1.27 Long-LivedChronometers

1.27
Long-Lived ChronometersM. Wadhwa
Arizona State University,Tempe, AZ, USA

1.27.1 INTRODUCTION
1 1.27.1.1 Basic Principles 1 1.27.1.2 Application to Meteorites and Planetary Materials: A Historical Perspective 2
1.27.2 CHONDRITES AND THEIR COMPONENTS
3 1.27.2.1 Formation Ages of Chondritic Components 3
1.27.2.1.1 Calcium-, aluminum-rich inclusions
3 1.27.2.1.2 Chondrules 5
1.27.2.2 Ages of Secondary Events Recorded in Chondrites
5 1.27.2.2.1 Aqueous alteration 6 1.27.2.2.2 Thermal metamorphism 6 1.27.2.2.3 Shock metamorphism 7
1.27.3 DIFFERENTIATED METEORITES
8 1.27.3.1 Primitive Achondrites: Timing of Incipient Differentiation on Planetesimals 8 1.27.3.2 Basaltic and Other Achondrites: Timing of Asteroidal Differentiation and Cataclysm 9
1.27.3.2.1 Crust-formation timescales from chronology of achondrites and their components
9 1.27.3.2.2 Global differentiation timescales based on whole-rock isochrons and initial 87Sr/86Sr 11 1.27.3.2.3 Inner solar system bombardment history based on reset ages 13
1.27.3.3 Iron Meteorites and Pallasites: Timescales of Core Crystallization on Planetesimals
14
1.27.4 PLANETARY MATERIALS
16 1.27.4.1 Timing of Lunar Differentiation and Cataclysm from Chronology of Lunar Samples 16
1.27.4.1.1 Lunar differentiation history
16 1.27.4.1.2 Lunar bombardment history 17
1.27.4.2 Timescales for the Evolution of Mars from Chronology of Martian Meteorites
18
1.27.5 CONCLUSIONS
19 1.27.5.1 A Timeline for Solar System Events 19 1.27.5.2 Outlook and Future Prospects 20
REFERENCES
21
1.27.1 INTRODUCTION

1.27.1.1 Basic Principles

Long-lived radioactive isotopes, definedhere as those that have half-lives in excess of afew hundred million years, have been utilized forchronology since the early part of the twentiethcentury. The decay of a radioactive (‘‘parent’’)isotope involves its spontaneous transformation,sometimes through other intermediate radioiso-topes, into a stable (‘‘daughter’’) isotope at a rateproportional to the number of atoms of the

1

radioisotope at any given time, such that

P ¼ P0e�lt ð1Þ

where P is the number of atoms of the parentisotope remaining at present, P0 the initialabundance of the parent isotope at the timeof isotopic closure, t the time elapsed since iso-topic closure (e.g., crystallization age for a rock)and l the decay constant. Equation (1) maybe rewritten in terms of the abundance of the

2 Long-Lived Chronometers

radiogenic daughter isotope (D*) as follows:

Dn ¼ Pðelt � 1Þ ð2Þ

However, since the total number of atoms ofthe daughter isotope (D) is the sum of the ra-diogenic (D*) and the nonradiogenic (D0) com-ponents,

D ¼ D0 þ Pðelt � 1Þ ð3Þ

Normalizing to a stable isotope of the daughterelement (Ds),

D=Ds ¼ D0=Ds þ P=Ds ðelt � 1Þ ð4Þ

As such, the slope in an isochron plot for along-lived chronometer (i.e., where D/Ds isplotted versus P/Ds) is given by (elt�1), fromwhich the age (t) may be determined.

The past several decades have seen significantimprovements in the precision and accuracy ofchronological information based on the decayof long-lived radioisotopes. These have resultedparticularly from advances in the mass spectro-metric techniques for measurement of isotoperatios and better constriants on the relevant de-cay constants. Chronometers based on the de-cay of radioisotopes essentially date the time ofisotopic closure following a chemical event thatfractionated the parent element from thedaughter element. Assuming that parent/daugh-ter isotope ratios can be determined accuratelyand precisely and that the decay constant isknown, meaningful age information based onsuch chronometers may only be obtained if:(1) there was complete equilibration of the iso-topic composition of the daughter element priorto fractionation of the parent element from thedaughter element; and (2) there has been nodisturbance of isotope systematics following theisotopic closure event that is to be dated.

1.27.1.2 Application to Meteorites andPlanetary Materials: A HistoricalPerspective

Clair Patterson’s analyses of terrestrial andmeteoritic lead isotopic compositions (Patterson,

Table 1 Long-lived radioisotopes used fo

Radioisotope Daughter isotope Ref

40K 40Ar, 40Ca87Rb 87Sr147Sm 143Nd176Lu 176Hf187Re 187Os190Pt 186Os232Th 208Pb235U 207Pb238U 206Pb

1955, 1956) heralded the modern age of isotopechronology. He obtained a 207Pb/206Pb age fromthree stony meteorites of 4.5570.07Ga andsuggested that this represented the time offormation of the solar system and the Earth.Since that time, (1) advances in analytical in-strumentation (allowing more precise isotopicratio measurements), (2) more accurately deter-mined decay constants, and (3) more appro-priate sample selection have led to increasinglyrefined and precise estimates of this age. Bychance, changes in these three factors have com-pensated one another in such a way that half acentury later, Patterson’s initial estimate of theage of the solar system still agrees with the cur-rent best estimate of this age. The 207Pb/206Pbsystematics in the refractory calcium-, alumi-num-rich inclusions (CAIs), believed to beamong the first solids formed in the early his-tory of the solar system, have been utilized toprovide an estimate of the (minimum) age of thesolar system. As will be discussed in more detailin the section below, the most recent analyses oflead-isotope systematics in CAIs from theEfremovka carbonaceous (CV3) chondrite yielda highly precise age of 4,567.170.2Ma (Amelinet al., 2002, 2006).

The various long-lived radioisotopes thathave thus far been used for chronologicalinvestigations of meteorites and their compo-nents are given in Table 1. Among these, theones that have been most commonly applied arethe 40K–40Ar, 87Rb–87Sr, 147Sm–143Nd,and 235,238U,232Th–207,206,208Pb chronometers.These have mostly been used for determiningthe crystallization and secondary alteration(e.g., by shock metamorphism) ages of variousclasses of meteorites. For the same meteorites,different chronometers may date differentevents in their histories, depending on the geo-chemical behaviors of the parent and daughterelements and their ease of equilibration. Forexample, while the 40K–40Ar system in mostbasaltic eucrites is partially or totally resetas a result of shock metamorphism at 3.4–4.1Ga (Bogard, 1995), the 147Sm–143Nd ages

r chronological studies of meteorites.

erence stable isotope Half-life (109 years)

36Ar 1.2786Sr 48.8

144Nd 106177Hf 35.7188Os 41.6188Os 489204Pb 14.01204Pb 0.704204Pb 4.469

Chondrites and their Components 3

of several samples belonging to this class ofmeteorites still reflect their crystallization atB4.5Ga.

Of all the long-lived chronometers applied tometeorites so far, the combined 235,238U–207,206 Pbsystems provide the highest time resolution. Thisis so because the combination of two chronom-eters that involve the same parent and daughterelements effectively allows the determination of atime ‘‘t’’ (or a 207Pb–206Pb age) without havingto measure the parent/daughter elemental ratioand based only on the isotopic composition(207Pb/206Pb ratio) of the daughter element, whichcan be very precisely measured. Moreover, therelatively short half-life of 235U compared tothe other radioisotopes in Table 1 implies that,following a parent/daughter fractionation event,the 207Pb/206Pb ratio evolves rapidly over geolo-gic timescales, thereby allowing sub-Myr timeresolution. The 207Pb–206Pb age for a sample caneither be a single-stage model age, which is deter-mined by subtracting an assumed isotopic com-position for ‘‘common Pb’’ (which includes theinitial Pb and any extraneous Pb of terrestrialor extraterrestrial origin) from the measuredcomposition, or an isochron age. The latter isobtained from a regression of the data for mul-tiple samples, or components of a sample, on aPb–Pb isochron plot (i.e., 207Pb/206Pb versus204Pb/206Pb) to obtain the purely radiogenic207Pb/206Pb ratio (i.e., the intercept of thisisochron plot) from which an age is calculated.As long as it is reasonable to assume that allsamples plotted on a Pb–Pb isochron plot sharedthe same common lead component, the isochronmethod of calculating the age is the preferableone since no assumption of a common lead com-position need be made.

Although much valuable chronological in-formation is now being obtained from chrono-meters based on the decay of short-livedradionuclides that were present in the earlysolar system (see Chapter 1.16), long-livedchronometers (particularly those based on the235,238U–207,206Pb systems) provide the onlymeans of anchoring the relative ages providedby the extinct chronometers to an absolutetimescale. In this review, an overview is pre-sented of the chronological constraints thathave been obtained so far for events occurringin the early history of the solar system based onlong-lived radionuclides. Although results fromearlier studies are briefly summarized, the focusof this review will be on more recent reports(i.e., those published within the last decadeor so) and their implications. For additionaldetails on previous studies of early solar systemchronology based on both long- and short-livedradionuclides, the reader is referred to severalexcellent reviews (e.g., Wasserburg, 1985;

Tilton, 1988; Podosek and Nichols, 1997; Carl-son and Lugmair, 2000; Kita et al., 2005;Chapter 1.16).

1.27.2 CHONDRITES AND THEIRCOMPONENTS

1.27.2.1 Formation Ages of ChondriticComponents

1.27.2.1.1 Calcium-, aluminum-rich inclusions

CAIs are refractory millimeter- to centime-ter-sized objects found in primitive chondritemeteorite groups. They are thought to repre-sent some of the first solids that formed in thesolar protoplanetary disk. The earliest lead-iso-tope studies of CAIs (Chen and Tilton, 1976;Tatsumoto et al., 1976) indicated that thesewere indeed ancient objects that formed in theearliest history of the solar system, close to4.56Ga. Subsequently, Chen and Wasserburg(1981) reported the lead-isotope compositionsof several CAIs from the Allende carbonaceous(oxidized CV3) chondrite. Considering themost radiogenic of these samples and regres-sing these data through the Canyon Diablolead-isotope composition (assumed here as theinitial lead composition for the solar system),these authors reported an age of 4.559Ga forAllende CAIs. However, if all of the data forCAIs from Chen and Wasserburg (1981) aretaken together, they fall along a single lineararray in a Pb–Pb isochron plot that (althoughit does not pass through the Canyon Diablolead-isotope composition, implying that theseCAIs contain a common lead componentwith a composition distinct from this) yields a207Pb–206Pb age of 4,56678Ma (Tera andCarlson, 1999) (Figure 1). Following thiswork, U–Pb analyses of several otherAllende CAIs gave a consistent, but more pre-cise, 207Pb–206Pb age of 4,56672Ma (Gopelet al., 1991; Allegre et al., 1995). In more recentyears, several studies have demonstrated theimportance of the removal of common lead forobtaining high-precision 207Pb–206Pb ages formeteorites and their components (e.g., Lugmairand Galer, 1992; Amelin et al., 2002, 2005). Inparticular, using extensive acid leaching toremove the common lead component, Amelinet al. (2002) obtained a precise 207Pb–206Pb ageof 4,567.270.6Ma for two CAIs (E49 andE60) from the Efremovka carbonaceous (re-duced CV3) chondrite (Figure 2). Additionalanalyses of the E60 CAI using step-leachingand 202Pb–205Pb double-spike in combinationwith the results reported by Amelin et al. (2002)have yielded the most precise 207Pb–206Pb ageof 4,567.170.2Ma (Amelin et al., 2006).

0.9333

0.7467

1.12

+_

207 P

b/20

6 Pb

0 0.04 0.08 0.12204Pb/206Pb

PAT

Age = 4.566 Ga0.008

C-1

3529-40A3529-40BEgg-5A

Terr Pb

Egg-5B

Egg-3B

Egg-4

Egg-1 & Egg-6WA-5

Figure 1207Pb–206Pb isochron diagram of the

data of Chen and Wasserburg (1981) for AllendeCAIs. The data define a single array that corre-sponds to a 207Pb–206Pb age of 4.566Ga. PAT¼Pbisotope composition of Canyon Diablo troilite.Reproduced by permission of Elsevier from Tera

and Carlson (1999).

EfremoAge = 4,5

MSW

204Pb

207 P

b/20

6 Pb

0.000 0.001 0.0020.620

0.624

0.628

0.632

0.636

0.640

0.644

0.648

Age = 4,567.4 ± 1.1 MaMSWD = 1.09

Efremovka CAI E60

Figure 2207Pb–206Pb isochron diagram for acid-washed f

weighted average of the 207Pb–206Pb ages obtained for thPb-isotope data for the six most radiogenic analyses of acidReproduced by permission of American Association for th


However, E60 is a relatively rare type of CAI(forsterite-bearing Type B; Amelin et al., 2002),and it is unclear whether its age (the most pre-cisely defined though it is) is indeed represen-tative of that of the more common CAI types.Nevertheless, at present, this represents the bestestimate for time of formation of the earliestsolids in the solar nebula and, therefore, thebest estimate of the minimum age of thesolar system.

The 87Sr/86Sr ratio has also been used as atracer for the formation time of CAIs. In thisapproach, a formation time interval is esti-mated based on the measured initial 87Sr/86Srratio of a particular sample with a low Rb/Srratio (such as a CAI) and the time that istaken to evolve to this composition from a lessradiogenic strontium-isotope composition(such as the starting composition inferred forthe solar nebula) in an environment with agiven Rb/Sr ratio. The high Rb/Sr ratio in thesolar nebula (Anders and Grevesse, 1989;Chapter 1.03) implies that the 87Sr/86Sr ratiowould increase rapidly in material evolving insuch an environment until a major Rb/Sr frac-tionation event (such as CAI formation) definesthe initial 87Sr/86Sr ratio of the object formedduring this event. Comparison of the initial87Sr/86Sr ratios for solar system materials canpotentially resolve time differences of the orderof a million years or so. The antiquity of CAIsis indicated by their extremely unradiogenic

vka CAI E4967.17 ± 0.70 Ma

D = 0.88

Acfer 059 chondrulesAge = 4,564.66 ± 0.63 Ma

MSWD = 0.51

/206Pb

0.003 0.004 0.005

ractions from two Efremovka CAIs (E49 and E60); theese two CAIs is 4,567.270.6Ma. Also shown are the-washed chondrules From the CR chondrite Acfer 059.e Advancement of Science from Amelin et al. (2002).

0.0000 0.0008 0.0016 0.0024204Pb/206Pb

207 P

b/20

6 Pb

0.638

0.634

0.630

0.626

0.622

0.618

Gujba

Chondrules 3, 4, 5

Age = 4,562.68 ± 0.49 MaMSWD = 1.3

Figure 3 207Pb–206Pb isochron diagram for threeGujba chondrules. Reproduced by permission ofNature Publishing Group from Krot et al. (2005).


strontium isotopic compositions. A CAI (D7)from Allende has the lowest reported 87Sr/86Srratio of any solar system material (Gray et al.,1973). However, there is some complexity inthe strontium isotopic composition of CAIssince other Allende inclusions analyzed byGray et al. (1973) and Podosek et al. (1991)have 87Sr/86Sr ratios that are slightly higherthan in D7 (translating to a time span of up toB3Myr between these and D7). Furthermore,more recent analyses of the strontium-isotopecomposition of the D7 inclusion are alsoslightly higher than the initial 87Sr/86Sr ratiothat was reported for this CAI by Gray et al.(1973). Potential factors affecting this issuemay be the presence of nucleosynthetic anom-alies in the strontium-isotope composition orthe disturbance of the strontium isotopic com-position by secondary events or by contamina-tion. In this regard, nucleosynthetic anomaliesin strontium isotopes have been reported inCAIs that have been shown to record otherfractionation and unknown nuclear (FUN) ef-fects (Papanastassiou and Wasserburg, 1978;Loss et al., 1994).

1.27.2.1.2 Chondrules

Chondrules are sub-millimeter to centimeter-sized ferromagnesian silicate spherules foundin chondrites. Although, in detail, there areseveral hypotheses for the exact mechanisminvolved in chondrule formation (currently, thetwo leading ones being the X-wind and theshock-wave models; see Ciesla, 2005 and refer-ences therein): they are generally considered tohave resulted from transient heating events inthe solar nebula. The earliest lead-isotope studyof chondrules was performed on those sepa-rated from the Allende chondrite and gave anage of 4,560767Ma (Chen and Tilton, 1976).The low precision of this date was due to therelatively large unradiogenic Pb component inthese chondrules and small spread in the207Pb/206Pb ratios. More recently, using ex-tensive leaching procedures to remove the com-mon lead component, the chondrules from acarbonaceous chondrite belonging to the CRgroup were precisely dated at 4,564.770.6(Amelin et al., 2002) (Figure 2), which indi-cates that these were formed 2.470.6Myr afterCAIs. Lead-isotope compositions of chondrulesfrom Allende have also been recently analyzedand yield an older age of 4,566.771.0Ma(Amelin et al., 2004), which overlaps withinthe uncertainties with the 207Pb–206Pb age ofCAIs. As such, chondrules from these primitivechondrite groups define ages that suggest thatthey began forming almost contemporaneouslywith CAIs and continued to form for at least

another 2–3Myr afterwards. Recently, Amelinet al. (2005) reported a 207Pb–206Pb age of4,562.771.7Ma from pyroxene-rich chondrulesand chondrule fragments from the RichardtonH5 equilibrated ordinary chondrite. Given theequilibrated nature of this chondrite, theseauthors argued that this age was the minimumage for the formation of chondrules in thissample (possibly corresponding to the time ofcessation of lead loss).

Chondrules from the metal-rich CB chondritesGujba and Hammadah al Hamarah 237 gave207Pb–206Pb ages of 4,562.770.5Ma (Figure 3)and 4,562.870.9Ma, respectively (Krot et al.,2005). It had previously been suggested thatchondrules in the CB chondrites originated as aresult of an energetic impact between large plan-etesimals (Rubin et al., 2003) and these relativelyyoung ages have been used in support of thishypothesis (Krot et al., 2005).

1.27.2.2 Ages of Secondary Events Recordedin Chondrites

Chondritic meteorites record a variety of sec-ondary alteration processes, including aqueousalteration, thermal and shock metamorphism,and brecciation. These events spanned a longtime interval, beginning almost close to the for-mation of the solar system at B4.56Ga andextending over a period of several millions ofyears thereafter. The following is a discussion ofthe timescales for these different secondaryalteration processes inferred from variouslong-lived chronometers, although the bounda-ries between these processes are not always welldefined (e.g., an energetic impact event on aplanetesimal may result in shock metamorphismas well as thermal processing).


1.27.2.2.1 Aqueous alteration

Aqueous alteration of chondrites and theircomponents is thought to have occurred in avariety of settings, including the solar nebulaand on accreted parent bodies (Zolensky andMcSween, 1988; Brearley, 2006; Chapter 1.09).The CI carbonaceous chondrites contain sev-eral secondary mineral phases (particularly car-bonates and sulfates) that provide a record ofaqueous alteration on their parent body. Thecarbonates are a good candidate for datingusing the initial 87Sr/86Sr approach sincethis mineral is typically characterized by lowRb/Sr ratios. Macdougall et al. (1984) reportedthe strontium isotopic compositions ofseveral carbonate separates (dolomite andbreunnerite) from the Orgueil CI chondrite.These authors showed that Orgueil carbonateshave a range of 87Sr/86Sr ratios, with thelowest one being similar to the most primitiveSr-isotope composition measured in AllendeCAIs (Gray et al., 1973; Podosek et al., 1991).This implies that the onset of aqueousalteration on the CI chondrite parent body oc-curred essentially contemporaneously with itsformation.

1.27.2.2.2 Thermal metamorphism

Most chondrites have experienced somedegree of thermal metamorphism, defined hereas alteration resulting from heating at temper-atures in the range of 400–1,000 1C at lowlithostatic pressure for extended time periods(McSween et al., 1988; Huss et al., 2006). Cal-cium phosphates in chondritic meteorites areminor but uranium-rich secondary mineralsthat were formed during this thermal process-ing of the parent bodies of these meteorites,most likely by oxidation of phosphorus-richmetal (Perron et al., 1988). As such, U–Pb sys-tematics in secondary phosphates from chond-rite groups that have experienced differentdegrees of metamorphic equilibration canprovide constraints on the timescales involvedin thermal processing of planetesimals follow-ing their accretion. The first studies of U–Pbsystematics in phosphates from a chondritewere performed on the LL6 equilibrated ordi-nary chondrite Saint-Severin (Manhes et al.,1978; Chen and Wasserburg, 1981). The resultsfrom these studies are consistent with a laterinvestigation of phosphates from this same me-teorite (Gopel et al., 1994) and together yield a207Pb–206Pb age of 4,55876Ma. In addition toSaint-Severin, Gopel et al. (1994) reportedU–Pb systematics in phosphates from 14 otherequilibrated ordinary chondrites belonging tothe H4, H5, H6, L5, L6, LL5, and LL6 groups.

For the H chondrites, these authors noted acorrelation between the 207Pb–206Pb ages of thephosphates and their degree of metamorphism.Specifically, 207Pb–206Pb model ages (with atypical precision of 71Ma) ranged from4,563Ma for the Ste. Marguerite H4 chondriteto 4,504Ma for the Guarena H6 chondrite,thereby indicating that thermal processing ofthe H chondrite parent body(ies) extendedover a period of B60Myr. The 207Pb–206Pbage of phosphate fractions from the Richard-ton H5 chondrite (4,550.772.6Ma; Amelinet al., 2005) falls within this time span. Inthe case of the L chondrites, 207Pb–206Pb modelages for their phosphates ranged from 4,543 to4,511Ma, while for the LL chondrites theseages ranged from 4,557 to 4,536Ma. These agesalso suggest that thermal metamorphism of theL and LL parent bodies had extended for tensof millions of years in the early history of thesolar system.

Although less precise than the U–Pb chro-nometer, the 39Ar–40Ar technique has also beenapplied toward constraining the duration ofthermal processing of chondrite parent bodies.Some equilibrated (but unshocked) ordinarychondrites show a range of 39Ar–40Ar agesfrom B4.5 toB4.4Ga (with a typical precisionof 730Ma) (Turner et al., 1978; Hohenberget al., 1981) which, to first order at least, iscomparable to the duration indicated by the U–Pb systematics in phosphates from the ordinarychondrites. The initial 87Sr/86Sr method(described in the previous section) may addi-tionally be used to assess the duration of ther-mal processing of chondrite parent bodies. Twoexamples of the application of this approach(i.e., to phosphates from Beaver Creek andGuarena chondrites) toward obtaining chron-ological information regarding formationof secondary phosphates are illustrated inFigure 4. Time intervals of tens of millions ofyears are obtained based on the initial 87Sr/86Srisotopic compositions reported for ordinarychondrite phosphates (Wasserburg et al., 1969;Manhes et al., 1978; Brannon et al., 1988; Po-dosek and Brannon, 1991) and assumingevolution from a primitive strontium isotopicvalue similar to the average value for Allendeinclusions (Gray et al., 1973; Podosek et al.,1991) and equilibration of strontium isotopeson the whole-rock scale. This is again broadlyconsistent with the duration of metamorphicevents as indicated by the 207Pb–206Pb and39Ar–40Ar ages discussed above. However, eventhough all three of these chronometers seem tobe indicating generally similar timescales oftens of millions of years for the thermal proc-essing on the chondrite parent bodies, whenspecific ages from these three chronometers are

Interval (Ma)100 80 60 40 20 0

0.6985

0.6990

0.6995

87S

r/86

Sr

�ALL

0.7000

0.7005

0ALL

BabI 5

10

15

20Nebular evolution

Guarena

4.564.524.48

Age (Ga)

phosphate

Beaver Creek

Bulk chondriteevolution

phosphate

Figure 4 Two examples (phosphates from theBeaver Creek and Guarena chondrites) illustratingthe application of the initial 87Sr/86Sr approach to-ward obtaining chronological constraints. ALL¼initial 87Sr/86Sr ratio for the D7 Allende CAI withthe most primitive composition (Gray et al., 1973).Steeper solid line shows evolution of 87Sr/86Sr in anebular environment (87Rb/86SrB1.5); shallowersolid line shows evolution with bulk chondritic87Rb/86Sr (B0.75). Shaded gray bands are extra-polations based on the measured Rb–Sr systematicsin phosphates from the Beaver Creek and Guarenachondrites (the width of the bands indicating theuncertainties). As shown here, the time interval be-tween formation of Allende CAIs and formation ofsecondary phosphates in these two chondrites (i.e.,indicated on the x-axis by the points of the intersec-tion of the gray bands with the bulk chondriteevolution line) is tens of millions of years. Repro-duced by permission of Meteoritics and Planetary

Science from Podosek and Brannon (1991).


considered in detail (as was done by Gopelet al., 1994), there does not appear to be anycorrelation between them. This could be indi-cative of complex histories of ordinary chond-rite parent bodies following their initialaccretion, which might affect these threechronometers differently.

Finally, 87Rb–87Sr investigations of chond-rules additionally indicate that thermal process-ing of chondrite parent bodies extented toB4.4Ga. Chondrules from the Richardtonordinary equilibrated (H5) chondrite yielda 87Rb–87Sr age of 4.4570.03Ga (Evensenet al., 1979). Studies of chondrules from the Al-lende carbonaceous chondrite also indicate thatthe 87Rb–87Sr system was significantly affectedby late thermal processing (Gray et al., 1973;Tatsumoto et al., 1976; Shimoda et al., 2005). Inparticular, Shimoda et al. (2005) suggest thatthe mesostasis-rich chondrules (considered tobe most susceptible to disturbance, particularly

in terms of alkali elements such as rubidium)would best record this late processing. Assuch, consideration of only the mesostasis-rich chondrules from their study and fromprevious investigations (Gray et al., 1973;Tatsumoto et al., 1976) yields a 87Rb–87Sr ageof 4.3670.08Ga. Within errors, this age issimilar to that obtained for the Richardtonchondrules.

1.27.2.2.3 Shock metamorphism

Impacts between solar system bodies haveplayed an important role in their evolutionaryhistories and many chondritic meteorites pre-serve a record of these events (Stoffler et al.,1988). The timing of impact events affectingchondrite parent bodies have been determinedpredominantly with the 39Ar–40Ar datingmethod (e.g., Turner, 1969; Bogard et al.,1976; Bogard and Hirsch, 1980; Kaneoka,1981; Stephan and Jessberger, 1988; Kringet al., 1996; Grier et al., 2004), although insome cases other isotope systems such as Rb–Sr, Pb–Pb, and Sm–Nd (roughly in that orderof susceptibility to disturbance) are also af-fected by such events. Most experimental workon the effects of shock alone on various iso-topic chronometers indicates that shock pres-sures up to B60GPa are usually insufficient toreset these chronometers (Deutsch and Scharer,1994). However, if the shock event is associatedwith thermal annealing, it can easily affect someisotopic systems (particularly the K–Ar system,but also Rb–Sr) (e.g., Nyquist et al., 1991). Inthe particular case of the 39Ar–40Ar datingmethod, if the shocked sample is insufficientlyheated or if it cooled rapidly, there may be onlya small amount of the radiogenic 40Ar lost fromthe sample and thus the K–Ar system wouldonly be partially reset. On the other hand, if theshock is accompanied by extended thermal an-nealing (and also if the sample had small grainsizes facilitating diffusional loss of argon), thenthe K–Ar system may be almost totally reset bythis event (at least in some minerals of ashocked sample that may be more susceptibleto being reset). In any case, the 39Ar–40Ar ageof the shock event is determined by using astepwise temperature release of argon, whichhelps to separate the K–Ar chronologies ofdifferent minerals of the shocked sample.

Bogard (1995) summarized the impact agesof various chondrite classes, that are mostlybased on the 39Ar–40Ar method. Most chondri-tic meteorites have 39Ar–40Ar ages that areyounger than B1.3Ga (Figure 5). These youngages are considered to be reflecting relativelyfew impact events and, in fact, the peak at

0

2

00.1 0.3 0.5 0.7 0.9 1.1 1.3

4

6

8

10

12

14

0 0.4 0.8 1.2 1.6 2 2.4

Ar−Ar Age in Ga

Plateau Intercept

Ar−Ar Age in Ga

L-chon

H-chon

LL-chonN

umbe

r

Num

ber

2.8 3.2 3.6 4.0 4.2

2

4

6

8

10

12

14

Figure 5 Histograms of impact-reset 39Ar–40Ar ages of ordinary chondrites, with plotted age interval of0.1Ga. Inset shows an expanded scale figure with data for chondrites having impact ages of o1.3Ga. In mainfigure, ages defined by a significant age plateau (‘‘plateau’’) are in black while those defined primarily from theage intercept of a diffusion profile (‘‘intercept’’) are in gray. Reproduced by permission of Meteoritics and

Planetary Science from Bogard (1995).


B0.5Ga for the L class of ordinary chondritesin the histogram shown in Figure 5 is thoughtto be the result of a single large impact thatcatastrophically disrupted the parent body ofthese meteorites (Haack et al., 1996a). At leastB4 distinct impact events are required toaccount for the Ar–Ar ages of most chond-rites. The B0.3 and B0.5Ga events are thebest defined and affected the L and H ordinarychondrite parent bodies. Additional events atB0.9Ga (affecting the L and H parent bodies)and at B1.2Ga (affecting the LL parent body)are also indicated.

The Ar–Ar and Rb–Sr ages of some chond-rites, however, indicate significantly older im-pact ages of B3.5–4.0Ga (e.g., Keil et al.,1980; Stephan and Jessberger, 1988; Nakamuraet al., 1994). These ages most likely reflect thetime of heavy bombardment experienced bybodies in the inner solar system. This event hasalso been recorded in the impact-reset ages ofmany achondrites (see Section 1.27.3.2.3) andlunar samples (see Section 1.27.4.1.2).

1.27.3 DIFFERENTIATED METEORITES

1.27.3.1 Primitive Achondrites: Timing ofIncipient Differentiation onPlanetesimals

Primitive achondrites, such as acapulcoites,lodranites, winonaites, and brachinites, areconsidered to be the products of the earliest

stages of melting and igneous processing onplanetesimals (see Mittlefehldt et al., 1998and references therein). The acapulcoites andlodranites are thought to be the residualproducts of partial melting of chondritic pre-cursors (Mittlefehldt et al., 1996; McCoy et al.,1997a, b). 147Sm–143Nd systematics determinedby Prinzhofer et al. (1992) for Acapulco gave avery old age (4.6070.03Ga), which the authorsinterpreted as the time of recrystallization im-mediately following its formation event. Ura-nium–lead systematics in phosphates fromAcapulco give a 207Pb–206Pb model age of4,55772Ma (Gopel et al., 1992, 1994), indica-ting that this achondrite formed approximatelyB10Myr after the formation of CAIs (at4,567.170.2Ma; Amelin et al., 2002, 2006).More recently, a 207Pb–206Pb isochron age forAcapulco phosphates and mixed grain frac-tions of 4,556.5270.78Ma (or 10.670.8Myrafter CAI formation) has been reported (Am-elin and Pravdivtseva, 2005; Amelin et al.,2006). This 207Pb–206Pb age for Acapulco ismarginally younger, but much more precise,than the 147Sm–143Nd age. McCoy et al. (1996)have argued that the older 147Sm–143Nd agecould be due to disturbance during extensivelater metamorphism experienced by this mete-orite. The Divnoe meteorite is an ultramaficprimitive achondrite whose relationship withother primitive achondrite groups is as yet un-clear (Petaev et al., 1994; Weigel et al., 1996).This meteorite also has an old 147Sm–143Nd age

Differentiated Meteorites 9

of 4.6270.07Ga. Although uncertainties arelarge (Bogdanovski and Jagoutz, 1996), theyoungest formation time indicated by this147Sm–143Nd age is B17Myr after CAI forma-tion. The above discussion shows that the onsetof melting on some planetesimals, as evidencedby primitive achondrites such as Acapulco andDivnoe, occurred within B10–20Myr of thebeginning of the solar system.

0.629

0.628

0.627

0.626

0.625

0.6240.0000 0.0004 0.0008 0.0012

207 P

b/20

6 Pb

204Pb/206Pb

MSWD = 1.5 (n = 7)4,566. 18 ± 0.14 Ma

Figure 6 207Pb–206Pb isochron diagram for frag-ments of Sahara 99555 and NWA 1296 angritesand acid-washed pyroxene from Sahara 99555.Reproduced by permission of Nature Publishing

Group from Baker et al. (2005).

1.27.3.2 Basaltic and Other Achondrites:Timing of Asteroidal Differentiationand Cataclysm

1.27.3.2.1 Crust-formation timescales fromchronology of achondrites and theircomponents

Primary crystallization ages of individualmembers of achondrite groups such as theangrites and noncumulate eucrites, which rep-resent basaltic rocks that formed in asteroidalnear-surface environments, provide constraintson the timing of silicate differentiation andcrust formation on planetesimals during theearly history of the solar system.

Angrites are a small group of mineralogicallyunique basalts composed mostly of Ca–Al–Ti-rich pyroxenes (fassaite), olivine and anorthiticplagioclase (see Mittlefehldt et al., 1998 and ref-erences therein; Chapters 1.05 and 1.11).147Sm–143Nd systematics in Angra dos Reis(ADOR) and LEW 86010 (LEW) are well-behaved and give old crystallization ages be-tween 4.5370.04 and 4.5670.04Ga (Lugmairand Marti, 1977; Wasserburg et al., 1977; Jacob-sen and Wasserburg, 1984; Lugmair and Galer,1992; Nyquist et al., 1994). 147Sm–143Nd sys-tematics have also been determined in the morerecently discovered angrite D’Orbigny, and, de-spite some disturbance evident in the plagio-clase, possibly due to late metamorphism and/orterrestrial weathering, are generally consistentwith earlier results for ADOR and LEW(Nyquist et al., 2003a; Tonui et al., 2003). Themost precise estimate of the crystallization ageof the angrites is offered by their 207Pb–206Pbsystematics. A 207Pb–206Pb isochron defined byLEW minerals gives an age of 4,558.273.4Ma,concordant with the highly precise model age of4,557.870.5Ma obtained from the extremelyradiogenic lead compositions in the pyroxenesof ADOR and LEW (Lugmair and Galer,1992). Preliminary model ages derived fromthe lead-isotope compositions of the D’Orbignypyroxenes appeared to be in agreement withthose derived from ADOR and LEW pyroxenes(Jagoutz et al., 2003). However, a recent reeval-uation of these data by Jagoutz and colleagueshas resulted in a somewhat older age of

4,56371Ma for D’Orbigny (G. W. Lugmair,personal communication). This revised age is inagreement with the results from a more recentstudy that reported a 207Pb–206Pb model ageof 4,563.870.6Ma for D’Orbigny (Zartmanet al., 2006). Finally, Baker et al. (2005)reported a highly precise and extremely ancient207Pb–206Pb isochron age of 4,566.1870.l4Mafor the Sahara 99555 angrite (Figure 6).Given the 207Pb–206Pb age for CAIs of4,567.170.2Ma (Amelin et al., 2002, 2006), thissuggests that basalts were forming on the sur-face of the angrite parent body within B1Myrof CAI formation. The above discussion showsthat the 207Pb–206Pb ages of the angrites span atime interval of almost B8Myr, the youngest(LEW and ADOR) having an age of 4,558Maand the oldest (Sahara 99555) being 4,566Ma.

Like the angrites, the noncumulate eucritesare pyroxene–plagioclase rocks. However,there are significantly greater numbers ofknown noncumulate eucrites than there areangrites. Recent high-precision oxygen-isotopedata of Wiechert et al. (2004) demonstrate thatmost noncumulate eucrites along with the cu-mulate eucrites, diogenites, and howarditeshave uniform D17O (within 70.02%) and thuslie on a single mass fractionation line, consist-ent with their origin on a single parent body.Therefore, these basalts are the most numerouscrustal rocks available from any single solarsystem body other than the Earth andthe Moon. A handful of the noncumulate euc-rites (in particular Ibitira, but possibly alsoCaldera, Pasamonte, and ALHA 78132) haveoxygen-isotope compositions distinct from theothers, implying either that these samplesoriginated on different parent bodies or thatisotopic heterogeneity was preserved on the


Howardite–Eucrite–Diogenite (HED) parentbody (Wiechert et al., 2004). Unlike angrites(which did not undergo any significant degreeof recrystallization or metamorphism), thenoncumulate eucrites appear to record a pro-tracted history of extensive thermal processingon their parent body subsequent to theiroriginal crystallization. As a result, many ofthe chronometers investigated in these samplesappear to record varying degrees of disturbancefrom secondary thermal events. Nevertheless,there are several lines of evidence that suggestthat these basalts crystallized very early in thehistory of the solar system. Although typicallycharacterized by large uncertainties, the147Sm–143Nd ages of several of these samplessuch as Chervony Kut (4,580730Ma; Wad-hwa and Lugmair, 1995), Juvinas(4,560780Ma; Lugmair, 1974), Pasamonte(4,5807120Ma; Unruh et al., 1977), PipliaKalan (4,570723Ma; Kumar et al., 1999), andYamato 792510 (4,570790Ma; Nyquist et al.,1997) are close to B4.56Ga. In some caseswhere 147Sm–143Nd systematics appear torecord ages younger than B4.56Ga, the 146Sm–142Nd systematics hint at earlycrystallization of basaltic eucrites such as Cal-dera (146Sm/144Sm¼ 0.007370.0011; Wadhwaand Lugmair, 1996) and Ibitira (146Sm/144SmB0.00970.001; Prinzhofer et al., 1992). Thismay be explained by a model proposed byPrinzhofer et al. (1992), according to which theapparent discrepancy between the long-lived147Sm–143Nd and the short-lived 146Sm–142Ndsystems can be interpreted in terms of a shortepisodic disturbance resulting in partial reequi-libration of the rare earth elements (REEs),predominantly between plagioclase (which hasvery low REE abundances) and phosphates(which are the primary REE carriers). Asshown by the modeling results of these au-thors, such a disturbance could partially resetthe 147Sm–144Nd isochron, without having aresolvable effect on the 146Sm–142Nd system.

The Rb–Sr chronometer also generally indi-cates ancient formation ages for the noncumu-late eucrites (e.g., Allegre et al., 1975; Nyquistet al., 1986). The most precise of the absolutechronometer, that is, the U–Pb system, appearsto have been affected by postcrystallizationevents and terrestrial Pb contamination in mostnoncumulate eucrites, recording mineral iso-chron ages in the range of 4,128–4,530Ma(Tatsumoto et al., 1973; Unruh et al., 1977;Galer and Lugmair, 1996; Tera et al., 1997).However, Ibitira whole-rock samples withradiogenic Pb isotopic compositions gave old207Pb–206Pb model ages of 4,55676Ma (Chenand Wasserburg, 1985) and 4,56073Ma(Manhes et al., 1987). Furthermore, a recent

determination of the 207Pb–206Pb mineral is-ochron for the Asuka 881394 eucrite yielded aprecise and ancient age of 4,566.570.3Ma(Amelin et al., 2006). This is only 0.670.4Myryounger than the time of CAI formation(4,567.170.2Ma; Amelin et al., 2002, 2006)and, as in the case of the Sahara 99555 angritediscussed earlier, indicates that crust formationon differentiated planetesimals occurred withinB1Myr of the formation of CAIs.

Basaltic noncumulate eucrites thus showclear evidence of having formed close toB4.56Ga in the crust of their parent planet-esimal. In contrast, cumulate eucrites, whichformed in the crust of the same parent planet-esimal as the noncumulate eucrites (Claytonand Mayeda, 1996; Wiechert et al., 2004), havesignificantly younger concordant Sm–Nd andPb–Pb ages, ranging from the oldest of4,456725Ma (Sm–Nd) and 4,484719Ma(Pb–Pb) for Moore County (Tera et al., 1997)to the youngest of 4,410720Ma (Sm–Nd;Lugmair et al., 1977) and 4,399735Ma (Pb–Pb; Tera et al., 1997) for Serra de Mage. Thus,Sm–Nd and Pb–Pb systematics in the cumulateeucrites indicate that isotopic closure occurredup to B150Myr after the noncumulate eucrites(Lugmair et al., 1977; Jacobsen and Was-serburg, 1984; Lugmair et al., 1991; Teraet al., 1997), possibly implying that activemagmatism persisted on the eucrite parentbody (EPB) for this extended period (Teraet al., 1997).

Since, as discussed above, the oldest basalticmeteorites formed within B10Myr of CAI for-mation (and some, specifically the angrite Sa-hara 99555 and the eucrite Asuka 881394,within only B1Myr), the decay of short-lived radionuclides such as 26Al and 60Feis the likely heat source for the early and ex-tensive differentiation experienced by their par-ent planetesimal. Energy sources that canaccount for later igneous activity (i.e., tens ofmillions of years after CAI formation) on smallplanetesimals are not obvious unless the cumu-late eucrites are the crystallization products ofimpact melting on the EPB. Alternatively (andperhaps more likely), since the cumulate euc-rites are slowly cooled rocks that possiblyformed deeper within the crust of their parentbody than noncumulate eucrites, the long-livedchronometers could be recording the long cool-ing times required to achieve subsolidus tem-peratures. This is supported by the modelingresults of Ghosh and McSween (1998), whichshow that, assuming reasonable parameters, itis possible to maintain temperatures in excessof the solidus temperature for basalt at a depthof B100 km for over B100Myr in a Vesta-sized planetesimal.

0.0000.6988

0.6990

0.6992

0.6994

0.6996

0.6998

0.7000

0.7002

0.7004Noncumulate eucrites

T = 4.548 ± 0.058 Ga

Juvinas PxBereba

Jonzak

Nuevo Laredo

Sioux County

Pomozdino

Juvinas

JuvinasPlag

Ibitira

I (87Sr/86Sr) = 0.698925 ± 14

0.005 0.01087Rb/86Sr

87S

r/86

Sr

0.015 0.020

Figure 7 87Rb–87Sr whole-rock isochron for non-cumulate eucrites. All data points indicate whole-rocks samples or clasts, with the exception of twomineral fractions from Juvinas (Px¼ pyroxene;Plag¼ plagioclase) that are also plotted here. Dataare from various literature sources given in Smoliar(1993). Adapted by permission of Meteoritics and

Planetary Science from Smoliar (1993).


Besides the angrites and the eucrites, thereare other types of achondrites, such as aubritesand ureilites, whose origins are somewhatenigmatic but which were nevertheless formedin the crusts of extensively differentiated aster-oidal bodies. There are very few chronologicalinvestigations of the aubrites, which are essen-tially monomineralic rocks composed ofcoarse-grained enstatite. U–Th–Pb and Sm–Nd isotope systematics in the ureilites generallyindicate early formation, although there areapparent complications resulting from laterdisturbance during a metasomatic event onthe ureilite parent body and/or by recent ter-restrial contamination (Goodrich and Lugmair,1995; Torigoye-Kita et al., 1995a, b).

1.27.3.2.2 Global differentiation timescalesbased on whole-rock isochrons andinitial 87Sr/86Sr

While an internal mineral isochron canprovide constraints on the timing of formationof an individual achondrite in the crust of aplanetesimal, a whole-rock isochron of a par-ticular achondrite group can provide limits onthe timing of the major fractionation event thatestablished the parent/daughter element ratioin the whole rocks (which could have pre-datedthe formation of an individual sample). De-pending on the geochemical behavior of theparent and daughter elements, this major par-ent/daughter fractionation recorded in thewhole rocks may reflect a global silicate frac-tionation event (possibly associated with crys-tallization of a magma ocean) that establishedthe source characteristics for these rocks, or itcould simply reflect crystal fractionation from aparental melt, which resulted in the formationof these rocks. Whole-rock Rb–Sr isochronsfor the basaltic eucrites established early onthat Rb–Sr fractionation in the mantle sourcereservoir of these achondrites occurred close toB4.6Ga (Papanastassiou and Wasserburg,1969; Birck and Allegre, 1978). Smoliar (1993)evaluated all available Rb–Sr data for the euc-rites and obtained a whole-rock isochron age of4.5570.06Ga for the noncumulate eucrites(Figure 7). A whole-rock 147Sm–143Nd iso-chron for 18 noncumulate and cumulate euc-rites was reported by Blichert-Toft et al. (2002)and gave a relatively young age of4,464775Ma (Figure 8). Since the slope ofthis isochron is controlled by the cumulateeucrites (which have the most fractionatedwhole-rock Sm/Nd ratios), the authors inter-preted this to imply that cumulate eucrites wereformed B100Myr after solar system forma-tion. This is supported by the relatively young

207Pb–206Pb ages and 147Sm–143Nd internalisochron ages of the cumulate eucrites (see ear-lier discussion in this section). Patchett andTatsumoto (1980) reported the first whole-rock 176Lu–176Hf isochron for the eucrites. Atthe time, the half-life of 176Lu was not well-constrained and these authors assigned anage of 4.55Ga for this whole-rock isochronand thus estimated a half-life of 35.3Gyr(corresponding to a decay constant for 176Luor l176Lu of 1.96� 10�11 yr�1). A more recentstudy of 176Lu–176Hf systematics in wholerocks of eucrites by Blichert-Toft et al. (2002)reported an errorchron corresponding toan age of 4,604739Ma for the noncumu-late eucrites (Figure 9). The cumulate eucrites,on the other hand, defined a 176Lu–176Hf iso-chron with an age of 4,470722Ma (Figure 9),indistinguishable from their 147Sm–143Ndwhole-rock age. Based on these whole-rock147Sm–143Nd and 176Lu–176Hf systematics inthe eucrites, these authors suggested that cu-mulate eucrites were formed B100Myr afterthe noncumulate eucrites, while the latter wereformed close to the beginning of the solar sys-tem. In their study, Blichert-Toft et al. (2002)assumed a l176Lu value of B1.93� 10�11 yr�1

(similar to that suggested by Patchett andTatsumoto, 1980). However, studies of terres-trial samples on the one hand and meteoriticsamples on the other indicate that the half-lifeof 176Lu may be either 37.2Gyr (l176Lu

B1.86� 10�11 yr�1; Scherer et al., 2001) or

0.5142

0.5140

0.5138

0.5136

0.5134

0.5132

0.5130

0.5128

0.5126

0.5124

0.51220.185 0.195

Pasamonte

Caldera

0.205 0.215 0.225 0.235 0.245 0.255 0.265147Sm/144Nd

143 N

d/14

4 Nd

Talampaya Serra de MagèMoore county

BSE

Basaltic eucrites

Cumulate eucrites

Not regressed

T = 4,464 ± 75 Ma

MSWD = 1.26

(143Nd/144Nd)i = 0.50680 ± 0.00010

Nagaria

Moama

Figure 8 147Sm–143Nd whole-rock isochron for eucrites. Reproduced by permission of Elsevier from Blichert-Toft et al. (2002).

0.0200.2810

0.2820

0.2830

0.2840

0.2850

0.2860

0.2870

0.2880

0.2890

0.2900

0.2910

0.030

CalderaMoore County

Serra de Magé

TalampayaNagaria

MoamaBasaltic eucritesCumulate eucrites

0.040 0.050 0.060176Lu/177Hf

176 H

f/177 H

f

0.070 0.080 0.090 0.100 0.110

(176Hf/177Hf)i = 0.27966 ± 0.00002

Figure 9 176Lu–176Hf whole-rock isochron for the eucrites. A statistically significant isochron cannot beobtained if all data are considered together. If only noncumulate eucrites are considered (with the exception ofone sample, Palo Blanco Creek) an errorchron corresponding to an age of 4,604739Ma is obtained. If onlythe three cumulate eucrites Moama, Moore County, and Serra de Mage are regressed together, they yield anisochron corresponding to an age of 4,470722Ma. Reproduced by permission of Elsevier from Blichert-Toft

et al. (2002).


34.9Gyr (l176LuB1.98� 10�11 yr�1; Bizzarroet al., 2003), respectively. Amelin and Davis(2005) have shown that this apparent discrep-ancy cannot be accounted for by a possiblebranched decay of 176Lu. Although other pos-sible explanations were evaluated by theseauthors, none were considered to be plausible.As such, the half-life of 176Lu still remains

uncertain, thereby limiting the usefulness of the176Lu–176Hf chronometer at the present time.

Finally, the ancient ages for the establish-ment of the highly volatile-depleted mantlesource reservoirs of the angrites and theeucrites on their respective parent bodies canbe inferred from their initial 87Sr/86Sr ratios.Papanastassiou and Wasserburg (1969) first


estimated the initial 87Sr/86Sr of the EPB (ba-saltic achondrite best initial or BABI). At thetime, this was the most primitive known stron-tium isotopic composition. Subsequently, how-ever, an even more primitive strontium isotopiccomposition was reported for an Allende CAI(ALL) (Gray et al., 1973). Assuming that theinitial 87Sr/86Sr ratio at the beginning of thesolar system is represented by the average in-itial 87Sr/86Sr ratio measured in Allende CAIs(Gray et al., 1973; Podosek et al., 1991), andthat subsequent evolution of radiogenic stron-tium occurred in an environment with solarRb/Sr ratios, the initial 87Sr/86Sr ratios of theangrites (Lugmair and Galer, 1992; Nyquistet al., 1994, 2003a; Tonui et al., 2003) translateto an age difference of B4Myr between CAIformation and the timing of Rb/Sr depletionevent that established the angrite source char-acteristics. The very low initial 87Sr/86Sr ratiosof the eucrites similarly indicate that the vola-tile depletion characterizing the EPB may haveoccurred early in the history of the solar system(Carlson and Lugmair, 2000, and referencestherein). A re-evaluation of the strontium-iso-tope data for the eucrites by Smoliar (1993)shows that whole-rock Rb–Sr isochrons for thenoncumulate and the cumulate eucrites defineslightly, but resolvably, different initial87Sr/86Sr ratios (that are both distinctly lowerthan the eucrite initial, BABI, previously de-fined by Papanastassiou and Wasserburg,1969). In fact, the initial 87Sr/86Sr ratio forthe cumulate eucrites proposed by Smoliar(1993) is, within errors, similar to that for theangrites (e.g., Lugmair and Galer, 1992),suggesting that the volatile depletion in theirsources was established at similar times (possi-bly B4Myr after CAI formation; see above).However, the slighter higher initial 87Sr/86Srratio defined by the noncumulate eucrites ispotentially problematic since the simplest in-terpretation would be that their source evolvedwith a chondritic Rb/Sr ratio for B4Myr longer (and is thus younger) than that ofthe cumulate eucrites, further implyingthat these two types of eucrites originated ondistinct parent bodies (Smoliar, 1993). This isinconsistent with recent high-precision oxygen-isotope data (Wiechert et al., 2004) that suggestthat the cumulate and noncumulate eucrites(with the possible exception of Ibitira) origin-ated on a common parent planetesimal. Asdiscussed by Carlson and Lugmair (2000), apossible explanation could be that the severevolatile depletion on the EPB did not occur in asingle-step process, but rather took placeover the course of its accretionary and earlydifferentiation history. Subsequently, the proc-ess of magma ocean crystallization may have

resulted in a slighter higher Rb/Sr ratio in thesource of the noncumulate eucrites comparedto that of the cumulate eucrites, therebyresulting in the higher initial 87Sr/86Sr ratio in-dicated by the whole-rock isochron for thenoncumulate eucrites.

1.27.3.2.3 Inner solar system bombardmenthistory based on reset ages

As discussed previously for chondritic mete-orites, the 39Ar–40Ar technique has also beenwidely applied toward dating the thermal his-tories of achondritic meteorites, and particu-larly for determining the ages of impact eventson their parent bodies (e.g., Podosek and Hun-eke, 1973; Bogard et al., 1990; Takeda et al.,1994; Bogard and Garrison, 2003). As is thecase for severely shocked chondritic samples,other isotopic chronometers such as Rb–Sr,Pb–Pb, and Sm–Nd in some differentiated ac-hondrites that have experienced postshockthermal annealing record varying degrees ofdisturbance (e.g., Unruh et al., 1977; Nyquistet al., 1990). The impact ages of most achond-rites (particularly the HED meteorites) fallwithin a relatively narrow time interval ofB3.4–4.1Ga (Bogard, 1995; Bogard and Gar-rison, 2003) (Figure 10). As discussed in theearlier sections, these achondrites have crystal-lization ages that are close to B4.56Ga and sothe age distribution shown in Figure 10 may bereasonably interpreted to reflect the timing ofthermal metamorphism on the HED parentbody resulting from large impacts (which areconsidered to be the most plausible heat sourcefor these late events). Most of the HEDs arebrecciated rocks that preserve clear textural,mineralogical, and chemical evidence of shockmetamorphism (e.g., Stoffler et al., 1988; Met-zler et al., 1995), further supporting the aboveinterpretation.

Based on the age distribution shown in Figure10, it has been suggested that there was apeak in the impactor flux (i.e., a cataclysm) onthe HED parent body at B3.7Ga, whichdecreased sharply and tailed down to agesslightly younger than B3.4Ga (Bogard, 1995;Bogard and Garrison, 2003). Therefore, the39Ar–40Ar ages of the HED meteorites suggestthat the region where their parent body residedin the inner solar system experienced a period ofheavy bombardment during the time intervalextending from B4.1Ga until at least B3.4Ga.The peak at B4.48Ga in Figure 10 has beenexplained by Bogard and Garrison (2003)in terms of a single large impact that excavatedthese eucritic meteorites from their parentbody.

13

12

11

10

9

8

7

6

5

4

3

2

1

03 3.2 3.4 3.6 3.8 4 4.2 4.4

Num

ber

Age in Ga

Eucriteages

Rb−Sr & Pb−Pb

Approx. Ar−Ar

Precise Ar−Ar

Figure 10 Histogram of impact-reset ages of theeucrites, with plotted age interval of 0.1Ga. In thecase of 39Ar–40Ar ages, those reported with uncer-tainties (‘‘precise Ar–Ar’’) are shown in black whileless precise ones reported without uncertainties(‘‘approximate Ar–Ar’’) are shown in gray. Rb–Srand Pb–Pb ages o4.3Ga are shown as the stipledareas. Reproduced by permission of Elsevier from

Bogard and Garrison (2003).


1.27.3.3 Iron Meteorites and Pallasites:Timescales of Core Crystallization onPlanetesimals

Limits on the timescales involved in metalliccore formation and crystallization on planet-esimals may be obtained from chronologicalinvestigations of iron-rich meteorites, such asmagmatic irons (that represent the cores of dif-ferentiated asteroidal bodies) and pallasites (con-sidered to have formed near the core–mantleboundary). Once the metal has segregated intothe core of a planetesimal, the timescalesinvolved in the crystallization of this metal maybe constrained by isochrons based on bulk sam-ples and mineral phases of magmatic iron me-teorites and pallasites. In recent years,precise Re–Os isochrons have been obtainedfor various groups of the iron meteorites (Figure11). Nevertheless, one of the main problems af-fecting the ability to obtain accurate and preciseages based on such isochrons has been the large(B73%) uncertainty in the measured decayconstant of 187Re (Lindner et al., 1989). RecentRe–Os studies (e.g., Smoliar et al., 1996; Horanet al., 1998; Chen et al., 2002) have assumed amore precise value for the 187Re decay constant

under the assumption that the Re–Os system at-tained closure in the IIIAB irons almost con-temporaneously with the formation of theangrites ADOR and LEW at 4,557.870.5Ma(Lugmair and Galer, 1992). Indeed, the Mn–Crmodel ages for the IIIAB irons (Hutcheon andOlsen, 1991; Hutcheon et al., 1992; Sugiura andHoshino, 2003) and these angrites (Nyquistet al., 1994; Lugmair and Shukolyukov, 1998)are approximately coincident within B75Myrand, therefore, this assumption may be valid atthis level of uncertainty.

Assuming then that the IIIAB magmaticirons attained closure of the Re–Os system at4,558Ma, most iron meteorite groups give re-latively old Re–Os ages (i.e., older thanB4.5Ga) (Shen et al., 1996; Smoliar et al.,1996; Horan et al., 1998; Cook et al., 2004).There is an apparent discrepancy, however, inthe Re–Os systematics reported in the IVAmagmatic irons. While Shen et al. (1996) reporta Re–Os age for the IVA irons that is60745Myr older than for IIAB irons, the dataof Smoliar et al. (1996) give an age of4,456725Ma, significantly younger than otheriron meteorite groups, which these authors at-tributed to later disturbance of the Re–Os sys-tem. As Horan et al. (1998) have pointed out,while the Re–Os isotopic compositions of mostof the IVA irons analyzed by Smoliar et al.(1996) do lie on the 4,456725Ma isochron,some (specifically, Duel Hill and BushmanLand) are consistent with the older age re-ported by Shen et al. (1996). It is possible thatthe Re–Os isotope systematics in different IVAirons are variably disturbed, perhaps by proc-esses such as breakup and reassembly of theparent planetesimal (a process that has beeninvoked to explain the range of metallographiccooling rates of the IVA irons, e.g., Haacket al., 1996b). This may have resulted in theresetting of the Re–Os system in some IVAirons but not in others.

As noted earlier, Re–Os ages reported sofar use a 187Re decay constant that was deter-mined by assuming that the IIIAB isochronshould give the same age as the U–Pb age of theangrites, so the accuracy of these ages is only asgood as the validity of this assumption. Never-theless, the age range indicated by Re–Osisochrons is largely independent of the precisevalue of half-life, so the results for variousmagmatic iron meteorite groups suggest thatcore crystallization (or more specifically,Re–Os isotopic closure) in planetesimalsspanned a period of B30Myr (although therelatively large errors certainly allow this timeinterval to be significantly narrower than this).This time interval for core crystallization (theprocess that is most likely to have produced

0.4 0.5 0.6 0.70.15

0.14

0.13

IIIA Irons

0.4 0.5 0.6 0.7

4

2

0

−2

−4

0.4 0.5 0.6 0.7187 O

s/18

8 Os

0.15

0.14

0.13

(a)

(b)

IVA Irons

0.4 0.5 0.6 0.7

8

4

0

−4

−8

0.4 0.6 0.8 1.00.18

0.16

0.14

0.12

IIA Irons

0.4 0.6 0.8 1.0

4

2

0

−2

−4

�

� �

0.30 0.35 0.40

0.126

0.124

0.122

0.120

IVB Irons

0.30 0.35 0.40

4

2

0

−2

−4

�

187Re/188Os

187Re/188Os

0.140

0.136

0.132

0.128

0.124

187Os188Os

+2

0

−2

−4

�

IA isochron:T = 4,537 ± 21 MaIo = 0.09556 ± 16MSWD = 1.0

0.3 0.4 0.5 0.6

187Re/188Os

IA irons - isochron set

IA isochron

IIA isochron

Sh CD Se

0.3 0.4 0.5 0.6

Other IAB-IIICD irons

Figure 11 187Re–187Os isochrons for meteorites from (a) the magmatic IIIA (4,557712Ma), IIA (4,53778Ma),IVA (4,456726Ma; open diamonds indicate three samples that were omitted from the regression), and IVB(4,527729Ma) groups and (b) the nonmagmatic IAB-IIICD (4,537721Ma) groups (labeled data points are:CD—Canyon Diablo, Se—Seelasgen, and Sh—Shrewsbury). Ages are calculated assuming a decay constant of1.666� 10�11 yr�1. Insets show the deviation in parts per 104 (i.e., in e units) of the data points from the best-fitline; these deviations are calculated relative to the IIA isochron (shown as the horizontal line in each of the insets).(a) Reproduced by permission of American Association for the Advancement of Science from Smoliar et al. (1996).

(b) Reproduced by permission of Elsevier from Horan et al. (1998).


Re–Os fractionation among the different mem-bers of a particular group of the magmaticirons) is distinct from that for core formation(or metal-silicate segregation) on the ironmeteorite parent bodies. The latter has been

inferred (predominantly from the short-lived182Hf–182W chronometer) to have occurredover a relatively short time period of only afew Myr after CAI formation (e.g., Markowskiet al., 2006).

0.4

0.12

0.13

0.14

0.15

0.16

0.6

0.3−20

−10

� PAL 0

10

0.4

FM

TM

ESS

B

GM

0.5 0.6 0.7 0.8

0.8187Re/188Os

187 O

s/18

8 Os

Pallasites

Slope = 0.0775 ± 0.0008T = 4.50 ± 0.04 Ga (� = 1.66 x 10−11a−1)I = 0.09599 ± 0.00046

Brenham

Springwater

Otinapa

Marjalahti

Finmarken

Thiel MtnsEagle station

Newport

Glorieta Mtn

Chen et al. (2002)Shen et al. (1998)

Figure 12 187Re–187Os isochron diagram for pallasite meteorites. Inset shows deviations from the best-fit line(dPAL) versus

187Re/188Os ratios. Reproduced by permission of Elsevier from Chen et al. (2002).


Re–Os systematics in pallasites indicate thatthey may be younger than iron meteorites byB60Myr (Figure 12). However, this appar-ently young age may be due to later reequili-bration of the Re–Os system (Chen et al.,2002). Cook et al. (2004) recently reported thefirst high-precision 190Pt–186Os isochrons forthe IIAB and IIIAB magmatic irons, and esti-mated ages for these meteorite groups of4,323780Ma and 4,325726Ma, respectively.These ages are somewhat younger than Re–Osages for iron meteorites, and these authorssuggested that this discrepancy could reflect anerror in the decay constant for 190Pt.

1.27.4 PLANETARY MATERIALS

Besides the Earth, the Moon and Mars arethe only other planetary bodies from whichthere are samples currently available forchronological investigations. In the caseof the Moon, there are the samples that werereturned by the Apollo and Luna missions, aswell as B40 distinct meteorites of basaltic andfeldspathic compositions that are thought tohave originated from a variety of terrains onthe Moon (Chapter 1.21). In the case of Mars,there are currently about three dozen or sodistinct meteorites of mafic and ultramaficcompositions thought to have originated fromthis planet (Chapter 1.22). The followingprovides a brief summary of the chronologyof these samples and the inferred differentia-tion and evolutionary histories of the Moon

and Mars. More detailed discussions onthese topics may be found in the Chapters1.21 and 1.22 in this volume.

1.27.4.1 Timing of Lunar Differentiation andCataclysm from Chronology of LunarSamples

1.27.4.1.1 Lunar differentiation history

Lunar samples returned by the Apolloand Luna missions as well as the lunar mete-orites broadly fall within two compositionalcategories: mare basalts and feldspathic (high-lands) rocks. The mare basalts are a geochem-ically diverse group and are comprised ofhigh-titanium (9–14wt.% TiO2), low-titanium(1–5wt.% TiO2), and very low-titanium(o1wt.% TiO2) basaltic samples. The feld-spathic rocks are also geochemically andpetrologically diverse and are composed of avariety of pristine igneous rocks as well as poly-mict breccias. Among the pristine igneoussamples are the ferroan anorthosites andmagnesium-rich rocks. The latter are composedof subgroups of magnesian-suite, alkali-suite,granites/felsites, and KREEP basalt and quartzmonzodiorite rocks.

Among known lunar samples, the ones thathave yielded the oldest crystallization ages arethe ferroan anorthosites, which are thus in-ferred to be remnants of the earliest-formedlunar crust (Hanan and Tilton, 1987; Carlsonand Lugmair, 1988; Borg et al., 1999a; Alibertet al., 1994; Norman et al., 2003). These

Planetary Materials 17

ages have been determined mostly using theSm–Nd chronometer (in one case, with theU–Pb chronometer), and span a time intervalfrom B4.29Ga70.06 (Borg et al., 1999a) toB4.56Ga70.07 (Alibert et al., 1994). Normanet al. (2003) have argued that plagioclase inthese anorthosites may have been subject tolater disturbance by impact metamorphism andif only the mafic minerals are considered, theseyield an Sm–Nd age for four ferroan anortho-sites of 4.4670.04Ga. This is then likely to bethe best estimate of the crystallization age ofthe oldest lunar crustal samples.

The crystallization ages of the magnesium-rich highlands rocks (determined predomi-nantly with Sm–Nd, but also with the Rb–Srand U–Pb chronometers) are also generally an-cient. Among these, the oldest (B4.1–4.5Ga)are the magnesian-suite norites, troctolites, anddunites, some of which may have formed con-temporaneously with the ferroan anorthosites.These were followed by the alkali-suite (B4.0–4.3Ga), granites/felsites (B3.8–4.3), KREEPbasalts (B3.8–4.0Ga), and quartz monzodio-rite rocks (B4.3Ga) (see Nyquist and Shih,1992; Papike et al., 1998; Snyder et al., 2000;and references therein).

The formation ages of the mare basalts andlunar pyroclastic glasses based on the Sm–Nd,Rb–Sr, and U–Pb chronometers are summa-rized in table 4 of Chapter 1.21. Among theoldest dated basaltic lunar material arethe mare-like clasts in highland breccias fromthe Apollo 14 landing site that have ages as oldas B4.2Ga (Taylor et al., 1983). Most marevolcanism, however, postdated the period ofheavy bombardment at B3.9Ga (see discussionof impact ages of lunar samples in the followingsection). Thus far, one of the youngest marebasalts to be dated is the KREEP-rich basalticlunar meteorite NWA 773 that has 39Ar–40Arand Sm–Nd ages of B2.9Ga (Fernandes et al.,2003; Borg et al., 2004). Fernandes et al. (2003)reported a similarly young 39Ar–40Ar age ofB2.8Ga for another lunar meteorite, NWA032, an unbrecciated mare basalt, and suggestedthat this also represented the time of crystalli-zation for this basalt. Figure 16 of Chapter 1.21summarizes the crystallization ages of thevariety of lunar samples discussed here.

The Moon is thought to have undergonemajor global differentiation early in its historythat resulted in the formation of the earliestcrust (represented by the feldspathic highlandsrocks) as well as the mantle source reservoirs ofthe basaltic lunar samples. The timing of thislunar global differentiation has been inferredfrom a variety of methods. The crystallizationages of the feldspathic highlands rocks, par-ticularly the ferroan anorthosites, discussed

earlier have been used to estimate a minimumage of B4.3 to B4.5Ga for this event (e.g.,Carlson and Lugmair, 1988; Borg et al., 1999a;Norman et al., 2003). Estimates for the forma-tion of the KREEP component in the lunarmantle (thought to represent the residuumfrom Moon-wide differentiation; e.g., Wood,1972; Warren and Wasson, 1979) also placelimits on the timing of this event. This was es-timated to be at B4.6Ga from the Rb–Srmodel age of lunar soils (Papanastassiou et al.,1970), B4.42Ga from U–Pb systematics (Teraand Wasserburg, 1974), and B4.36Ga fromSm–Nd model ages of KREEP samples (Lug-mair and Carlson, 1978). Nyquist and Shih(1992) estimated an average value from vari-ous Rb–Sr model ages for KREEP to be4.4270.14Ga. They suggested this as the bestvalue for the timing of lunar global differenti-ation, with the uncertainty reflecting thepossibility that this event did not occur at asharply defined time and that some Rb–Srfractionation may have occurred during thepetrogenesis of KREEP basalts. Finally, com-bined 147Sm–143Nd and 146Sm–142Nd systemat-ics for lunar basaltic samples indicate that thesources of these basalts were established in thelunar mantle B200Myr after the beginning ofthe solar system (Nyquist et al., 1995b;Rankenburg et al., 2006), a timescale that isconsistent with the others discussed above.

1.27.4.1.2 Lunar bombardment history

The flux of impactors on the Moon as afunction of time is a topic that is highly debatedand is of great interest since it has implicationsfor the bombardment history of the Earth,which may in turn have played a role in theevolution of the Earth’s atmosphere and in theinitiation and evolution of life on this planet. Adistinct spike in the bombardment history ofthe Moon or a ‘‘lunar cataclysm’’ at B3.9Gawas first explicitly hypothesized on the basis of39Ar–40Ar, U–Pb, and Rb–Sr systematics inrocks from the Apollo 15, 16, and 17 landingsites that appeared to have been reset or dis-turbed at this time (Tera et al., 1974). Subse-quent studies of 39Ar–40Ar ages of the Apolloand Luna impact melt rocks further demon-strated the lack of ages significantly older thanB4.0Ga (e.g., Dalrymple and Ryder, 1991,1993, 1996; Swindle et al., 1991), thus support-ing this hypothesis. However, since the Apolloand Luna samples represent only a very smallproportion of the lunar crust, they may bedominated by the effects of a few large impacts.Lunar meteorites provide a potential source ofa more random, possibly less biased, selection


of material from the Moon. Recent 39Ar–40Arinvestigations of impact melt clasts fromseveral lunar meteorites have yielded ages thatare typically o3.9Ga (Fernandes et al., 2000;Cohen et al., 2000, 2005; Gnos et al., 2004).Although these studies support the cataclysmhypothesis in that there are no ages older thanB4.0Ga, they also indicate that, rather thandropping sharply to a nearly constant rate (ashad been suggested by some studies of theApollo samples; e.g., Guinness and Arvidson,1977; Bogard et al., 1994), the impact flux wasdeclining gradually in the B3 or so billionyears following the enhanced period of bom-bardment at B3.9Ga. A recent 39Ar–40Arstudy of glass spherules from the Apollo 14soils also suggests a gradual decline in the im-pact flux (by a factor of B2–3) between B3.5andB0.5Ga (Culler et al., 2000), followed by amarked enhancement in the impact flux (by afactor of B3–5) within the last B400–500Myr.Another similar study of glass spherules fromApollo 12 soils (Levine et al., 2005) was con-sistent with, but did not require a recent in-crease in the bombardment rate. However, thisinterpretation has been questioned (Horz,2000), and remains one of the controversial is-sues related to the lunar bombardment historythat has yet to be resolved. A more thoroughdiscussion of the impact ages of a variety oflunar samples obtained with various chrono-meters based on long-lived radionuclides (par-ticularly the 39Ar–40Ar method), including theimplications for the bombardment history ofthe Moon, has been provided in Chapter 1.21.

1.27.4.2 Timescales for the Evolution of Marsfrom Chronology of MartianMeteorites

The martian meteorites represent a variety ofvolcanic to subvolcanic as well as plutonicigneous rocks that are broadly categorized,based on their petrologic and geochemicalcharacteristics, into four groups: the basalticand lherzolitic shergottites, the clinopyroxeniticnakhlites, the dunitic chassignites, and theorthopyroxenite ALH84001 (Chapter 1.22).Chronological investigations of these martianmeteorites and their implications for theevolution of Mars have been thoroughly re-viewed by Nyquist et al. (2001) and Borg andDrake (2005). As such, only a brief summary ispresented here.

Figure 13 summarizes the absolute ages forvarious events in martian history based on theradiometric dating of the martian meteoritesand their components. The oldest formationage of B4.5Ga is yielded by a 147Sm–143Nd

internal isochron for the ALH84001 or-thopyroxenite (Jagoutz et al., 1994; Nyquistet al., 1995a). All other dated martian meteor-ites have internal isochron ages (based mostlyon the Sm–Nd and Rb–Sr chronometers) thatare younger than B1.3Ga, with the youngestsamples being only B170Ma. Based on theirgeochemical features (particularly the trace-and minor-element zonations in their minerals;e.g., Jones, 1986), these ages are generally in-terpreted to reflect the crystallization ages ofthese samples, thus indicating that Mars maystill be a geologically active planet. A recentstudy has suggested that the young ages of theshergottites may in fact reflect the timing ofsecondary alteration, and that their formationages are close toB4.0Ga (Bouvier et al., 2005).This interpretation requires that the late-stageminerals (i.e., phosphates) in these shergottitesbe affected by this secondary alteration event.However, studies of trace-element microdistri-butions have shown that these phosphates wereformed by closed system fractional crystalliza-tion of the shergottite parent melts and thatthey preserve their original igneous composi-tions (Wadhwa et al., 1994). As such, at thepresent time, the interpretation of the Sm–Ndand Rb–Sr internal isochrons ages of theshergottites as their crystallization ages is themost likely. Therefore, as discussed by Borgand Drake (2005) and summarized in Figure13, there is evidence from the martian meteor-ites for igneous events on Mars at 4,3007130,1,327729, 57577, 474711, 33279, and17472Ma.

Besides primary igneous events, the martianmeteorites also provide evidence for secondaryevents on Mars. A 39Ar–40Ar age in the range ofB3.8–4.1Ga for the ALH84001 orthopyroxenitehas been interpreted as possibly reflecting thetime of heavy bombardment on Mars co-incident with a similar event on the Moon (Ashet al., 1996; Turner et al., 1997). Moreover, car-bonates in this rock, thought to be precipitated asa result of interaction with an aqueous fluid in anear-surface environment, have also been datedat B3.9Ga based on Rb–Sr and U–Pb system-atics (Borg et al., 1999b). In addition to the car-bonates in ALH84001, there are other productsof near-surface aqueous alteration in the othermartian meteorites, in particular, a hydrous min-eral (iddingsite) in the nakhlites (e.g., Bunch andReid, 1975; Treiman et al., 1993) and a variety ofsecondary minerals (including carbonates, sul-fates, and clays) in the shergottites (e.g., Goodinget al., 1988, 1990). Based on Rb–Sr and K–Arsystematics in iddingsite-rich fractions (Shihet al., 1998, 2002; Swindle et al., 2000), it hasbeen argued that this hydrous secondary phasewas formed at B600–700Ma. The only age

Salts shergottites (0−175 Ma)Iddingsite nakhlites (633 ± 23 Ma)

Carbonates ALH84001 (3,929 ± 37 Ma)

Shergotty (165 ± 11 Ma)

Zagami (169 ± 7 Ma)LA1 (170 ± 7 Ma)NWA856 (170 ± 19 Ma)

174 ± 2 Ma

332 ± 9 Ma

1,327 ± 39 Ma

EET79001A (173 ± 10 Ma)Y793605 (173 ± 14 Ma)

EET79001B (173 ± 3 Ma)ALH77005 (177 ± 6 Ma)LEW88516 (178 ± 9 Ma)NWA1056 (185 ± 11 Ma)

Y980459 (290 ± 40 Ma)QUE94201 (327 ± 10 Ma)NWA1195 (348 ± 19 Ma)

DaG 476 (474 ± 11 Ma)Dhofar 019 (575 ± 7 Ma)

Nakhla (1,260 ± 70 Ma)NWA998 (1,290 ± 50 Ma)Y000593 (1,310 ± 30 Ma)Lafayette (1,320 ± 50 Ma)Chassigny (1,362 ± 62)Gov. Valad. (1,370 ± 20 Ma)

ALH84001 (4,500 ± 130 Ma)Silicate differentiation (4,526 ± 21 Ma)

Core segregation (4,556 ± 1 Ma)LEW86010; silicate differentiation reference (4,558 ± 0.5 Ma)

CAI (solar system formation reference) (4,567 ± 0.6 Ma)

0 1,000 2,000

Age (Ma)

3,000 4,000 4,657

Figure 13 Absolute ages of events in Mars’ history based on radiometric dating of martian meteorites andtheir components. Crystallization ages of the shergottites (B170–575Ma; filled circles, open squares, filledtriangle, and open diamond), the nakhlites (B1.3Ga; open circles), and the orthopyroxenite ALH84001(B4.5Ga) are shown. The ages of aqueous alteration events recorded by secondary products in the martianmeteorites are shown as the open triangles. Also shown are the ages of global differentiation events (coreformation and silicate differentiation) based on the 147,146Sm–143,142Nd and 182Hf–182W systematics of the

martian meteorites. Reproduced with permission from Borg and Drake (2005).

Conclusions 19

constraint on the secondary alteration productsin the shergottites may be obtained from consid-eration of the crystallization ages of the partic-ular samples (B175Ma) in which these arefound; as such, it is estimated that these second-ary minerals were formed at some timeo175Ma. Based on the abundances of altera-tion products in the martian meteorites and theirdiscrete formation time intervals, Borg andDrake (2005) argued that aqueous fluids werepresent episodically, and not continuously, in thenear-surface environment on Mars.

Finally, whole-rock Rb–Sr and U–Pb sys-tematics of the martian meteorites indicate thatglobal silicate differentiation on Mars occurredclose to B4.5Ga (Shih et al., 1982; Chen andWasserburg, 1986). Using an approach similarto that applied by Nyquist et al. (1995b) tothe lunar basaltic samples (i.e., combined

147Sm–143Nd and 146Sm–142Nd systematics), amore precise estimate of 4,525720Ma for thetiming of major silicate differentiation on Marsthat established the source reservoirof the shergottites has been made (Borget al., 2003; Foley et al., 2005). The short-lived146Sm–142Nd and 182Hf–182W chronometersfurther indicate that the nakhlite mantle sourcemay have been established contemporaneouslywith that of the shergottites, or it may have pre-dated this event by a few tens of millions ofyears (Foley et al., 2005).

1.27.5 CONCLUSIONS

1.27.5.1 A Timeline for Solar System Events

From the application of various chronome-ters based on the long-lived radionuclides to


meteoritic and planetary materials, the follow-ing may be inferred as the sequence of eventsthat occurred in the history of the solar system:

1. The earliest solids to form in the solarprotoplanetary disk were the refractoryCAIs, which formed at 4,567.170.2Ma.This represents the minimum age of thesolar system.

2. Chondrules from CV and CR chondritesbegan forming more or less contempora-neously with CAIs (4,566.771.0Ma forAllende chondrules), and continued toform for at least another 2–3Myr.

3. Accretion and differentiation of planet-esimals also began almost contemporane-ously with CAIs, with basalts forming onthe surfaces of these bodies well withinB1Myr of CAI formation. The accretionprocess for planetesimals is also likely tohave continued for a few million years(with at least some chondrite parent bodiespossibly being assembled a few millionyears after CAIs).

4. Energetic collisions between the accretedplanetary embryos resulted in the forma-tion of impact-generated chondrules sev-eral millions of years after CAI formation.Specifically, chondrules from the CBchondrites, which are thought to resultfrom such a process, were formed B5Myrafter CAIs.

5. Thermal processing of accreted undifferen-tiated planetesimals began early (e.g., atB4,563Ma for the parent body of the Ste.Marguerite H4 chondrite), but this processcontinued for tens of millions of years afterthe beginning of the solar system. Meta-morphism on chondrite parent bodiesis likely to have extended for at leastB60Myr after CAI formation. Aqueousalteration of the CI chondrite parent bodyis likely to have begun almost as soon as itwas accreted.

6. Basaltic melts continued to be generated ondifferentiated planetesimals for a period ofB10Myr after solar system formation. Inthe case of the EPB, these basalts were thensubjected to a complex and protracted (las-ting tens of millions of years) history ofthermal processing, most likely resultingfrom large impacts on the surface of theEPB. Based on the ages of the cumulateeucrites, it is inferred that either igneousactivity may have lasted for B100–150Myr on the EPB or isotopic systemsin these samples record slow cooling atdepth in the EPB.

7. The iron-nickel cores of some differentiatedplanetesimals began crystallization and

solidification within B10Myr of CAI for-mation, but this process on other differen-tiated parent bodies may have extended foranother tens of millions of years.

8. Global silicate differentiation on the Moonand Mars and establishment of mantlesource reservoirs for lunar and martianbasalts occurred B200Myr and withinB50Myr after the beginning of the solarsystem, respectively.

9. The earliest crustal (highlands) rocks onthe Moon was formed at B4.5Ga. Theoldest basaltic lunar materials have agesof B4.2Ga, although most mare basaltscrystallized after B3.9. Mare volcanismcontinued at least till B2.9Ga.

10. Crystallization ages of martian meteoritesrange from B4.5Ga for the ALH84001orthopyroxenite to B170Ma for someshergottites, suggesting that magmatic ac-tivity on Mars began early and may stillcontinue. Aqueous alteration events re-corded in the martian meteorites are esti-mated to be have occurred at B3.9Ga,600–700Ma, and o175Ma. These discreteformation times suggest that aqueous fluidswere present episodically, and not contin-uously, in the near-surface environment onMars.

11. Members belonging to several groups ofmeteorites (eucrites and chondrites) andplanetary materials (lunar samples and themartian meteorite ALH84001) record im-pact ages that are consistent with a periodof heavy bombardment in the inner solarsystem that peaked close to B4Ga. Peaksin the impact ages at o1.3Ga for thechondrites and close to B4.5Ga for theeucrites may point toward a few large im-pact events that occurred on their parentbodies.

1.27.5.2 Outlook and Future Prospects

Among the long-lived chronometers, theone based on the 235,238U–207,206Pb systemsprovides the best precision. With extensive lea-ching procedures to effectively remove the com-mon lead component and use of a double spiketo improve the analytical precision, it is nowpossible to attain a precision better than70.5Myr on 207Pb–206Pb ages for radiogenicsamples. As such, more future studies of earlysolar system chronology are likely to focus onthe utilization of this chronometer as the abso-lute anchor for other chronometers based onshort-lived radionuclides, which also have thepotential for providing sub-Myr time resolutionfor events occurring in the earliest history of the

References 21

solar system. Nevertheless, there are numerousassumptions and complexities involved in theutilization of chronometers based on the short-lived radionuclides, and the U–Pb system is notan appropriate chronometer for many types ofmeteoritic materials (e.g., those that have re-latively unradiogenic lead-isotope composi-tions). As such, obtaining the absolute ages ofa variety of meteorites and their componentsusing various long-lived chronometers will re-main a high priority for understanding theirformation histories in the early solar system,particularly if future developments allow higherprecision and accuracy for these absolute ages.These future developments would include ana-lytical advances in mass spectrometric tech-niques for isotope ratio measurements and theprecise and accurate determination of decayconstants of some of the long-lived radionucl-ides. Furthermore, there are as yet unresolvedissues that are likely to be only addressedthrough chronological studies based on thelong-lived radionuclides. In particular, betterconstraints on the impact flux in the inner solarsystem will only be possible through additionaland more extensive studies of the impact ages(mostly from 39Ar–40Ar, but possibly also usingother chronometers such as 87Rb–87Sr and235,238U–207,206Pb) of a variety of meteoriticand planetary materials.

REFERENCES

Allegre C. J., Birck J.-L., Fourcade S., and Semet M. P.(1975) Rubidium-87/strontium-87 age of Juvinas basalticachondrite and early igneous activity in the solar system.Science 187, 436–438.

Allegre C. J., Manhes G., and Gopel C. (1995) The age ofthe Earth. Geochim. Cosmochim. Acta 59, 1445–1456.

Alibert C., Norman M. D., and McCulloch (1994) An an-cient Sm–Nd age for a ferroan noritic anorthosite clastfrom lunar breccia 67016. Geochim. Cosmochim. Acta 58,2921–2926.

Amelin Y. and Davis W. J. (2005) Geochemical test forbranching decay of 176Lu. Geochim. Cosmochim. Acta 69,465–473.

Amelin Y., Ghosh A., and Rotenberg E. (2005) Unravelingthe evolution of chondrite parent asteroids by preciseU–Pb dating and thermal modeling. Geochim. Cosmo-chim. Acta 69, 505–518.

Amelin Y., Krot A. N., Hutcheon I. D., and Ulyanov A. A.(2002) Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science 297, 1678–1683.

Amelin Y., Krot A. N., and Twelker E. (2004) Pb isotopicages of the CB chondrite Gujba, and the duration of thechondrule formation interval. Geochim. Cosmochim. Acta68, A759.

Amelin Y. and Pravdivtseva O. (2005) U–Pb age of theAcapulco phosphate: testing the calibration of the I–Xechronometer. Meteorit. Planet. Sci 40, A16.

Amelin Y., Wadhwa M., and Lugmair G. (2006) Pb-isotopicdating of meteorites using 202Pb–205Pb double-spike: com-parison with other high-resolution chronometers. In LunarPlanet. Sci. XXXVII, #1970. The Lunar and Planetary In-stitute, Houston (CD-ROM).

Anders E. and Grevesse N. (1989) Abundances of theelements: meteoritic and solar. Geochim. Cosmochim.Acta 53, 197–214.

Ash R. D., Knott S. F., and Turner G. (1996) A 4Gyrshock age for a martian meteorite and implications forthe cratering history of Mars. Nature 380, 57–59.

Baker J., Bizzarro M., Wittig N., Connelly J., and HaackH. (2005) Early planetesimal melting from an age of4.5662Gyr for differentiated meteorites. Nature 436,1127–1130.

Birck J.-L. and Allegre C. J. (1978) Chronology and chem-ical history of the parent body of the basaltic achondritesstudied by the 87Rb–87Sr method. Earth Planet. Sci. Lett.39, 37–51.

Bizzarro M., Baker J. A., Haack H., Ulfbeck D., andRosing M. (2003) Early history of Earth’s crust–mantlesystem inferred from hafnium isotopes in chondrites.Nature 421, 931–933.

Blichert-Toft J., Boyet M., Telouk P., and Albarede F.(2002) 147Sm–143Nd and 176Lu–176Hf in eucrites and thedifferentiation of the HED parent body. Earth Planet.Sci. Lett. 204, 167–181.

Bogard D. D. (1995) Impact ages of meteorites: a synthesis.Meteoritics 30, 244–268.

Bogard D. D. and Garrison D. H. (2003) 39Ar–40Ar agesof eucrites and thermal history of asteroid 4 Vesta.Meteorit. Planet. Sci. 38, 669–710.

Bogard D. D., Garrison D. H., Jordan J. L., andMittlefehldt D. W. (1990) 39Ar–40Ar dating of mesoside-rites: evidence for major parent disruption o4Ga ago.Geochim. Cosmochim. Acta 54, 2549–2564.

Bogard D. D., Garrison D. H., Shih C.-Y., and NyquistL. E. (1994) 39Ar–40Ar dating of two lunar granites—theage of Copernicus. Geochim. Cosmochim. Acta 58, 3093–3100.

Bogard D. D. and Hirsch W. C. (1980) 39Ar–40Ar dating,Ar diffusion properties, and cooling rate determinationsof severely shocked chondrites. Geochim. Cosmochim.Acta 44, 1667–1682.

Bogard D. D., Husain L., and Wright R. J. (1976)40Ar–39Ar dating of collisional events in chondriteparent bodies. J. Geophys. Res. 81, 5664–5678.

Bogdanovski O. and Jagoutz E. (1996) Sm–Nd system inthe Divnoe meteorite. In Lunar Planet. Sci. XXVII. TheLunar and Planetary Institute, Houston, pp. 129–130.

Borg L. and Drake M. J. (2005) A review of meteoriteevidence for the timing of magmatism and of surface ornear-surface liquid water on Mars. J. Geophys. Res. 110,E12S03.

Borg L., Norman M., Nyquist L., Bogard D., Snyder G.,Taylor L., and Lindstrom M. (1999a) Isotopic studies offerroan anorthosite 62236: a young lunar crustal rockfrom a light rare-earth-element-depleted source. Geo-chim. Cosmochim. Acta 63, 2679–2691.

Borg L. E., Connelly J. N., Nyquist L. E., Shih C.-Y.,Wiesmann H., and Reese Y. (1999b) The age of thecarbonates in martian meteorite ALH84001. Science 268,90–94.

Borg L. E., Nyquist L. E., Wiesmann H., Shih C.-Y., andReese Y. (2003) The age of Dar al Gani 476 and the dif-ferentiation history of the martian meteorites inferred fromtheir radiogenic isotope systematics. Geochim. Cosmochim.Acta 67, 3519–3536.

Borg L. E., Shearer C. K., Asmerom Y., and Papike J. J.(2004) Prolonged KREEP magmatism on the Moonindicated by the youngest dated lunar igneous rock.Nature 432, 209–211.

Bouvier A., Blichert-Toft J., Vervoort J. D., and Albarede F.(2005) The age of SNC meteorites and the antiquity of themartian surface. Earth Planet. Sci. Lett. 240, 221–233.

Brannon J. C., Podosek F. A., and Lugmair G. W. (1988)Initial 87Sr/86Sr and Sm–Nd chronology of chondriticmeteorites. Proc. 18th Lunar Planet. Sci. Conf. 555–564.


Brearley A. J. (2006) The action of water. In Meteorites andthe Early Solar System Vol. II (eds. D. Lauretta andH. Y. McSween Jr.). University of Arizona Press,Tucson, AZ (in press).

Bunch T. E. and Reid A. M. (1975) The nakhlites. 1:Petrology and mineral chemistry. Meteoritics 10, 303–315.

Carlson R. W. and Lugmair G. W. (1988) The age of fer-roan anorthosite 60025: oldest crust on a young Moon?Earth Planet. Sci. Lett. 90, 119–130.

Carlson R. W. and Lugmair G. W. (2000) Timescalesof planetesimal formation and differentiation basedon extinct and extant radionuclides. In The Originof the Earth and Moon (eds. R. M. Canup andK. Righter). University of Arizona Press, Tucson, AZ,pp. 25–44.

Chen J. H., Papanastassiou D. A., and Wasserburg G. J.(2002) Re–Os and Pd–Ag systematics in Group IIABirons and in pallasites. Geochim. Cosmochim. Acta 66,3793–3810.

Chen J. H. and Tilton G. R. (1976) Isotopic lead investigat-ions on the Allende carbonaceous chondrite. Geochim.Cosmochim. Acta 40, 635–643.

Chen J. H. and Wasserburg G. J. (1981) The isotopiccomposition of uranium and lead in Allende inclusionsand meteoritic phosphates. Earth Planet. Sci. Lett. 5,21–51.

Chen J. H. and Wasserburg G. J. (1985). U–Th–Pb isotopicstudies on meteorite ALHA 81005 and Ibitira. In LunarPlanet. Sci. XVI. The Lunar and Planetary Institute,Houston, pp. 119–120.

Chen J. H. and Wasserburg G. J. (1986) Formation agesand evolution of Shergotty and its parent planet fromU–Th–Pb systematics. Geochim. Cosmochim. Acta 50,955–968.

Ciesla F. J. (2005) Chondrule forming processes—anoverview. In Chondrites and the Protoplanetary Disk(eds. A. N. Krot, E. R. D. Scott, and B. Reipurth), ASPConference Series. Astronomical Society of the Pacific,San Francisco, vol. 341 pp. 558–587.

Clayton R. N. and Mayeda T. K. (1996) Oxygen isotopestudies of achondrites. Geochim. Cosmochim. Acta 60,1999–2017.

Cohen B. A., Swindle T. D., and Kring D. A. (2000) Sup-port for the lunar cataclysm hypothesis from lunar me-teorite impact melt ages. Science 290, 1754–1756.

Cohen B. A., Swindle T. D., and Kring D. A. (2005)Geochemistry and 40Ar–39Ar geochronology of impact-melt clasts in feldspathic lunar meteorites: implicationsfor lunar bombardment history. Meteorit. Planet. Sci. 40,755–777.

Cook D. L., Walker R. J., Horan M. F., Wasson J. T., andMorgan J. W. (2004) Pt–Re–Os systematics of groupIIAB and IIIAB iron meteorites. Geochim. Cosmochim.Acta 68, 1413–1431.

Culler T. S., Becker T. A., Muller R. A., and Renne P. R.(2000) Lunar impact history from 40Ar/39Ar dating ofglass spherules. Science 287, 1785–1788.

Dalrymple G. B. and Ryder G. (1991) 40Ar/39Ar ages of sixApollo 15 impact melt rocks by laser step heating.Geophys. Res. Lett. 18, 1163–1166.

Dalrymple G. B. and Ryder G. (1993) 40Ar/39Ar age spec-tra of Apollo 15 impact melt rocks by laser step heatingand their bearing on the history of lunar basin formation.J. Geophys. Res. 98, 13085–13095.

Dalrymple G. B. and Ryder G. (1996) Argon-40/argon-39age spectra of Apollo 17 highlands breccia samples bylaser step heating and the age of the Serenitatis basin.J. Geophys. Res. 101, 26069–26084.

Deutsch A. and Scharer U. (1994) Dating terrestrial impactevents. Meteoritics 29, 301–322.

Evensen N. M., Carter S. R., Hamilton P. J., O’Nions R. K.,and Ridley W. I. (1979) A combined chemical-petrological

study of separated chondrules from the Richardton me-teorite. Earth Planet. Sci. Lett. 42, 223–236.

Fernandes V. A., Burgess R., and Turner G. (2000) Laserargon-40–argon-39 age studies of Dar al Gani 262 lunarmeteorite. Meteorit. Planet. Sci. 25, 1355–1364.

Fernandes V. A., Burgess R., and Turner G. (2003)40Ar–39Ar chronology of lunar meteorites NorthwestAfrica 032 and 773. Meteorit. Planet. Sci. 38, 555–564.

Foley C. N., Wadhwa M., Borg L. E., Janney P. E., HinesR., and Grove T. L. (2005) The early differentiationhistory of Mars from 182W–142Nd isotope systematics inthe SNC meteorites. Geochim. Cosmochim. Acta 69,4557–4571.

Galer S. J. G. and Lugmair G. W. (1996) Lead isotopesystematics of noncumulate eucrites. Meteorit. Planet.Sci. 31, A47–A48.

Ghosh A. and McSween H. Y. M., Jr. (1998) A thermalmodel for the differentiation of Asteroid 4 Vesta, basedon radiogenic heating. Icarus 134, 187–206.

Gnos E., Hofmann B. A., Al A-K., Lorenzetti S., EugsterO., Whitehouse M. J., Villa I. M., Jull A. J. T., EikenbergJ., Spettel B., Krahenbuhl U., Franchi I. A., and Green-wood R. C. (2004) Pinpointing the source of a lunarmeteorite: implications for the evolution of the Moon.Science 305, 657–659.

Gooding J. L., Aggrey K. E., and Muenow D. W. (1990)Volatile compounds in shergottites and nakhlite meteor-ites. Meteoritics 25, 281–289.

Gooding J. L., Wentworth S. J., and Zolensky M. E. (1988)Calcium carbonate and sulfate of possible extraterrestrialorigin in the EETA 79001 meteorite. Geochim. Cosmo-chim. Acta 52, 909–915.

Goodrich C. A. and Lugmair G. W. (1995) Stalking theLREE-enriched component in ureilites. Geochim. Cosmo-chim. Acta 59, 2609–2620.

Gopel C., Manhes G., and Allegre C. J. (1991) Constraintson the time of accretion and thermal evolution of chond-rite parent body. Meteoritics 26, 338.

Gopel C., Manhes G., and Allegre C. J. (1992) U–Pb studyof the Acapulco meteorite. Meteoritics 27, 226.

Gopel C., Manhes G., and Allegre C. J. (1994) U–Pb sys-tematics of phosphates from equilibrated ordinarychondrites. Earth Planet. Sci. Lett. 121, 153–171.

Gray C. M., Papanastassiou D. A., and Wasserburg G. J.(1973) The identification of early condensates from thesolar nebula. Icarus 20, 213–239.

Grier J. A., Kring D. A., Swindle T. D., Rivkin A. S.,Cohen B. A., and Britt D. T. (2004) Analyses of thechondritic meteorite Orvino (H6): insights into theorigins and evolution of shocked H chondrite material.Meteorit. Planet. Sci. 39, 1475–1493.

Guinness E. A. and Arvidson R. E. (1977) On the con-stancy of the lunar cratering flux over the past3.3� 109 yr. Proc. 8th Lunar Sci. Conf. 3475–3494.

Haack H., Farinella P., Scott E. R. D., and Keil K. (1996a)Meteoritic, asteroidal and theoretical constraints on the500Ma disruption of the L chondrite parent body. Icarus119, 182–191.

Haack H., Scott E. R. D., Love S. G., Brearley A. J.,and McCoy T. J. (1996b) Thermal histories of IVAstony-iron meteorites: evidence for asteroid fragmentat-ion and reaccretion. Geochim. Cosmochim. Acta 60,3103–3113.

Hanan B. B. and Tilton G. R. (1987) 60025: Relict ofprimitive lunar crust? Earth Planet. Sci. Lett. 84, 15–21.

Hohenberg C. M., Hudson B., Kennedy B. M., andPodosek F. A. (1981) Noble gas retention chronologiesfor the St Severin meteorite. Geochim. Cosmochim. Acta45, 535–546.

Horan M. F., Smoliar M. I., and Walker R. J. (1998) 182Wand 187Re–187Os systematics of iron meteorites: chronol-ogy for melting, differentiation, and crystallization inasteroids. Geochim. Cosmochim. Acta 62, 545–554.

References 23

Horz F. (2000) Time variable cratering rates? Science 288,2095a.

Huss G. R., Rubin A. E., and Grossman J. N. (2006)Thermal metamorphism in chondrites. In Meteorites andthe Early Solar System Vol. II (eds. D. Lauretta andH. Y. McSween Jr.). University of Arizona Press,Tucson, AZ (in press).

Hutcheon I. D. and Olsen E. J. (1991) Cr isotopic compo-sition of differentiated meteorites: a search for 53Mn.In Lunar Planet. Sci. XXII. The Lunar and PlanetaryInstitute, Houston, pp. 605–606.

Hutcheon I. D., Olsen E., Zipfel J., and Wasserburg G. J.(1992) Chromium isotopes in differentiated meteorites:evidence for 53Mn. In Lunar Planet. Sci. XXIII. The Lu-nar and Planetary Institute, Houston, pp. 565–566.

Jacobsen S. B. and Wasserburg G. J. (1984) Sm–Nd iso-topic evolution of chondrites and achondrites, II. EarthPlanet. Sci. Lett. 67, 137–150.

Jagoutz E., Jotter R., Kubny A., Varela M. E., Zartman R.,Kurat G., and Lugmair G. W. (2003) Cm?–U–Th–Pbisotopic evolution of the D’Orbigny angrite. Meteorit.Planet. Sci. 38, A81.

Jagoutz E., Sorowka A., Vogel J. D., and Wanke H. (1994)ALH84001: alien or progenitor of the SNC family?Meteoritics 29, 478–479.

Jones J. H. (1986) A discussion of isotopic systematics andmineral zoning in the shergottites: evidence for a 180m.y.igneous crystallization age. Geochim. Cosmochim. Acta50, 969–977.

Kaneoka I. (1981) 40Ar–39Ar ages of Antarctic meteorites:Y-74191, Y-75258, Y-7308, Y-74450, and ALH-765.Proc. 6th NIPR Symp. Antarct. Met. 250–263.

Keil K., Fodor R. V., Starzyk P. M., Schmidt R. A.,Bogard D. D., and Husain L. (1980) A 3.6 b.y.-oldimpact melt rock fragment in the Plainview chondrite:implications for the age of the H-group chondrite parentbody regolith formation. Earth Planet. Sci. Lett. 51,235–247.

Kita N. T., Huss G. R., Tachibana S., Amelin Y., NyquistL. E., and Hutcheon I. D. (2005) Constraints on theorigin of chondrules and CAIs from short-lived and long-lived radionuclides. In Chondrites and the ProtoplanetaryDisk (eds. A. N. Krot, E. R. D. Scott, and B. Reipurth),ASP Conference Series. Astronomical Society of thePacific, San Francisco, vol. 341, pp. 558–587.

Kring D. A., Swindle T. D., Britt D. T., and Grier J. A.(1996) Cat Mountain: a meteoritic sample of an impact-melted asteroid regolith. J. Geophys. Res. 101, 29353–29371.

Krot A. N., Amelin Y., Cassen P., and Meibom A. (2005)Young chondrules in CB chondrites from a giant impactin the early solar system. Nature 436, 989–992.

Kumar A., Gopalan K., and Bhandari N. (1999)147Sm–143Nd and 87Rb–87Sr ages of the eucrite PipliaKalan. Geochim. Cosmochim. Acta 63, 3997–4001.

Levine J., Becker T. A., Muller R. A., and Renne P. R.(2005) 40Ar/39Ar dating of Apollo 12 impact spherules.Geophys. Res. Lett. 32, L15201.

Lindner M., Leich D. A., Russ G. P., Bazan J. M., andBorg R. J. (1989) Direct determination of the half-life of187Re. Geochim. Cosmochim. Acta 53, 1597–1606.

Loss R. D., Lugmair G. W., Davis A. M., and MacPhersonG. J. (1994) Isotopically distinct reservoirs in the solarnebula: isotopic anomalies in Vigarano meteorite inclu-sions. Astrophys. J. 436, L193–L196.

Lugmair G. W. (1974) Sm–Nd ages: a new dating method.Meteoritics 9, 369.

Lugmair G. W. and Carlson R. W. (1978) The Sm–Ndhistory of KREEP. Proc. 9th Lunar Planet. Sci. Conf.689–704.

Lugmair G. W. and Galer S. J. G. (1992) Age and isotopicrelationships among angrites Lewis Cliff 86010 and Angrados Reis. Geochim. Cosmochim. Acta 56, 1673–1694.

Lugmair G. W., Galer S. J. G., and Carlson R. W. (1991)Isotope systematics of cumulate eucrite EET-87520.Meteoritics 26, 368.

Lugmair G. W. and Marti K. (1977) Sm–Nd–Pu timepiecesin the Angra dos Reis meteorite. Earth Planet. Sci. Lett.35, 273–284.

Lugmair G. W., Scheinin N. B., and Carlson R. W. (1977)Sm–Nd systematics of the Serra de Mage eucrite. Meteo-ritics 10, 300–301.

Lugmair G. W. and Shukolyukov A. (1998) Early solarsystem timescales according to 53Mn–53Cr systematics.Geochim. Cosmochim. Acta 62, 2863–2886.

Macdougall J. D., Lugmair G. W., and Kerridge J. F.(1984) Early solar system aqueous activity: Sr isotopeevidence from the Orgueil CI meteorite. Nature 307,249–251.

Manhes G., Gopel C., and Allegre C. J. (1987) High res-olution chronology of the early solar system based onlead isotopes. Meteoritics 22, 453–454.

Manhes G., Minster J. F., and Allegre C. J. (1978)Comparative U–Th–Pb and Rb–Sr study of the Saint-Severin amphoterite: consequences for early solar systemchronology. Earth Planet. Sci. Lett. 39, 14–24.

Markowski A., Quitte G., Kleine T., and Halliday A. N.(2006) Tungsten isotopic composition of iron meteorite:chronological constraints vs. cosmogenic effects. EarthPlanet. Sci. Lett. 214, 1–15.

McCoy T. J., Keil K., Clayton R. N., Mayeda T. K.,Bogard D. D., Garrison D. H., Huss G. R., HutcheonI. D., and Wieler R. (1996) A petrologic, chemical, andisotopic study of Monument Draw and comparisonwith other acapulcoites: evidence for formation by in-cipient partial melting. Geochim. Cosmochim. Acta 60,2681–2708.

McCoy T. J., Keil K., Clayton R. N., Mayeda T. K., BogardD. D., Garrison D. H., and Wieler R. (1997a) A pet-rologic and isotopic study of lodranites: evidence of earlyformation as partial melt residues from heterogeneousprecursors. Geochim. Cosmochim. Acta 61, 623–637.

McCoy T. J., Keil K., Muenow D. W., and Wilson L.(1997b) Partial melting and melt migration in theacapulcoite–lodranite parent body. Geochim. Cosmo-chim. Acta 61, 639–650.

McSween H. Y., Jr., Sears D. W. G., and Dodd R. T.(1988) Thermal metamorphism. In Meteorites and theEarly Solar System (eds. J. F. Kerridge and M. S.Matthews). University of Arizona Press, Tucson, AZ, pp.102–113.

Metzler K., Bobe K. D., Palme H., Spettel B., and StofflerD. (1995) Thermal and impact metamorphism on theHED parent asteroid. Planet. Space Sci. 43, 499–525.

Mittlefehldt D. W., Lindstrom M. M., Bogard D. D.,Garrison D. H., and Field S. W. (1996) Acapulco-and Lodran-like achondrites: petrology, geochemistry,chronology and origin. Geochim. Cosmochim. Acta 60,867–882.

Mittlefehldt D. W., McCoy T. J., Goodrich C. A., andKracher A. (1998) Non-chondritic meteorites fromasteroidal bodies. In Planetary Materials (ed. J. J.Papike). Rev. Mineral., Mineralogical Society of America,Washington, DC, vol. 36, pp. 4-1–4-195.

Nakamura N., Morikawa N., Hutchison R., Clayton R. N.,Mayeda T., Nagao K., Misawa K., Okano O., YamamotoK., Yanai K., and Matsumoto Y. (1994) Trace elementand isotopic characteristrics of inclusions in the Yamatoordinary chondrites Y-75097, Y-793241, and Y-794046.Proc. 7th NIPR Symp. Antarct. Met. 125–143.

Norman M. D., Borg L. E., Nyquist L. E., and BogardD. D. (2003) Chronology, geochemistry, and petrologyof a ferroan noritic anorthosites clast from Descartesbreccia 67215: clues to the age, origin, structure, andimpact history of the lunar crust. Meteorit. Planet. Sci.38, 645–661.


Nyquist L. E., Bansal B., Wiesmann H., and Shih C.-Y.(1994) Neodymium, strontium and chromium isotopicstudies of the LEW 86010 and Angra dos Reis meteoritesand the chronology of the angrite parent body. Meteori-tics 29, 872–885.

Nyquist L. E., Bansal B. M., Wiesmann H., and Shih C.-Y.(1995a) ‘‘Martians’’ young and old: Zagami andALH84001. In Lunar Planet. Sci. XXVI. The Lunarand Planetary Institute, Houston, pp. 1065–1066.

Nyquist L. E., Bogard D. D., Garrison D. H., Bansal B.,Wiesmann H., and Shih C.- Y. (1991). Thermal resettingof radiometric ages. In Lunar Planet. Sci. XXII. TheLunar and Planetary Institute, Houston, pp. 985–988.

Nyquist L. E., Bogard D. D., Shih C.-Y., Greshake A.,Stoffler D., and Eugster E. (2001) Ages and geologichistories of martian meteorites. Space Sci. Rev. 96,105–164.

Nyquist L., Bogard D., Takeda H., Bansal B., WiesmannH., and Shih C.-Y. (1997) Crystallization, recrystal-lization, and impact-metamorphic ages of eucritesY792510 and Y791186. Geochim. Cosmochim. Acta 61,2119–2138.

Nyquist L. E., Bogard D. D., Wiesmann H., Bansal B.,Shih C.-Y., and Morris R. M. (1990) Age of a eucriteclast from the Bholghati howardite. Geochim. Cosmo-chim. Acta 54, 2195–2206.

Nyquist L. E. and Shih C.-Y. (1992) The isotopic recordof lunar volcanism. Geochim. Cosmochim. Acta 56, 2213–2234.

Nyquist L. E., Shih C. Y., Wiesmann H., and Mikouchi T.(2003a). Fossil 26Al and 53Mn in D’Orbigny and Sahara99555 and the timescale for angrite magmatism. In LunarPlanet. Sci. XXXIV, #1388. The Lunar and PlanetaryInstitute, Houston (CD-ROM).

Nyquist L. E., Takeda H., Bansal B. M., Shih C.-Y.,Wiesmann H., and Wooden J. L. (1986) Rb–Sr andSm–Nd internal isochron ages of a subophitic basaltclast and a matrix sample from the Y75011 eucrite.J. Geophys. Res. 91, 8137–8150.

Nyquist L. E., Wiesmann H., Bansal B., Shih C.-Y., KeithJ. E., and Harper C. L. (1995b) 46Sm–142Nd formationinterval for the lunar mantle. Geochim. Cosmochim. Acta59, 2817–2837.

Papanastassiou D. A. and Wasserburg G. J. (1969) Initialstrontium isotopic abundances and the resolution ofsmall time differences in the formation of planetaryobjects. Earth Planet. Sci. Lett. 5, 361–376.

Papanastassiou D. A. and Wasserburg G. J. (1978) Stron-tium isotopic anomalies in the Allende meteorite.Geophys. Res. Lett. 5, 595–598.

Papanastassiou D. A., Wasserburg G. J., and Burnett D. S.(1970) Rb–Sr ages of lunar rocks from the Sea of Tran-quility. Earth Planet. Sci. Lett. 8, 1–19.

Papike J. J., Ryder G., and Shearer C. K. (1998) Lunarsamples. In Planetary Materials (ed. J. J. Papike). Rev.Mineral., Mineralogical Society of America, Washington,DC, vol. 36, 5-1–5-234.

Patchett P. J. and Tatsumoto M. (1980) Lu–Hf total-rockisochron for the eucrite meteorites. Nature 288, 571–574.

Patterson C. C. (1955) The 207Pb/206Pb ages of some stonemeteorites. Geochim. Cosmochim. Acta 7, 151–153.

Patterson C. C. (1956) Age of meteorites and the Earth.Geochim. Cosmochim. Acta 10, 230–237.

Perron C., Bourot-Denise M., Marti K., Kim S., andCrozaz G. (1988) The metal–phosphate connection inchondrites. Meteoritics 27, 275.

Petaev M. I., Baruskova L. D., Lipschutz M. E., WangM.-S., Arsikan A. A., Clayton R. N., and Mayeda T. K.(1994) The Divnoe meteorite: petrology, chemistry,oxygen isotopes and origin. Meteoritics 29, 182–199.

Podosek F. A. and Brannon J. C. (1991) Chondritechronology by initial 87Sr/86Sr in phosphates. Meteori-tics 26, 145–152.

Podosek F. A. and Huneke J. C. (1973) Argon 40–argon 39chronology of four calcium-rich achondrites. Geochim.Cosmochim. Acta 37, 667–684.

Podosek F. A. and Nichols R. H., Jr. (1997) Short-livedradionuclides in the solar nebula. In Astrophyical Impli-cations of the Laboratory Study of Presolar Materials(eds. T. J. Bernatowicz and E. Zinner). American Insti-tute of Physics, Woodbury, pp. 617–647.

Podosek F. A., Zinner E. K., MacPherson G. J., LundbergL. L., Brannon J. C., and Fahey A. J. (1991) Correlatedstudy of initial 87Sr/86Sr and Al/Mg isotopic systematicsand petrologic properties in a suite of refractory inclu-sions from the Allende meteorite. Geochim. Cosmochim.Acta 55, 1083–1110.

Prinzhofer A., Papanastassiou D. A., and Wasserburg G. J.(1992) Samarium–neodymium evolution of meteorites.Geochim. Cosmochim. Acta 56, 797–815.

Rankenburg K., Brandon A. D., and Neal C. R. (2006)Neodymium isotope evidence for a chondritic composi-tion of the Moon. Science 312, 1369–1372.

Rubin A. E., Kallemeyn G. W., Wasson J. T., Clayton R.N., Mayeda T. K., Grady M., Verchovsky A. B., EugsterO., and Lorenzetti S. (2003) Formation of metal andsilicate globules in Gujba: a new Bencubbin-like meteo-rite fall. Geochim. Cosmochim. Acta 67, 3283–3298.

Scherer E., Munker C., and Mezger K. (2001) Calibrationof the lutetium–hafnium clock. Science 293, 683–687.

Shen J. J., Papanastassiou D. A., and Wasserburg G. J.(1996) Precise Re–Os determinations and systematics ofiron meteorites. Geochim. Cosmochim. Acta 60, 2887–2900.

Shih C.-Y., Nyquist L. E., Bogard D. D., McKay G. A.,Wooden J. L., Bansal B., and Wiesmann H. (1982)Chronology and petrology of young achondrites, Sher-gotty, Zagami and ALHA77005: late magmatism on ageologically active planet. Geochim. Cosmochim. Acta 46,2323–2344.

Shih C.-Y., Nyquist L. E., Reese Y., and H. Wiesmann(1998) The chronology of the nakhlite Lafayette: Rb–Srand Sm–Nd isotopic ages. In Lunar Planet. Sci. XXXIV,#1145. The Lunar and Planetary Institute, Houston(CD-ROM).

Shih C.-Y., Wiesmann H., Nyquist L. E., and Misawa K.(2002) Crystallization age of Antarctic nakhlite Y000593:further evidence for nakhlite launch pairing. Antarct.Meteorites XXVII, 151–153.

Shimoda G., Nakamura N., Kimura M., Kani T.,Nohda S., and Yamamoto K. (2005) Evidence from theRb–Sr system for 4.4Ga alteration of chondrules in theAllende (CV3) parent body. Meteorit. Planet. Sci. 40,1059–1072.

Smoliar M. I. (1993) A survey of Rb–Sr systematics ofeucrites. Meteoritics 28, 105–113.

Smoliar M. I., Walker R. J., and Morgan J. W. (1996)Re–Os ages of group IIA, IIIA, IVA, and IVB ironmeteorites. Science 271, 1099–1102.

Snyder G. A., Borg L. E., Nyquist L. E., and Taylor L. A.(2000) Chronology and isotopic constraints on lunarevolution. In Origin of the Earth and Moon (eds.R. Canup and K. Righter). University of Arizona Press,Tucson, AZ, pp. 361–395.

Stephan T. and Jessberger E. K. (1988) 40Ar–39Ar ages oftypes 3 and 4, L and H chondrites from Antarctica.Meteoritics 23, 373–377.

Stoffler D., Bischoff A., Buchvald V., and Rubin A. E.(1988) Shock effects in meteorites. In Meteorites and theEarly Solar System (eds. J. F. Kerridge and M. S.Matthews). University of Arizona Press, Tucson, AZ, pp.165–202.

Sugiura N. and Hoshino H. (2003) Mn–Cr chronologyof five IIIAB iron meteorites. Meteorit. Planet. Sci. 38,117–143.

Swindle T. D., Spudis P. D., Taylor G. J., Korotev R. L.,and Nichols R. H. (1991) Searching for Crisium basin

References 25

ejecta—chemistry and ages of Luna 20 impact melts.Proc. 21st Lunar Planet. Sci. Conf. 167–181.

Swindle T. D., Treiman A. H., Lindstrom D. L., BurklandM. K., Cohen B. A., Grier J. A., Li B., and Olsen E. K.(2000) Noble gasses in iddingsite from the Lafayette me-teorite: evidence for liquid water on Mars in the last fewhundred million years. Meteorit. Planet. Sci. 35, 107–115.

Takeda H., Morei H., and Bogard D. D. (1994) Mineralogyand 39Ar–40Ar age of an old pristine basalt: thermal his-tory of the HED parent body. Earth Planet. Sci. Lett.122, 183–194.

Tatsumoto M., Knight R. J., and Allegre C. J. (1973) Timedifferences in the formation of meteorites as determinedfrom the ratio of lead-207 to lead-206. Science 180, 1279–1283.

Tatsumoto M., Unruh D. M., and Desborough G. A.(1976) U–Th–Pb and Rb–Sr systematics of Allende andU–Th–Pb systematics of Orgueil. Geochim. Cosmochim.Acta 40, 617–634.

Taylor L. A., Shervais J. W., Hunter R. H., Shih C. Y.,Nyquist L., Bansal B., Wooden J., and Laul J. C. (1983)Pre-4.2 AE mare basalt volcanism in the lunar highlands.Earth Planet. Sci. Lett. 66, 33–47.

Tera F. and Carlson R. W. (1999) Assessment of thePb–Pb and U–Pb chronometry of the early Solar System.Geochim. Cosmochim. Acta 63, 1877–1889.

Tera F., Carlson R. W., and Boctor N. Z. (1997) Radio-metric ages of basaltic achondrites and their relation tothe early history of the solar system. Geochim. Cosmo-chim. Acta 61, 1713–1731.

Tera F., Papanastassiou D. A., and Wasserburg G. J.(1974) Isotopic evidence for a terminal lunar cataclysm.Earth Planet. Sci. Lett. 22, 1–21.

Tera F. and Wasserburg G. J. (1974) U–Th–Pb systematicson lunar rocks and inferences about lunar evolution andthe age of the moon. Proc. 5th Lunar Sci. Conf. 1571–1599.

Tilton G. R. (1988) Age of the solar system. In Meteoritesand the Early Solar System (eds. J. F. Kerridge and M. S.Matthews). University of Arizona Press, Tucson, AZ, pp.249–258.

Tonui E. K., Ngo H. H., and Papanastassiou D. A. (2003)Rb–Sr and Sm–Nd study of the D’Orbigny angrite. InLunar Planet. Sci. XXXIV, #1812. Lunar and PlanetaryInstitute, Houston (CD-ROM).

Torigoye-Kita N., Misawa K., and Tatsumoto M. (1995a)U–Th–Pb and Sm–Nd isotopic systematics of the Goal-para ureilites: resolution of terrestrial contamination.Geochim. Cosmochim. Acta 59, 381–390.

Torigoye-Kita N., Tatsumoto M., Meeker G. P., and YanaiK. (1995b) The 4.56 Ga age of the MET 78008 ureilite.Geochim. Cosmochim. Acta 59, 2319–2329.

Treiman A. H., Barrett R. A., and Gooding J. L. (1993)Preterrestrial aqueous alteration of the Lafayette (SNC)meteorite. Meteoritics 28, 86–97.

Turner G. (1969) Thermal histories of meteorites by the39Ar–40Ar method. In Meteorite Research (ed. P. M.Milliman). Springer-Verlag, New York, pp. 407–417.

Turner G., Enright M. C., and Cadogan P. H. (1978) Theearly history of chondrite parent bodies inferred from40Ar–39Ar ages. Proc. 9th Lunar Planet. Sci. Conf.989–1025.

Turner G., Knott S. F., Ash R. D., and Gilmour J. D. (1997)Ar–Ar chronology of the martian meteorite ALH84001:evidence for the timing of early bombardment of Mars.Geochim. Cosmochim. Acta 61, 3835–3850.

Unruh D. M., Nakamura N., and Tatsumoto M. (1977)History of the Pasamonte achondrite: relative suscepti-bility of the Sm–Nd, Rb–Sr, and U–Pb systems tometamorphic events. Earth Planet. Sci. Lett. 37, 1–12.

Wadhwa M. and Lugmair G. W. (1995) Sm–Nd systematicsof the eucrite Chervony Kut. In Lunar Planet. Sci. XXVI.The Lunar and Planetary Institute, Houston, pp. 1453–1454.

Wadhwa M. and Lugmair G. W. (1996) Age of the eucrite‘‘Caldera’’ from convergence of long-lived and short-lived chronometers. Geochim. Cosmochim. Acta 60,4889–4893.

Wadhwa M., McSween Jr H. Y., and Crozaz G. (1994)Petrogenesis of the shergottite meteorites inferred fromtheir minor and trace element microdistributions. Geo-chim. Cosmochim. Acta 58, 4213–4229.

Warren P. H. and Wasson J. T. (1979) The origin ofKREEP. Rev. Geophys. Space Phys. 17, 73–88.

Wasserburg G. J. (1985) Short-lived nuclei in the early solarsystem. In Protostars and Planets II (eds. D. C. Black andM. S. Matthews). The University of Arizona Press,Tucson, AZ, pp. 703–737.

Wasserburg G. J., Papanastassiou D. A., and Sanz H. G.(1969) Initial strontium for a chondrite and the determi-nation of a metamorphism or formation interval. EarthPlanet. Sci. Lett. 7, 33–43.

Wasserburg G. J., Tera F., Papanastassiou D. A., andHuneke J. C. (1977) Isotopic and chemical investiga-tions on Angra dos Reis. Earth Planet. Sci. Lett. 35,294–316.

Weigel A., Eugster O., Koeberl C., and Krahenbuhl U. (1996)Primitive differentiated achondrite Divnoe and its relation-ship to brachinites. In Lunar Planet. Sci. XXVII. The Lunarand Planetary Institute, Houston, pp. 1403–1404.

Wiechert U. H., Halliday A. N., Palme H., and Rumble D.(2004) Oxygen isotope evidence for rapid mixing of theHED meteorite parent body. Earth Planet. Sci. Lett. 221,373–382.

Wood J. (1972) Fragments of terra rock in the Apollo 12soil samples and a structural model of the Moon. Icarus16, 461–501.

Zartman R. E., Jagoutz E., and Bouring S. A. (2006) Pb–Pbdating of the D’Orbigny and Asuka 881371 angrites anda second absolute time calibration of the Mn–Cr chro-nometer. In Lunar Planet. Sci. XXXVII, #1580. The Lu-nar and Planetary Institute, Houston (CD-ROM).

Zolensky M. and McSween H. Y., Jr. (1988) Aqueousalteration. In Meteorites and the Early Solar System (eds.J. F. Kerridge and M. S. Matthews). University ofArizona Press, Tucson, AZ, pp. 114–143.

r 2007, Elsevier Ltd. All rights reserved. Treatise On Geochemistry

No part of this publication may be reproduced, stored in a retrieval system or

transmitted in any form or by any means, electronic, mechanical, photocopying,

recording or otherwise, without prior written permission of the Publisher.

ISBN (set): 0-08-043751-6

pp. 1–25

1.27 Long-LivedChronometers

Documents

Transcript of 1.27 Long-LivedChronometers