.001 4 ± 1Vesta accretion Earliest age 87Rb-87Sr Halliday and Porcelli (2001) <4.563 ± 0.002 >4 ± 2Vesta differentiation Silicate-metal 182Hf-182W Lee and Halliday (1997) 4.56* 10*Vesta differentiation Silicate-silicate 53Mn-53Cr Lugmair and Shukulyokov (1998) 4.5648 ± 0.0009 1 ± 2Vesta differentiation Silicate-metal 182Hf-182W Quitté et al. (2000) 4.550 ± 0.001* 16 ± 1*Vesta differentiation Silicate-metal 182Hf-182W Kleine et al. (2002) 4.563 ± 0.001 4 ± 1Vesta differentiation Silicate-metal 182Hf-182W Yin et al. (2002b) 4.564 3Early eucrites Asuka 881394 235/238U-207/206Pb Wadhwa et al. (2005), M. Wadhwa 4.5665 ± 0.0009 0.7 ± 1.1

(personal communication, 2005)Early eucrites Noncumulate eucrites 182Hf-182W Quitté and Birck (2004) 4.558 ± 0.003 9 ± 3Early eucrites Chervony Kut 53Mn-53Cr Lugmair and Shukulyokov (1998) 4.563 ± 0.001 4 ± 1Angrite formation Angra dos Reis and 235/238U-207/206Pb Lugmair and Galer (1992) 4.5578 ± 0.0005 9 ± 1

LEW 86010Mars accretion Youngest age 146Sm-142Nd Harper et al. (1995) ≥4.54 ≤30Mars accretion Mean age 182Hf-182W Lee and Halliday (1997) 4.560* 6*Mars accretion Youngest age 182Hf-182W Lee and Halliday (1997) ≥4.54* ≤30*Mars accretion Youngest age 182Hf-182W Halliday et al. (2001b) ≥4.55* ≤20*Mars accretion Youngest age 182Hf-182W Kleine et al. (2002) ≥4.55 <13 ± 2Mars accretion Youngest age 182Hf-182W This study >4.566 <1Earth accretion Mean age 235/238U-207/206Pb Halliday (2000) 4.527 to 4.562 15 to 40Earth accretion Mean age 182Hf-182W Yin et al. (2002b) 4.556 ± 0.001 11 ± 1Earth accretion Mean age 235/238U-207/206Pb Halliday (2004) 4.550 ± 0.003 17 ± 3Moon formation Best estimate of age 235/238U-207/206Pb Tera et al. (1973) 4.47 ± 0.02 100 ± 20Moon formation Best estimate of age 235/238U-207/206Pb Carlson and Lugmair (1988) 4.44 to 4.51 60 to 130

and 147Sm-143NdMoon formation Best estimate of age 182Hf-182W Halliday et al. (1996) 4.47 ± 0.04* 100 ± 40*Moon formation Best estimate of age 182Hf-182W Lee et al. (1997) 4.51 ± 0.01* 55 ± 10*Moon formation Earliest age 182Hf-182W Halliday (2000) ≤4.52* ≥45*Moon formation Earliest age 87Rb-87Sr Halliday and Porcelli (2001) <4.556 ± 0.001 >11±1Moon formation Best estimate of age 182Hf-182W Lee et al. (2002) 4.51 ± 0.01* 55 ± 10*Moon formation Best estimate of age 182Hf-182W Kleine et al. (2002) 4.54 ± 0.01 30 ± 10Moon formation Best estimate of age 182Hf-182W Yin et al. (2002b) 4.546 29Moon formation Best estimate of age 182Hf-182W Halliday (2004) 4.52 ± 0.01 45 ± 10Lunar highlands Ferroan anorthosite 60025 235/238U-207/206Pb Hanan and Tilton (1987) 4.50 ± 0.01 70±10Lunar highlands Ferroan anorthosite 60025 147Sm-143Nd Carlson and Lugmair (1988) 4.44 ± 0.02 130 ± 20Lunar highlands Norite from breccia 15445 147Sm-143Nd Shih et al. (1993) 4.46 ± 0.07 110 ± 70Lunar highlands Ferroan noritic anorthosite 67016 147Sm-143Nd Alibert et.al. (1994) 4.56 ± 0.07 10 ± 70Earliest crust Jack Hills zircons 235/238U-207/206Pb Wilde et al. (2001) 4.44 ± 0.01 130 ± 10

*Based on early solar system initial abundances now thought incorrect.

CAIs = calcium-aluminum-rich refractory inclusions. UOC = unequilibrated ordinary chondrite.


results in better agreement between the Hf-W and Mn-Crwhole-rock ages of the eucrites and also between the Hf-W and 235/238U-207/206Pb age for CAIs (Kleine et al., 2005a).

4. Most recently, the 182Hf-182W chronometer has beencross-calibrated onto an absolute timescale by determiningan internal isochron for CAIs (Kleine et al., 2005a). Thispermits direct comparison with 235/238U-207/206Pb and 26Al-26Mg data.

These calibrations are tremendously important as thefield develops in the direction of increasing age resolution.For example, for many years there was a debate aboutwhether 26Al/27Al variations in early solar system objectsreflected time differences, variable resetting, or differencesin the abundance of 26Al in different settings. The lattermodel was fueled by more recent models of 26Al productionclose to the Sun (Lee et al., 1998) and the detection of for-merly live 10Be in CAIs (McKeegan et al., 2000). However,since 10Be is clearly spallogenic and excess 10B (from thedecay of 10Be) has been detected in CAIs lacking evidenceof excess 26Mg, the vast majority of 26Al cannot be spallo-genic (Marhas et al., 2002). The close agreement between26Al-26Mg and 235/238U-207/206Pb chronometry, as well as thevery precise initial abundance of (26Al/27Al)BSSI recentlydetermined for bulk CAIs (Bizzarro et al., 2004, 2005), ap-pears to demonstrate that any variation in initial abundanceis very small.

Although such issues mandate further effort to achieveclarity, there is no question that we now know the absoluteage of certain key objects extremely well and can get a goodidea of how quickly they formed.

5. CHONDRITES, ACHONDRITES, IRONS,AND EARLIEST PLANETESIMALS

The most precise and accurate current estimate of theabsolute age of the solar system is taken to be the 235/238U-207/206Pb age of Efremovka CAIs of 4.5672 ± 0.0006 Ga(Amelin et al., 2002). The most widely held view is thatCAIs did form within the earliest solar system. The reasonfor believing this is that they are the earliest objects yetdated, have the highest abundances of short-lived nuclides,and have roughly chondritic proportions of highly refractoryelements and isotopic compositions that are, broadly speak-ing, similar to those of average solar system.

When and how CAIs formed relative to the collapse ofthe portion of the molecular cloud that formed the Sun isunknown. The most precise estimate of the duration of CAIformation is provided by the new high-precision 26Al-26Mgdata produced by Bizzarro et al. (2004, 2005), who obtainedan identical apparent initial 26Al abundance for all CAIsmeasured, implying a time interval of perhaps as little as50,000 years for Allende CAI formation. These measure-ments need to be extended to CAIs from other meteoritesbut the preliminary data are consistent with a very short dur-ation of CAI production. As discussed above, there are ap-parent discrepancies between this result and the Mg-isoto-pic data of Young et al. (2005), who found a higher apparent

26Al/ 27Al for bulk CAIs than the canonical value based onminerals. It is currently unclear whether this difference isreal. Either way, the resolution of these chronometers is get-ting to the point at which incredible detail should be re-vealed about the earliest solar system over the coming yearsas these problems are resolved.

How chondrules form is also unknown and highly con-troversial. The most widely accepted theory is that theyform within the dusty circumstellar disk from which theplanetesimals formed. A variety of studies have now shownthat chondrule formation overlapped with CAI production,but was protracted, extending to roughly 2 m.y. after thestart of the solar system (Russell et al., 1996; Galy et al.,2000; Bizarro et al., 2004). Whether all these objects wereformed in a similar manner is unclear. Some may have beenproduced as a result of molten planetesimal impacts (Zook,1981; Sanders, 1996). Another “planetary” model for chon-drule formation involves bow shocks around eccentric as-teroids driven by jovian resonances (Weidenschilling et al.,1998). Similarly, Lugmair and Shukolyukov (2001) arguethat early planetesimals would melt rapidly and that colli-sions among them would release melt that ultimately formschondrules.

The question then arises as to how well we know howquickly the first planetesimals formed. From dynamic con-siderations it is thought that planetesimals formed within afew times 105 yr (e.g., Kortenkamp et al., 2000). Opticallydense circumstellar disks appear to form very early and lastfor about a million years or more. Some dusty disks likeHR 4796A and Beta Pictoris are an order of magnitude older(Schneider et al., 1999). The presence of these disks doesnot imply that planetesimals or planets have not yet formed.It is assumed that they are already formed and are presentin the midplane.

Chondrites themselves, while judged to be undifferenti-ated, cannot have formed very early because they containa variety of CAIs, chondrules, and presolar grains that musthave become mixed together some time after chondruleformation. Certain kinds of presolar grains could not havesurvived the temperatures and conditions of chondrule for-mation. Therefore, a scenario for chondrite-parent-body for-mation must involve mixing of CAIs, chondrules, and pre-solar grains into a particular region where chondrite parentbodies were able to form by sticking or local gravitationalcollapse of these small objects. The 26Al/27Al (Bizzarro etal., 2004, 2005; Young et al., 2005) and high 60Fe/56Fe(Mostefaoui et al., 2004) now thought to have character-ized the earliest solar system should have produced moltenand differentiated planetesimals at an early stage. The veryexistence of chondrites provides evidence of small-bodyaccretion of undifferentiated material millions of years afterthe start of the solar system. As such, their existence in theinner solar system is dynamically problematic.

Isotopic data for differentiated meteorites provide clearevidence that some of the earliest planetesimals formedwithin the first 10 m.y. of the solar system. Exactly howearly is a matter of some uncertainty. The primary lines of


evidence come from short-lived nuclides and in particular26Al-26Mg, 53Mn-53Cr, and 182Hf-182W. For example, Cher-vony Kut is an unbrecciated eucrite with relatively high 60Fe(Shukolyukov and Lugmair, 1993), lending confidence to theview that it is a less-disturbed sample well suited for defin-ing early solar system timescales. The 53Mn-53Cr data forChervony Kut can be compared with the data for the an-grites Angra dos Reis and LEW 86010 (Table 2). These ob-jects are very precisely dated by 235/238U-207/206Pb and, al-though relatively young, they permit mapping of the Mn-Crtimescale for eucrites to an absolute age. On this basis Cher-vony Kut formed within 5 m.y. of the start of the solar sys-tem, allowing for the uncertainties in Pb-Pb chronometryof the angrites and CAIs (Lugmair and Shukolyukov, 1998).The eucrite whole-rock Mn-Cr isochron, thought to reflectplanetary differentiation, defines an earlier age no more than4 m.y. after the start of the solar system (Table 2).

Despite these lines of evidence that differentiated objectsformed early, the Sr-isotopic data for eucrites are difficultto explain as the primary planetesimals that supposedlyformed within the first few hundred thousand years of thesolar system. There is a difference between the initial Sr-isotopic compositions of CAIs (Gray et al., 1973; Podoseket al., 1991) on the one hand and eucrites and angrites(Wasserburg et al., 1977; Lugmair and Galer, 1992) on theother. This cannot be readily explained unless the eucriteparent body formed more than 2 m.y. after the start of thesolar system (Halliday and Porcelli, 2001). The basis forthis is as follows. All these objects are strongly depleted inmoderately volatile Rb relative to refractory Sr and, as such,there is no possibility of the Sr-isotopic composition of theparent body or its mantle changing significantly by decayof 87Rb. Yet the Sr-isotopic difference requires a significantperiod of time in a high Rb/Sr environment. The highestRb/Sr environment possible is the Rb/Sr of the solar nebula(0.3). The minimum time required to generate the differencein Sr-isotopic composition then defines the earliest timethese objects could have formed. This time difference is atleast 2 m.y. (Halliday and Porcelli, 2001).

Basaltic achondrites are relatively rare and it could wellbe that in the harsh energetic processes associated with ac-cretion and destruction of asteroids the silicate portions ofmany primary planetesimals have been destroyed. A bettersample may be found among the many iron meteorites (Hal-liday, 2003). It has long been known from Pb-isotopic dataas well as 107Pd-107Ag chronometry (Chen and Wasserburg,1996) that iron meteorites are early objects. However, thePb-isotopic data for iron meteorites appear to be offset fromthe initial compositions defined by CAIs for reasons thatare not well understood (Tera and Carlson, 1999; Carlsonand Lugmair, 2000) and the 107Pd-107Ag system is not wellcalibrated against other chronometers (Chen and Wasser-burg, 1996). Alternative chronometers such as 129I-129Xeand 53Mn-53Cr probably date cooling rather than the pri-mary formation of these objects.

The most powerful chronometer of the accretion andmetal segregation processes associated with iron meteorite

formation is 182Hf-182W. The timing of metal segregationis well defined by the W-isotopic composition because Hfand W are both refractory and early solar system parentbodies can be safely assumed to have chondritic Hf/W (Leeand Halliday, 1996; Horan et al., 1998) (Fig. 2). As suchthe parent-body isotopic evolution is predictable and the W-isotopic composition of the iron meteorite can define thetime of metal separation to extremely high precision usingmodern MC-ICPMS techniques (Halliday, 2003, Hallidayet al., 2003; Lee and Halliday, 2003). The W-isotopic datafor iron meteorites are strikingly unradiogenic, indicatingthat they formed within a few million years of each other atthe start of the solar system (Lee and Halliday, 1996; Horan

Fig. 2. The W-isotopic compositions of iron meteorites provideunequivocal evidence of early metal segregation. Assuming theparent bodies had chondritic relative proportions of refractory Hfand W, one can calculate a model time difference between thetimes of formation of different metals. The most comprehensivedataset published so far is that of Horan et al. (1998) shown here.The vertical line is the best estimate of average solar system basedon data for carbonaceous chondrites (Kleine et al., 2002, 2004a).


et al., 1998). The absolute age can be calculated from cross-calibrations of the Hf-W-isotopic system such as that pro-vided by Ste. Marguerite, discussed above. New studies nowshow that the W-isotopic compositions of iron meteoritesyield resolvable differences at very high precision (Lee andHalliday, 2003; Markowski et al., 2004, 2006; Schersten etal., 2004; Kleine et al., 2005a; Lee, 2005).

The least-radiogenic 182W/184W reported to high precisionis ε182W = –4.4 ± 0.2 (the deviation in parts per 104 rela-tive to bulk silicate Earth), obtained for Tlacotopec (Quittéand Birck, 2004). However, this low value may in part re-flect burnout of W isotopes caused by the interaction withthermal neutrons produced during a prolonged exposure tocosmic rays (Masarik, 1997; Leya et al., 2003). Several ironmeteorites, including Tlacotopec, have very long cosmic-ray exposure ages of hundreds of millions of years. As such,their W-isotopic compositions are not expected to be pris-tine. The production rate of thermal neutrons and, hence,the cosmogenic effects on W isotopes vary with depth in ameteorite. Therefore, different W-isotopic data may be ob-tained from different aliquots of the same meteorite. Toobtain really precise estimates of the timescales of metalsegregation it may be necessary to measure exposure agesand W-isotopic compositions on the same aliquots and thento apply a correction (Markowski et al., 2006).

Even if the least-radiogenic W in iron meteorites is onlyas low as εW = –3.7, other problems are apparent. The Hf-W data for Allende CAIs (Kleine et al., 2005a) define a(182W/184W)BSSI that is higher than ε182W = –3.7, apparentlysuggesting that some iron meteorites formed at least a fewhundred thousand years before CAIs. This is contrary to thereigning paradigm that CAIs are the first solids that formedin the solar nebula. Challenging this paradigm will requirea precise quantification of the 182W-burnout effects in ironmeteorites (Markowski et al., 2006). It will also requiredemonstration that the Hf-W systematics of CAIs have notbeen preferentially reset. Given the presence of W com-pounds as Fremdlinge in Allende CAIs and the demonstra-tion that these are secondary features (Armstrong et al.,1985, 1987; Hutcheon et al., 1987), one is left with somelevel of uncertainty as to whether there has been preferen-tial disturbance of the W isotopes. To counter this, it seemsunlikely that the CAI isochron is reset for the followingreasons [a detailed discussion of this issue can be found inKleine et al. (2005a)]: (1) The isochron is based on mineralseparates from one CAI and bulk analyses from two otherCAIs. (2) Resetting of the Hf-W system would require mo-bilization of radiogenic W from silicates into metal. Thediffusion of W from silicates into metal, however, appearsto require temperatures in excess of 600°C (Kleine et al.,2005b), which were not achieved on the Allende parent as-teroid. The formation of scheelite and powellite as second-ary products from Fremdlinge as well as the mobilizationof W from Fremdlinge to surrounding silicates would notcause resetting of the isochron. (3) The 182Hf-182W age de-rived from the CAI isochron (calculated relative to the Ste.Marguerite H chondrite) is 4568.0 ± 1.7 Ma, which is iden-

tical to the U-Pb age for Efremovka CAIs (Amelin et al.,2002).

Recently, very precise W-isotopic data have been ob-tained for a large range of magmatic and nonmagmaticirons. By correcting for the maximum possible cosmogeniceffect, it can be demonstrated that some magmatic irons seg-regated within <0.5 m.y. of the start of the solar system, asdefined by W-isotopic data for Allende CAIs (Markowskiet al., 2006).

6. MARTIAN METEORITES ANDEARLIEST PLANETARY

DIFFERENTIATION PROCESSES

Martian meteorites provide the most powerful constraintson the timescales of formation of another terrestrial planet.This is not just of interest to comparative planetology. Inmany respects, Mars represents an example of how Earthand other terrestrial planets may have first started. There isevidence that a Mars-sized object sometimes referred to as“Theia” (the Greek goddess who was the mother of Selene,the goddess of the Moon) was responsible for the Moon-forming giant impact (Canup and Asphaug, 2001). Therealso is evidence that certain features of the composition ofTheia were similar to those of Mars (Halliday, 2004; Halli-day and Porcelli, 2001).

The constraints from short-lived nuclide data for mar-tian meteorites are dominated by the data for four isotopicsystems: 129I-129Xe, 146Sm-142Nd, 182Hf-182W, and 244Pu-136Xe (Table 1). Although all these data provide a similarpicture of rapid development of Mars, the comparisonsbetween W and Nd have proved most interesting. The lat-est compilation of W-isotopic data is shown in Fig. 3. It canbe seen that all W-isotopic compositions are now resolv-able from that of the silicate Earth, with radiogenic valuesin the range ε182W = 0–3. In general terms the shergottitestend to be closer to 0 (Kleine et al., 2004a; Foley et al.,2005) whereas the nahklites are closer to 3 (Lee and Halli-day, 1997). The original W-isotopic data (Lee and Halliday,1997; Kleine et al., 2004a) show a weakly defined relation-ship with the original ε142Nd (Harper et al., 1995; Jagoutzand Jotter, 2000; Jagoutz et al., 2003) that provided evi-dence that the early processes that affected each systemwere somehow correlated. Recently this has been ques-tioned (Foley et al., 2005) on the basis of new data, as dis-cussed in detail below. The processes fractionating Hf/Ware partial melting and core formation (Lee and Halliday,1997; Halliday and Lee, 1999), whereas Sm/Nd respondsonly to partial melting because both parent and daughterelements are lithophile. Lee and Halliday (1997) proposedthat this reflects contemporaneous partial melting and coreformation. However, this was based on the assumption thatthe least-radiogenic W was chondritic, as was the least-ra-diogenic Nd. Now it is known that chondrites do not haveε182W = 0 (Lee and Halliday, 1995, 1996) but ratherε182W ~ –2 (Lee and Halliday, 2000a; Kleine et al., 2002;Schoenberg et al., 2002; Yin et al., 2002b). It has been pro-


posed that there is a small (~20 ppm) offset in chondriticε142Nd as well (Boyet and Carlson, 2005). However, thisis of second-order importance for the comparison with Wisotopes. The least-radiogenic W-isotopic composition mea-sured thus far for Mars is significantly offset to higher thanchondritic values (Fig. 3). Kleine et al. (2002, 2004a) haveexplained this in terms of a period of core formation priorto silicate melting, whereas the correlation with ε142Nd re-flects the effects of silicate melting only.

More precise 182W and 142Nd data on previously stud-ied and additional martian meteorites have now been usedto argue that no such overall Nd-W correlation exists (Foleyet al., 2005). The analytical uncertainties on ε182W areconsiderably improved in the Foley et al. study. Most pre-viously published W data agree with the new data withinthe stated uncertainties except for the early report of ε182W ~2 for EETA 79001 (Lee and Halliday, 1997), which is notreplicated by both of the more recent studies (Kleine et al.,2004; Foley et al., 2005). There are also apparent discrepan-cies between the ε142Nd of two Saharan martian meteorites,DaG 476 and SAU 051 (Jagoutz and Jotter, 2000; Jagoutzet al., 2003; Foley et al., 2005). The most precise Nd- and

W-isotopic data available as of August 2005 are plotted inFig. 4. The nakhlites form a group with relatively uniformand radiogenic Nd and W (ε182W ~ 3). Chassigny is slightlyless radiogenic. The shergottites are less radiogenic in W,as already noted. However, with the inclusion of the latestdata for the two Saharan shergottites DaG 476 and SAU 051(Foley et al., 2005), there is a broad spread in ε142Nd from–0.3 to radiogenic values of ~+1.0, like those of the nakh-lites, and offset from the main correlation. The addition ofthe new data for the Saharan meteorites therefore changesthe apparent distribution somewhat. In Fig. 4 both the newFoley data and the Jagoutz ε142Nd data are plotted to illus-trate the effect. The reason for this apparent discrepancy isnot certain although, as discussed by Foley et al., the lowerε142Nd value reported by Jagoutz et al. (2003) in the Sa-haran martian meteorites could result from the presence ofa terrestrial weathering component. Until this discrepancyis clarified it is perhaps premature to venture an explanationfor the apparent offset of these two Saharan Nd data fromFig. 3. Tungsten-isotopic data for martian meteorites reveal a sig-

nificant offset relative to the newly defined average solar systemat εW = –1.9. Data from Lee and Halliday (1997), Kleine et al.(2004a), and Foley et al. (2005). In general terms the sequence ofpapers is accompanied by a dramatic improvement in precision.Where the same meteorite has been analyzed in more than onestudy, only the most precise measurements are shown here and inthe following figures. The three studies show generally good agree-ment but the imprecise value for EETA 79001 in Lee and Halliday(1997) differs significantly from the more precise measurementsreported by Kleine et al. (2004a) and Foley et al. (2005). Assum-ing that these more precise recently determined values are likelyto be correct, the data obtained for shergottites are relatively uni-form and distinct from nakhlites, as pointed out by Kleine et al.and Foley et al.

Fig. 4. The W-isotopic compositions of martian meteorites dis-play a broad relationship with their corresponding ε142Nd result-ing from decay of formerly live 146Sm. Data from Harper et al.(1995), Jagoutz et al. (2003), and Foley et al. (2005). The latest142Nd data for the two Saharan meteorites (Foley et al., 2005)differ from those reported by Jagoutz et al. (2003). The reasonsfor these discrepancies are unclear and both sets of values areshown here. Foley et al. (2005) point out that inclusion of theirnew Saharan data seriously reduces the evidence for a relation-ships between Nd- and W-isotopic effects and the proposal thatmetal segregation and silicate melting overlapped in time (Lee andHalliday, 1997). The fact that some samples have W that is moreradiogenic than chondritic, while still possessing nearly chondriticNd, provides evidence that a component of core formation pre-ceded large-scale silicate differentiation (Kleine et al., 2002,2004a). Note the new values for chondritic W (Kleine et al.,2004a) and chondritic Nd (Boyet and Carlson, 2005). The concen-trations of W (in ppb) are shown in parentheses.


the main trend. Foley et al. (2005) note that there still existsa correlation between the ε182W and ε142Nd values for theshergottites, as can be seen from Fig. 4, but it has a very dif-ferent (shallower) slope.

A concern with some of the Saharan samples is withmobility of certain trace elements. Some such samples haveanomalous trace-element compositions, thought to reflectpartial terrestrial contamination or leaching. As shown inFig. 4, the meteorites DaG 476 and SAU 051 have extremelydepleted W concentrations (Kleine et al., 2004) — amongthe most depleted yet recorded from shergottites and nahk-lites (Lodders, 1998). The Nd/W ratios are perfectly nor-mal (Kleine et al., 2004; Dreibus et al., 2003). Therefore,whether this depletion reflects leaching/dissolution of in-compatible matrix elements or is a primary igneous featureis unclear. There is no question from the bulk compositionsthat there has been enrichment in some fluid mobile ele-ments (e.g., Ba) and so leaching and exchange of either Ndor W may have also occurred. Whether this has anythingto do with the isotopic data distribution or the interlabora-tory differences is also unclear at present.

The W-isotopic compositions of martian meteorites arenot correlated with chemical indices of magmatic fractiona-

tion (Halliday and Lee, 1999). This is not unexpected be-cause martian meteorites are mainly young and their sourceshave experienced several billion years of evolution since thevery early processes that produced the W-isotopic effects.However, the first, imprecise, W-isotopic compositions doshow a trend with Ba/W (Halliday et al., 2001b). This isless convincing with more precise measurements (Fig. 5a),as discussed below. Barium and W are equally incompatiblein mantle melting and their relative proportions in terres-trial basalts are more or less uniform (Newsom et al., 1986,1996). The Ba/W ratios of terrestrial basalts dominantly re-flect the amount of W depletion in the mantle caused by coreformation. Assuming this is also true of silicate partitioningon Mars, one can use the variation in Ba/W as a proxy fordifferent degrees of metal segregation. The implication isthat the incompatible trace elements in martian basalts carrya record of very early metal segregation processes on theplanet. This at first seems quite remarkable. On Earth sucheffects have largely been homogenized by billions of yearsof mantle convection. The oldest incompatible-element iso-topic heterogeneities in the present-day convecting mantle,as preserved in modern basalt magmas, are <2 b.y. old. Eventhe subcratonic lithospheric mantle is overwhelmingly dom-

Fig. 5. Tungsten-isotopic compositions of martian meteorites do not correlate with measured Hf/W but do correlate with model sourceHf/W determined from the chondritic value of (a) Hf/Ba and the sample’s Ba/W and (b) Hf/Th and the sample’s Th/W. Assuming Hf/W is only fractionated by core formation and not silicate partial melting, these data can be modeled in terms of a progressive coreformation model. However, it is very well established that W is much more incompatible than Hf during partial melting of the mantle(Newsom et al., 1986, 1996; Halliday and Lee, 1999). Note therefore that metal segregation probably only accounts for a fraction ofthe W-isotopic heterogeneity. The present-day W-isotopic compositions are shown for a range of martian reservoirs that have suffereddifferent degrees of W depletion by core formation over different timescales. Each model curve deploys exponentially increasing Hf/W with time and a time constant (m.y.–1) as shown. There is no particular basis for assuming that W depletion did change exponen-tially with time. Other models show a similar general curvature and the timescales calculated (see Fig. 6) are similar. However, aportion of the apparent rapidity of differentiation could reflect the unknown magnitude of increase in Hf/W resulting from silicatepartial melting. See text for details.


inated by heterogeneity that is ≤3 b.y. Traces of very early(>4.4 Ga) ε142Nd heterogeneity have been found in the an-cient sources of early Archean rocks from West Greenland(Boyet et al., 2003; Caro et al., 2003). Therefore, Earth hadsmall mantle isotopic heterogeneities in the early Archeanthat have since been homogenized. It has recently beenproposed that the silicate Earth, as sampled by magmas, isoffset from chondrites in ε142Nd by 20 ± 14 ppm (Boyet andCarlson, 2005). Whether this reflects a complementary hid-den reservoir or a slightly nonchondritic silicate Earth isunclear. None of these effects compare with the large het-erogeneities preserved within the martian mantle from theearliest solar system.

The fact that the martian W-isotopic compositions relateto Ba/W provides evidence that the isotopic variations dodefine differences in the degree of W depletion related tocore formation and not just partial melting. The samplescarry a chemical as well as isotopic record of variations inthe degree of siderophile depletion. One can simply convertthe Ba/W to a model source Hf/W that would have beenproduced as a result of metal-silicate fractionation by mul-tiplying by the chondritic Hf/Ba ratio (0.044). This is il-lustrated (Fig. 5a) for all the most precise W-isotopic datafor which Ba/W ratios are available. Saharan meteorites arestrongly altered and as such it is not possible to use Hf/Baas a refractory lithophile multiplier in such samples. Un-fortunately, U, the commonly used alternative to Ba, is alsohighly mobile in Saharan meteorites. Thorium is probablythe best, although also imperfect, alternative and is usedhere to calculate a second set of model source Hf/W valuesfrom the measured Th/W and chondritic Hf/Th (4.0). Whenthe W-isotopic data are compared to these model source Hf/W ratios, a trend is evident, whichever multiplier is deployed(Fig. 5a,b). This cannot be explained by silicate partial melt-ing. It indicates that core formation was a primary factorin generating the variations in W-isotopic composition.

The scatter to the data distribution in Fig. 5 could becaused by one or more of four factors:

1. Alteration. The different symbols in Fig. 5 refer toSaharan (diamonds), Antarctic (squares), and other (circles)martian meteorites. The model Hf/W values calculated forthe two Saharan meteorites, using Th/W, define the fullspread in shergottite values (Fig. 5b) and should thereforebe treated with some caution.

2. Analytical uncertainties and sample heterogeneityissues. The source Hf/W calculated either way has beenassigned a somewhat arbitrary uncertainty of ±20%. How-ever, the trace-element compositions are generally deter-mined on different aliquots from those used for measuringisotopic compositions. Given the variability in trace-elementconcentrations reported by different workers, this may bean underestimate of the real uncertainties. (See, for example,C. Meyer’s Mars Meteorite Compendium of available dataat http://www-curator.jsc.nasa.gov/curator/antmet/mmc/mmc.htm.) Furthermore, some of the apparent variability inBa/W for the same sample could be related to these ele-ments being predominantly cited in different minor phases

that could be distributed heterogeneously on the scale ofthe typical sample amounts used by various workers.

3. Partial melting. A second effect may have fraction-ated the source Hf/W, resulting in changes in W-isotopiccomposition. The variability in ε142Nd (Fig. 4) indicates thatsome silicate partial melting was consanguineous with coreformation. This would have produced source variations inHf/W that are not determinable with the approach utilizedhere.

4. Variable siderophile-element depletion in time andspace. Given the preservation of heterogeneity in the W-isotopic composition and degree of W depletion, there is nobasis for assuming that metal segregation proceeded at aconstant rate in all portions of Mars.

This last feature can be modeled quite easily. This isimportant because clearly the W-isotopic compositions werenot generated from a single primitive mantle reservoir withuniform Hf/W as is commonly deployed to calculate amodel age of core formation. The data distributions shownin Fig. 5a,b are plotted along with model curves, each ofwhich defines the predicted present-day W-isotopic com-positions of a reservoir that became isotopically heteroge-neous by undergoing variable degrees of metal segregation,hence W depletion, with a particular time constant, λ, i.e.,

180Hf/184W = (180Hf/184W)BSS × eλt

where (180Hf/184W)BSS is the parent/daughter ratio of thebulk solar system (or chondrites), λ is the time constant, andt is time. The model assumes that Mars was a chondriticobject that developed various such heterogeneous reservoirs,which differed in the rate at which they underwent W deple-tion because of core formation. If the W depletion developsrapidly the Hf/W ratio increases early and the reservoirdevelops W that is relatively radiogenic for a given Hf/W.

The longevity defined by the locus of each of these dif-ferent time constant curves is shown in Fig. 6. It can be seenthat the timescales that are required to generate martianmeteorite compositions like that of Nakhla could in fact bequite long (>10 m.y.) because, despite having radiogenic W,the model Hf/W is also very high. Unless other effects (1through 3 above) are responsible, the nakhlite source ap-pears to have carried on segregating metal over long (107 yr)timescales. The steepest line in Fig. 5 defines a very fast,effectively instantaneous differentiation rate (λ = 1 m.y.–1)and several of the martian meteorites plot along this lineor slightly to the left with low Hf/W for a given W-isoto-pic composition. The differentiation timescales for gener-ating the W-isotopic compositions along this low Hf/W lineare incredibly fast — around 1 m.y., as shown for Zagami(Fig. 6).

Coeval silicate partial melting effects mean that these Hf/W values could both over- and underestimate the real sourceHf/W values. This could explain why some samples haveslightly radiogenic W for a given model Hf/W and plot tothe left in Fig. 5, for example. That is, the source was alsodepleted by partial melting generating higher Hf/W than


accounted for with the approach adopted here that assumesthat all the W-isotopic effects result from core formation.However, this does not appear to apply to Zagami and Sher-gotty (Fig. 4). These meteorites yield chondritic Nd-isoto-pic compositions; there is no evidence of fractionation as aresult of silicate partial melting. Further combined W- andNd-isotopic data and high-quality trace-element data on thesame and representative sample aliquots are needed toevaluate this more generally. Model attempts based on thelithophile behavior of Nd have been utilized by Kleine etal. (2004a) and Foley et al. (2005). The deduced timescalesfor martian mantle reservoirs tend to be similar. Brandonet al. (2000) have also presented a broadly similar storyfrom Re-Os systematics. Marty and Marti (2002) similarlyhave presented the case for very rapid (<35 m.y.) differenti-ation based on Xe-isotopic systematics but with a more pro-tracted degassing history for the nakhlite source. All theseapproaches yield approximately the same view of Mars, thatthe accretion and primary differentiation were exceedinglyrapid.

Nevertheless, the Hf-W data provide unique evidencethat core formation proceeded over a range of timescalesthat in some cases were quite long (107 yr). Other meteor-ites such as Zagami appear to record more rapid timescales(106 yr). This in turn implies that accretion of Mars wasearly and rapid, on the order of 1 m.y. Such timescales are

similar to those predicted from runaway growth (Weiden-schilling et al., 1997). There is little evidence for late-stageplanetary-scale collisions affecting and effecting the growthof Mars.

7. THE AGE OF THE MOON

The discovery of 182W variability in lunar samples hasprovided the most powerful constraint on the age of theMoon (Lee et al., 1997, 2002; Leya et al. 2000; Kleine etal., 2005c). The W-isotopic compositions of lunar samplesare offset to radiogenic values relative to chondrites. At first,it appeared these effects were relatively large such that arather simple model age calculation could be applied. How-ever, the most enhanced εW values are now known to bethe product of cosmogenic effects on 181Ta, which produces182Ta, the intermediate decay product of 182Hf decay (Leyaet al., 2000; Lee et al., 2002). The isotopic variations appearto be at most 1 or 2 ε units (Lee et al., 2002). However,variations within the lunar mantle, even of this magnitude,restrict the age of the Moon to the first 30 to 55 m.y. of thesolar system (Kleine et al., 2005c; Yin et al., 2002b; Halli-day, 2003, 2004), in stark contrast to the limited constraintsbased on the oldest dated rocks (Wasserburg et al., 1977b)(Table 2).

The determination of the Hf-W age of the Moon dependson knowing the magnitude of the W-isotopic effect pro-duced by decay of 182Hf within the Moon itself (Halliday,2004). For this it is necessary to know (1) the initial Hf-and W-isotopic compositions of the solar system, (2) theinitial W-isotopic composition of the Moon, (3) the present-day W-isotopic composition of the lunar mantle, and (4) theHf/W of the lunar mantle.

The initial 182Hf/180Hf of the solar system is now fairlywell established (Kleine et al., 2002, 2005a; Schoenberg etal., 2002; Yin et al., 2002b). The initial ε182W of the solarsystem based on the W-isotopic compositions of CAIs andchondrites is established to be –3.5 (Kleine et al., 2002,2005a; Yin et al., 2002b); nevertheless, as discussed earlier,there is still some degree of uncertainty because of the ap-parently less-radiogenic W-isotopic compositions of someiron meteorites.

The initial W-isotopic composition of the Moon is atpresent unknown. The argument has been made (Halliday,2003, 2004) that the large number of samples with W-iso-topic compositions close to ε182W = 0 most likely meansthat this approximates a common composition from whichthe Moon started. This could well be wrong; however, it isunlikely to be far wrong. Clearly the initial composition ofthe Moon cannot be higher than its least-radiogenic value.It is unlikely to be much lower either since the current the-ory for the origin of the Moon involves a giant impact. TheMoon is thought to have been derived from the silicatemantles of the proto-Earth and the impacting planet Theia.All such silicate reservoirs studied thus far have ε182W ≥–0.5 (Quitté et al., 2000) and unless the Moon formed very

Fig. 6. Timescales implied by the model development of Hf/Wwith time in the martian mantle assuming all the W-isotopic ef-fects are the result of siderophile depletion associated with coreformation. Two extreme examples are highlighted. The W deple-tion of the Nakhla source would have taken more than 10 m.y. todevelop, whereas the time necessary to form the source of Zagamiis <1 m.y.


early (<20 m.y.), there is no reason why it would have alower value.

Most of the present-day W-isotopic compositions of lu-nar samples corrected for cosmogenic effects lie in therange ε182W = 0–1. Most cosmogenic corrections are veryapproximate. Some values are unequivocally higher (Leeet al., 2002). Lunar sample 15555 with no cosmogeniceffect yields a value of 1.30 ± 0.39. From the current data-base it would appear that the average composition of thelunar mantle is unlikely to be greater than ε182W = 1, but itcould conceivably be this high if 15555 is especially rep-resentative of the lunar mantle. Therefore, a likely scenariois an average radiogenic increase of about 0.5–1.0 (Halli-day, 2004). A more conservative estimate of the amount ofradiogenic increase would be that it is less than 2.0 ε182Wunits.

The fact that there are endemic, indigenous W-isotopicvariations at all provides clear evidence that they were pro-duced by radioactive decay within the lunar mantle; theycannot be residual from Earth or Theia given the inevitablemixing associated with the high temperatures of accretionand early convective overturn. The initial W-isotopic com-position of the Moon does reflect the mix of componentsfrom these parent planets, however, and provides informa-tion on their history, as described below.

The Hf/W ratio of the lunar mantle has been estimatedmost recently and comprehensively by Jones and Palme(2000) to be 25.2. (Note that this value changes to 26.5when using more up-to-date CI abundances; the estimateis based on U/W = 1.93 for the lunar mantle and dependson the Hf/U used for chondrites.) This is similar to previ-ously used estimates (Halliday, 2003, 2004).

The age of the Moon can be deduced from these com-bined estimates as shown in Fig. 7 and would appear to lie inthe range 40–50 m.y. (Halliday, 2004; Kleine et al., 2004b).The age can be refined more closely once more preciseestimates of the amount of radiogenic growth on the Moonare known. This estimate is later than the 29 m.y. proposedby Kleine et al. (2002) and the value of 30 m.y. calculatedby Yin et al. (2002b). These are model ages relative to achondritic reservoir and are the equivalent of assuming thatthe Moon formed in a single step from an undifferentiatedchondritic reservoir. We now know the W-isotopic compo-sition of such a reservoir very well (Kleine et al., 2002;Schoenberg et al., 2002; Yin et al., 2002b). However, theMoon has a relatively low uncompressed density comparedwith that of the terrestrial planets, and this has long beentaken as evidence of an origin from a silicate-rich precursor,i.e., a planetary mantle. Therefore, current models of lunarformation (e.g., Canup and Asphaug, 2001) involve a colli-sion between Earth and a Mars-sized planet that is alreadydifferentiated into silicate and metal. As such, the prior W-isotopic evolution of these parent planetary mantles will havecontributed to the W-isotopic composition recorded in lunarbasalts. A model age relative to chondrites is the equivalentof assuming that none of this had happened. It provides a

useful upper limit, that is, the earliest likely time that theMoon can have formed, given its W-isotopic composition.

Recently, Kleine et al. (2005c) have reported W-isoto-pic data for metals separated from lunar samples that pro-vide what appears to be the best constraint on the exact ageof the Moon. The significance of these measurements liesin the fact that the metal has very low Hf/W and Ta/W suchthat the W-isotopic composition is a real initial ratio for theigneous rocks studied whatever their age and exposure. Bycomparing these W-isotopic compositions with independentestimates of the Hf/W ratio of the magma sources, one canderive a differentiation age for the Moon. On this basisKleine et al. derive an age for lunar differentiation of 40 ±10 m.y. This is the most precise estimate yet obtained forthe age of crystallization of the lunar magma ocean.

Whatever the exact age of the Moon, there are clear im-plications for these results:

1. First, there is no question that the Moon formed laterelative to other small objects. Therefore, models for its for-mation must explain this. The fission and late impact theo-ries would seem best suited. Capture and coaccretion do notpredict a late formation.

2. A significant time gap can now be identified betweenthe age of crystallization of the lunar magma ocean deducedfrom 182Hf-182W and the precise ages deduced from 235/238U-207/206Pb and 147Sm-143Nd chronology for early lunar rocksand lunar differentiation events, previously thought to re-

Fig. 7. The age information that can be derived from current W-isotopic data for the Moon at the present time is strongly modeldependent. The effects on the calculated age of the Moon of notknowing the average amount of radiogenic, as opposed to inher-ited or cosmogenic, 182W are shown. The increase in εW withinthe Moon as a function of decay of primordial 182Hf, as can bededuced from currently available data, is probably ≤1 ε unit. Thedata are therefore consistent with an age of the Moon (hence gi-ant impact) of about 40 to 50 m.y. after the start of the solar sys-tem. The different curves are based on differing values of theε182WBSSI, which directly affects the calculated (182Hf/180Hf)BSSI.All calculations assume λ182Hf = 0.077 × 10–6 yr–1, Hf/WMOON =22, and ε182WBSS = –1.9.


late to the formation and crystallization of the lunar magmaocean.

3. Assuming the Moon was formed in a giant impactwhen the proto-Earth was ~90% formed, its age providesthe single most important piece of independent evidencethat can be used with the meteorite reference frame to cali-brate Earth’s growth history, as explained next.

8. METEORITES AND THEGROWTH OF EARTH

As explained above, the W-isotopic data for iron mete-orites provide evidence of very rapid accretion and coreformation of earliest planetesimals in the inner solar sys-tem. Similarly, the data for Mars are hard to explain unlessaccretion and earliest core formation were extremely rapid(<1 m.y.). The isotopic data for Earth, however, providestrong support for protracted accretion over tens of millionsof years. This is consistent with dynamic simulations thatpredict timescales for terrestrial planet accretion that are onthe order of a few tens of millions of years (Safronov, 1954;Wetherill, 1980, 1986; Agnor et al., 1999; Canup and Agnor,2000; Chambers, 2001a,b, 2004). These simulations can beextremely sensitive to the amount of nebular gas present(Agnor and Ward, 2002; Kominami and Ida, 2002). Theiraccuracy for describing the real Earth needs to be testedrigorously, hence the central importance of isotopic ap-proaches.

The three most powerful and effective techniques for de-termining the growth rate of Earth are 235/238U-207/206Pb,244Pu/129I-136/129Xe, and 182Hf-182W. These yield differingage estimates, and this in turn may provide insights into theprocesses that are likely to have accompanied Earth accre-tion. As the first 182Hf-182W data became available there ap-peared to be excellent agreement with the protracted time-scales for planetary accretion and atmospheric loss as de-duced from 92Nb-92Zr (Münker et al., 2000; Jacobsen andYin, 2001), 129I/244Pu-129/136Xe (Porcelli and Pepin, 2000),and 235/238U-207/206Pb chronometry (Halliday, 2000). Thetimescales deduced from 97Tc-97Mo (Yin and Jacobsen,1998) and 107Pd-107Ag (Carlson and Hauri, 2001) chro-nometry were shorter.

Now it is clear that some of these constraints are not verystrong because some of the critical Mo, Zr, and Ag dataappear to be incorrect or misinterpreted (Dauphas et al.,2002a,b, 2004; Chen et al., 2004; Lee and Halliday, 2003;Schönbächler et al., 2002, 2003; Woodland et al., 2004).For example, the anomalies in Mo are hard to distinguishfrom nucleosynthetic effects and different groups have re-ported significantly different results for some of the isotopeseven when the same normalization procedure is deployed(Yin et al., 2002a; Becker and Walker, 2003; Dauphas et al.,2002a,b, 2004; Chen et al., 2004; Lee and Halliday, 2003).The reasons for these discrepancies are unclear. Dauphaset al. (2004) reported that his Mo-isotopic data correlatedwith effects for Ru reported by Chen et al. (2003), appar-ently confirming their validity as nucleosynthetic hetero-geneities. The latest results for Mo reported by Chen et al.

(2004) reveal differences relative to previously publishedresults. They confirm only some of the nucleosynthetic ef-fects and find a decoupling between p- and r-process com-ponents. The results no longer correlate with Ru so well.

The apparent consistency between 235/238U-207/206Pb,244Pu/129I-136/129Xe, and 182Hf-182W chronometry may alsobe incorrect. The early W-isotopic data for chondrites (Leeand Halliday, 1995, 1996) are wrong by about 180–200 ppm(Kleine et al., 2002; Schönberg et al., 2002; Yin et al.,2002b). The new 182Hf-182W timescales are shorter andappear to be inconsistent with 235/238U-207/206Pb timescales(Halliday, 2003, 2004; Wood and Halliday, 2005). Even thelatest 244Pu/129I-136/129Xe timescales appear to be too long(Porcelli et al., 2001). In detail, therefore, it now appearsas though the chronometers are reflecting different chemi-cal fractionations that took place over distinct timescales,or that the degree to which the real processes are accuratelysimulated by the isotopic models differ between elements(Halliday, 2003, 2004; Wood and Halliday, 2005).

The 235/238U-207/206Pb and 182Hf-182W chronometers bothrely on determining the timing of fractionation during coreformation. Both are usually utilized in one of two ways:(1) A model age can be calculated as the time that the sili-cate Earth was last in total isotopic equilibrium, since thetime when its silicate (high U/Pb and Hf/W) and metal (lowU/Pb and Hf/W) reservoirs developed distinct compositions.This can be thought of as defining an age of instantaneouscore formation in a completely formed planet. Alternatively,it can be thought of as catastrophic reequilibration of sili-cate and metal reservoirs during some major overturn mix-ing event. Finally, it can be thought of as simply the age ofthe planet itself assuming the core formed simultaneously.There is considerable evidence against any of these repre-senting the way Earth formed and developed in its earlieststages. (2) A mean age for accretion can be calculated asthe inverse of the time constant for exponentially decreas-ing growth of Earth, assuming that, as it grew, the accretedmaterial isotopically equilibrated with the primitive mantleand continuously segregated further core material in pres-ent-day proportion to Earth’s mass (Jacobsen and Harper,1996). There seems little question that this is more realis-tic (Halliday, 2004).

The first of these models represents the standard proce-dure adopted for dating Earth or core formation (Allègreet al., 1995; Galer and Goldstein, 1996; Lee and Halliday,1995, 1996; Halliday and Lee, 1999). However, the “event”defined probably never occurred as such. The second ofthese models is more sophisticated and useful and is basedon the formalism first proposed by Jacobsen and Wasser-burg 1979). It was explored for 182Hf-182W by Harper andJacobsen (1996) and Jacobsen and Harper (1996) and fur-ther explored in a series of models by Halliday et al. (1996,2000), Halliday and Lee (1999), and Halliday (2000). Theseall predate the correct determination of the isotopic com-positions of chondrites (Kleine et al., 2002, 2004a; Lee andHalliday, 2000a; Schoenberg et al., 2002; Yin et al., 2002b).Recent studies have made extensive use of this approach(and the correct reference compositions) to obtain an ac-


curate 182Hf-182W accretion timescale for Earth (Yin et al.,2002b; Halliday, 2003, 2004; Kleine et al., 2004b). Thesame idea was developed and exploited for 235/238U-207/206Pbby Halliday (2000, 2003, 2004).

When the results for 182Hf-182W and 235/238U-207/206Pb arecompared there appears to be a discrepancy (Halliday, 2003,2004; Wood and Halliday, 2005). The 235/238U-207/206Pb ac-cretion rates appear to be more protracted than those de-termined from 182Hf-182W (Fig. 8). One possible explana-tion to consider is that it relates to the differences in the ratesof metal-silicate equilibration of refractory W vs. volatile Pbduring the accretion process itself (Halliday, 2003, 2004).The possibility that there had been a lack of equilibrationduring accretion was discussed by Halliday (2001), but atthat stage it appeared that the W-isotopic composition ofthe silicate Earth was identical to that of chondrites suchthat equilibration had to be the general rule (Halliday, 2001).Now that a resolvable difference has been established be-tween the silicate Earth and chondrites, the question arisesas to whether this is caused by faster accretion, or lack ofisotopic equilibration, or both (Halliday, 2004). The onlyway to test this is with some independent assessment of theage of the Moon, or the giant impact.

The age of the Moon provides the only firm independentconstraint on the accretion rate because it is thought to bethe byproduct of the last major growth phase of Earth(Canup and Asphaug, 2001). The later the age of the Moon,the more disequilibrium is needed to explain the radiogenicW-isotopic composition of the silicate Earth. Using a win-dow of 45 ± 5 m.y. (Fig. 7) as the most realistic for the giantimpact (Halliday, 2003, 2004; Kleine et al., 2004b, 2005c)allows a model accretion curve to be constructed usinglarger-sized impacts with time, culminating in the giant im-pact (Fig. 9). From this accretion curve one can deduce thelevel of W-isotopic equilibration that would be requiredassuming a general level of disequilibrium throughout theaccretion history or disequilibrium during the giant impactalone (Fig. 10). Assuming the latter case it can be seen that

Fig. 9. Change in mass fraction of Earth as a function of timein the model used in Halliday (2004) illustrated with an accre-tion scenario calculated from a giant impact at 50 m.y. after thestart of the solar system. The approximated mean life of accre-tion (τ) is the time taken to achieve 63% growth. Both this andthe timing of each increment are calculated from the timing ofthe giant impact (tGI) to achieve an overall exponentially decreas-ing rate of growth for Earth broadly consistent with dynamic simu-lations. The smooth curves show the corresponding exponentiallydecreasing rates of growth. The step function curves define thegrowth used in the isotopic calculations. Growth of Earth is mod-eled as a series of collisions between differentiated objects. Theoverall rate of accretionary growth of Earth may have decreasedin some predictable fashion with time, but the growth events wouldhave become more widely interspersed and larger. Therefore, themodel simulates further growth by successive additions of 1% Mobjects from 1% to 10% of the current mass, then by 2% objectsto 30%, and then by 4% objects to 90%. The Moon-forming gi-ant impact is modeled to take place when Earth was ~90% of itscurrent mass and involved an impactor planet Theia that was ~10%of the (then) mass of Earth. Therefore, in the model the giantimpact at 45 m.y. after the start of the solar system contributes afurther 9% of the current Earth mass. There is evidence againstlarge amounts of accretion after the giant impact. Although epi-sodic relative to more conventional continuous core formationmodels (Jacobsen and Harper, 1996; Halliday et al., 1996, 2000;Halliday and Lee, 1999; Halliday, 2000), the model is still smoothrelative to accretion simulations (e.g., Agnor et al., 1999; Cham-bers, 2001a,b, 2004).

Fig. 8. Calculated values for Earth’s mean life of accretion (τ)and time of the giant impact (TGI) given in million years as de-duced from the different estimates of the Pb-isotopic compositionof the bulk silicate Earth (BSE) (Doe and Zartman, 1979; Davies,1984; Allègre et al., 1988; Zartman and Haines, 1988; Allègre andLewin, 1989; Kwon et al., 1989; Galer and Goldstein, 1991; Liewet al., 1991; Kramers and Tolstikhin, 1997; Kamber and Coller-son, 1999; Murphy et al., 2003) using the type of accretion modelshown in Fig. 9 (see Halliday, 2004). Note that, while all calcu-lations assume continuous core formation and total equilibrationbetween accreted material and the BSE, the accretion is punctu-ated, as predicted from the planetesimal theory of planetary ac-cretion. This generates more protracted calculated timescales thanthose that assumer smooth accretion (Halliday, 2004). The 238U/204Pb values assumed for Earth = 0.7 (Allègre et al., 1995). Eventhe Hf-W timescales using the exact same style of model are sig-nificantly shorter. The mean life model ages given in Table 2 arefor a smooth accretion model that yields less-protracted results(Halliday, 2000) and are for the average (Halliday, 2004) of themost recent (Kramers and Tolstikhin, 1997; Kamber and Coller-son, 1999; Murphy et al., 2003) of these Pb-isotopic estimates.


the amount of equilibration of Theia’s core with the silicateEarth required to explain its W-isotopic composition waslimited to <60% during a giant impact at 45 m.y. and couldhave been as little as 40% (cf. Halliday, 2004).

This point is made primarily to illustrate the importanceof considering equilibration when determining model iso-topic age constraints such as the accretionary mean life. Ifthe giant impact involved a larger impactor, for example, a0.15 instead of 0.10 M object, then a greater proportionof Earth would have been accreted at a later time. There-fore, the radiogenic W in the silicate Earth would implygreater amounts of disequilibrium during accretion. The sizeof the impactor affects the energy that is released. As such,it can be tuned in dynamic simulations to produce an Fe-depleted Moon at the right distance with the correct angu-lar momentum. However, a critical issue is the equation ofstate, which defines the amount of energy released for agiven sized impactor, and which is poorly known for high-pressure materials such as perovskite, even though advancesare being made. A closer integration of isotopic modeling

with dynamic simulations and experimental and theoreticalmineral physics should help in this respect.

Another variable that is becoming critical is the exactcomposition of Earth. We now have excellent W-isotopicdata for chondrites (Kleine et al., 2004a), assumed to rep-resent the bulk Earth. This is unlikely to be far wrong. How-ever, it now becomes important to know the exact Hf/W ofthe total Earth. If Earth has a Mg/Fe ratio that is lower thanchondritic (Halliday et al., 2001a; H. Palme, personal com-munication, 2003), then it probably has a Hf/W ratio that isalso subchondritic. The effect on W-isotopic models dependson when this Hf/W depletion occurred. If it is an early fea-ture it requires that the amount of disequilibrium during ac-cretion was greater. Even the chondritic Hf/W ratios are notso well known. New high-precision isotope-dilution data byKleine et al. (2004a) are, on average, slightly lower thanthe previously utilized “standard” data (Newsom, 1995). Al-though the data agree within error, these uncertainties areof sufficient magnitude that they now significantly limitattempts to produce high-resolution model age and accre-tion rate calculations. Sample size effects have plagued Lu-Hf and Sm-Nd studies of chondrites and new Hf-W stud-ies of large representative samples would be worthwhile.

A Moon-forming impact at 45 m.y. does little to explainthe discrepancy with the rates determined from the Pb-iso-topic composition of Earth (Fig. 8). Assuming any one ofthese 11 estimates is approximately correct, the most likelyexplanation is that there was a change in U/Pb in the bulksilicate Earth during accretion. This might have been causedby changes in partitioning or volatile loss (Halliday, 2004).A recent study by Wood and Halliday (2005) proposes a newversion of the former explanation. Tungsten and Pb maywell segregate into planetary cores at different stages. Tung-sten is more siderophile than chalcophile, whereas Pb dis-plays the opposite behavior. During planetary accretion anddifferentiation a point will be reached where metal segre-gates as a result of the removal of ferric iron into perovskite.The metallic iron removes W at an early stage. However,Pb remains in the silicate Earth until the upper mantle ofEarth becomes oxidized as a result of the giant impact dis-rupting perovskite in the lower mantle. The S added fromthe impactor, which appears to have been rich in chalcophileelements (Yi et al., 2000), results in the formation of sul-fides as Earth cools. These will strip the silicate Earth ofits Pb and, if sufficiently abundant, will accumulate at thebase of the magma ocean, eventually sinking to the core(Wood and Halliday, 2005). Therefore, there appears to bea logical explanation for why the Pb-isotopic system recordsa more protracted timescale than the Hf-W system. Coolingof Earth’s lower mantle should have been very fast afterthe giant impact (Solomatov, 2000). However, the uppermantle from which the sulfide segregated could have cooledmuch more slowly (up to 108 yr). The earliest calculatedtimescales for removing Pb in this fashion are about 30 m.y.after a 45 m.y. giant impact based on the most recent esti-mate for the Pb-isotopic composition of the BSE (Wood andHalliday, 2005). Temperatures of roughly 3000 K are re-quired to segregate sulfide in this fashion. The latest mod-

Fig. 10. An illustration of the effect on calculated Hf-W time-scales for Earth’s formation of incomplete mixing and equilibra-tion of impacting core material. This plot shows the true time ofthe giant impact that generates ε182WBSE of zero as a function ofvarious levels of incomplete mixing of the impacting core mate-rial with the BSE. The lower curves are for disequilibrium duringthe giant impact alone. The upper curves correspond to disequi-librium during the entire accretion process up to and includingthe giant impact. The different curves for different Hf/WBSI (theHf/W in the silicate portion of the impactor) are also shown. Allcurves are calculated with ε182WBSSI = –3.5. Assuming an age forthe Moon in the range 40–50 m.y. after the start of the solar sys-tem, it would seem likely that there was significant isotopic dis-equilibrium during the giant impact (Halliday, 2004).


els of the giant impact (Canup, 2004) raise the tempera-ture of Earth to about 7000 K. Therefore, broadly speak-ing, Earth cooled at more than 100 K per m.y. after the giantimpact. This would correspond to the time required to coolthe upper portion of the mantle to the temperature at whichsulfide stripped Pb from the mantle.

9. TIME-INTEGRATED CHEMICALCOMPOSITIONS AND THE VOLATILE

BUDGETS OF EARLY PLANETESIMALSAND TERRESTRIAL PLANETS

The radiogenic isotopic compositions of early solar sys-tem objects can be used to define time-averaged parent/daughter ratios of precursor materials. This can be exploredfor any isotopic system and provides useful insights intothe paleocosmochemistry of the early solar system. Thereare four systems to which this can be very usefully appliedto constrain early chemical evolution during accretion.

9.1. Time-integrated Uranium/Lead

Lead-isotopic compositions provide simultaneous con-straints on the age of an object and the parent/daughterratios (238U/204Pb or primary “µ”) of the precursor materi-als. For objects like Earth and Mars the present-day Pb-iso-topic compositions of the silicate reservoirs are overwhelm-ingly dominated by the time-integrated µ since accretionand core formation. As such, these data say little about anyearlier history. However, several lunar rocks formed earlyand sampled the Pb-isotopic compositions of their precur-sor reservoirs. Lead is moderately volatile, and most lunarrocks contain so little inherited Pb that the initial composi-tion is hard to resolve (Tera et al., 1973). It has been pro-posed that this may represent the early silicate Earth (Galer,1993). However, an early high U/Pb silicate Earth is noteasy to reconcile with the present Pb-isotopic compositionof the silicate Earth (Halliday, 2004). Furthermore, somelunar rocks carry initial Pb-isotopic compositions that arerelatively unradiogenic (Meyer et al., 1975; Hanan and Til-ton, 1987; Torigoye-Kita et al., 1995) and provide evidencethat the Moon formed from material that was not so de-pleted in Pb. In fact, the primary µ values are broadly simi-lar to that of the bulk silicate Earth (Halliday et al., 1996).Assuming that the Moon sampled a mixture of materialfrom the proto-Earth and Theia, there is no clear evidencefor any great difference in the magnitude of volatile deple-tion in these precursor materials compared with the presentEarth. This raises the issue of how the strong depletion in204Pb in the Moon was introduced. Presumably it was aproduct of the giant impact.

9.2. Time-integrated Rubidium/Strontium

The Rb-Sr system also provides an indication of themagnitude of volatile-element depletion (Halliday and Por-celli, 2001) of precursor materials. The differences betweenthe Sr-isotopic compositions of early objects like eucrites,

angrites, and lunar rocks relative to CAIs provides strongevidence that some process causes loss of Rb during accre-tion (Halliday and Porcelli, 2001). This process cannot frac-tionate K isotopes (Humayun and Clayton, 1995), and par-tial evaporation would therefore appear unlikely. However,elemental fractionations unaccompanied by isotopic effectscan in fact occur as a result of partial evaporation or con-densation, as long as environmental conditions change suf-ficiently slowly to allow for thermodynamic equilibrium(Richter, 2003). In the case of early planetesimals like theangrite and eucrite parent bodies, it is thought that there wasvery early melting, differentiation, and volcanism predatingthe formation of the rocks sampled as achondrites (Lugmairand Shukolyukov, 1998). It is conceivable that such an earlyglobal differentiation was accompanied by extensive out-gassing and almost complete loss of moderately volatile ele-ments. The volatile budgets would then be added by sub-sequent accretion of new material. In the case of the Moon,however, it would appear more likely that the major loss(>90%) of Rb was associated with the giant impact (O’Neill,1991a,b).

9.3. Time-integrated Hafnium/Tungsten

The Hf/W of a silicate reservoir is a function of the par-titioning of W during core formation but may also be ren-dered heterogeneous by silicate partial melting (Hallidayand Lee, 1999). The metal/silicate partition coefficient forW is strongly dependent on oxygen fugacity (Schmitt et al.,1989; Walter et al., 2000). It presumably is also affected byvolatile contents and the pressure and temperature of coreformation (Righter and Drake, 1996, 1999; Righter et al.,1997; Walter et al., 2000). The lack of depletion of W inthe martian mantle has long been taken as providing evi-dence of more oxidizing conditions thought to be linked tothe greater abundances of moderately volatile elements (fora review, see Halliday et al., 2001b). The fact that the W-isotopic compositions recorded in martian meteorites areonly 200–500 ppm more radiogenic than chondritic despitebeing generated very early is entirely consistent with thislack of depletion. In the case of the silicate Earth the fact thatthe W-isotopic composition also is within 200 ppm of chon-dritic despite a Hf/W that is an order of magnitude greaterthan chondritic is explained by protracted accretion of chon-dritic material equilibrating with the silicate Earth. How-ever, the fact that the initial W-isotopic composition of theMoon is close to chondritic is more problematic (Halliday,2004). The Moon has a high Hf/W, and it is thought that itwas mainly formed from the silicate portions of Theia, withsubordinate contributions from the proto-Earth. If Theia wasonly a Mars-sized object it should have formed relativelyfast like Mars itself, as discussed above. However, with sucha high Hf/W it would then be expected to have producedhighly radiogenic W by the time of the giant impact. Instead,the Moon started with εW ~ 0 (Halliday, 2004; Kleine et al.,2005c), within ~200 ppm of chondritic values. Assumingthe Moon was mainly derived from Theia, there are fourpossible explanations: (1) The bulk silicate Theia had its


composition modified by large impacts shortly before thegiant impact (Halliday et al., 2000), which is somewhat un-testable but not impossible. (2) Theia formed very slowly,and would need to be accreting at a rate much slower thanEarth, Mars, or Vesta to have such unradiogenic W. (3) Thematerial that formed the Moon equilibrated with quite largeamounts of metal; to achieve an initial W-isotopic composi-tion of εW ~ 0 in the Moon the amount of metal requiredis roughly 10% of the total material, which is well in excessof that predicted in giant impact simulations; (4) The sili-cate portions of Theia and/or the proto-Earth had a Hf/Wthat was close to chondritic; W partitioning changes as afunction of oxygen fugacity (Walter et al., 2000), and theimplication would be that the combination of silicate ma-terial that formed the Moon was significantly more oxidiz-ing than the present silicate Earth.

The last of these scenarios seems the most likely. Whenthe time-integrated Hf/W for the source of the Moon’s con-stituents (generally thought to be dominated by the bulksilicate Theia) is compared with the time-integrated Rb/Sr(Halliday and Porcelli, 2001), the combined compositionis strikingly similar to that of the martian mantle (Halliday,2004), lending support for the idea that Mars represents atype of normal protoplanet that was an essential buildingblock in the formation of larger terrestrial planets. In fact,the composition plots at the extreme end of an array ofknown compositions for terrestrial planets and differenti-ated planetesimals (Fig. 11), providing evidence that lossof moderately volatile elements is somehow linked to themetal/silicate partitioning of W.

Mars has a higher FeO/Fe ratio than Earth, which meansthat the average oxidation state of Mars is higher, so thiswould explain why W was more lithophile (Walter et al.,2000). The Earth’s upper mantle has a higher ferric ironcontent and thus oxygen fugacity than Mars. However, thisis thought to be due to the self-oxidation process of Earth’smantle associated with growth of the core. Once Earth hadestablished a certain mass, greater than that of Mars, per-ovskite would have become stable in the lower mantle.Therefore, during accretion and despite its low FeO/Fe,Earth started producing ferric iron that pushed up the oxy-gen fugacity (Wood and Halliday, 2005). This would nothave happened on Mars because perovskite is only just sta-ble at the base of the martian mantle.

The striking differences between the present composi-tion of the Moon and the time-integrated Rb/Sr, U/Pb, andHf/W of Theia provide evidence that a major compositionalchange was effected by the giant impact. In some respectsthis is no surprise because the energy was enormous. How-ever, little has been established about the physical chemis-try of such accretion processes and remarkably little isknown about the feasibility of losing volatile elements fromthe debris of a giant impact.

9.4. Time-integrated Manganese/Chromium

The depletion of Earth’s and the Moon’s inventories ofsome volatile elements must have occurred earlier than the

giant impact and probably predates planetary formation.Models for the collapse of the solar nebula and accretionof a planetary disk predict high temperatures (1500 K) inthe inner solar system (Boss, 1990). Therefore, it is likelythat the inner solar system became depleted in volatile el-ements before accretion of sizable bodies. The Earth andMoon share similar Cr-isotopic compositions, distinct fromthose of chondrites (Lugmair and Shukolyukov, 1998), con-sistent with an early depletion in (more volatile) Mn (Halli-day et al., 1996; Cassen and Woolum, 1997). The magni-tude of the difference in Cr-isotopic composition betweenthe BSE and some chondritic meteorites (about 0.5 ε53Cr/52Cr units) would be consistent with the depletion occur-ring within about 3.5 m.y. after the formation of the mate-rial in Allende, given the differences between the estimatedMn/Cr ratios of the total Earth (~0.18) and chondrites(~0.40) and assuming an initial 53Mn/55Mn ratio of 4.4 ×10–5. It should be emphasized that the interpretation of Mn-Cr data is controversial and complex (Lugmair and Shukol-yukov, 1998; Birck et al., 1989). The primary hypothesisproposed by Lugmair and Shukolyukov (1998) is that thevariations in Cr-isotopic composition simply reflect a ra-dial gradient in the 53Mn distribution resulting from incom-plete mixing of material injected into the disk from a nearbystellar source that synthesized this and other nuclides.

Isotope geochemistry therefore provides limited evidenceto support the widely accepted and almost certainly correctview that some volatile-element depletion in the planets was

Fig. 11. Hafnium/tungsten appears to be negatively correlatedwith Rb/Sr in the primitive mantles of inner solar system plan-etesimals and planets. A possible explanation for this is that theloss of moderately volatile elements was linked to loss of othervolatiles during planetary collisions such that the mantles changedfrom more oxidizing to more reducing. The Moon is an extremeexample of this. The fact that the mixture of material from theproto-Earth and Theia that is calculated to have formed the Moonis so like Mars provides evidence that such volatile-rich objectsmay have been common in the inner solar system during the earlystages of planetary accretion. See Halliday and Porcelli (2001)and Halliday (2004) for further details.


very early. Other lines of isotopic evidence indicate thatsome changes in volatile abundance were late and somemay have been related to energetic accretion processes likethe Moon-forming giant impact.

10. CONCLUSIONS AND OUTLOOK

Isotope geochemistry of meteorites has been central tothe determination of the ages, rates, and mechanisms of ac-cretion of the terrestrial planets. Not only do isotopic stud-ies of meteorites provide information on the average solarsystem and hence a reference for models of accretion anddifferentiation, meteorites provide important clues aboutfirst planetesimals and the formation of the earliest plane-tary embryos from which the rest of the solar system wasbuilt. The development of the Hf-W chronometer has hada bigger effect on this area of science than any other as-pect of isotope geochemistry. With the excellent referencedataset now available for carbonaceous chondrites and thenewly determined highly precise value for the 182Hf decayconstant (Vockenhuber et al., 2004), the precision and ac-curacy of Hf-W chronometry has taken major steps forward.

At the time of this writing, several important conclusionscan be drawn from Hf-W and other isotopic data about howthe inner solar system was built (Fig. 12):

1. Iron meteorites appear to provide the best bet for sam-ples of the first planetesimals. The ubiquitous unradiogenicW, sometimes lower in 182W than any other solar systemsample measured so far, provides strong evidence that theyformed early.

2. The exact timing of metal segregation as representedby iron meteorites requires improved understanding of cos-mogenic effects, which tend to decrease 182W and will bepresent in many iron meteorites with long exposure ages.By applying a maximum correction for cosmogenic effectsit appears that some iron meteorites formed within <1 m.y.of the start of the solar system, as defined by the W-isoto-pic composition of Allende CAIs (Markowski et al., 2006).

3. It is also essential to provide better cross calibrationof chronometers. These will be important as improvementsin mass spectrometric techniques will provide increasingage resolution.

4. There exist conflicts between the interpretations ofdifferent isotopic data for angrites and eucrites. The Sr-iso-topic data are hard to explain unless the parent body formed>2 m.y. after the start of the solar system.

5. The W-isotopic data for Mars are best explained ifparts of the martian mantle differentiated very rapidly. Thecurrent data are consistent with accretion and differentiationof some reservoirs within less than 1 m.y. Other reservoirsappear to have developed over longer timescales (>10 m.y.).However, these timescales assume that other factors suchas alteration, analytical and sampling issues, and silicatedifferentiation have not affected the inferred Hf/W ratiosin the source reservoirs of the martian meteotites. Furtherstudies that combine high-quality trace-element determina-tion with new W and Nd measurements are needed to con-firm this.

6. The age of the Moon is best defined by Hf-W datathat provide strong evidence of an origin from a giant im-pact, in the time range 30–50 m.y. after the start of the solarsystem, the exact age depending on the model deployed.

7. The age of the Moon provides an important, indepen-dent, and unique constraint on the accretion rate of Earthassuming it defines the last major phase adding roughly10% of the present mass of Earth.

8. The W-isotopic composition of the silicate Earth isconsistent with this timescale for accretion. The small ex-cess of 182W in the bulk silicate Earth relative to chondritesis explicable if the initial W-isotopic composition of the so-lar system is ε182W < –3.5, or if there was some level ofdisequilibrium between incoming metal and the BSE dur-ing accretion.

9. The difference between W- and Pb-isotopic estimatesfor the rates of accretion of Earth cannot be explained bydisequilibrium during accretion. Either these estimates forthe average Pb-isotopic composition of the bulk silicateEarth are in error or there has been an additional process thathas fractionated U/Pb at a relatively late stage during Earthaccretion.

10. Time-integrated parent/daughter ratios provide lim-ited evidence that some volatile-element depletion in theplanets was very early. Other lines of isotopic evidence indi-cate that some changes in volatile abundance were late andpossibly related to energetic accretion processes like theMoon-forming giant impact. It is essential to understandhow planetary chemical compositions are achieved given

Fig. 12. The current best estimates for the timescales over whichvery early inner solar system objects and the terrestrial planetsformed. The approximated mean life of accretion (τ) is the timetaken to achieve 63% growth at exponentially decreasing rates ofgrowth. The dashed lines indicate the mean life for accretion de-duced for Earth based on W and Pb isotopes (Halliday, 2000,2003, 2004; Kleine et al., 2002; Yin et al., 2002b). The earliestage of the Moon assumes separation from a reservoir with chon-dritic Hf/W (Kleine et al., 2002; Yin et al., 2002b). The best esti-mates are based on the radiogenic ingrowth deduced for theinterior of the Moon (Halliday, 2003, 2004; Kleine et al., 2005c).See Table 2 for details of other sources.


the evidence from isotopic compositions that these havechanged over time.

Although the development of Hf-W has resulted in majorprogress in quantifying the accretion of the terrestrial plan-ets, further progress is essential in certain critical areas.

Acknowledgments. We are very grateful to S. Jacobsen, C.Münker, and an anonymous reviewer for their comments on anearlier version of this chapter. M. Wadhwa kindly provided accessto the unpublished martian meteorite data of N. Foley and co-workers. B. Wood is thanked for discussions on mantle oxidationand siderophile-element partitioning during core formation. Weare deeply indebted to D. Lauretta and M. Wadhwa for editorialadvice and for their patience in giving us time to produce and re-vise this paper in the midst of many complications that intervened.

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