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1.07
Chondritesand Their ComponentsE.R.D. Scott and A.N. Krot
University of Hawai’i atManoa, Honolulu, HI, USA

1.07.1 INTRODUCTION
2 1.07.1.1 What Are Chondrites? 2 1.07.1.2 Why Study Chondrites? 2 1.07.1.3 Historical Views on Chondrite Origins 3 1.07.1.4 Chondrites and the Solar Nebula 4
1.07.2 CLASSIFICATION AND PARENT BODIES OF CHONDRITES
4 1.07.2.1 Chondrite Groups, Clans, and Parent Bodies 4 1.07.2.2 Ordinary Chondrites 7 1.07.2.3 Carbonaceous Chondrites 7 1.07.2.4 Enstatite Chondrites 7
1.07.3 BULK COMPOSITION OF CHONDRITES
8 1.07.3.1 Cosmochemical Classification of Elements 8 1.07.3.2 Chemical Compositions of Chondrites 8 1.07.3.3 Isotopic Compositions of Chondrites 10 1.07.3.4 Oxygen-Isotopic Compositions 10
1.07.4 METAMORPHISM, ALTERATION, AND IMPACT PROCESSING
11 1.07.4.1 Petrologic Types 11
1.07.4.1.1 Type 1 and 2 chondrites
11 1.07.4.1.2 Type 3 chondrites 12 1.07.4.1.3 Type 4–6 chondrites 13
1.07.4.2 Thermal History and Modeling
14 1.07.4.3 Impact Processing of Chondritic Asteroids 14
1.07.5 CHONDRITIC COMPONENTS
15 1.07.5.1 Calcium- and Aluminum-Rich Inclusions 15
1.07.5.1.1 Comparison of CAIs from different chondrite groups
16 1.07.5.1.2 Oxygen-isotopic compositions of CAIs 17
1.07.5.2 Forsterite-Rich Accretionary Rims around CAIs
19 1.07.5.3 Amoeboid Olivine Aggregates 20
1.07.5.3.1 Mineralogy and petrology of AOAs
21 1.07.5.3.2 Trace elements and isotopic composition of AOAs 22 1.07.5.3.3 Origin of amoeboid olivine aggregates 25
1.07.5.4 Aluminum-Rich Chondrules
25 1.07.5.4.1 Isotopic and trace element studies of aluminum-rich chondrules 25 1.07.5.4.2 Origin of aluminum-rich chondrules 27 1.07.5.4.3 Relict CAIs and AOAs in ferromagnesian chondrules 29 1.07.5.4.4 Relict chondrules in igneous CAIs and igneous CAIs overgrown by chondrule material 31
1.07.5.5 Refractory Forsterite-Rich Objects
33 1.07.5.6 Chondrules 34
1.07.5.6.1 Chondrule textures, types, and thermal histories
34 1.07.5.6.2 Compositions of chondrules 36 1.07.5.6.3 Compound chondrules, relict grains, and chondrule rims 37 1.07.5.6.4 Closed-system crystallization 38 1.07.5.6.5 Open-system crystallization 38 1.07.5.6.6 Chondrules that formed on asteroids 39
1.07.5.7 Metal and Troilite
40 1.07.5.7.1 CR chondrite metallic Fe,Ni 41 1.07.5.7.2 CH and CB chondrite metallic Fe,Ni 42 1.07.5.7.3 Troilite 43
1

Ed Scott

Note

Scott E. R. D. and Krot A. N. (2007) Chondrites and their components. In Meteorites, Comets and Planets (ed. A. M. Davis) Chapter 1.07, Treatise on Geochemistry Update 1, Elsevier, online at http://www.sciencedirect.com/science/referenceworks/9780080437514 Publication numbers: HIGP #1460 SOEST #6882

2 Chondrites and Their Components

1.07.5.8 Matrix Material
43 1.07.5.8.1 CI1 chondrites 45 1.07.5.8.2 CM2 chondrite matrices 46 1.07.5.8.3 CR2 chondrite matrices 46 1.07.5.8.4 CO3 chondrite matrices 46 1.07.5.8.5 CK3 chondrite matrices 50 1.07.5.8.6 CV3 chondrite matrices 50 1.07.5.8.7 Matrix of ungrouped C chondrites, Acfer 094, and Adelaide 51 1.07.5.8.8 H–L–LL3 chondrite matrices 51 1.07.5.8.9 K3 chondrite matrix (Kakangari) 52 1.07.5.8.10 Heavily altered, matrix-rich lithic clasts 52 1.07.5.8.11 Origins of matrix phases 53
1.07.6 FORMATION AND ACCRETION OF CHONDRITIC COMPONENTS
55
1.07.7 HEATING MECHANISMS IN THE EARLY SOLAR SYSTEM
55 1.07.7.1 Nebular Shocks 56 1.07.7.2 Jets and Outflows 56 1.07.7.3 Impacts on Planetesimals 58
ACKNOWLEDGMENTS
58
REFERENCES
59
1.07.1 INTRODUCTION

1.07.1.1 What Are Chondrites?

Chondrites are meteorites that provide thebest clues to the origin of the solar system. Theyare the oldest known rocks—their componentsformed during the birth of the solar systemca. 4,567Ma—and their abundances of non-volatile elements are close to those in the solarphotosphere. Chondrites are broadly ultramaficin composition, consisting largely of iron,magnesium, silicon, and oxygen. The mostabundant constituents of chondrites are chond-rules, which are igneous particles that crystal-lized rapidly in minutes to hours. They arecomposed largely of olivine and pyroxene,commonly contain metallic Fe,Ni and are0.01–10mm in size. Some chondrules arerounded as they were once entirely molten butmany are irregular in shape because they wereonly partly melted or because they accretedother particles as they solidified. Chondritesthemselves were never molten. The definition ofa chondrite has expanded recently with thediscovery in Antarctica and the Sahara Desertof extraordinary meteorites with chondrules10–100mm in size, and chondrites so rich inmetallic Fe,Ni that they were initially classifiedas iron meteorites with silicate inclusions. Thus,in meteoritics, as in other fields of planetaryscience, new discoveries sometimes requiredefinitions to be modified.

Chondrites are so diverse in their miner-alogical and textural characteristics that it isnot possible to describe a typical chondrite. Weshow one with diversely textured chondrulesincluding prominent, esthetically pleasing,rounded chondrules (Figure 1a), and anotherwith more uniformly textured chondrules(Figure 1b). Owing to the high abundance of

rounded or droplet chondrules in the so-calledordinary chondrites (Figure 1a), studies of theorigin of chondrules have commonly beenbased on these chondrites.

Chondrites contain diverse proportions ofthree other components: refractory inclusions(0.01–10vol.%), metallic Fe,Ni (o0.1–70%), andmatrix material (1–80%). Refractory inclusionsare tens of micrometers to centimeters in dimen-sions, lack volatile elements, and are the productsof high-temperature processes including conden-sation, evaporation, and melting. Two types arerecognized: calcium- and aluminum-rich inclu-sions or Ca–Al-rich inclusions (CAIs), and am-oeboid olivine aggregates (AOA). CAIs arecomposed of minerals such as spinel, melilite,hibonite, perovskite, and Al–Ti-diopside,whichare absent in other chondritic components (seeChapter 1.08). Amoeboid olivine aggregatesconsist of fine-grained olivine, Fe,Ni metal,and a refractory component largely composedof aluminum-diopside, anorthite, spinel, andrare melilite. Grains of metallic Fe,Ni occur in-side and outside the chondrules as grains up to amillimeter in size and, like the chondrules andrefractory inclusions, formed at high tempera-tures. Matrix material is volatile-rich, and fine-grained (10 nm–5mm) and forms rims on othercomponents and fills the interstices betweenthem. Chondrite matrices have diverse miner-alogies: most are disequilibrium mixtures of hy-drated and anhydrous silicates, oxides, metallicFe–Ni, sulfides, and organic material and con-tain rare presolar grains.

1.07.1.2 Why Study Chondrites?

Goldschmidt, Suess, and Urey showed thatchondrites provide the best estimates for themean abundances of condensable elements inthe solar system. These estimates were essential

POP POP

Mg K� PCA91082, CR

POII

POI

POI1 mm(a)

Mg K�

POP

PP

1 mm

C

BO

Tieschitz, H/L3.6

RP

POI

POII

(b)

Figure 1 Maps showing magnesium concentrationsin two chondrites: (a) PCA91082, a CR2 carbon-aceous chondrite, and (b) Tieschitz, an H/L3.6ordinary chondrite. In CR chondrites, as in mostcarbonaceous chondrites, nearly all chondrules haveporphyritic textures and are composed largely offorsterite (white grains), enstatite (gray), and metal-lic Fe,Ni (black). The subscripts show type I chond-rules, which are common, and type II, which areFeO-rich and rare in this chondrite. Tieschitz, likeother ordinary chondrites, is composed of all kindsof chondrules with diverse FeO concentrations. Keyto chondrule types: BO, barred olivine; C, crypto-crystalline, PO, porphyritic olivine; POP, porphyriticolivine-pyroxene; PP, porphyritic pyroxene; RP,radial pyroxene. These maps were made with an

electron microprobe from Mg Ka X-rays.

Introduction 3

for developing theories for the formation ofelements in evolved stars (see Chapters 1.01and 1.03). Presolar grains (see Chapter 1.02)provide additional clues to nucleosynthesis andthe subsequent growth of circumstellar grains.Chondrules, metal grains, refractory inclusions,and matrix materials formed under very diverseconditions in the solar system and appear tooffer insights into processes that occurredduring the formation of the Sun and planetsfrom a collapsing cloud of interstellar dust andgas. The rocks themselves provide clues to the

geological processes including impact processesthat affected asteroids over 4.5Ga. Studies ofchondrites help us to match chondrite groupswith asteroid classes, to understand the originand evolution of the asteroid belt, the nature ofplanetesimals that accreted into terrestrialplanets, and the reason for the dearth of plan-etary material between Mars and Jupiter. Fi-nally, chondrite studies help us to understandthe physical and mineralogical structure of un-melted asteroids and to assess what should bedone about rogue near-Earth objects thatthreaten Earth.

In this chapter, we focus on the insightsthat chondrites offer into the earliest stages inthe formation of asteroids and planetesimalsin the solar nebula—the protoplanetary diskof dust and gas that evolved into the plane-tary system. Other chapters review many otheraspects of chondrite studies. Since the firstedition of the Treatise on Geochemistry, therehave been considerable advances in understan-ding the chronology of chondrule and CAIformation by reconciling data from short- andlong-lived isotopes (see Kita et al., 2005 andChapter 1.16), possible mechanisms responsiblefor the oxygen-isotopic anomalies in chond-rites, and in the use of detailed isotopic meas-urements to distinguish and date nebula andparent body alteration effects. A surprise tomany is the discovery that planetesimals al-ready existed in the nebula when chondruleswere forming, and that in some cases at least,chondrules may have formed from preexistingbodies (Kleine et al., 2005; Hevey and Sanders,2006; Krot et al., 2005a). Comets too are notas pristine as once believed. Chondritic porousinterplanetary dust particles (IDPs), like pris-tine chondrite matrices, contain abundantforsterite and enstatite grains that lack largeisotopic anomalies, showing that most silicatesin comets were thermally processed in thesolar nebula (Messenger et al., 2003; Woodenet al., 2005).

1.07.1.3 Historical Views on ChondriteOrigins

At the beginning of the space age in the1960s, there were two views about the origin ofchondrules and the variations in the concen-trations of the most volatile elements in chond-rites. Ringwood, Urey, and others invokedvolcanism, impacts, and other planetary proc-esses, whereas Wood, Larimer, and Andersinvoked processes in the solar nebula (Wood,1963; Larimer, 1967; Larimer and Anders, 1967).In the early 1970s, studies of CAIs by Grossmanand others appeared to strongly favor an origin


for chondritic components as condensates fromthe solar nebula (Grossman and Larimer, 1974).However, this simple picture was disrupted bydiscoveries of ubiquitous isotopic anomalies inCAIs and the complexity of the mineralogicaland isotopic record in chondritic components,which seemed inconsistent with the standardmodel of a hot, monotonically cooling solarnebula. In addition, astronomical evidence andtheoretical studies suggested that the solarnebula was not hot enough to vaporize silicateswhere chondrites formed. The solar nebula wasenvisaged as a place where there were dynamic,energetic, dust-rich zones with constantlychanging dust/gas ratios and fluctuating hightemperatures. ‘‘Mineral material caught in thismaelstrom went through multiple cycles ofmelting, evaporation, recondensation, crystal-lization, and aggregation’’ (Wood, 1988).

Since the mid-1980s, much progress has beenmade in understanding chondrites, though theorigin of their components has not beenresolved. Discoveries of new kinds of carbon-aceous chondrites have shown that the Allendemeteorite, which is the most-studied chondrite,is not as pristine as once believed. Its compo-nents were modified in asteroids (e.g., Krotet al., 1995, 1998b) or the solar nebula (e.g.,Palme and Fegley, 1990; Weisberg and Prinz,1998). Detailed isotopic and chemical data forCAIs and chondrules from ion microprobestudies have been invaluable in understandinghow and where chondritic components werealtered, but have not yet resolved their origin.Most authors now accept that CAIs and mostchondrules formed in some kind of solar nebulaenvironment, and several promising modelshave been developed, but some workers arguethat planetesimals were essential for the forma-tion of chondrules (e.g., Sanders, 1996; Lugmairand Shukolyukov, 2001; Hsu et al., 2000). Tometeoriticists like Wood, who argued in 1963that chondrites were aggregates of solar nebularcondensates, it has seemed that chondrite re-search has stagnated and that little progress hasbeen made in developing plausible theories forthe origins of chondrules, CAIs, and chondrites(Kerr, 2001). However, most find the excitementof studying presolar grains, solar nebula parti-cles, and annual batches of new types of meteo-rites from Antarctica to be irresistible (Bischoff,2001).

1.07.1.4 Chondrites and the Solar Nebula

Chondrites used to be considered as productsof a ‘‘minimum’’ solar nebula (containing justenough material to make the planets) thataccreted from components formed at a single

nebula location. However, it now appears thatchondritic components formed at various loca-tions over a period of several million yearsduring which the nebula mass would havedecreased by an order of magnitude or more.Refractory inclusions probably formed close tothe protosun at an early stage when the diskwas most active and accretion rates onto theprotosun were highest (Wood, 2000b, 2004).Chondrules probably formed over several mil-lion years in localized heating events near theiraccretion site with the youngest and oldestchondrules forming and accreting by differentmechanisms (see Kita et al., 2005; Scott andKrot, 2005b; Krot et al., 2005a). However,some chondrules, such as CAIs, may haveformed at the inner edge of the disk prior toinjection into the asteroid belt by bipolar out-flows (Liffman and Brown, 1996a, b; Shu et al.,1996, 2001). Alternatively, Alexander et al.(2001) have argued that CAIs and chondrulesboth formed in localized heating events in theasteroid belt.

Astronomical studies of star-forming regionshave provided considerable insights into theearly evolution and dynamic environment ofthe Sun and its protoplanetary disk andpossible origins for components in chondritesand comets (Reipurth, 2005; Alexander, 2005).However, the 71 difference between the plane-tary and solar angular momentum vectors(Hutchison et al., 2001), the multiplicity ofprotostellar systems, and the discovery ofnumerous ‘‘hot Jupiters’’ around other starsprovide additional reminders of the complexityof disk evolution.

1.07.2 CLASSIFICATION AND PARENTBODIES OF CHONDRITES

1.07.2.1 Chondrite Groups, Clans, andParent Bodies

Many thousands of chondrites have beenclassified into 15 groups: about 15 other chond-rites do not fit comfortably into these groupsand are called ungrouped (Table 1). Thirteen ofthese groups comprise three classes: carbon-aceous (CI, CM, CO, CV, CR, CH, CB, andCK), ordinary (H, L, and LL), and enstatite(EH and EL). The K and R chondrites do notbelong to the three classes (see Chapter 1.05).Chondrites are also classified into petrologictypes 1–6, which indicate the extent of aster-oidal processing (see table 2 of Chapter 1.05).Type 3 chondrites are the least metamorphosedand least altered; type 6 are the most meta-morphosed. Type 1 chondrites, which lackchondrules, are composed almost entirely of

Table 1 Abundances of refractory inclusions, chondrules, metallic Fe,Ni, and matrix and other key properties of the chondrite groups.

Group Refract. lith./Mgrel. CI a

CAI and AOAðvol:%Þ

Chondruleaverage diameter

ðmmÞ

Chondrulesðvol:%Þb

Metalðvol:%Þ

Matrixðvol:%Þc

Fall frequencyd

ð%ÞExamples

CarbonaceousCI 1.00 o0.01 none o5 o0.01 95 0.5 Ivuna, OrgueilCM 1.15 5 0.3 20 0.1 70 1.6 MurchisonCO 1.13 13 0.15 40 1–5 30 0.5 OrnansCV 1.35 10 1.0 45 0–5 40 0.6 Vigarano,

AllendeCR 1.03 0.5 0.7 50–60 5–8 30–50 0.3 RenazzoCH 1.00 0.1 0.02–0.09 B70 20 5 0 ALH 85085CBa 1.0 o0.1 B5 40 60 o5 0 BencubbinCBb 1.4 o0.1 B0.5 30 70 o5 0 QUE 94411CK 1.21 4 0.8 15 o0.01 75 0.2 Karoonda

OrdinaryH 0.93 0.01–0.2 0.3 60–80 8 10–15 34.4 DhajalaL 0.94 o0.1 0.5 60–80 3 10–15 38.1 KhoharLL 0.90 o0.1 0.6 60–80 1.5 10–15 7.8 Semarkona

EnstatiteEH 0.87 o0.1 0.2 60–80 8 o0.1–10 0.9 Qingzhen, AbeeEL 0.83 o0.1 0.6 60–80 15 o0.1–10 0.8 Hvittis

OtherK 0.9 o0.1 0.6 20–30 6–9 70 0.1 KakangariR 0.95 o0.1 0.4 440 o0.1 35 0.1 Rumuruti

Sources: Scott et al. (1996); other data from Weisberg et al. (1996, 2001), Rubin (2000), Krot et al. (2002a), Kimura et al. (2002), and Bischoff et al. (1993).aMean ratio of refractory lithophiles relative to magnesium, normalized to CI chondrites. b Includes chondrule fragments and silicates inferred to be fragments of chondrules. c Includes matrix-rich clasts, which accountfor all matrix in CH and CBb chondrites (Greshake et al., 2002). dFall frequencies based on 918 falls of differentiated meteorites and classified chondrites (Grady, 2000).


minerals that formed during aqueous altera-tion. Type 2 chondrites are partly altered.

Elemental ratios are commonly used for com-paring chondrite compositions instead of con-centrations, as some chondrites have excessiveamounts of one component such as water ormetallic Fe,Ni, for example, that dilute elementsin other components. The elements silicon andmagnesium are commonly used for normalizinglithophiles; nickel is used for comparing side-rophile elements. CI chondrites, which have thehighest concentrations of volatile elements, areused as the reference standard because theyprovide the best match for the composition ofthe solar photosphere, neglecting the incom-pletely condensed elements hydrogen, helium,carbon, nitrogen, oxygen, and the noble gases(see Chapter 1.03).

Ratios of refractory and moderately volatilelithophile (figure 3 in Chapter 1.05) or siderophileelements (Figure 2) provide the best criteria forclassifying chondrites. On these plots, the rangewithin each group is a small fraction of the totalrange shown by all chondrites. Each group has aunique chemical composition implying that itsmembers formed in a separate location. The CBgroup, which is split into two subgroups withrather different properties, is an exception. Un-fortunately, many newly discovered, unusualchondrites have not been chemically analyzed,so their classification and the compositionallimits of newly discovered groups are not welldefined. Two other properties are useful forclassifying chondrites and understanding theircompositional diversity and genetic relation-ships: bulk oxygen-isotopic compositions (see

18

16

14

12

10

83 4 5 6 7 8 9

EH

EL

Ga/Ni × 104 atom ratio

Cl

CM

LL

L

CO

CV

H

Ir/N

i × 1

06 at

om r

atio

Figure 2 Plot of Ga/Ni versus Ir/Ni showing bulkcompositions of chondrites in nine groups. Eachgroup is well resolved. The proportions of thesesiderophiles are not correlated with other chemicalproperties showing that the chemical variations inchondrites are complex. Reproduced by permissionof Verlag der Zeitschrift fur Naturforschuang from

Scott and Newsom (1989).

Chapter 1.06) and the distribution of ironamong silicate, metal, and sulfide phases (figure7 of Chapter 1.05). The latter requires classicalmethods of rock analysis (Jarosewich, 1990),which are no longer applied to chondrite. Manychondrite groups can also be distinguished byother parameters such as the sizes and propor-tions of their components and the minerals intheir chondrules, CAIs, or other components.See Brearley and Jones (1998) (hereafter ab-breviated B&J), for a detailed account ofchondrite mineralogy.

The chemical, mineralogical, and isotopicproperties of chondrites distinguish closelyrelated groups, which constitute ‘‘clans.’’ Forexample, chondrules in the H, L, and LLgroups are chemically and isotopically similar,though their average properties are different,and these three groups form a clan. Four otherclans are recognized: the EH–EL, the CM–CO,the CV–CK, and the CR–CH–CB.

Cosmic-ray exposure ages suggest thatchondrites in many groups were exposed tospace as meter-sized objects in a limitednumber of impacts (Chapter 1.13). For exam-ple, about one-third of the LL chondrites wereexposed to cosmic rays 15Ma and half theH chondrites have exposure ages of 7Ma (Grafand Marti, 1994, 1995). Three or four largecollisions appear to have generated two-thirdsof all H chondrites (Wieler, 2002). Since allpetrologic types are found in the 7 and 33Mapeaks, it is possible that all H chondrites comefrom a single body. Some other groups showdistinctive impact ages or shock properties,suggesting that they were involved in a singleimpact. For example, many L chondrites, whichare strongly shocked, have radiometric ages ofB0.5Ga (Bogard, 1995). These features suggestthat each group comes from one, or possibly afew bodies.

We do not know for certain that two chond-rite groups did not originate from a single bodybut it is rather unlikely (except possibly for theEH and EL chondrites). Evidence from mete-orites, impact modeling of asteroid collisions,and asteroid density determinations suggeststhat all asteroids, except the few largest, werewell fragmented and mixed by collisions beforemeteoroids were removed. Impacts also weldrock fragments from different locations intocoherent rocks that reach Earth as meteorites.Thus, if CV and CO chondrites, for example,had formed in a single body, we might expectto find meteorites containing both types ofmaterial. Their absence makes it unlikely thatone body ever contained both groups. Never-theless, regolith breccias contain a few volumepercent of foreign clasts showing that smallamounts of material from different sources

Classification and Parent Bodies of Chondrites 7

were mixed together. The close proximity of theorbits of an EL6 and an H5 chondrite, whichmight imply that they have a common parent(Spurny et al., 2003), appears to be coincidental(Pauls and Gladman, 2005).

Given the extraordinary diversity of chemicaland mineralogical properties of the 15 chondritegroups and the rapid rate at which new groupsare being identified (only nine were listed byWasson, 1985), we should expect many moreunusual types to be discovered. Our sampling ofchondrites is also biased toward tough rocksthat can survive the journey to Earth, and wecannot expect these to be representative sam-ples. Thus, it is important to study componentsin all kinds of chondrites if we need to under-stand how they formed, not just those in themost common chondrite groups or those thatare present in the largest falls, like Allende.

1.07.2.2 Ordinary Chondrites

Since ordinary chondrites account for B80%of all meteorite falls (Table 1), it was oncebelieved that their parent bodies were commonin the main asteroid belt. However, spectralstudies show that most asteroids are dark andfeatureless (C and related types), and most ofthe brighter S types are spectrally unlike ordi-nary chondrites. H group chondrites probablycome from one or more S-type asteroids, pos-sibly 6 Hebe (Burbine et al., 2002). Ordinarychondrites are probably rare in the main part ofthe asteroid belt (Meibom and Clark, 1999),though a significant fraction of the near-Earthasteroids including Eros are probably made ofordinary chondritic material (Binzel et al., 2002;Chapman, 2004). A few ordinary chondritesincluding Tieschitz (Figure 1a) do not fit com-fortably into the H, L, and LL groups and maybe derived from separate bodies.

1.07.2.3 Carbonaceous Chondrites

Some carbonaceous chondrites are rich incarbon (CI and CM chondrites have 1.5–6%carbon), but others are not. Carbonaceouschondrites are now defined on the basis of theirrefractory elemental abundances, which equal orexceed those in CI chondrites. Carbonaceouschondrites are derived from very diverse aster-oids, which probably formed in very differentlocations. The parent bodies of CI and CMchondrites are highly altered, yet the parentbodies of CH and CB chondrites are less alteredthan all other chondrite bodies. Young et al.(1999) infer from oxygen-isotopic compositionaldata that CI, CM, and CV chondrites couldhave been derived from different zones in a

single, aqueously altered body. However, bulkchemical differences between these groups indi-cate fractionation during nebular processes, notaqueous alteration (see below), and the compo-nents in CM and CV chondrites are quitedifferent.

Most carbonaceous chondrites are thoughtto come from the low-albedo, C-type asteroids,which are the most abundant type between2.7 and 3.4AU (Bell et al., 1989), CM chond-rites may be derived from an altered C-likeasteroid called G-type (Burbine et al., 2002).The ungrouped C chondrite, Tagish Lake, hasbeen linked with D asteroids, which appear todominate the asteroid population beyond 4AU(Jones et al., 1990; Hiroi et al., 2001). CBchondrites have only very minor amounts ofphyllosilicates and may come from W-typeasteroids (‘‘wet-M’’ asteroids). Lodders andOsborne (1999) discuss the possibility that CMand CI chondrites could be derived from asmall fraction of comets that evolve into near-Earth objects after losing volatiles. A smallfraction of near-Earth asteroids are derivedfrom Jupiter-family comets, and Lodders andOsborne (1999) suggest that they are supplyingus with CI and CM chondrites. Gounelle et al.(2006) also favor a cometary origin for Orgueilmeteorite based on its inferred orbit. However,Campins and Swindle (1998) argue that come-tary meteorites, if they exist, probably resemblecarbon-rich, unaltered chondrites withoutchondrules. The extremely fragile nature ofchondritic porous IDPs, which are probablyderived from comets (Rietmeijer, 1998, 2002;Chapter 1.26), suggests that larger cometarymeteoroids would not survive atmospheric en-try (Messenger et al., 2006).

Although most comets clearly accreted farbeyond the asteroid belt, similarities betweenthe matrices of the most primitive carbon-aceous chondrites and the chondritic porousIDPs (Section 1.07.5.8) and the discovery thatsome C-type asteroids show cometary activity(Hsieh and Jewitt, 2006) have blurred the dis-tinction between comets and asteroids.

1.07.2.4 Enstatite Chondrites

The genetic relationships among E chondritesare poorly understood, in part because of poorsampling but also because the cause of theirchemical and mineralogical diversity is obscure.EH and EL chondrites may come from one(Kong et al., 1997) or two bodies (Keil, 1989).Although there are compositional discontinui-ties between the EH and EL groups (e.g., in goldand aluminum), Kong et al. concluded thatelemental concentrations appear to vary


continuously through the sequence EH4, 5,EH3, EL3, and EL6. Oxygen-isotopic composi-tions of EH and EL chondrites show a largeoverlap, but do not resolve the issue (Newtonet al., 2000). On the basis of their noble gasanalyses, Patzer and Schultz (2002b) advancedan entirely different proposal that ‘‘EH3 andEL3 chondrites are derived from one body andEH4–6 and EL4–6 from a second.’’ They foundthat E3 and E4–6 chondrite types have differenttrapped noble gas components (called ‘‘Q-gases’’), and inferred that this difference couldnot be attributed to asteroidal processing. Inaddition, about one-third of the E3 chondritescontain solar noble gases, which are absentamong E4–6 chondrites. Unfortunately, the cos-mic-ray exposure age data do not yet providedefinitive constraints. The 60-odd E chondriteshave a wide range of exposure ages, 0.07–70Mawith broad peaks as their precision isB20% (cf.10% for ordinary chondrites) (Patzer andSchultz, 2002a; Okazaki et al., 2000). However,the data are not inconsistent with a single Echondrite body, as Kong et al. (1997) proposed.E chondrites may be derived from one or moreM-type asteroids (Burbine et al., 2002), whichare concentrated in the inner asteroid belt at2–3AU (Bell et al., 1989).

1.07.3 BULK COMPOSITION OFCHONDRITES

1.07.3.1 Cosmochemical Classification ofElements

The chemical fractionations observed amongchondrites and the compositions of manychondritic components are best understoodin terms of quenched equilibrium betweenphases in a nebula of solar composition (Palme,2001; Chapters 1.03 and 1.15). The equilibriummodel assumes that minerals condensed from, orequilibrated with, a homogeneous solar nebula atdiverse temperatures. Isotopic variations amongchondrites and their components show that thisassumption is not correct and detailed petrologicstudies have identified relatively few chondriticcomponents that resemble equilibrium nebularproducts. Nevertheless, the equilibrium modelis invaluable for understanding the chemicalcomposition of chondrites and their componentsas the solar nebular signature is etched deeplyinto their chemistry and mineralogy.

In a cooling nebula at a pressure of 10�3 atm,calcium, aluminum, and titanium condense asoxides in the temperature range 1,800–1,450Kalong with refractory trace elements like therare earth elements (REEs) and platinum groupmetals (figure 2 of Chapter 1.15). Between

1,450 and 1,350K, magnesium and silicon con-dense as forsterite and enstatite and iron con-denses as metallic Fe,Ni along with cobalt.Between 1,250 and 650K, the alkalis and mod-erately volatile siderophiles condense and at650K, iron reacts to form FeS. Below 650K,the highly volatile elements like the halogensand inert gases condense. The division of theelements into refractories (1,800–1,450K),major elements (1,350–1,250K), moderatelyvolatile elements (1,250–650K), and highlyvolatile elements (o650K) is the key to under-standing the chemical variations among chond-rites and their components. Note that only threephases condense entirely from the gas phase—Al2O3, Mg2SiO4, and Fe,Ni—all other mineralsform by reaction between solids and gas.

At 10�3 atm, FeO concentrations in silicatesare minimal above 650K and all condensedphases are solid. If the total pressure increases,condensation temperatures rise, but the sequencein which minerals condense does not changesignificantly except that silicate melts becomestable at pressures above B10�2 atm. If conden-sation occurs in systems that have been enrichedin dust relative to gas, FeO concentrations in sil-icates become appreciable at high temperatures(up to B1,100K for 103-fold dust enrichments),and liquids are also stabilized (Wood and Hashi-moto, 1993; Ebel and Grossman, 2000). Pres-sures and temperatures at the midplane of thenebula vary considerably as the solar nebulaevolves, decreasing as the accretion rate falls(Wood, 2000a; Woolum and Cassen, 1999).Additional insights into chondrite propertiescan be gained from fractional condensationmodels in which condensates are continuallysequestered (Petaev and Wood, 2000).

1.07.3.2 Chemical Compositions ofChondrites

A large number of chondrites have beenanalyzed with instrumental and radiochemicalneutron activation analysis usingB250–350mgsamples of chondrites (Kallemeyn and Wasson,1981; Wasson and Kallemeyn, 1988; Kalle-meyn et al., 1991, 1994, 1996). Samples of thissize can also be analyzed for many elementswith comparable accuracy using X-ray fluores-cence analysis (Wolf and Palme, 2001). Meancompositions of chondrite groups are given byLodders and Fegley (1998).

Chemical variations among chondrites arecommonly attributed to the accretion of dif-ferent proportions of five components: therefractory component, magnesium silicates,metallic Fe,Ni, moderately volatile elements,and highly volatile elements (see chapters in

Bulk Composition of Chondrites 9

Kerridge and Matthews, 1988; Palme, 2000).In addition, variations in oxygen fugacityare invoked to explain the diverse distributionof iron among metallic and silicate phases(figure 7 of Chapter 1.05). In general, chemicalfractionation patterns are simple: refractoryelements, for example, are uniformly enrichedrelative to CI chondrites in carbonaceous chond-rites by factors of 1.0–1.4, whereas they are de-pleted in ordinary and enstatite chondrites byfactors of 0.8–0.95 (Figure 3). Moderately vola-tile elements decrease in abundance with dec-reasing condensation temperature. (Note thatenstatite chondrites do not follow these rules asclosely as the other groups.) Although the chem-ical fractionation patterns for most groups ofchondrites are known (Wasson and Kallemeyn,1988), their origins have proved to be surpri-singly elusive. In particular, the relative impor-tance of localized processes that madechondrules and CAIs and nebula-wide thermalgradients is controversial (see Cassen, 2001a):

1. Refractory elements. All the elements thatcondense above 1,450K are precisely thosefound in CAIs, and there is some correla-tion between the bulk concentrations ofrefractory elements and refractory inclusions(Table 1). This suggests that the fraction-ation of refractory elements among chond-rites is simply caused by addition or loss ofCAIs. However, this is incorrect as the pro-portion of the refractory elements that is

1.4

1.2

1.0

0.9

0.8

0.7

0.6

0.5

0.4

0.3

Al Sc Ca Lu Sm La Yb

Gro

up: C

I abu

ndan

ce r

atio

CO

CV

CM H

L

EH

EL

Figure 3 Mean abundances of lithophile elements normof increasing volatility in seven chondrite groups. Refraenriched in CO, CM, and CV chondrites and depletedelements, which condense below magnesium and silicofractionations are related in poorly understood ways to th

permission of The Royal Society from

present in CAIs is generally small, except forCVs, where it approaches B50%. Chond-rules, which are much more abundant andhave approximately CI-like levels of refrac-tories relative to silicon, account for most ofthe refractory elements in chondrites. Notethat the depletions of refractory elementsin E and O chondrites require that theirchondrules formed from a condensate thatwas depleted in refractories.

2. Magnesium silicates. Forsterite and ens-tatite, which are the major minerals inchondrules, form in the temperature range1,450–1,350K, the latter by reaction bet-ween forsterite and gaseous SiO. Thus, equi-librium condensation can produce a range ofmolar enstatite/forsterite ratios of 0–1.2(figure 2 of Chapter 1.15), provided thatthe condensates are sequestered from the gasover a range of temperatures. To reach thecomposition of E chondrites, which arealmost pure enstatite, would require removalof forsterite condensate. A correlation be-tween the Mg/Si and refractory/Si ratios ofthe chondrite groups suggests that refracto-ries were partly associated with forsterite(Larimer and Wasson, 1988).

3. Metallic Fe,Ni. which condenses over thesame temperature range as the magnesiumsilicates, is closely associated with them inchondrules. At 10�5 atm, Fe,Ni condenses atslightly lower temperatures than forsterite,but above 10�4 atm there is a reversal. Some

Eu V Mg Si Cr Mn K Na

Si-normalized

alized to CI chondrites and silicon arranged in orderctories (elements condensing above V) are uniformlyin H, L, and EH chondrites. Moderately volatile

n, are all depleted relative to CI chondrites. Thesee formation of CAIs and chondrules. Reproduced byWasson and Kallemeyn (1988).


groups such as the CV3 chondrites did notsuffer metal/silicate fractionation (figure 2 ofChapter 1.15), but the EH–EL, H–L–LL,and CR–CH–CB clans clearly did. Fe/Si ra-tios are B(2–5)�CI levels in CH and CBchondrites because metal was enriched re-lative to silicate; EL, L, and LL chondriteswere depleted in metallic Fe,Ni.

4. Moderately volatile elements. The gradualdepletion of moderately volatiles with dec-reasing condensation temperature can beattributed to loss of fine volatile-rich dust(Wasson and Kallemeyn, 1988), incom-plete condensation due to isolation of con-densed phases (Palme et al., 1988; Cassen,2001b), or to evaporation prior to or duringchondrule formation (see, e.g., Young, 2000;Lugmair and Shukolyukov, 2001; see alsoChapter 1.15).

5. Highly volatile elements may have failed toaccrete from the nebula because of highambient temperatures during accretion orinefficient accretion of volatile-rich dust,or they may simply have been lost duringasteroidal metamorphism.

1.07.3.3 Isotopic Compositions of Chondrites

Major and minor planetary bodies areessentially uniform isotopically (e.g., Palme,2001). The great majority of elements are iso-topically uniform to within B0.01%. Theexceptions are found for oxygen (discussedbelow), effects due to short-lived isotopes(Chapter 1.16), slightly larger effects (per mil orless) due to anomalies in 54Cr and 50Ti in chond-rites and their components (e.g., Podosek et al.,1997; Niemeyer, 1988), and isotopically anoma-lous components in chondrites: interstellargrains, certain rare CAIs called FUN inclusions(containing fractionated and unidentified nuclearisotopic effects) and some hibonite inclusions(see Chapter 1.08). The isotopic anomalies ininterstellar grains are orders of magnitude largerthan those in refractory grains in chondritesand the latter probably reflect incompletehomogenization of the former (e.g., Ott, 1993,2001). Depending on one’s viewpoint, theisotopic anomalies provide evidence thatvaporization of presolar materials was limited(Taylor, 2001) or that isotopic homogenizationwas remarkably complete (Palme, 2001).

1.07.3.4 Oxygen-Isotopic Compositions

The discovery of large and ubiquitous oxygen-isotopic anomalies in refractory inclusions

in carbonaceous chondrites and smaller anom-alies in chondrules transformed the study ofchondrites (see Clayton, 1993; Chapter 1.06). Itgalvanized the search for presolar grains,provided another way to classify meteorites,and offered numerous constraints on the originof chondrites and their constituents and sub-sequent asteroidal processing. In particular, theoxygen-isotopic anomalies appeared to eliminatethe conviction that the meteorites and innerplanets were derived from a hot, gaseous, andhomogeneous nebula (e.g., Begemann, 1980).However, the origin of the oxygen anomaliesremains a mystery, presolar grains were notdiscovered by searching for the source of theanomalous oxygen (Ott, 1993), and the oxygenanomalies are now known to be associated withnebular condensates, as well as evaporative res-idues (see Scott and Krot, 2001; Krot et al.,2002b). The recent introduction of ion and lasermicroprobe techniques has produced an explo-sion of new oxygen-isotopic analyses of mineralsin CAIs and chondrules, which have been espe-cially valuable in distinguishing asteroidal andnebula processes (e.g., Choi et al., 1998; Yuri-moto et al., 1998; Yurimoto and Wasson,2002).

Differences between the bulk compositions ofchondrites, planets and asteroids can beattributed to accretion from different batchesof CAIs, chondrules and other components,which are spread along 16O variation lines on thestandard three-isotope plot. The preservation ofoxygen-isotopic anomalies shows that there werenumerous oxygen reservoirs for the manufactureof chondrules and CAIs that were quite separate.

The sourceof theanomalieshasbeenattributedto nucleosynthetic differences between batchesof presolar materials (Wasson, 2000), mass-independent partitioning among oxygen-bearingmolecules in the nebular gas (Thiemens, 1999),and photochemical self-shielding of CO at the in-ner edgeof the solarnebula (Clayton, 2002), in thepresolar molecular cloud (Yurimoto and Kura-moto, 2002),or in the surfaceof thenebula (Lyonsand Young, 2005). Whatever the source, the 16Ovariations in CAIs and chondrules imply that theanomalies can be understood in terms of mixingof 16O-rich and 16O-poor components in the solarnebula. The anomalies could have been preservedby enriching 16O-rich dust relative to gas prior tovaporization (Cassen, 2001b; Scott and Krot,2001). However, in the self-shielding model, the16O-poor component was anomalously heavy iceand the nebula was initially 16O-rich, like theprotosun. Luck et al. (2003) inferred that 16Oexcesses were correlated with 63Cu-isotopicanomalies, but the reason for this is not clear. Itis hoped that the NASA Genesis mission willdetermine the oxygen-isotopic compositionof the

Metamorphism, Alteration, and Impact Processing 11

solar wind and establish the origin of the oxygenanomalies (Wiens et al., 1999; Clayton, 2002).

1.07.4 METAMORPHISM, ALTERATION,AND IMPACT PROCESSING

All chondrites were modified in some wayby geological processes in asteroids operatingover 4.5Gyr. If we want to understand howthe components in chondrites were formed, wemust understand how chondritic materials weremodified in asteroids. Three processes affectedchondrites: aqueous and hydrothermal altera-tion, thermal metamorphism, and impacts.

The most important heat source for thealteration and metamorphism of chondriteswas heat from the decay of 26Al, and to a muchlesser extent, 60Fe (Chapter 1.16). If chondritescome from the first generation of planet-esimals that formed, one could exclude 26Alas a heat source for melting planetesimals(Kunihiro et al., 2004a). However, 182Hf–182Wdating and thermal modeling show that anearlier generation of bodies accreted o0.5Myrafter CAIs formed and melted due to the short-lived radioistopes forming cores that are thesources of most iron meteorites (Kleine et al.,2005; Bizzarro et al., 2005; Schersten et al., 2006).Evidence that 26Al was adequately mixed in thesolar nebula comes from the consistent chronol-ogy provided by 26Al–26Mg and Pb–Pb dating(Zinner and Gopel, 2002; Kita et al., 2005).

Impacts also provide kinetic energy but forasteroids smaller than a few hundred kilome-ters in size, they are not an effective heat sourcefor global metamorphism (Keil et al., 1997).For alternative views, see Rubin (1995, 2005),who favors heating by asteroidal impacts, andKurat (1988), who argues for nebular heatingprocesses. Electrical induction heating of aster-oids by plasma outflow from the protosun isnot currently favored (see Scott et al., 1989).Arguments against the concept of hot accretionwithout radiometric heating (Hutchison et al.,1979) were given by Haack et al. (1992) andRubin and Brearley (1996).

Alteration and metamorphism used to beconsidered as separate processes: the formeraffecting carbonaceous bodies and the latteroperating on enstatite and ordinary chondrites.However, it is now clear that fluids of somekind were present during thermal processing ofvirtually all chondrite classes.

1.07.4.1 Petrologic Types

Chondrites are classified into petrologic(or petrographic) types 1–6 to provide a guide

to the extent and nature of asteroidal process-ing. This classification scheme, which has beeninvaluable (Anders and Kerridge, 1988), never-theless provides only a very rough guide as ourideas about asteroidal processes have changedsignificantly since the scheme was first estab-lished, and because one number cannot sum-marize all the effects of complex asteroidalprocessing. As a result, the classification isnot rigorous and may be ambiguous or evenmisleading.

The original division of chondrites by VanSchmus and Wood (1967) into six petrologictypes was based on two underlying assump-tions that are now known to be incorrect:(1) that type 1 chondrites are the least alteredchondrites that best represent the mineralogyof the original materials that accreted intoasteroids, and (2) that types 2–6 representincreasing degrees of thermal metamorphism ofmaterial that mineralogically resembled type 1chondrites. Although CI1 chondrites are‘‘chemically’’ most pristine as they are compo-sitionally similar to the solar composition,almost all of their ‘‘minerals’’ formed duringaqueous alteration. Type 3 chondrites are clos-est mineralogically to the original materialsthat accreted into asteroids (McSween, 1979).

An additional problem with the division intosix petrologic types is that it ignores the roleplayed by impacts in modifying and mixingchondritic material. Many chondrites are brec-cias of materials with diverse alteration ormetamorphic histories. Thus the petrologictype may provide no information about themetamorphic history of the whole chondrite,only an approximate guide to the history ofmost constituents.

The criteria used for assigning petrologictypes were first given by Van Schmus and Wood(1967) and have not changed significantly since(table 3 of Chapter 1.05). Table 2 shows thenumbers of classified chondrites in petrologictypes 1–6 within each group. The unrepresentedcells in Table 2 may reflect poor sampling (e.g.,in the poorly represented K group). However inmany cases, it is likely that the missing typeswere volumetrically insignificant. It is very un-likely, for example, that the ordinary chondriteparent bodies ever contained significantamounts of type 1 and 2 material.

1.07.4.1.1 Type 1 and 2 chondrites

There is much ambiguity about the classifi-cation of type 1 and 2 chondrites that are notclearly associated with the major groups, CIand CM. Different workers favor different cri-teria, so that type 1, for example, may mean a

Table 2 Numbers of classified chondrites in petrologic types 1–6 by group.

1 2 3 4 5 6 Total no.

CarbonaceousCI 5 5CM 4 44 48CO 31 31CV 1 35 36CR 1 13 1 15CH 7 7CBa 3 3CBb 2 2CK 2 13 6 1 23

Ordinarya

H 187 1,371 3,319 1,784 6,661L 316 415 1,220 4,053 6,004LL 71 64 419 406 960

EnstatiteEH 18 9 6 2 35EL 8 0 2 19 29

OtherK 2 2R 3 2 14

9 R3–6

Sources: Based largely on lists in Grady (2000) with additional data from Krot et al. (2002a) for CR, CH, and CB and Kallemeyn et al. (1996)and Bischoff (2000) for R chondrites. A small number of chondrites that are classified as intermediate types (e.g., 3–4 and 4–5) and mixedbreccias (e.g., types 3–5) were omitted except for the R chondrites, which are mostly breccias of type 3–5 or 3–6 material.aNumbers of H, L, and LL chondrites are italicized as they have not been corrected for probable pairings: the number of individual meteoritesis probably B2–4 times smaller than the number of specimens.


chondrite with major minerals similar to thosein CI1, or carbon-rich like CI1, or simply thatall the chondrules have been almost completelyaltered. Since the original CI and CM chond-rites were studied (see Zolensky and McSween,1988), CM1 chondrites have been described(Zolensky et al., 1997), as well as several CI orCM chondrites that were metamorphosed afteralteration but have not been assigned a pet-rologic type (B&J, 1998, pp. 70–71, 240–244).

CR2 chondrites, which contain abundantFe,Ni, are the least altered type 1 and 2 chond-rites and offer exceptional insights into thenature of nebular components. Renazzo wasthe first to be recognized, but other CR2s in-cluding PCA 91082 (Figure 1a) appear to bethe least altered (Wood, 1967; Krot et al.,2002c). GRO 95577, a type 1 chondrite, is themost altered CR (Weisberg and Prinz, 2000).Since there are only two known CR1 and CR3chondrites, the term ‘‘CR chondrite’’ is com-monly used as a synonym for CR2. Similarly,CM is used for CM2.

1.07.4.1.2 Type 3 chondrites

Type 3 chondrites show very diverse miner-alogies and in several groups they have beensubdivided. CO3, H3, L3, and LL3 chondrites,

which are divided into subtypes 3.0–3.9, showcorrelated variations in mineralogy and concen-trations of carbon and highly volatile elementsthat can largely be attributed to metamorphism.Subtypes were first defined for ordinary chond-rites based largely on their thermoluminescencesensitivity, which is a measure of plagioclasegrowth in chondrule mesostases (Sears andHasan, 1987; Sears et al., 1991, 1995a). Manyother properties vary systematically withprogressive metamorphism and have been usedto assign subtypes, for example, chemicalchanges in chondrule olivines (Figure 4), thechemical homogenization of olivines, FeO/(FeOþMgO) ratios in matrices (Huss et al.,1981), loss of trapped noble gases and loss ofcarbon (Sears and Hasan, 1987), chromiumlevels in metallic Fe,Ni grains (Scott and Jones,1990), and growth of ferroan olivine in amoe-boid–olivine aggregates (Chizmadia et al.,2002). Especially in low-subtype chondrites,there is evidence that hydrous fluids promotedchemical equilibration and alteration in chond-rule mesostases (B&J, 1998; Tomeoka and Itoh,2004), CAIs (Russell et al., 1998), and matrices(Section 1.07.5.7). With increasing metamor-phism, fluids were lost from the parent bodies.For CV3 chondrites, the effects of metamorphicalteration are more complex: subtypes based onthermoluminescence sensitivities appear to be

CaO

(w

t.%)

0.5

0.4

0.3

0.2

0.1

0

FeO (wt.%)

0 5 10 15 20 25 30

4−6

3.3−3.4

3.23.0−

3.5−3.6

3.0−3.2

3.3−3.43.7−3.9

3.5− 3.6

10 20 30Fa (mol.%)

LL chondrites

Type IA chondrules

Type IIA chondrules

3.7−3.9

Figure 4 Plot showing effects of metamorphism onthe mean CaO and FeO concentrations in olivine intwo types of chondrules in LL3 chondrites (Scottet al., 1994). Type 3 chondrites can be readily sub-divided into subtypes using these parameters ormany others (e.g., Sears and Hasan, 1987). Repro-duced by permission of Elsevier from Scott et al.

(1994).

Metamorphism, Alteration, and Impact Processing 13

less useful (Guimon et al., 1995; Krot et al.,1995), and the degree of crystallinity of carbon-aceous material appears to be a more reliableguide (Bonal et al., 2006). The latter parametersuggests that Allende was heated above type 3.6levels and should no longer be considered as apristine chondrite.

Although type 3 chondrites are the leastaltered and metamorphosed chondrites, only afewof themare now thought to be relatively pris-tine samples, and all of these have componentsthat were tainted by their asteroidal environm-ent. Among the most pristine are the CBb3 andCH3 chondrites, Vigarano and Leoville (CV3),ALHA77307 (CO3.0), Semarkona (LL3.0),Kakangari (K3), and two ungrouped carbon-aceous chondrites, Acfer 094 and Adelaide. Toaid in identifying the least metamorphosedchondrites, subtypes 3.0 and 3.1 have beendivided into 3.00 to 3.15 on the basis of Cr2Oconcentrations in chondrule olivines thatcontain42 wt.%FeO (Grossman andBrearley,2005). These data confirm that Acfer 094 andALHA77307 are among the least metamor-phosed chondrites. Chondrules in type 3 chond-rites in the ordinary, K, CO, and CV groupswere probably heated to around 400–600 1Cfor 4106 years, as they contain kamacite andtaenite that equilibrated at these temperatures(Keil, 2000).However, the CH3 andCB3 chond-rites contain martensite that has not exsolved tokamacite and taenite and has retained chemicalzoning patterns that were established prior toaccretion (Meibom et al., 1999, 2000; Krot et al.,2002c). Electron microscopic studies of themetal show that CH3 chondrites were not

heated above 300 1C for more than a year andprobably experienced much lower peak temper-atures (Reisener et al., 2000).

1.07.4.1.3 Type 4–6 chondrites

Type 4–6 ordinary chondrites, which arealso called equilibrated chondrites, are thoughtto represent an isochemical, metamorphicsequence: types 4 and 5 were heated to 500–800 1C and type 6 to 800–1,000 1C (see Keil,2000). However, Wlotzka (2005) found similarolivine/Cr-spinel temperatures of 700–800 1C fortype 4–6 ordinary chondrites and suggests thatthese temperatures correspond to peak meta-morphism. The most noticeable metamorphiceffects are due to grain coarsening and growthof feldspar, chromite and phosphate grains.Some type 4–6 chondrites are not simple meta-morphic rocks: they are fragmental or regolithbreccias of rock fragments that were metamor-phosed prior to assembly (Rubin, 1990). In thesecases, the petrologic type may only be a measureof the metamorphic history of most material inthe rock. In addition, Rubin (2002, 2003, 2004)infers that practically all type 5 and 6 ordinarychondrites were shocked and later annealed andthat many experienced multiple episodes ofshock and impact-induced annealing. Withineach group of ordinary chondrites, there aresmall but systematic increases in the FeO con-centrations of olivines and pyroxenes with inc-reasing petrologic. These are probably due tooxidation of metal by traces of water vaporduring metamorphism (McSween and Labotka,1993). Fluid inclusions are present in halite intwo H chondrite regolith breccias and may havebeen derived from asteroidal fluids or icy pro-jectiles (see Rubin et al., 2002).

The criteria for classifying types 4–6 werebased solely on the properties of H, L, and LLtypes 3–6 chondrites and most cannot be usedfor asteroidally processed carbonaceous andenstatite chondrites. For these chondrites, thetype is based almost entirely on the degree towhich the chondrules are delineated. If thechondrules have been obscured by recrystalli-zation during prolonged heating, the petrologictype is a good measure of the metamorphichistory. However, for the EH4–6 and EL4–6chondrites, evidence for simple thermal meta-morphism is lacking. Rapid cooling ratesrecorded by sulfides and numerous examplesof impact heating suggest that impact process-ing may have been more important for enstatitechondrites than heating by decay of 26Al(Zhang et al., 1996; Rubin et al., 1997; Linand Kimura, 1998). R chondrites experiencedubiquitous brecciation and although some have


been classified as R4 (or R3), all are probablybreccias of type 3–5 or 3–6 material (Bischoff,2000).

1.07.4.2 Thermal History and Modeling

To understand geological processes onchondritic bodies and to estimate their sizeswe need to investigate the thermal history ofunshocked or only lightly shocked samples. Thesimplest metamorphosed chondrites are theordinary chondrites, for which there are threekinds of constraints: (1) mineral equilibrationtemperatures (e.g., McSween and Patchen,1989); (2) radiometric ages that date coolingbelow isotopic closure in minerals usingisotope systems such as 207Pb–206Pb, 40K–39Ar,and 87Rb–87Sr (e.g., Trieloff et al., 2003);and (3) cooling-rate determinations based onannealing of Pu fission tracks (Lipschutz et al.,1989) and metallographic cooling rates (Tayloret al., 1987). The cooling rates determinedby the metallographic technique, which aremostly 1–103 1CMyr�1 in the temperature rangeof 400–500 1C, are compatible with cooling ratesin the range 100–500 1C determined by thefission track technique. A third technique basedon Fe–Mg ordering in orthopyroxene givescooling rates in the range 340–480 1C thatare systematically higher by several orders ofmagnitude (Folco et al., 1997), probably be-cause metamorphism was insufficient to reequil-ibrate the ordering imposed during chondruleformation (Artioli and Davoli, 1994). Coolingrates can also be estimated from radiometric agedata for mineral grains that close isotopically atdifferent temperatures (Bogard, 1995; Gangulyand Tirone, 2001; Amelin et al., 2005).

Metamorphic ages of ordinary chondritesfrom Pb–Pb ages of phosphates in seven H4–6chondrites range from 4.50 to 4.56Ga (Gopelet al., 1994), while their Ar–Ar ages range from4.45 to 4.53Ga (Trieloff et al., 2003). The twoH4 chondrites, which have metallographiccooling rates of 103 1CMyr�1, have the oldestages while the H6 chondrites with cooling ratesof B10 1CMyr�1 have younger ages. Thenegative correlation between age and meta-morphic type for these seven chondrites iscompatible with an onion-shell model in whichburial depth correlates with petrologic type.However, Pb–Pb ages for two ordinary type 3chondrites are significantly younger than thephosphate ages for the H6 chondrites and arenot readily reconciled with this model (Gopelet al., 1994), and the selected H4 chondritesmay not be representative (Wlotzka, 2005).Metallographic cooling rates for larger num-bers of unshocked, ordinary chondrites show

no correlation between cooling rate and pet-rologic type (Taylor et al., 1987) suggestingthat if the petrologic types were once arrangedsequentially in layers, the bodies must havebeen catastrophically fragmented and reassem-bled by impacts during cooling from peakmetamorphic temperatures (Grimm, 1985).Regolith breccias show extreme variations inthe cooling rates of their metal grains from 1 to103 1CMyr�1 implying that the ordinary chond-rite bodies were scrambled after metamorphismso that material from all depths could be com-bined together (Williams et al., 1999).

Estimates of the sizes of the ordinary chond-rite parent bodies can be derived from thermalmodels as the timescale for cooling is controlledby conduction for dry asteroids 410–20 km insize (McSween et al., 2002). However, the ther-mal conductivity is very sensitive to the poros-ity, which can be reduced by sintering andincreased by impacts. Layers of regolith o1 kmthick can drastically reduce the cooling rate ofasteroids and increase near-surface thermalgradients (Haack et al., 1990). Akridge et al.(1998) argue that H chondrites are derivedfrom depths of o10 km from a 100 km radiusasteroid and that material at greater depthshas not been sampled. However, Bennett andMcSween (1996) infer complete sampling ofbodies with radii of B80–95 km for the H andL chondrite parent bodies.

For volatile-rich carbonaceous chondrites likeCI and CM chondrites, constraints on thermalhistories are derived from carbonate ages andoxygen-isotopic data. The former indicate thatalteration began soon after CAI formation andlasted B20Myr (Endress et al., 1996; Brearleyet al., 2001). Oxygen-isotopic compositions ofcarbonates provide model-dependent tempera-tures under 50 1C (see McSween et al., 2002).Modeling metamorphism of wet asteroidsrequires more complex models to track fluidflow, chemical and mineralogical changes duringmetamorphism, and exothermic serpentinizat-ion reactions (e.g., Wilson et al., 1999; Cohenand Coker, 2000; Young, 2001; Young et al.,2003). McSween et al. (2002) suggest that ther-mal buffering by ice and water prevented high-temperature metamorphism in carbonaceouschondrite bodies. However, some CV and CKchondrites were clearly heated above 500 1C.

1.07.4.3 Impact Processing of ChondriticAsteroids

Impacts fragmented, mixed, modified, melted,devolatilized, and lithified chondritic material.Asteroids were impacted throughout their his-tory: when they accreted, during alteration and

Chondritic Components 15

metamorphism, and during the 4.5Ga aftertheir parent bodies were heated. For reviews ofimpact processing of ordinary chondrites, seeStoffler et al. (1991); carbonaceous chondrites,Scott et al. (1992); and enstatite chondrites,Rubin et al. (1997). Experimental studies ofimpact metamorphism of ordinary chondriteswere made by Schmitt (2000) and Langenhorstet al. (2002), CV chondrites by Nakamura et al.(2000), and CM chondrites by Tomeoka et al.(1999). Meteorite evidence for impact proc-essing of chondritic and other asteroids wasreviewed by Keil et al. (1994) and Scott(2002). High-pressure minerals in shock veinsin chondrites also provide clues to mineralstability at depth in the Earth (Stoffler, 1997).

Advocates for impact heating of chondritespoint to Portales Valley, a breccia of H6 clastsembedded in a matrix of once-molten metallicFe,Ni, which cooled slowly to form a Wid-manstatten pattern (Rubin et al., 2001). Gaffeyand Gilbert (1998) suggest that molten metal onthe H parent body was derived from impact-formed melt sheets or residues of metal-richprojectiles and that the IIE irons were alsoderived from this melt (see section 1.12.2.5).However, the origin of Portales Valley and thelink with IIE irons remain obscure as asteroidsare thought to be too small for impacts to gene-rate melt sheets (Keil et al., 1997), and PortalesValley may have been heated radiogenicallyprior to the impact (Ruzicka et al., 2005).

1.07.5 CHONDRITIC COMPONENTS

Below we review the geochemical propertiesand possible origins of chondritic componentsfocusing on what appear to be primary featuresin the least altered type 2–3 chondrites. Therich diversity of chondritic components even inunaltered chondrites and the vast differencesbetween chondrite groups are key features ofchondrites. Several comprehensive reviews ofchondrites and their components have beenpublished since Kerridge and Matthews (1988).These include a comprehensive review ofchondrite mineralogy (B&J, 1998), edited bookson chondrules (Hewins et al., 1996), chondrites(Krot et al., 2005c), and constraints on earlysolar system processes (Lauretta and McSween,2006), and several reviews on chondrules(Hewins, 1997; Rubin, 2000; Jones et al.,2000a; Zanda, 2004; Connolly and Desch,2004) and CAIs (Ireland and Fegley, 2000;Chapter 1.08).

First we review CAIs and their forsterite-richaccretionary rims, then amoeboid olivine in-clusions and aluminum-rich chondrules, whichare intermediate in composition between CAIs

and chondrules, and finally matrix, which is acomplex mixture of many ingredients.Although chondrules are volumetrically moreimportant, we start with CAIs as these formedfirst and their origins are better constrainedby chemical and isotopic data. Chemical andisotopic variations in chondrules are muchmore modest than those in CAIs but there areenough related features and components withintermediate compositions to suggest thatchondrules cannot be understood in isolationfrom CAIs and other components.

1.07.5.1 Calcium- and Aluminum-RichInclusions

There are two types of refractory inclusions:calcium- and aluminum-rich inclusions (thissection) and amoeboid olivine aggregates (Sec-tion 1.07.5.3). Since the mineralogy, chemistry,and isotope chemistry of refractory inclusionswere reviewed by MacPherson et al. (1988),many new analyses have been made of CAIs inCV, CM, CO, CR, CH, CB, ordinary, andenstatite chondrites that provide importantconstraints on physicochemical conditions,time, and place of CAI formation. CAIs areaddressed in detail in Chapter 1.08, the role ofcondensation and evaporation in their forma-tion in Chapter 1.15, and their clues to earlysolar system chronology in Chapter 1.16.

CAIs are the oldest objects in chondritesexcluding presolar grains (Amelin et al., 2002)and probably formed during the most energeticphase of protosolar disk evolution (e.g., Irelandand Fegley, 2000; Wood, 2000b, 2004). Theyare composed of the minerals corundum, hibo-nite, grossite, perovskite, melilite, spinel,Al–Ti-diopside, anorthite, and forsterite, whichare predicted to condense from a coolinggas of solar composition at temperatures41,200–1,300K and total pressure of 10�5 bar(e.g., Ireland and Fegley, 2000; Ebel and Gross-man, 2000). Some CAIs are thought to becondensates because they have irregularshapes, fluffy textures, and minerals that matchthe sequence of mineral formation (Figure 5),and trace-element chemistry (e.g., Simon et al.,2002). The strongest argument in favor of con-densation origin of some CAIs comes fromtheir group II REE pattern, which is a highlyfractionated pattern with much lower abun-dances of the heavy REEs Gd–Er and Lu,exhibited by B30% of all CAIs (Palme andBoynton, 1993; Ireland and Fegley, 2000). Thisvolatility-controlled pattern requires the priorremoval of the ultrarefractory elements. How-ever, the preservation of large isotopic anom-alies (e.g., 48Ca, 50Ti) in platy hibonites

Adelaide

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Figure 5 Backscattered electron images of refractoryinclusions with clear condensation signatures: (a) CAIin Adelaide (unique carbonaceous chondrite), (b) CAIin Efremovka (CV3), and (c) AOA in Acfer 094(unique carbonaceous chondrite). (a) Corundum (cor)is replaced by hibonite (hib), which is corroded bygrossite (grs). Reproduced by permission of Universityof Arizona, r The Meteoritical Society from Krotet al. (2001a). The inferred sequence of mineral crys-tallization is consistent with the predicted equilibriumcondensation sequence at total pressure of r10�5 barand CI chondritic dust/gas enrichment relative to solarcomposition of B1,000 (Simon et al., 2002). (b) Re-placement of melilite by a fine-grained mixture ofspinel and diopside, suggesting the following gas–solidreaction: Ca2Al2SiO7 þ 3SiO(g)þ 3Mg(g)þ 6H2O(g)¼ 2CaMgSi2O6þMgAl2O4þ 6H2(g). After Krotet al. (2004c). (c) Replacement of forsterite by low-calcium pyroxene (px), suggesting the following reac-tion Mg2SiO4þ SiO(g)þH2O(g)¼Mg2Si2O6þH2(g).Adapted by permission of Elsevier from Krot et al.

(2004e).


(PLACs) in CM chondrites and their lack of26Al, plus mass fractionation effects andsmaller nuclear anomalies in FUN inclusions(see Chapter 1.08), favor evaporation as thepredominant process for other CAIs. CAIsmay represent mixtures of solar nebula con-densates and refractory residue materialsformed by evaporating material of presolarorigin (Ireland and Fegley, 2000).

Because of the diversity of CAIs, there is nouniversal classification scheme for all groups.The best studied CAIs from CV chondrites aredivided into type A (melilite–spinel-rich), typeB (composed of pyroxene, melilite, spinel,7plagioclase), forsterite-bearing type B, typeC (melilite-poor, pyroxene–plagioclase-rich),and fine-grained spinel-rich (e.g., MacPhersonet al., 1988). Type A CAIs are subdivided intofluffy and compact subtypes. In other chondritegroups, CAIs are typically classified accordingto their modal mineralogy (e.g., corundum-,grossite-, and hibonite-rich). In CV chondrites,most coarse-grained CAIs were probably oncemolten. However, in other chondrite groups,CAIs were probably not molten, though defi-nitive compositional data for comparison withmelt–solid distribution coefficients are typicallynot available.

1.07.5.1.1 Comparison of CAIs from differentchondrite groups

A comparison of CAIs from carbonaceous,ordinary and enstatite groups (Table 3) showsthat most inclusion types can be found in manychondrite groups, and that the most commonvarieties of CAIs are spinel-, pyroxene-, andmelilite-rich (e.g., Lin et al., 2006). However,the relative proportions of CAI types and theirmean sizes, which are correlated with chond-rule size (May et al., 1999), differ among thechondrite groups. For example, type B CAIsare almost exclusively found in CV chondrites;PLACs occur mostly in CM chondrites; gros-site-rich, hibonite-rich, and aluminum-diopsidespherules are common only in CH chondrites;CAIs in enstatite chondrites are rich in spineland, in many cases, hibonite. There are alsomineralogical and isotopic differences betweenCAIs of the same mineralogical type: forexample, (1) melilite in compact type A CAIsfrom CV chondrites is more akermanitic thanthat in compact melilite-rich CAIs from COand CR chondrites (Weber and Bischoff, 1997;Russell et al., 1998; Aleon et al., 2002a) and (2)the grossite-rich and hibonite-rich CAIs in CRchondrites are 16O-rich and have a canonical

Table 3 Distribution of types of CAIs, AOAs, accretionary rims, and aluminum-rich chondrules in type 2and 3 chondrites.

Object CM CO CV CR CH CB CK Acfer 094 Adelaide H–L–LL EH–EL

Melilite-rich (type A) r c c c c c — c c r —Al,Ti-px-rich (type B) — — c r — r r — — — —Fo-type B CAIs — — r — — r — — — — —Plagioclase-pyroxene-rich (type C) — r c r — — — r — — —Hibonite7spinel-rich c c c r c c — r c r rGrossite-rich — r — r c r — c — — —Spinel7pyroxene-rich c c c c c c r c c r rPyx-hibonite spherules r r — — c c — r — — rFo-rich accretionary rims — — c — — — — — — — —AOAs c c c c c r c c c r —Aluminum-rich chondrules c c c c c c c c c c c

Source: MacPherson et al. (1988) with additional data: CHs, Kimura et al. (1993); COs, Russell et al. (1998); CMs, MacPherson et al. (1984),Simon et al. (1994); CRs, Aleon et al., (2002a), Weisberg et al. (1993), Kallemeyn et al. (1994); CKs, Weber and Bischoff (1997), McSween(1977a), MacPherson and Delaney (1985), Kallemeyn et al. (1991), Keller et al. (1992), Noguchi, (1993); H-L-LLs, Bischoff and Keil (1984),Russell et al. (2000), Kimura et al. (2002), Itoh et al. (2004a); ECs, Bischoff et al. (1985), Guan et al. (2000), Fagan et al. (2000); CBs, Krotet al. (2001a); Acfer 094, Adelaide, Krot (unpublished).Note: r, rare; c, common; —, absent.


26Al/27Al ratio (Marhas et al., 2002; Aleonet al., 2002a), whereas those in CH chondritesare 16O-rich, but generally lack radiogenic 26Mg(26Mg*) (MacPherson et al., 1989; Kimura et al.,1993; Weber et al., 1995; McKeegan et al.,2000a). The grossite- and hibonite-rich CAIs inCB chondrites are 16O-poor (Krot et al., 2001a);no Al–Mg-isotopic measurements have beenpublished yet.

1.07.5.1.2 Oxygen-isotopic compositions ofCAIs

In situ ion microprobe analyses of oxygen-isotopic compositions of individual minerals inCAIs in the chondrite groups provide very im-portant constraints on the origin of CAIs. CAIsfrom CV chondrites typically exhibit oxygen-isotope heterogeneity within an individualinclusion, with spinel, hibonite, and pyroxeneenriched and anorthite and melilite depleted in16O. At the same time, several coarse-grainedigneous CAIs containing 16O-rich and 16O-poormelilite, anorthite and Al–Ti-diopside, and Al–Ti-diopside with oxygen-isotope zoning havebeen described in the Allende and Efremovkameteorites (e.g., Kim et al., 1998; Yurimotoet al., 1998; Ito et al., 1999; Imai and Yurimoto,2000). Based on these observations and exper-imental data on oxygen-isotope self-diffusion inCAI minerals (Ryerson and McKeegan, 1994),Yurimoto et al. (1998) concluded that coarse-grained CAIs in CV chondrites wereinitially 16O-rich and subsequently experiencedoxygen-isotope exchange with 16O-poor nebulargas while partly molten during reheating. A few

CAIs in CV chondrites appear to have experi-enced multiple heating events in nebularenvironments that fluctuated from 16O-rich, to16O-poor, and back to 16O-rich (Yoshitakeet al., 2005). 16O-poor compositions of second-ary minerals (nepheline, sodalite, andradite, andhedenbergite) in CV CAIs (Cosarinsky et al.,2003) and oxygen-isotope heterogeneity in fine-grained spinel-rich CAIs that were probably notmelted suggest that some oxygen-isotopic he-terogeneity resulted from secondary alterationin asteroids (Aleon et al., 2002b; Fagan et al.,2002).

CAIs from CO chondrites (Figure 6) show acorrelation between oxygen-isotopic composi-tion and degree of thermal metamorphism:spinel, hibonite, melilite, diopside, anorthite,and olivine in most CAIs from CO3.0 chon-drites Colony and Y-81020 are uniformly16O-enriched, whereas melilite and secondarynepheline in CAIs from metamorphosed COchondrites Kainsaz (3.2) and Ornans (3.2) are16O-depleted relative to spinel and diopside.These observations favor oxygen-isotopeexchange during fluid–rock interaction in anasteroidal setting (Itoh et al., 2004a, b; Wassonet al., 2001). However, the 16O-poor composi-tions of some CAIs in CO3.0 chondrites (e.g.,Yurimoto et al., 2001) suggest that someisotopic exchange also occurred in the solarnebula in the presence of 16O-poor gas.

CAIs in type 2 and 3 CR, CH, CB, CO, CM,ordinary, and enstatite chondrites are typicallyisotopically uniform (within 3–4%): both16O-rich and 16O-poor CAIs have been found.Most CAIs in CR and CH chondrites are16O-rich (Figure 7); the 16O-poor CAIs in CR,

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Figure 6 Oxygen-isotopic compositions ofindividual minerals in CAIs from the CO chondritesY-81020, Colony, Kainsaz, and Ornans. Primaryminerals in CAIs from the least metamorphosed COchondrites Y-81020 (type 3.0) and Colony (3.0) areuniformly 16O-enriched, whereas CAIs from Kainsaz(3.2) and Ornans (3.3) tend to show oxygen-isotopicheterogeneity with spinel and high-calcium pyroxeneenriched in 16O and melilite and secondary nephelinedepleted in 16O. Based on these observations,Wasson et al. (2001) inferred that oxygen-isotopeexchange took place during thermal metamorphismand alteration in an asteroid. Data from Itoh et al.

(2004b) and Wasson et al. (2001).


CH, and CB chondrites show petrographicevidence for extensive melting (Aleon et al.,2002a; Krot et al., 2001a; McKeegan et al.,2006). CAIs in CM, ordinary, and enstatitechondrites are 16O-rich (D17Oo�20%). InCM chondrites, hibonites in spinel–hibonitespherules (SHIBs), which generally lack stable-isotopic anomalies in 48Ca and 49,50Ti, have

canonical (26Al/27Al)0 and (41Ca/40Ca)0 ratiosand group II REE patterns, whereas platyhibonites (PLACs), which generally have groupIII REE patterns, large (410%) stable-isotopicanomalies in 48Ca and 50Ti, and lack 26Al and41Ca, are similarly 16O-enriched. No obviouscorrelation between hibonite morphology andradiogenic 26Mg, Ca (d48Ca) or Ti (d50Ti)isotopic anomalies has been found (Sahijpalet al., 2000; Goswami et al., 2001).

To summarize, variations in oxygen-isotopiccompositions within a single CAI in CV andCO chondrites require isotopic exchange,however, the exact mechanism of exchangeremains poorly understood. Proposed mecha-nisms include gas–solid and/or gas–melt ex-change in the solar nebula (e.g., Clayton, 1993;Yurimoto et al., 1998, 2000, 2001; Krot et al.,2005c), and solid–fluid exchange in an aster-oidal setting (e.g., Wasson et al., 2001; Faganet al., 2002, 2004). The presence of 16O-poorCAIs with igneous textures in primitive CR,CH, and CB chondrites suggests that in thesecases oxygen-isotopic exchange occurred dur-ing melting. It remains unclear whether thisexchange occurred in the CAI- or chondrule-forming regions and whether this exchange waslate or early.

The oxygen-isotopic compositions ofchondritic components are generally assumedto reflect isotopic variations among the differ-ent locations within the meteorite-formingregions of the solar nebula. However, sinceCAIs in most chondrite groups have similar16O-rich isotopic compositions, it was inferredthat CAIs formed in a single, restricted nebularlocale and were then unevenly distributed tothe regions where chondrites accreted after sor-ting by size and type (McKeegan et al., 1998;Guan et al., 2000). Although formation in asingle restricted region appears to be generallyaccepted, sorting by size and type cannotexplain the observed variations in mineralogyand isotope chemistry of CAIs in differentchondrite groups. We suggest instead that CAIformation occurred multiple times in the CAI-forming region under diverse physicochemicalconditions (e.g., dust/gas ratio, peak heatingtemperature, cooling rate, and number of re-cyclings) and that CAIs were subsequently iso-lated from the hot nebular gas at variousambient nebular temperatures. The isolationcould be due to removal by stellar wind (e.g.,Shu et al., 1996, 2001). The 16O-rich and 16O-poor CAIs reflect either temporal variations inoxygen-isotopic compositions of the nebulargas in the CAI-forming region(s) (e.g., Yuri-moto et al., 2001) or radial transport of someCAI precursors into an 16O-poor gas prior totheir melting (see also Scott and Krot, 2001;

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Figure 7 (a, b) Oxygen-isotopic compositions of individual minerals in CAIs from CR chondrites. Most CAIsare 16O-rich and isotopically homogeneous; CAIs, which show petrographic evidence for extensive melting areless 16O-rich. Secondary phyllosilicates replacing 16O-rich melilite plot along terrestrial fractionation line. Datafrom Aleon et al. (2002a). (c, d) Oxygen-isotopic compositions of individual minerals in CAIs from CH chond-rites show a bimodal distribution. The most refractory CAIs, which are composed of grossite, hibonite, spinel,perovskite, and melilite, are enriched in 16O; aluminum-diopside rims around these CAIs, pyroxene-spinel CAIs,and two melilite-rich CAIs surrounded by a forsterite layer are 16O-poor. Data from McKeegan et al. (2006).(e, f) Oxygen-isotopic compositions of individual minerals in CAIs from CB chondrites Hammadah al Hamra

237 and QUE 94411. The CAIs are rather homogeneous and 16O-poor. Data from Krot et al. (2001c).


Aleon et al., 2002a; Krot et al., 2005c;McKeegan et al., 2006).

1.07.5.2 Forsterite-Rich Accretionary Rimsaround CAIs

Olivine-rich, accretionary rims around CAIs,which have grain sizes 45–20mm, were first de-scribed in the altered CV chondrite Allende byMacPherson et al. (1985). In this meteorite, theaccretionary rims have four layers that differ intexture and mineralogy. The innermost layerconsists of either pyroxene needles, olivine,clumps of hedenbergite, and andradite or olivine‘‘doughnuts’’ (i.e., crystals with central cavities).The next two layers outward both containolivine plates and laths. The final layer separa-ting accretionary rims from the Allende matrixcontains clumps of andradite and hedenbergitesurrounded by salitic pyroxene needles. Nephe-line and Fe,Ni-sulfides are common constituents

in all layers. Based on the observed disequilib-rium mineral assemblages, abundant euhedralcrystals with pore space between them, and rimthicknesses controlled by the underlyingtopography of their host inclusions, MacPher-son et al. (1985) concluded that the layeredaccretionary rims in Allende are aggregates ofgas–solid condensates which reflect significantfluctuations in physicochemical conditions inthe solar nebula and grain/gas separation proc-esses. Hiyagon (1998) analyzed the oxygen-iso-tope compositions of minerals in accretionaryrims around two type A CAIs from Allende.Olivines in both accretionary rims are charac-terized by 16O-rich-isotopic compositions(d18O¼ –28%, d17O¼ –30%), whereas wollasto-nite, andradite, and hedenbergite show muchheavier oxygen-isotopic compositions (d18Ofrom þ 2% to þ 10%, d17O from –3% toþ 4%).In the reduced CV chondrites, which are less

altered than Allende, accretionary rims are

Efremovka, E38

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Figure 8 (a) Combined elemental map in magnesium (red), calcium (green), and Al Ka (blue) X-rays of a typeB CAI E38 from Efremovka with an accretionary rim (AR); (b)–(d) backscattered electron images fromregions marked in elemental map. This CAI consists of melilite (mel), aluminum, titanium-diopside (fas),spinel (sp), and anorthite (an) and is surrounded by a multilayered Wark–Lovering rim (WLR) composedof \spinel, perovskite (pv), anorthite, aluminum-diopside, and forsterite, and a thick outer forsterite-richaccretionary rim (red in (a)). The accretionary rim consists of forsterite (ol) and refractory components

composed of spinel, aluminum-diopside, and anorthite.


not layered and consist of coarse-grained(20–40 mm), anhedral forsterite (Fa1–8 versusFa5–50 in Allende), Fe,Ni-metal nodules, andfine-grained refractory components composedof aluminum-diopside, anorthite, spinel, and,in some cases, forsterite (Krot et al., 2001f,2002b). These rims surround different types ofCAIs: types A, B, and fine-grained spinel-rich(Figure 8). Forsterite grains in the accretionaryrims are 16O-enriched similar to spinel, alumi-num-diopside, and hibonite of the host CAIs(Figure 9).

Krot et al. (2001f, 2002b) concluded thatforsterite-rich rims formed by accretion of high-temperature condensates in the CAI-formingregions. The rims were emplaced after high-tem-perature events experienced by the CAIs, suchas melting and formation of Wark–Lovering

rims (see Wark and Boynton, 2001), which musttherefore have occurred in the CAI-formingregion as well. Andradite, wollastonite, he-denbergite, and fayalitic olivine in the Allenderims formed by alteration under oxidizing con-ditions in the presence of 16O-poor fluid (Krotet al., 1995, 1998a, b, c, 2001e).

1.07.5.3 Amoeboid Olivine Aggregates

AOAs are irregularly shaped objects with finegrain sizes (5–20mm) that occupy up to a fewpercent of type 2 and 3 carbonaceous chondrites(Table 3). AOAs are similar in size to co-accreted chondrules and CAIs being mostly100–500mm in size in CO (Chizmadia et al.,

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Figure 9 Oxygen-isotopic compositions of mineralsin the Efremovka CAI E104 (a) and Vigarano CAI1623-2 (b) and their forsterite-rich accretionary rims(AR) and Wark–Lovering rims (WLR) showing thatboth formed in an 16O-rich reservoir. Data from

Krot et al. (2002b).


2002), up to 5mm in size in CV (Komatsu et al.,2001), and typically o500mm in CR chon-drites (Aleon et al., 2002a). In the least alteredchondrites, they are porous aggregates andtheir mineralogy matches that expected forhigh-temperature nebular condensates (Krotet al., 2004d). Unlike CAIs and chondrules,AOAs do not appear to show mineralogical andisotopic differences between groups and thusprovide an excellent guide to the alteration his-tory of the meteorite and its constituents.

1.07.5.3.1 Mineralogy and petrology of AOAs

AOAs, which were first described inthe heavily altered Allende CV3 chondrite

(Grossman and Steele, 1976), are best observedin the least altered chondrites, CO3.0, CR2,CH, CB chondrites, Adelaide, and Acfer 094(Krot et al., 2004d and references therein).Here they consist of anhedral, fine-grained(1–20 mm) forsterite (Fao1–3), Fe,Ni-metal,and a refractory component composed of alu-minum-diopside, spinel, anorthite, and raremelilite, or CAIs of similar mineralogy. AOAsin CO3.0 chondrites contain B8 vol.% metallicFe,Ni: troilite is rare or absent (Chizmadiaet al., 2002). The refractory objects in AOAsare characterized by significant variations insize (1–250 mm), shape (irregular, rounded), dis-tribution (uniform, heterogeneous), modalmineralogy (spinel-rich, spinel-poor), and abun-dance (rare, abundant) (Figure 10). Forsteritegrains typically contain numerous pores and tinyinclusions of aluminum-diopside. However,some AOAs show triple junctions betweenforsterite grains and coarse-grained shells sur-rounding finer-grained cores, indicating high-temperature annealing (Komatsu et al., 2001).About 10% of AOAs contain low-calcium pyr-oxene replacing forsterite at grain boundariesin AOA peripheries (Figure 5c), or rimmingthe AOA (Krot et al., 2004e). Although mostAOAs show no clear evidence of being melted,those with low-Ca pyroxene rims are compactrather than porous and may have experi-enced small degrees of melting. Some AOAs ap-pear to contain CAIs that have been melted(Figure 10).

Forsterite grains in AOAs in pristine chond-rites show diverse compositions. Forsteritegrains in Adelaide AOAs that lack low-Capyroxene are nearly pure (FeOo1wt.%) withonly small amounts of Cr2O3 (o0.3wt.%) andMnO (o0.07 wt.%) whereas those in pyroxene-bearing AOAs are richer in FeO (1–4wt.%),Cr2O3 (0.2–0.6wt.%), and MnO (up to0.8wt.%) (Krot et al., 2004d). Mn and Cr con-centrations in AOAs in a CO3.0 chondrite(Yamato 81020) appear to be higher in frac-tured or peripheral regions and lowest in AOAswith high CaO concentrations (0.2–0.8wt.%)(Sugiura et al., 2004, 2006), This suggests thatboth Mn and Cr were introduced by gas–solidreactions after Ca-rich forsterite condensed,consistent with thermodynamic models (e.g.,Weisberg et al., 2004).

Secondary nepheline, sodalite and ferrousolivine are common in AOAs in the alteredand metamorphosed CV and CO chondrites,but are absent in AOAs from the unalteredchondrites. Olivines in AOAs from CV chond-rites are generally more fayalitic than inthe unaltered chondrites, with fayalite con-centrations increasing from Leoville and Vigar-ano to Efremovka to Allende (Figure 11).

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Figure 10 Combined elemental map in magnesium (red), calcium (green), and Al Ka (blue) X-rays (a) and BSEimages (b)–(d) of AOAs in the reduced CV chondrite Efremovka. They consist of fine-grained forsteritic olivines(Ol), Fe,Ni-metal (Met), and CAIs composed of aluminum-diopside (cpx), anorthite (an), spinel (sp), andrare melilite (mel); melilite is replaced by anorthite. Although AOAs show no clear evidence of being melted,some CAIs inside them appear to have experienced melting. Regions outlined in (a) and (b) are shown in detail

in (b)–(d).


This sequence correlates with the degree of sec-ondary alteration and thermal metamorphismexperienced by CV chondrites (e.g., Krot et al.,1995, 1998a, b, c, 2001e; Komatsu et al., 2001).In CO3 chondrites, fayalite contents are wellcorrelated with petrogic subtype (Chizmadiaet al., 2002).

1.07.5.3.2 Trace elements and isotopiccomposition of AOAs

Trace element concentrations have beenmeasured in 10 AOAs from Allende (Grossmanet al., 1979). Most AOAs have unfractionatedabundances of refractory lithophile and side-rophile elements (2–20�CI); two AOAs analy-zed have group II REE patterns.

Oxygen-isotopic analysis of primary min-erals (forsterite, aluminum-diopside, spinel,

and anorthite) in AOAs from CM, CR, CB,and unmetamorphosed CO chondrites areuniformly enriched in 16O, whereas anorthite,melilite, and secondary minerals (nepheline andsodalite) in AOAs from CVs and metamor-phosed COs are 16O-depleted to various degrees(Hiyagon and Hashimoto, 1999a; Imai andYurimoto, 2000, 2001; Itoh et al., 2002;Krot et al., 2001f, 2002b; Fagan et al., 2004)(Figure 12). Melilite grains in AOAs from thereduced CV3 chondrite Efremovka are 16O-enriched relative to those from Allende, sugges-ting exchange during alteration (Fagan et al.,2004). Imai and Yurimoto (2000, 2001)reported two generations of olivines in AllendeAOAs: primary forsteritic olivines, which are16O-rich (D17OB–20%), and secondary fayali-tic olivines, which are 16O-poor (D17OB–5%)and probably grew on the parent asteroid in thepresence of 16O-poor fluid. Grain-boundary

Freq

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)

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Figure 11 Compositions of olivine in AOAs from the reduced CV chondrites Vigarano (a), Leoville (b),Efremovka (c), oxidized CV chondrite Allende (d), CR chondrites (e), CH and CB chondrites (f), uniquecarbonaceous chondrites Adelaide (g), and Acfer 094 (h). Olivines in CRs, Adelaide, and Acfer 094 aremagnesium-rich compared to olivines in CV AOAs. Fayalite contents in olivines from CV AOAs increase inthe order Leoville and Vigarano, Efremovka, Allende; this is correlated with the degree of secondary alterationand thermal metamorphism experienced by CV chondrites. Data for Allende AOAs are from Hashimoto and

Grossman (1987); and for Efremovka, Leoville, and Vigarano AOAs from Komatsu et al. (2001).


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fo

sp

focpxansp

cpx

sp

nphphyl

an

ol

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(a) (b)

px

Figure 12 Ion microprobe analyses of oxygen-isotopic compositions of individual minerals in AOAs fromCR2 (a), CV3 (b), CM2 (c), and CO3 (d) chondrites. Key: fo, forsterit; cpx, Al-diopside; an, anorthite;sp, spinel; fas, Al–Ti-diopside; phyl, phyllosilicates; nph, nepheline; fo relict, relict AOA forsterite in chond-rules; pig, chondrule pigeonite. Data for CR AOAs from Aleon et al. (2002a); CV AOAs are from Imaiand Yurimoto (2000, 2001), Fagan et al. (2004), Hiyagon and Hashimoto (1999a), and Krot et al. (2002b);CM AOAs from Hiyagon and Hashimoto (1999b); CO AOAs from Itoh et al. (2002); CB AOA from Krot

et al. (2001c).


low-Ca pyroxene in AOA peripheries in prim-itive carbonaceous chondrites is 16O-rich(D17Oo–20% ) whereas low-Ca pyroxene igne-ous rims on AOAs are 16O-poor (D17OB�5% )(Krot et al., 2005b). Thus AOAs appear tohave been modified in two different nebularenvironments, in addition to the final, pre-sumably asteroidal environment where Na-richphases were introduced.

Three AOAs in the CO3.0 chondrite Y-81020that were studied for Al–Mg systematics showed

26Mg excesses corresponding to initial 26Al/27Al

ratios of (2.770.7)� 10�5, (2.971.5)� 10�5,and (3.271.5)� 10�5. Comparable excesses of26Mg were not detected in the AOA-bearingchondrule (Itoh et al., 2002). The lower26Al/27Al ratios in AOAs compared to a canon-ical value of B5� 10�5 found in most CAIs(e.g., MacPherson et al., 1995) suggest thateither AOAs formed 0.1–0.5Ma after CAIs orthat 26Al was heterogeneously distributed in theCAI-forming region (Itoh et al., 2002).


1.07.5.3.3 Origin of amoeboid olivineaggregates

The mineralogy, petrology, 16O-rich compo-sitions, depletion in moderately volatileelements, such as manganese, Cr, and Na,and the general absence of low-calcium pyrox-enes associated with pristine AOAs stronglysuggest that AOAs are aggregates of grains offorsterite, aluminum-diopside, spinel, anort-hite, and Fe,Ni-metal that condensed in thenebula from an 16O-rich gaseous reservoir andaggregated with refractory objects (Krot et al.,2004d, e, 2005b; Weisberg et al., 2004).Mineralogical and chemical similarities bet-ween AOAs and forsterite-rich accretionaryrims suggest that AOAs formed contempora-neously with these rims. Some refractoryobjects inside AOAs were melted prior toaggregation. AOAs subsequently experiencedhigh-temperature annealing and solid-staterecrystallization with only very minor meltingand nebular alteration resulting in formationof anorthite and, in some cases, low-calciumpyroxene. The AOAs formed in the CAI-forming region(s) and they were either largelyabsent from chondrule-forming region(s) whenchondrules formed, or the scales of the chond-rule-forming events were small and their fre-quency low. Their preservation can beattributed to removal from the CAI-formingregion prior to formation of low-calcium pyr-oxene by gas–solid reactions and/or kinetic in-hibition of these reactions. Nevertheless, AOAshave preserved a record of nebular processingin 16O-rich and 16O-poor environments andmay have been important precursor materialsin the formation of several kinds of igneous,coarse-grained chondritic components such asrefractory forsterite-rich objects and type Ichondrules.

1.07.5.4 Aluminum-Rich Chondrules

Aluminum-rich chondrules are a diversegroup of objects with igneous textures and410wt.% Al2O3 (Bischoff and Keil, 1984; Husset al., 2001; MacPherson and Huss, 2005; Guanet al., 2006), which appear to have distinctorigins from ferromagnesian chondrules (Sec-tion 1.07.5.5). Their mineralogy, bulk chemicalcompositions, oxygen, and Al–Mg isotopesystematics suggest a close link betweenCAIs and aluminum-rich chondrules. Usingthe CaO–MgO–Al2O3–SiO2 phase diagram(Figure 13), Huss et al. (2001) subdivided alu-minum-rich chondrules in ordinary chondrites

according to their bulk composition and miner-alogy into three groups: ‘‘plagioclase-rich,’’which plot in the anorthite (þ spinel) field onthe liquidus and contain abundant euhedral ca-lcic plagioclase crystals, ‘‘olivine-rich,’’ whichplot in the forsterite (þ spinel) field and com-monly contain large euhedral olivine pheno-crysts, and ‘‘glassy,’’ which plot near thecorundum–forsterite side of the ternaryand are characterized by abundant transparentNa–Al-rich and calcium-poor glass that enclosesa phenocryst assemblage of olivine, pyroxene,and spinel. In carbonaceous chondrites,however, Krot and Keil (2002) and Krotet al. (2002a) found a continuum between theplagioclase-rich and olivine-rich aluminum-rich chondrules: glassy aluminum-richchondrules are absent (Figure 14). Because al-uminum-rich chondrules in carbonaceouschondrites all have similar mineralogy (magne-sian low-calcium pyroxene and forsterite phe-nocrysts, Fe,Ni-metal nodules, interstitialanorthitic plagioclase, Al–Ti–Cr-rich low-calcium and high-calcium pyroxenes, and crys-talline mesostasis composed of silica, anorthite,and high-calcium pyroxene), Krot and Keil(2002) and Krot et al. (2002a) called them allplagioclase-rich chondrules, irrespective of mo-dal mineralogy.

Because the aluminum-rich chondrules arecompositionally most similar to CAIs andoften contain aluminum-rich phases with highAl/Mg ratios (plagioclase and, occasionally,glass) which make them suitable for Al–Mgisotope studies, these chondrules received spe-cial attention (e.g., Kring and Boynton, 1990;Sheng et al., 1991; Hutcheon and Jones, 1995;Russell et al., 1996, 2000; Hutcheon et al.,1994, 2000; Marhas et al., 2002; Srinivasanet al., 2000a, b; Huss et al., 2001; Krot andKeil, 2002; Krot et al., 2001b, 2002a, c; Guanet al., 2006).

1.07.5.4.1 Isotopic and trace element studies ofaluminum-rich chondrules

Oxygen-isotopic data for seven aluminum-richchondrules from ordinary chondrites (Russellet al., 2000) are continuous with the ordinarychondrite ferromagnesian chondrule fieldabove the terrestrial line, but extend it to more16O-enriched values (d18O¼ –15.771.8%,d17O¼ –13.572.6%) along a mixing lineof slope¼ 0.8370.09 (Figure 15a). Two chond-rules exhibit significant internal-isotopicheterogeneity indicative of partial exchange witha gaseous reservoir (Russell et al., 2000).

LaCa2SiO4 Mw Mo

Per+L

Mw+L

Mo+LMe+L

La+L

Fo+L

FoMg2SiO4

En

Pyx+L

Cord+LSaph+L

Mull+LAn

An+L

Cor+L

Hib+L

Grs+L

Geh

Hib Cor

Al2O3Al-rich chondrules

OCsCVsAdelaideCRsAOAs in Acfer 094, CRs

Figure 13 Bulk compositions of aluminum-rich chondrules plotted in the system Al2O3–Mg2SiO4–Ca2SiO4

and projected from spinel (MgAl2O4). Mineral abbreviations: An, anorthite; Grs, grossite; Cor, corundum;Cord, cordierite; Di, diopside; Fo, forsterite; Geh, gehlenite; Hib, hibonite; La, larnite; L, liquid; Mel, melilitesolid solution; Mull, mullite; Mo, monticellite; Mw, merwinite; Per, periclase; Pyx, pyroxene solid solution;Saph, sapphirine. Data from Bischoff and Keil (1984); Sheng et al. (1991); Huss et al. (2001); Krot and Keil

(2002); and Krot et al. (2002c, 2004a).


Porphyritic aluminum-rich chondrules are con-sistently 16O-rich relative to nonporphyriticones, suggesting that degree of melting was akey factor during nebular exchange (Russellet al., 2000).

Oxygen-isotopic compositions of alumi-num-rich chondrules from CR and CH chond-rites (Krot et al., 2006a) are plotted in Figure15b. Three out of six chondrules analyzed ex-hibit large internal-isotopic heterogeneity,whereas others are isotopically uniform; allchondrules have porphyritic textures. In con-trast to aluminum-rich chondrules in ordinarychondrites, oxygen-isotopic heterogeneity isdue to the presence of relict CAIs (Krot andKeil, 2002).

Aluminum–magnesium isotope measure-ments of aluminum-rich chondrules in ordi-nary and carbonaceous chondrites weresummarized by Huss et al. (2001). About 10–15% of all aluminum-rich chondrules studiedshow 26Mg*. The inferred (26Al/27Al)0 ratio ofthese chondrules is B(0.5–2)� 10�5, which isconsistent with the values found in ferromagne-sian chondrules from type 3.0–3.1 unequili-brated ordinary chondrites (Kita et al., 2000;

McKeegan et al., 2000b; Mostefaoui et al.,2000, 2002).

Correlated in situ analyses of the oxygen-and magnesium-isotopic compositions of Al-rich chondrules from unequilibrated enstatitechondrites show that among 11 Al-rich chond-rules measured for 26Al–26Mg systematics, onlyone chondrule contains 26Mg* correspondingto the initial 26Al/27Al ratio of (6.872.4)� 10�6

(Guan et al., 2006). The oxygen-isotopic com-positions of five Al-rich chondrules on a three-oxygen-isotope diagram define a line of slopeB0.670.1 and overlap with the field defined byferromagnesian chondrules in enstatite chond-rites but extend to more 16O-rich compositionswith a range in d18O of about 12% (figure 3 inGuan et al., 2006).

The REEs in most aluminum-rich chond-rules from carbonaceous chondrites studiedare unfractionated (Kring and Boynton, 1990;Weisberg et al., 1991). However, some havehighly fractionated REE patterns, similar tovolatility-controlled group II and ultrarefracto-ry patterns seen in refractory inclusions (Kringand Boynton, 1990; Misawa and Nakamura,1988, 1996).

50 µm

chd

CAI

sp

pl

met

px

300 µm cpx

opx

met

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ol

Mg K� Si K�

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sp

20 µm

pl

met

plsil

pv

pl

an

p

(d)

(g) (h)

(f)

(c)(b)Mg K�

(a)

Figure 14 Backscattered electron images (a, c), combined elemental map in magnesium (red), calcium (green)and Al Ka (blue) X-rays (b, d, g), and elemental maps in magnesium (e, h) and Si Ka X-rays (f, i) of theCAI-bearing aluminum-rich chondrule #2 in the CH chondrite Acfer 182 (a–c) and plagioclase-bearingchondrules #1 (d–f) and #17 (g–i) in the CR chondrites EET92042 and PCA91082, respectively. (a–c) RelictCAI consists of irregularly-shaped spinel (sp) core surrounded by anorthite (pl); tiny perovskite (pv) inclusionsoccur in spinel. The chondrule portion consists of low-calcium pyroxene (opx), high-calcium pyroxene (cpx),Fe,Ni-metal nodules (met) and silica-rich crystalline mesostasis composed of silica (sil), anorthite and high-calcium pyroxene. Region outlined in (b) is shown in detail in (c). White veins (w) in (a) and (c) are veins ofterrestrial weathering products. (d–f) Chondrule #1 consists of anorthite poikilitically enclosing spinelgrains, which are relict, high-calcium pyroxene, low-calcium pyroxene, Fe,Ni-metal nodules, minor forsteriteand silica-rich crystalline mesostasis composed of silica, anorthite, and high-calcium pyroxene. (g–i)Chondrule #17 consists of forsteritic olivine, low-calcium pyroxene, anorthitic plagioclase, high-calciumpyroxene, Fe,Ni-metal nodules (black), and silica-rich crystalline mesostasis composed of anorthite, silica,and high-calcium pyroxene. Low-calcium pyroxene occurs as elongated grains overgrown by high-calciumpyroxene in the chondrule cores and as phenocrysts poikilitically enclosing forsterite grains in chondrule

peripheries.


1.07.5.4.2 Origin of aluminum-rich chondrules

MacPherson and Russell (1997) and Mac-Pherson and Huss (2005) discussed severalpossible models for the origin of aluminum-rich chondrules including impact melting of

aluminous (evolved) parent body surface ma-terial, mixing of aluminous components withferromagnesian chondrule precursors prior tochondrule formation, and volatilization. Theyconcluded that (1) no single model appearscapable of explaining the entire spectrum of

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Chainpur 14-1, spChainpur 14-1, pxChainpur 14-1, olSharps 10-1, spSharps 10-1, ol+px+plSharps 10-1, nph+plSharps 10-1, ol

Chainpur 10-1, gl, spChainpur 14-2, mixChainpur 20-1, ol, px, spInman 1-1, mixQuinyambie 5-2

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Acfer 182 #2, spAcfer 182 #2, pxAcfer 182 #2, plEET92042 #1, spEET92042 #1, opxEET92042 #1, cpxEET92042 #1, plEET92147 #1, spEET92147 #1, cpxEET92147 #1, pl

GRA95229 #9, pl, cpx, opxMAC87320 #1, ol, pl, opxGRA95229 #7, pl, cpx, ol

(b)

Al-rich chondrules in CRs,without relict grains

Al-rich chondrules in OCs,homogeneous

Al-rich chondrules in CRs & CHs, with relict grains

Al-rich chondrules in OCs, heterogeneous

Figure 15 Oxygen-isotopic compositions of individual minerals in (a) aluminum-rich chondrules fromordinary chondrites (Russell et al., 2000) and (b) CR carbonaceous chondrites (Krot et al., 2006a).Abbreviations: cpx, clinopyroxene; gl, glass; nph, nepheline; ol, olivine; opx, orthopyroxene; pl, plagioclase;

px, pyroxene; sp, spinel.



aluminum-rich chondrule bulk compositions,(2) aluminum-rich chondrules are not inter-mediate in an evolutionary sense betweenknown CAIs and chondrules, and (3) alumi-num-rich chondrules are more closely related tonormal chondrules, either by addition of ananorthite-like component to chondrule precur-sors or possibly, in some cases, by volatilizationof chondrule or chondritic material of unusualoxygen-isotopic composition. Russell et al.(2000), MacPherson and Huss (2005), andGuan et al. (2006) concluded that ‘‘if alumi-num-rich chondrules were mixtures of fer-romagnesian chondrules and CAI material,their bulk chemical compositions would re-quire them to exhibit larger 16O enrichmentsthan we observe. Therefore, aluminum-richchondrules are not simple mixtures of thesetwo components.’’

However, Krot and Keil (2002) and Krotet al. (2001b, 2002a) showed that B15% of al-uminum-rich chondrules from different car-bonaceous chondrites are associated with relictCAIs (Figure 14). These relict CAIs are min-eralogically similar to CAIs in AOAs and con-sist of spinel, anorthite, 7aluminum-diopside,and 7forsterite; most of them were extensivelymelted during chondrule formation. Krot andcolleagues inferred that aluminum-rich chond-rules formed by melting of the ferromagnesianchondrule precursors (magnesian olivine andpyroxenes, Fe,Ni-metal) mixed with the refrac-tory materials, including relict CAIs, composedof anorthite, spinel, high-calcium pyroxene,and forsterite. This mixing explains the relativeenrichment of anorthite-rich chondrules in 16Ocompared to typical ferromagnesian chond-rules (Figure 12), and the group II REE pat-terns of some of the aluminum-rich chondrules(Misawa and Nakamura, 1988, 1996; Kringand Boynton, 1990).

1.07.5.4.3 Relict CAIs and AOAs inferromagnesian chondrules

Mineralogical observations indicate thatCAIs outside chondrules show no clear evid-ence for being melted during a chondrule-forming event (Krot et al., 2002a, b, c, 2005d).Relict CAIs inside ferromagnesian chondrulesare exceptionally rare. Only six relict CAIshave been described in ferromagnesian chond-rules: in the unique carbonaceous chondritesAcfer 094 and Adelaide, the CH chondritesAcfer 182 and Patuxent Range 91546, the CVchondrite Allende, and the H3.4 ordi-nary chondrite Sharps (Bischoff and Keil,

1984; Misawa and Fujita, 1994; Krot et al.,2006b). Relict CAIs are also present in AOAs(Itoh et al., 2002; Hiyagon, 2000). These rel-ict CAIs are very refractory, unlike those inaluminum-rich chondrules, being composed ofhibonite, grossite, spinel, melilite, perovskite,and platinum group metals (Figure 16). Therelict CAIs in chondrules are corroded by hostchondrule melts, but have largely preservedthe 16O-rich isotopic signatures of typicalCAIs (Figure 17a). The only relict AOAobserved inside a chondrule has also preservedits 16O-rich isotopic signature (Figure 17b).These observations indicate that some CAIswere present in the chondrule-forming regionsand were melted together with chondruleprecursors, suggesting that at least some CAIsformed before chondrules.

Magnesium isotope measurements of therelict CAIs and CAI- and AOA-bearing chond-rules typically show no resolvable 26Mg* (Krotet al., 2002a; Itoh et al., 2002), implying upperlimits for (26Al/27Al)I of o1� 10�5 ando5� 10�6. These data suggest either theseCAIs formed without 26Al or their Al–Mg sys-tems were reset during chondrule formation.If these CAIs formed initially with a canonical(26Al/27Al)I that was reset during chondrulemelting, the host chondrules must have formedat least B2Myr after the CAIs.

The only exception is the hibonite-richrelict CAI #9d from Adelaide (Figure 16d)that shows clear excesses of radiogenic 26Mg inhibonite that are correlated with 27Al/24Mgratio. The observed 26Mg* in hibonite togetherwith spinel define an isochron with (26Al/27Al)Iof (3.770.5)� 10�5. Melilite in this CAIshows no 26Mg*, while a mixture of hiboniteand melilite shows an intermediate 26Mg ex-cess that appears to plot on a mixing lineconnecting melilite and the most 26Mg*-enriched hibonite. Two measurements ofplagioclase in the host chondrule, whichwere made in the rim around the CAI, showno resolvable 26Mg* [(26Al/27Al)Io5� 10�6].Krot et al. (2006b) concluded that whenCAI #9d was incorporated into the surround-ing chondrule melt, diopside, melilite (oranorthite), and spinel in its Wark–Loveringrim sequence began reacting with the sur-rounding chondrule melt to produce a layer ofplagioclase, and melilite in the interior of theinclusion-exchanged magnesium with the sur-rounding chondrule melt. The newly formedplagioclase and the melilite inside the CAIno longer contain evidence of 26Mg*. Thisconstrains the timing of the chondrule meltingevent to be at least 2Myr after the CAI

Acfer 094 #17

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PAT91546 #16

cpx met

grs

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Y-81020 #1

(f)

px

Figure 16 Backscattered electron images of chondrules containing relict refractory inclusions: #17 in Acfer094 (a, b), #9d in Adelaide (c), #9 in Sharps (d), and #16 in PAT91546 (e). Relict CAI Acfer 094 (a, b) consistsof hibonite (hib) with inclusions of perovskite (pv), a Zr-bearing phase (Zr-ph), and Re-, Ir-, Os-bearing (PGE)nuggets. It is surrounded by a shell of fine-grained, iron-spinel (sp). The host chondrule consists of ferrousolivine (ol), plagioclase mesostasis (pl), Fe,Ni-sulfides (sf), Cr-spinel, and relict grains of forsteritic olivine (fo).The relict CAI in Adelaide (c) consists of hibonite, perovskite, and melilite (mel); it is surrounded by amagnesium-spinel rim. Its host chondrule consists of forsteritic olivine, low-calcium pyroxene, glassy me-sostasis, and kamacitic metal. The relict CAI in Sharps (d) consists of hibonite surrounded by iron-spineland a nepheline-like phase (nph). Its host chondrule consists of forsteritic olivine, low-calcium pyroxene,high-calcium pyroxene, anorthitic mesostasis, and kamacitic metal. The relict CAI in PAT91546 (e) has agrossite (grs) core surrounded by a spinel (sp) rim. This CAI is surrounded by an igneous chondrule-likematerial composed of plagioclase (pl), low-Ca pyroxene (opx), high-Ca pyroxene (cpx), olivine (ol), andFe,Ni-metal (met), mesostasis (mes) (after Krot et al., 2005d). (f) Relict AOA #1 in a type II chondrule from

the CO3.0 chondrite Y-81020 (after Yurimoto and Wasson, 2002).


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Relict chd A5enpgnCCAM

AllendeHost CAI ABC

anfasmelspCr-spdi

Relict chd ABCenol

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(c) (d)

Acfer 094CAI #17

Figure 17 Oxygen-isotopic compositions of chondrules with relict CAIs (a) and relict AOA (b), and type CCAIs with relict chondrules (c, d). (a) The relict CAIs are 16O-enriched compared to their host chondrules.(b) The core of the relict AOA (see Figure 16f) is 16O-enriched compared to the olivine phenocrysts andmesostasis of the host chondrule; olivine mantle overgrowing AOA has intermediate oxygen-isotopic com-positions. (c) Spinel and Al–Ti-diopside of the host CAI ABC are 16O-enriched; whereas other minerals are16O-depleted to various degrees. Relict olivine and low-Ca pyroxene grains have 16O-poor compositions. (d)Spinel of the CAI core is 16O-rich, whereas other minerals are 16O-depleted to various degrees. The Al–Ti-diopside grains in the core are less 16O-depleted compared to those in the mantle. Relict olivine and high-Capyroxene halo around it have 16O-poor compositions. Al–Ti-di, Al–Ti-diopside; an, anorthite; cpx, high-Capyroxene; di, diopside; mel, melilite; mes, mesostasis; ol, olivine; plag, plagioclase; px, low-Ca pyroxene; sp,

spinel. Data from Yurimoto and Wasson (2002) and Krot et al. (2005d).


formed, because 480% of 26Al had decayed.Thus, Adelaide CAI #9d and its surroundingchondrule provide clear evidence of a Z2Myrtime gap between the formation of the CAIand its incorporation into the chondrule,and this conclusion does not require an as-sumption of homogeneously distributed 26Alin the chondrule-forming region (MacPhersonet al., 1995).

1.07.5.4.4 Relict chondrules in igneous CAIsand igneous CAIs overgrown bychondrule material

Relict chondrules have been reported in threeigneous CAIs—compact type A CAI A5 fromY-81020, coarse-grained, igneous, anorthite-rich(type C) CAIs ABC, and TS26 from Allende(Itoh and Yurimoto, 2003; Krot et al., 2005d).


No relict chondrules have been found insideAOAs or nonigneous CAIs.

CAI A5 has a melilite-rich core surroundedby a very fine-grained mantle composed ofAl-diopside, spinel, and anorthite; the mantlecontains igneously zoned pyroxene grainsranging in composition from enstatite to augite(Figure 18). The melilite and pyroxene grainsare similarly 16O-depleted, whereas the mantleshows a wide compositional range and is16O-enriched (Figure 17d). Itoh and Yurimoto(2003) concluded that the pyroxene grains arerelict chondrule fragments incorporated intothe host CAI prior to its final melting in theCAI-forming region. Alternatively, the lack ofWark–Lovering rim layers around the hostCAI, the estimated fast cooling rate(4100 1Ch�1) during crystallization of theCAI mantle, and the wide range of oxygen-iso-topic compositions of the CAI minerals over-lapping with those of relict chondrulefragments, may indicate that this compoundCAI-chondrule object experienced late-stage,incomplete melting and oxygen-isotopic ex-change in an 16O-rich gaseous reservoir duringchondrule formation (Russell et al., 2005).

Type C CAI ABC in its peripheral zone con-tains a relict olivine-pyroxene fragment that iscorroded by Al-diopside and surrounded by ahalo of high-Ca pyroxene (Figures 19a and19b). Spinel and Al,Ti-diopside of the host CAIare 16O-enriched, whereas melilite, anorthite,olivine, and low-Ca pyroxene are 16O-depleted;Cr-spinel, high-Ca pyroxene, and Al,Ti-poordiopside have intermediate oxygen-isotopic

ol

CAI

20 µm

mes

mel

Figure 18 Backscattered electron image of the chondrutalline melilite core surrounded by a porous mesostaclinypyroxene (lighter area) and Al-rich glass (darker ardominated by enstatite with minor pigeonite and augite;

the mesostasis near the p

compositions (Figure 17c). The CAI anorthitehas a resolvable 26Mg* corresponding to the(26Al/27Al)I ratio of (4.771.4)� 10�6.

TS26 is a mineralogically zoned type C CAI(Figures 19c and 19d). The coarse-grained core issurrounded by a thick, fine-grained mantle con-taining abundant relict grains of forsteritic oli-vine and low-Ca pyroxene; the Wark–Loveringrim layers are absent. Spinel and Al,Ti-diopsideof the CAI core are 16O-rich, whereas melilite,anorthite, Al,Ti-poor diopside of the CAImantle, and relict olivine and low-Ca pyroxeneare 16O-depleted to various degrees (Figure 17c).Anorthite shows no resolvable 26Mg*; theinferred (26Al/27Al)I ratio is o2.5� 10�6.

The relict nature of the olivine–pyroxenefragments in the periphery of the CAIs doesnot necessarily mean that they predated theCAIs. Instead, the 16O-poor oxygen in theperiphery of the CAIs and resetting of their26Al–26Mg systematics, provide evidence thatthey experienced a late remelting event. Duringthis event, incorporation of the chondrule-like,olivine–pyroxene material could have occurred(Krot et al., 2005d). This interpretation isconsistent with some igneous CAIs being over-grown by chondrule-like material (Figures 19eand 19f). We note that the apparent discrep-ancy in interpretation of chondrule-bearingigneous CAIs (Itoh and Yurimoto, 2003; Krotet al., 2005d) is largely due to differences indefinitions of CAI-formation. According toItoh and Yurimoto (2003), all high-tempera-ture processes that affected CAIs should becalled CAI-forming events, whereas Krot et al.

tr pg mx

aug

en

chd

oI

tr

le-bearing CAI A5. The CAI consists of a polycrys-sis composed of a fine-grained mixture of Al-richea). The mesostasis encloses a pyroxene assemblagetroilite and Fe,Ni-metal (white spots) are scattered inyroxene assemblage.

ABC

TS26

93

spsod

200 µm

200 µm

500 µm 50 µm

px

sodnph

50 µmX300

20 µm 20 µmX850

20 µm 20 µm

cpx

di

sf

px

nphol

X700

di

cpxsf

px

ol

di

(a)di

di

fa+sod

(c)

core

sp

sp sp

an

an

an

(f)

fa+sod

mel

di

di

(e)

cpx

cpx

sod

olMantle

an

sec

Mg:Ca:Al

Mg:Ca:Al

Mg:Ca:Al

(d)

(b)

Figure 19 Combined elemental maps in Mg (red), Ca (green), and Al Ka (blue) X-rays (a, c, e) and back-scattered electron images (b, d, f) of the Allende type C CAIs associated with chondrule-like ferromagnesiansilicates. (a, b) CAI ABC. (c, d) CAI TS26. (e, f) CAI 93. Regions outlined in ‘‘a,’’ ‘‘c,’’ and ‘‘d’’ are shown indetail in ‘‘b,’’ ‘‘d,’’ and ‘‘e,’’ respectively. an, anorthite and anorthitic plagioclase; cpx, high-Ca pyroxene; di,diopside; mel, melilite; nph, nepheline; px, low-Ca pyroxene; sec, secondary sod, sodalite; sp, spinel.

Reproduced by permission of Astronomical Society of the Pacific from Russell et al. (2005).


(2005d) define CAI-forming events as thosewhich occurred in the CAI-forming regions,that is, regions with high ambient temperatures(41350K), relatively low pressure, 16O-richnebular gas, and approximately solar redoxconditions.

1.07.5.5 Refractory Forsterite-Rich Objects

A long-running controversy has existed overthe origin of refractory forsterites 0.1–0.5mm

in size in type 2 and 3 ordinary and carbon-aceous chondrites. Although most isolatedolivines are probably formed by disaggregat-ion of chondrules (McSween, 1977b; Roedder,1981), Steele (1986, 1989) argued that theextremely forsteritic grains (Fao2) in carbon-aceous chondrites, which have a blue cathodo-luminescence color, high concentrations ofCaO and Al2O3 (typically 0.5 and 0.25 wt.%,respectively) and comprise a few percent of allolivines, did not crystallize in chondrules.


Steele suggested that the refractory forsteritesmay have condensed from the nebula, formedrelict grains in chondrules, and were an impor-tant component of chondrule precursor mate-rial. Weinbruch et al. (2000) showed thatisolated refractory forsterites were more 16O-rich than olivines that crystallized in chond-rules and also favored a condensation origin.However, Jones et al. (2000b) showed thatthere are rounded objects composed largely ofa single large forsterite crystal with traces oflow-Ca pyroxene and very little mesostasis thathave compositions like those of the isolatedforsterites. Jones et al. (2000b) inferred thatthese objects were chondrules, that relict fors-terites from such objects were present in othertype I chondrules, and that refractory 16O-richforsterites were primitive nebular components.

Pack et al. (2004, 2005) found o0.01–0.4 vol.% luminescent refractory forsterites ina variety of ordinary, carbonaceous, and Rchondrites, and found no difference in theirchemical and isotopic compositions. Highly re-fractory forsterites with 40.6 wt.% CaO ando0.5 wt.% FeO are enriched in 16O with D17Obetween –4% and –10%, whereas moderatelyrefractory and nonrefractory olivines range be-tween –5% and þ 2%. They inferred that therefractory forsterites come from a single com-mon reservoir and that they formed in a red-ucing nebula gas of solar composition. Theoxygen-isotopic and chemical compositions ofthe refractory forsterites lie beyond those at theFeO-poor end of the sequence defined byolivine-rich, FeO-poor type IA chondrulesand pyroxene-rich type IB chondrules (seeJones, 1994; Section 1.07.5.6.1). Melt inclu-sions in refractory forsterites are not in equi-librium with the olivine hosts—either becauseforsterite crystallized outside its stability field(McSween, 1977b), or because the melt reactedmore readily with nebular gas than the crystal-lizing solids (Varela et al., 2005; Pack et al.,2004, 2005).

Refractory forsterite-rich objects are clearlya fascinating component of chondrites and mayprovide a possible link between refractoryinclusions, especially AOAs, and chondrules.Their 16O-enriched composition points to anearly solar nebula origin, though their forma-tion ages have not been determined. Rounded,refractory forsterite-rich objects may be akind of proto-type-I chondrules, as normaltype I chondrules show evidence for gas–solidreactions during cooling (Section 1.07.5.6.5).However, Libourel and Krot (2006) describedcoarse-grained forsteritic aggregates withgranoblastic textures inside type I chondrulesin CV chondrites and inferred that these igne-ous objects, which are also 16O-enriched

(Chaussidon et al., 2006), could not haveformed in the solar nebula and may be derivedfrom planetesimals that were disrupted. How-ever, if these planetesimals existed, they have notsupplied us with meteorites, as differentiatedmeteorites have D17O4–3% (Chapter 1.05).Whatever their origin, refractory 16O-enrichedforsterites appear to be an important compo-nent of type I chondrule precursor materials.

1.07.5.6 Chondrules

Chondrules are igneous-textured particlescomposed mainly of olivine and low-calciumpyroxene crystals set in a feldspathic glass ormicrocrystalline matrix. (Chondrules with 410wt.% Al2O3, which have closer links with CAIsare reviewed in Section 1.07.5.4.) Many chond-rules with FeO-poor silicates also contain me-tallic Fe,Ni grains, often clustered near theperiphery. Metal grains outside chondrules thatare of comparable sizes to those inside havegeochemical properties suggesting they wereonce ejected from or broken out of chondrules,or in the case of the CB group (where reducedchondrules lack metal), that metal and chond-rules formed in the same environment. A closelink between metal and reduced silicates is con-sistent with mineral–gas equilibration in thesolar nebula as forsterite, enstatite, and metal-lic Fe,Ni are stable over similar temperatureranges (figure 2 of Chapter 1.15). The miner-alogical, chemical, and isotopic properties ofchondrules have been reviewed most recentlyby Jones et al. (2005).

Links between asteroids and chondritessuggest that chondrules are abundant in theinner part of the asteroid belt and that theterrestrial planets probably formed fromchondritic materials. Thus, much of the massthat accreted in the inner solar system wouldhave been derived from chondrules. In theabsence of sticky organics, dry grains of silicatedo not readily adhere so that chondrule forma-tion may have been the process that triggeredaccretion in the inner solar system (e.g., Wood,1996a; Cuzzi et al., 2001; Scott, 2002). Under-standing the origin of chondrules has been amajor goal of meteorite studies since chondriteswere first described over 200 years ago.

1.07.5.6.1 Chondrule textures, types, andthermal histories

Chondrule textures are conveniently dividedinto two types: porphyritic, which have largecrystals of olivine and/or pyroxene in a fine-grained or glassy matrix, and nonporphyritic,which include those with cryptocrystalline,


radial-pyroxene, and barred-olivine textures(Figures 1 and 20). Table 4 shows how theproportions of different types vary considera-bly among the groups. Carbonaceous chond-rites tend to have the highest fraction ofporphyritic chondrules, but CH and CB groupsare exceptions. Mean sizes and abundances ofchondrules in each group are given in Table 1.

mes

metol

ol

mes

sf

200 µm

Dusty ol

(a) 400 µm

(b)

100 µm(c)

Figure 20 Backscattered electron images of threechondrules in Tieschitz (H/L3.6) chondrite. (a) TypeIA porphyritic chondrule composed largely of fors-teritic olivine (ol), mesostasis (mes), metallic Fe,Nidroplets (met) (white). Dusty olivines are relict, FeO-rich olivines that crystallized elsewhere and formedtiny metallic Fe,Ni particles when heated in thechondrule melt. (b) Type IIA porphyritic chondrulecontaining large euhedral, FeO-bearing, olivinephenocrysts, dark mesostasis (mes), and white sulfidedroplets (sf). (c) Nonporphyritic chondrule contain-ing fine pyroxene crystals that appear to radiate from

the upper edge of the chondrule.

Laboratory experiments constrain the initialconditions and thermal history required foreach textural type, but there are too manyvariables (e.g., initial state of chondrule, peaktemperature, time at peak temperature, andcooling rate) to make very quantitative state-ments (Lofgren, 1996; Hewins, 1997; Hewinset al., 2005). Porphyritic chondrules crystallizedfrom melts with many nuclei after incompletemelting of fine-grained, precursor materials(Lofgren, 1996; Connolly et al., 1998). Peaktemperature and duration of heating are linkedas porphyritic textures can be reproduced byheating fine-grained material to B100 1C abovethe liquidus temperature (generally 1,350–1,800 1C) for 5min, or just below the liquidustemperature for hours (Hewins and Connolly,1996; Lofgren, 1996). Nonporphyritic chond-rules crystallized from melts that were super-heated above their liquidus for long enough todestroy all solids that could act as nuclei. Forbarred olivine and radial pyroxene chondrules,nucleation was probably induced by collisionwith solid grains (Connolly and Hewins, 1995).Radial textured chondrules crystallized whilecooling at 5–3,000 1Ch�1, barred olivinechondrules at 500–3,000 1Ch�1, and porphyri-tic chondrules at 5–100 1Ch�1 (Desch andConnolly, 2002). Cooling rates above the liq-uidus may have been much higher; below thesolidus, chondrules cooled more slowly at ratesof 0.1–50 1Ch�1 (Weinbruch et al., 2001).

Porphyritic chondrules are classified into typeI, which are FeO-poor in unmetamorphosedchondrites, and type II, which are FeO-rich.Type I chondrules account for 495% of chond-rules in CO, CV, CR (Figure 1a), and CMchondrites and olivine-rich varieties were prob-ably abundant in the material that accreted toform the Earth (Hewins and Herzberg, 1996).Types I and II are both subdivided into A, whichare olivine-rich (490 modal percent olivine),and B, which are pyroxene-rich (Jones, 1990,1994, 1996a, b). Olivines in type IA chondrulesin unmetamorphosed chondrites are mostlyFa0.3–5 and Fa2–10 in IB; type II chondruleolivines are Fa10–20 in ordinary chondrites andup to Fa50 in carbonaceous chondrites.Rounded type I chondrules commonly havelow-calcium pyroxene rims, glassy mesostases,and grains of metallic Fe,Ni, whereas type IIchondrules lack these features, contain smallchromites and commonly have larger olivines(Figure 20; Scott and Taylor, 1983). Type I andII chondrules can be identified from these fea-tures in many type 2–6 chondrites so that meta-morphic and alteration effects can be tracked(McCoy et al., 1991; Hanowski and Brearley,2001). Type I chondrules also occur as finergrained, irregularly shaped aggregates, which

Table 4 Proportions of chondrule types in chondrite groups.

Type/group CM CO CV CR CH CBb CK H–L–LL EH EL R

PorphyriticOlivine 8 23 0.1Oliv-pyx. 69 48 4Pyroxene 18 10 77Granular oliv-pyx. o0.1 o0.1 3 1Total porphyritic B95 95 94 96–98 20 B1 499 84 82 B87 B92

NonporphyriticBarred olivine 2 6 1 4 0.1Radial pyx. 2 0.2 7 13Cryptocrystalline 1 0.1 79a 5 5Total nonporph. B5 5 6.3 2–4 80 B99 o1 16 18 B13 B8

Sources: Rubin (2000), Kallemeyn et al. (1994), Weisberg et al. (2001), Scott (1988).a Includes radial pyroxene and glass chondrules.

IAB,BIA

1

0.1

0.01

Abu

ndan

ce r

elat

ive

to C

l and

Si

Al Ca Ti Mg Si Cr Mn K Na

Figure 21 Elemental abundances in type IA chond-rules, which are olivine-rich and type IB and IABchondrules, which are pyroxene rich, from the Se-markona (LL3.0) chondrite normalized to CI chond-rites and silicon. Type IA chondrules are enriched inrefractories relative to IAB and IB. For moderatelyvolatile elements, all type I chondrules are depleted,with most IA showing larger depletions than typesIB and IAB. Reproduced by permission of Elsevier

from Jones (1994).


represent partly melted dust aggregates oraggregates of partly melted particles (e.g., Rubinand Wasson, 2005). In ordinary chondrites,which have very diverse textured chondrules,all chondrules having olivines with Fao10may be referred to as type I by some authors(Hewins, 1997).

Chondrules have also been classified usingtheir cathodoluminescence properties and thecompositions of olivines and mesostases (Searset al., 1995a). For unmetamorphosed chond-rites, groups A and B in this scheme correspondto types I and II. During metamorphism, thecathodoluminescence properties and chondrulegroup change. The merits of the two schemeswere reviewed by Scott et al. (1994) and Searset al. (1995b). Grossman and Brearley (2005)caution that the classification criteria based onthe compositions of chondrule mesostases needto be revised because of Na loss during electronprobe analysis.

1.07.5.6.2 Compositions of chondrules

Chemical compositions of chondrules havebeen determined from extracted samples usingneutron activation analysis and by in situ anal-ysis in polished sections using electron micro-probe and ion probe analysis (e.g., see Goodinget al., 1980; Grossman et al., 1988; Alexander,1995). Chondrules typically show flat refrac-tory abundances that are relatively close to themean chondrite value and abundances of mod-erately volatile elements that scatter morewidely about the mean. Type II chondrulesare relatively unfractionated with near-CIlevels of refractories and moderately volatileelements. However, type I chondrules showsystematic depletions of moderately volatileelements and a broader spread of refractoryabundances with the silicon-rich type IB chond-rules being poorer in refractories than the

silicon-poor type IA chondrules (Figure 21).The source of the fractionations in type Ichondrules is discussed below.

Oxygen-isotopic compositions of largeindividual chondrules suggest that chondrulesformed from several different reservoirs that werepoorer in 16O than the CAI source (Figure 22).Chondrules in ordinary chondrites generallyplot above the terrestrial fractionation line,enstatite chondrite chondrules plot on or nearthe line, and carbonaceous chondrite chond-rules lie below. Ion probe analyses for oxygenisotopes in olivine show that most ferromagne-sian chondrules have relatively homogeneouscompositions due to exchange between

6

4

2

0

−2

−4

−6

−8−4 −2 0 2 4 6 8

�18O (‰ rel. SMOW)

�17O

(‰

rel

. SM

OW

)

TF

CO

CVCM

CCAM

CR

EH

OCR

Figure 22 Oxygen-isotopic plot showing fields ofanalyses of large chondrules in chondrite groups.The data show that chondrules are derived frommany isotopically distinct reservoirs. Reproduced by

permission of Elsevier from Rubin (2000).


chondrule melt and nebula gas; but some type Ichondrules have 16O-rich forsterites that wereprobably present in precursor materials (seereferences in Scott and Krot, 2001; Jones et al.,2004; Krot et al., 2006a). Chondrules are iso-topically homogeneous in comparison to CAIs:50Ti anomalies (see Niemeyer, 1988) and massfractionation effects (e.g., Clayton et al., 1985;Misawa and Fujita, 2000; Alexander andWang, 2001; Alexander et al., 2000) are muchsmaller than those in many CAIs. Inferredinitial concentrations of 26Al in chondrules and207Pb/206Pb dating suggest that chondrulesformed 1–4Ma after CAIs (Kita et al., 2000;Mostefaoui et al., 2002; Amelin et al., 2002).See Gilmour (2002), Gilmour et al. (2006), Kitaet al. (2005), and Chapter 1.16 for reviews ofchondrule formation ages.

Effects of secondary processes in modifyingvolatile elements (Grossman et al., 2000, 2002)and oxygen isotopes in chondrules, especiallyin ordinary chondrites (Bridges et al., 1998;Franchi et al., 2001) require further study (seeChapter 1.09).

Trapped and cosmogenic noble gases inchondrules offer clues to their irradiation his-tory and formation but not definitive constraints(see Chapter 1.14). Trapped noble gases aregenerally absent or rare in chondrules, but signi-ficant amounts of trapped 36Ar, 84Kr, and 132Xefrom a so-called subsolar component were foundin chondrules in an EH3/4 chondrite (Okazakiet al., 2001). This component is intermediate in

composition between that of the Sun and themajor trapped component in type 1–3 chondrites(‘‘Q-type’’ gases) and occurs in E4–6 chondrites(Patzer and Schultz, 2002b), CR2, CH3, CBa3,and K3 chondrites (see Bischoff et al., 1993).Okazaki et al. (2001) inferred that solar windhad been implanted into chondrule precursormaterial near the protosun and then fractionatedduring melting to leave a subsolar pattern.However, subsolar noble gases are present inseveral other sites suggesting that the subsolargases were redistributed (Busemann et al., 2002).Further laser microprobe analyses of chondrulesand other components in the most pristinechondrites are clearly needed to establish whichchondrules, if any, acquired solar gases from thenebula or were irradiated by high-energy parti-cles near the inner edge of the disk.

1.07.5.6.3 Compound chondrules, relict grains,and chondrule rims

Chondrules in all groups except CB and CHshow evidence for collisions: with each otherwhile partly molten and with solid grains whilemolten and solid. Compound chondrules formby collisions of partly molten chondrules andtheir abundance gives some constraint on thenumber density of chondrules in the regionwhere they formed. Unfortunately, there arelarge uncertainties in chondrule speed andthe provenance of the joined chondrules,which make estimates of the chondrule densityvery uncertain. About 1–2% of chondrules inordinary and CV chondrites are compound(Wasson et al., 1995; Akaki and Nakamura,2005). In over half of these, the chondrules thatcollided have such similar compositions thatthey could have been derived from a singlelarge droplet. Other compound chondrulesmay have formed by part melting of a porousaggregate containing an embedded chondrule:relatively few compound chondrules probablyformed by random collisions. However, Cieslaet al. (2004b) infer that at least 5% of chond-rules are compound and that chondrulesformed where solids were concentrated by445 times above the canonical solar nebulavalue, possibly at the nebula midplane.

Foreign grains in chondrules appear to havebeen derived from other chondrules (Jones,1996b). Relict low-FeO olivines from type Ichondrules can be found in high-FeO, type IIchondrules andvice versa, indicating that chond-rule formation was a repeated process and thatmaterial was recycled through the chondrulefurnace (e.g., Kunihiro et al., 2004b, 2005a).

Chondrules may have two types of rims: fine-grained matrix rims and coarse-grained igneous


rims: both are missing in CH and CB chond-rites. Matrix rims vary in composition but thevariations are uncorrelated with chondrulecomposition (Section 1.07.5.7). However, thecompositions of igneous rims are related tothose of the host chondrule: type I chondruleshave igneous rims that are FeO poor, whiletype II chondrules have FeO-rich rims (Krotand Wasson, 1995). The igneous rims formedby collisions between chondrules and smallerpartly molten particles, or solid particles withdiverse oxygen-isotopic compositions that wereaccreted and later incompletely melted (Naga-shima et al., 2003).

From the abundances of igneous rims, com-pound chondrules, and relict grains and thedearth of unmelted, chondrule-sized dustballs,we infer that collisions between partly meltedobjects and remelting of aggregates from suchcollisions were important processes that ena-bled chondrules to form from fine particles.

1.07.5.6.4 Closed-system crystallization

Understanding whether chondrules actedas open or closed systems during chondruleformation is crucial for elucidating their origin(Sears et al., 1996). Type II chondrules appearto be good candidates for closed system crys-tallization, as chemical zoning profiles acrosstheir phenocrysts are consistent with fractionalcrystallization (Jones, 1990; Jones and Lofgren,1993). Type II chondrules contain significantconcentrations of sodium and other volatileelements that would have been readily lost ifthe chondrules had been molten for more thana few minutes (Yu and Hewins, 1998). Suchchondrules are commonly considered as arche-typal chondrules formed by closed-system mel-ting. Although most workers infer that type IIchondrules resulted from single heating events,Wasson and Rubin (2003) argue for multipleevents with low degrees of melting, whichwould help to explain how voltatiles wereretained.

Many workers once argued that all chond-rules formed by isochemical melting of precur-sor material and that the chemical fractionationsamong chondrules, except for metal loss, resultfrom processes that predate chondrule forma-tion (e.g., Gooding et al., 1980). Few workershold this view now, nevertheless it is commonlyheld that chemical fractionations prior tochondrule formation were much more impor-tant than any fractionation during chondruleformation, besides loss of metal (e.g., Gross-man, 1996). Thus, the standard model forchondrule formation envisages that volatile lossfrom chondrules was minimized by flash melting

of dust aggregates during melting events thatlasted seconds to minutes, or by high dust–gasratios or gas pressures.

The major difficulty with closed-system mod-els is that they do not explain why somechondrules are olivine-rich and others arepyroxene-rich. It is very difficult to see howfine-grained nebula dust could have been sortedinto high Mg/Si dustballs for olivine chond-rules and low Mg/Si dustballs for pyroxene-rich chondrules (e.g., Wood, 1996a). Althoughmatrix contains both phases, they are notsegregated into chondrule-sized volumes. Theonly plausible way to produce both olivine- andpyroxene-rich chondrules is by high-tempera-ture reactions between condensed material andnebular gas. New evidence, especially fromCR–CH–CB clan chondrites, supports thisthesis.

1.07.5.6.5 Open-system crystallization

Unlike the archetypal chondrules, type Ichondrules show a variety of evidence forvolatility-controlled chemical fractionationsthat may have occurred during chondruleformation. Hewins et al. (1997) studied micro-porphyritic and cryptoporphyritic type Ichondrules in Semarkona (LL3.0), which haveirregularly shaped interstitial metal grainsrather than the rounded droplets in mesostasiscommon in coarser type I chondrules. Theyfound an inverse correlation between grain sizeand abundance for the moderately volatileelements, sodium, potassium, and sulfur, aswell as iron and nickel (Figure 23). Hewinset al. inferred that volatilization increased withthe extent of melting of a fine-grained precursorhaving a chondritic composition. Alternatively,these very fine-grained chondrules may haveformed by aggregation of partly melted parti-cles that scavenged volatiles with an efficiencythat depended on their grain size. Despite thelarge increase in alkali concentrations fromtype IA to type IB chondrules, they do notshow any systematic variation in K isotopiccomposition (Alexander and Grossman, 2005).

Mesostases in some type IA chondrules inSemarkona show volatility-related chemicalzoning (Matsunami et al., 1993; Nagaharaet al., 1999). Mesostasis near the rims is higherin volatiles and lower in refractories thanmesostasis near the core. Matsunami et al.invoked recondensation and reduction proc-esses in the solar nebula. The chemical varia-tions across type I chondrules and those inmicroporphyritic and cryptoporphyritic type Ichondrules (Figure 23) mimic the fractionationpatterns shown by the bulk chemical data for

10

1

0.1

0.01

0.001Ca Al Ti Mg Si Cr Mn K Na Fe Ni P S

30−1,000

20−30

5−20

3−4

Cl

Chondrule grain size (µm)Abu

ndan

ce r

elat

ive

to C

I and

Mg

Figure 23 Elemental abundances in microporphyritic and cryptoporphyritic type I chondrules in Semarkonashowing an inverse correlation between grain size (in micrometer) and abundance for the moderately volatileelements (chromium to sulfur). Hewins et al. (1997) attribute this to volatilization during melting. Reproduced

by permission of National Institute of Polar Research from Hewins et al. (1997).


chondrules in the IA–IB sequence (Figure 21),suggesting single-process-controlled chemical

variations within and among type I chondrules.The concentration of pyroxene phenocrysts

around the periphery of many type I chond-rules was reproduced experimentally by Tis-sandier et al. (2002), who allowed melt to reactwith gaseous SiO during crystallization. Theyinferred that such gas–melt interactions wereprevalent during chondrule formation, andinvoked high dust–gas ratios to stabilize chond-rule melts in the nebula (Ebel and Grossman,2000). Further evidence for gas–melt interac-tions is provided by silicon-rich, igneous rimson many type I chondrules in CR2 chondrites(Krot et al., 2004b). These rims contain high-calcium and low-calcium pyroxene, glassymesostasis, Fe,Ni, and silica, with high con-centrations of chromium and manganese andlow concentrations of aluminum and calciumin high-calcium pyroxene (Figure 24). Meso-stasis zoning follows that in the ordinarychondrites with silicon, sodium, potassium,and manganese increasing and calcium, magne-sium, aluminum, and chromium decreasingfrom the olivine-rich core, through the pyrox-ene-rich rim to the silicon-rich rim. FeOremains nearly constant. Krot et al. (2004b)infer that gas–melt fractional condensation orgas–solid condensation followed by remeltingproduced these chemical fractionations. Isola-tion of early-formed forsterite condensates mayhave resulted naturally from crystallization, asforsterite would have equilibrated with thenebula more slowly than the melt.

Formation of chondrules from a gas–meltimpact plume, some by gas–melt condensation,

was proposed for chondrules in CBb chond-rites, which have unusual textures and compo-sitions (Krot et al., 2001d, 2006c). Thechondrules are cryptocrystalline or skeletalolivine, lack igneous or fine-grained rims andrelict grains, and do not form compound ob-jects (Figure 25). They clearly formed from to-tal melts in a dust-free environment. Somecryptocrystalline chondrules are embedded inzoned metal grains, which have compositionalprofiles consistent with nebular condensation(Krot et al., 2001d, 2006c). Chondrules haveuniform flat refractory element abundances butthe variation ((0.02–3)�CI) is far wider than inother chondrites (Figure 26). The depletions ofrefractory elements relative to silicon in thecryptocrystalline chondrules indicate that siliconor the refractory elements, or both, condensedinto chondrules or their precursors. The near-CIbulk abundances of refractory elements inthe chondrites favor closed-system fractionalcondensation. Moderately volatile elements inchondrules show large depletions (Figure 26)that are inversely correlated with condensationtemperature, as for bulk chondrites (Figure 3)and type I chondrules (Figure 21). Chondruletextures are consistent with condensation ofmelts.

1.07.5.6.6 Chondrules that formed on asteroids

Chondrites contain a small fraction ofchondrules that may have formed on asteroids.One rare type of chondrule (B0.1%) in type4–6 ordinary chondrite breccias is composedlargely of plagioclase (or mesostasis of plagio-clase composition) and chromite (Krot et al.,

250 250 µm

200 200 µm (d) 100 µm

POP

JSM−5900(e)

(c)(b)

JSM−5900 190190 30mm30mm200 mm200 mm

SIR

(a)

Mg K� Si K� Ni K�

Figure 24 X-ray elemental scanning maps using Mg Ka (a), Si Ka (b), and Ni Ka (c), plus backscatteredelectron images (d, e) of a typical type I porphyritic chondrule in the QUE 99177 CR2 chondrite, which has aforsterite-rich core, a porphyritic olivine–pyroxene rim (POP) and an outermost silica-rich rim (SIR). The NiKa map (c) shows that metallic Fe,Ni grains in the chondrule core are richer in nickel than those in the outerportions. The zoned structure of the chondrule can be explained by gas–melt fractional condensation or gas–

solid condensation followed by remelting (see Krot et al., 2003d).


1993). Krot and Rubin (1993) suggest thatthese chondrules may have formed by impactmelting as they find impact melts with similarcompositions inside shocked ordinary chond-rites and such chondrules are absent in type 3chondrites. In addition, some chromite-richchondrules contain chromite-rich aggregates,which appear to be fragments of equilibratedchondrites. Other possible impact productswere described in LL chondrites, which aremostly breccias, by Wlotzka et al. (1983). Theyfound microporphyritic potassium-rich clastswith highly variable K/Na ratios that aretotally unlike chondrules in type 3 chondrites.Large chondrules and clasts in Julesberg (L3)may also be impact melt products according toRuzicka et al. (1998).

1.07.5.7 Metal and Troilite

Two kinds of metal are found in chondrites:grains composedof refractory elements (iridium,osmium, ruthenium, molybdenum, tungsten,and rhenium), which condense along with the

refractory oxides above B1,600K at 10�4 atm,and grains composed predominantly of iron,cobalt, and nickel, which condense with forste-rite and enstatite atB1,350–1,450K (Campbellet al., 2005). The former are associated withCAIs (Palme and Wlotzka, 1976) and the latterwith chondrules, typically type I or FeO-poorchondrules (B&J, 1998, pp. 244–278). Unfortu-nately, few chondrites preserve a good record ofthe formation history of their metal grainsbecause subsequent low-temperature reactionscommonly formed oxides and sulfides and ther-mal metamorphism allowed kamacite to exsolvefrom taenite. Most refractory nuggets have beenstudied in Allende CAI, where rampant oxida-tion and sulfurization and redistribution ofvolatile tungsten and molybdenum compoundshave severely compromised the high-tempera-ture nebular record (Palme et al., 1994).

Most grains of metallic Fe,Ni and troiliteoutside chondrules were probably lost fromchondrules when they were molten. The sizes ofprominent grains are roughly correlated withchondrule size (Skinner and Leenhouts, 1993;Kuebler et al., 1999; Schneider et al., 2003),

mes

cpx

200 µm

cpx

px

ol

ol

mespx

(a)200 µm

ol

cpx

50 µm(b)

Ni Kα

150 µm(e)

CC

25 µm

cpx

ol

ol

(c)

(d)

(f)

CC

mes

mes

Figure 25 Combined elemental map in magnesium (red), calcium (green), and Al Ka (blue) X-rays (a, c) andbackscattered electron images (b, d) of two barred or skeletal olivine chondrules, and a combined X-ray map(e) and NiKa scanning map (f) of a zoned Fe,Ni metal grain with an enclosed cryptocrystalline chondrule (CC)in the CBb chondrite HaH 237. The chondrules, which contain forsteritic olivine (ol), low-calcium pyroxene(px), high-calcium pyroxene (cpx), and mesostasis (mes), lack rims and relict grains and clearly formed in adust-free environment from total melts. Reproduced by permission of the University of Arizona on behalf of

The Meteoritical Society from Krot et al. (2002c).


but were probably modified by shock aftercompaction. Grains of metallic Fe,Ni in mostunshocked type 3–6 chondrites provide arecord of slow cooling at B1–1,000KMyr�1

through the temperature range B550–350 1C,when kamacite and taenite ceased to equili-brate (Wood, 1967). In most type 2 and 3.0–3.3chondrites, metallic Fe,Ni grains typicallycontain concentrations of 0.1–1% chromium,silicon, and phosphorus, which are not foundin type 4–6 chondrites, and reflect high-temperature processing prior to accretion.

The most pristine metallic Fe,Ni grains arefound in the CR–CH–CB clan (Krot et al.,2002c). Most metal grains with 48% Ni lack

coexisting low-nickel kamacite and high-nickeltaenite showing that these chondrites were notheated above B300 1C for more than a year(Reisener et al., 2000). Metal grains are notpresent in most chondrules in CH chondrites,which are cryptocrystalline, and are entirelyabsent from chondrules in CB chondrites.

1.07.5.7.1 CR chondrite metallic Fe,Ni

CR2 chondrites contain 5–8 vol.% metallicFe,Ni grains in type I chondrules and betweenthem, which are not associated with troilite.Grains outside chondrules were probably once

1612

84

0.40.20.0

NiCoCr

Con

cent

ratio

n (w

t.%)

0 100 200 300 400 500

0 100 200 300 400 500

0 100 200 300 400 500

(a)

(b)

PdOsIrPt

3.53.02.52.01.51.00.5

Ele

men

t/Fe/

Cl

Ele

men

t/Fe/

Cl Mo

RuRh

1

2

3

4

5

Distance (µm)(c)

Figure 27 Concentration profiles of siderophile el-ements in a radially zoned Fe,Ni grain in the CBb

chondrite, QUE 94411: (a) electron microprobedata; (b, c) trace element data from laser ablationICPMS (Campbell et al., 2001). The nickel, cobalt,and chromium profiles can be matched by nonequi-librium nebular condensation assuming an enhanceddust–gas ratio of B36� solar, partial condensationof chromium into silicates, and isolation of 4% ofcondensates per degree of cooling (Petaev et al.,2001). Concentrations of the refractory siderophileelements, osmium, iridium, platinum, ruthenium,and rhodium, are enriched at the center of the grainby factors of 2.5–3 relative to edge concentrations,which are near CI levels after normalization to iron.Reproduced by permission of University of Arizonaon behalf of The Meteoritical Society from Krot

et al. (2002c).

1

0.1

0.01

0.001Ca AI Ti Mg Si Cr Mn Na

SOCC

Ele

men

ts/S

i/re

lativ

e to

CI

Figure 26 Bulk concentrations of lithophile ele-ments normalized to CI chondrites and silicon inskeletal-olivine and cryptocrystalline chondrules in-cluding cryptocrystalline inclusions in metallic Fe,Niin two CBb chondrites (HH 237 and QUE 94411).Shaded regions show compositional range of chond-rules in other carbonaceous chondrites, excludingthe CH group chondrules that are closely related toCB chondrules. The wide range of refractory abun-dances in CBb chondrules ((0.02–3)�CI levels)appears to reflect fractional condensation withskeletal-olivine chondrules condensing at highertemperatures than cryptocrystalline chondrules.Both types are more depleted in moderately volatileelements than other carbonaceous chondrites. After

Krot et al. (2002c).


associated with chondrules. Metal grains havebulk concentrations of 4–15wt.% Ni, solarCo/Ni ratios and up to B1% chromium andphosphorus (Weisberg et al., 1993). Grains atchondrule cores have higher nickel concentra-tions than those at the rims (Figure 24c) and allgrains are depleted in moderately volatile ele-ments (Connolly et al., 2001). Kong and Palme(1999) noted that Renazzo metal grains arecovered with layers of carbon, which appear tohave evolved at low temperatures. Two kinds ofprocesses have been invoked to account for theorigin and composition of the metal grains: con-densation as nebular solids or recondensationduring chondrule formation (Grossman andOlsen, 1974; Weisberg et al., 1993; Connollyet al., 2001) and silicate–metal equilibration andreduction during chondrule formation (Scottand Taylor, 1983; Zanda et al., 1994). Giventhat the silicate portions of type I chondrules inCR2 chondrites have retained a record ofcondensation processes, it seems likely that themetallic portions have also, and that both kindsof processes were involved in making CR metal.

1.07.5.7.2 CH and CB chondrite metallicFe,Ni

Between 0.01% and 1% of the metal grains inCH chondrites (Meibom et al., 1999, 2000) and

all the zoned metal grains in CBb chondrites(Meibom et al., 2001; Campbell et al., 2001)have radial chemical zoning patterns that areconsistent with formation as solid nebular con-densates above 1,200K (Figure 27). Fractionalcondensation models suggest that grains inQUE 94411 formed under a variety of oxida-tion conditions with enhanced dust–gas ratiosof 10–40 times solar prior to vaporiza-tion (Petaev et al., 2001). Kinetic growthmodels suggest growth timescales of 1–4 daysat cooling rates of a few degrees per hour andremoval from hot gas before temperatures fellbelow B1,200K (Meibom et al., 2001). The tinyfraction of appropriately zoned grains in CHchondrites might have condensed in convective


updrafts in the nebula and been transportedoutward to cooler regions by convection(Meibom et al., 2000). However, the large frac-tion of zoned grains in CBb chondrites requiresa more efficient transport mechanism such as anX-wind (Meibom et al., 2001; Shu et al., 2001).

CB chondrites also contain relatively homo-geneous metal grains with micrometer-sizedtroilite droplets. Large grains in CBa chondriteshave near-solar Co/Ni ratios and 0.1–1 wt.% Crand P and were among the first metallic grainsto be tagged as nebular condensates (Newsomand Drake, 1979; Weisberg et al., 1990). Mod-erately volatile siderophile elements gallium,germanium, arsenic, and antimony are depletedby factors of up to 103 from CI levels (Campbellet al., 2002). Condensation models suggest thatthe precursor gas was enriched in metals by afactor of 107 above solar levels prior to con-densation. Hypervelocity impacts between as-teroids or disruption of molten planetesimalsmight provide an appropriate environment forthe condensation of CBa metal as molten sulfur-bearing droplets (Campbell et al., 2002; Rubinet al., 2003; Krot et al., 2005a). However, thepresence of CAIs in CBa chondrites and theintimate mixture of zoned and sulfur-bearingmetallic Fe,Ni grains in CBb chondrites suggestthat components from diverse sources wereaccreted to form the CB chondrites.

1.07.5.7.3 Troilite

Troilite in chondrites may have condensedbelow B650K by reaction of nebular gas withmetallic Fe,Ni, or crystallized from Fe–Ni–Smelts in chondrules or during secondary proc-essing in asteroids. Whether any metal grains inchondrites have preserved nebular signatures isquestionable as sulfur is readily mobilizedduring metamorphism and shock (Laurettaet al., 1997). Criteria for identifying nebularsulfide condensates (Lauretta et al., 1998) andigneous sulfides (Rubin et al., 1999; Kong et al.,2000) have been proposed. Rubin et al. (1999)found that 13% of the chondrules in Semarkona(LL3.0) contain metal-troilite nodules with igne-ous textures inside phenocrysts. They inferredthat these chondrules were heated only briefly inthe nebula, otherwise sulfur would have beenlost. Lauretta et al. (2001) infer that troilite andfayalite rims on metal grains in Bishunpur(LL3.1) formed in a totally different way atB1,200K in a dust-enriched nebula.

1.07.5.8 Matrix Material

Matrix material is best defined as the opti-cally opaque mixture of mineral grains 10 nm

to 5mm in size that rims chondrules, CAIs, andother components and fills in the intersticesbetween them (Scott et al., 1988). Matrix grainsare generally distinguished from fragments ofchondrules, CAIs, and other components bytheir distinctive sizes, shapes, and textures.Minerals found in matrices include silicates,oxides, sulfides, metallic Fe,Ni, and espe-cially in type 2 chondrites, phyllosilicates, andcarbonates (Table 5). Matrices are broadlychondritic in composition though richer in FeOthan chondrules and have refractory abun-dances that deviate more from bulk chondritevalues (McSween and Richardson, 1977; Brear-ley, 1996; Bland et al., 2005; Huss et al., 2005).Matrix typically accounts for 5–50 vol.% of thechondrite (Table 1).

In the least altered chondrites, matrix rims,which are also called fine-grained rims, aresimilar in chemical composition and mineralogyto the interstitial matrix material, though matrixrims may be finer-grained and lack clasticmaterial. Compositions of rims on chondrulesand rims on CAIs appear indistinguishable.Rims are not uniform in width (Figure 28), butrim thickness is correlated with chondrule sizewithin a chondrite group (Metzler et al., 1992;Paque and Cuzzi, 1997). However, matrix rimsare absent in the CH, CB, and K chondritegroups. Clearly, matrix rims around chondriticcomponents were acquired after the componentsformed and before final lithification of thesecomponents into a rock. (Note that some rimsaround altered mineral grains appear to haveformed in situ by cementation of interstitialmaterial during aqueous alteration.) Mostauthors envisage that the chondritic compo-nents acquired their rims by accreting dust in thenebula (e.g., Morfill et al., 1998; Liffman andToscano, 2000; Cuzzi et al., 2001; Cuzzi, 2004),though some argue that rims form in asteroidalregoliths (e.g., Symes et al., 1998; Tomeoka andTanimura, 2000; Trigo-Rodriguez et al., 2006).Although chondrules and CAIs may have ac-quired transient, fluffy rims in the nebula, theidentical nature of well-compacted rims onobjects that formed at different times and placessuggests that rim acquisition and compactionmay have been delayed until planetesimalaccretion.

Millimeter-to-submillimeter lumps of matrix-rich material are present in many chondrites.Lumps that are mineralogically similar tonearby matrix occurrences are probablypieces of rims or aggregates of interstitialmaterial. However, heavily altered matrix-rich lumps called dark clasts with distinctivemineralogies probably have a different origin(Section 1.07.5.8.10). In CH and CB chond-rites, all the matrix material is present as dark

Table 5 Mineralogy of chondrite matrices.

Group Phases Reference

CarbonaceousCI Serpentine, saponite, ferrihydrite, magnetite, Ca–Mg

carbonate, pyrrhotite1

CM Serpentine, tochilinite, pyrrhotite, amorphous phase,calcite

1

CR2 Olivine, serpentine, saponite, minor magnetite, FeS,pentlandite, pyrrhotite, calcite

2

CO3.0: ALHA77307 Amorphous silicate, Fo30–98, low-calcium pyroxene,Fe,Ni metal, magnetite, sulfides

3

CO3.1–3.6 Fayalitic olivine, minor phyllosilicates, and ferric oxide 4CV3 reduced Fayalitic olivine, low-calcium pyroxene, low-nickel

metal, FeS5

CV3 oxidized:Bali-like Fayalitic olivine, phyllosilicate, fayalite, Ca–Fe

pyroxene, pentlandite, magnetite

6

Allende-like Fayalitic olivine, Ca–Fe pyroxene, nepheline,pentlandite

7

Ungrouped C:Acfer 094 Amorphous silicate, forsterite, enstatite, pyrrhotite,

ferrihydrite, traces of phyllosilicate

8

Adelaide Fayalitic olivine, amorphous silicate, enstatite,pentlandite, and magnetite

9

OrdinaryLL3.0: Semarkona Smectite, fayalitic olivine, forsterite, enstatite, calcite,

magnetite10

H, L, LL 3.1–3.6 Fayalitic olivine, amorphous material, pyroxene, albite,Fe,Ni metal

11

OtherK3: Kakangari Enstatite, forsterite, albite, ferrihydrite, troilite, Fe,Ni

metal12

References: 1, Zolensky et al. (1993); 2, Endress et al. (1994); 3, Brearley (1993); 4, Brearley and Jones (1998); 5, Lee et al., (1996); 6, Kelleret al., (1994); 7, Scott et al. (1988); 8, Greshake (1997); 9, Brearley (1991); 10, Alexander et al. (1989a); 11, Alexander et al. (1989b);12, Brearley (1989).


clasts: there are no matrix rims or interstitialmaterial.

According to Anders and Kerridge (1988),‘‘controversy permeates every aspect of thestudy of matrix material, from definition ofmatrix to theories of its origin.’’ Althoughcontroversy remains, major progress has beenmade in identifying matrix components andunderstanding their origin. Since matrix iscommonly richer in volatiles than other com-ponents, it is commonly referred to as the ‘‘low-temperature component’’ of chondrites. Butthis is a misnomer as it is composed of diversematerials that formed under different condi-tions. Prior to detailed TEM studies and theacquisition of mineralogical and isotopicevidence for aqueous alteration on asteroids,all matrix minerals in carbonaceous chond-rites including magnetite, carbonates, andphyllosilicates were thought to have condensedin the solar nebula (see McSween, 1979; seeChapter 1.09). However, matrix minerals inmost chondrites are now thought to be a com-plex mixture of presolar materials and ne-bular condensates that were mixed with fine

chondrule fragments and experienced aqueousalteration and metamorphism (e.g., Scott et al.,1988). However, nebular mechanisms for for-ming magnetite and phyllosilicates are stillbeing studied (Hong and Fegley, 1998; Cieslaet al., 2003). Matrix investigations are handi-capped by the extremely fine-grained nature ofthe constituents and the difficulty of distingui-shing primary and secondary mineralogicalfeatures (Brearley, 1996). The fine grain size(commonly as small as 50–100 nm), high po-rosity, and permeability of matrix ensure that itis more susceptible to alteration by aqueousfluids and metamorphism in asteroids thanother chondritic components.

Most of the matrix data that we review arechemical and mineralogical as there is littleisotopic information about individual matrixgrains. Bulk oxygen-isotopic compositions formatrix samples commonly differ from thoseof associated chondrules (e.g., Scott et al.,1988), but lack of oxygen-isotope data forsamples of key chondrites and specific matrixcomponents severely limits inferences aboutmatrix origins.

CAI

(a)

(b)

AOA

chd

JSM−53001610 0100

100 JSM-5900

Figure 28 Backscattered electron images showingfine-grained matrix rims on (a) Ca–Al-rich inclusion(CAI), and (b) an amoeboid–olivine aggregate(AOA) in ALHA77307. The light-gray rims containFeO-rich silicate material and are crossed by darkcracks; arrows mark the outer edge of the rims.


Other reviews of chondrite matrices haveaddressed aqueous alteration in the matrices ofcarbonaceous chondrites (Buseck and Hua,1993; see also Chapter 1.09), matrix miner-alogies and origins (Scott et al., 1988; Husset al., 2005), possible relations between matrixand chondrules (Brearley, 1996), the diversemineralogy of matrices in different chondritegroups (B&J, 1998, pp. 191–244) and relationsbetween matrices of primitive chondrites andcometary silicates (Nuth et al., 2005). Here wefocus on geochemical and mineralogicalfeatures in type 1–3 chondrites that offer thebest insights into the origin of the componentsof matrix and chondrites and the relationshipbetween matrix and chondrules. Unfortunately,well-characterized matrix material from E3 andR3 chondrites is lacking.

Some important constituents of matrixmaterial are discussed elsewhere: presolargrains (see Chapter 1.02), carbon and organicphases (see Chapter 1.10), and noble gases (seeChapter 1.14). The abundances and nature ofthese constituents in chondrite matrices provide

valuable clues to the thermal and alterationhistory of specific matrix samples.

1.07.5.8.1 CI1 chondrites

CI1 chondrites contain neither chondrulesnor refractory inclusions and have generallybeen considered as virtually pure matrix mate-rial. They are composed almost entirely ofphyllosilicates, oxides, sulfides, carbonates, andother minerals (Table 5) that were once widelythought to be nebular condensates but are nowknown to be products of aqueous alteration(B&J, 1998, p. 192). CI1 chondrites contain65wt.% of serpentine–saponite intergrowths,which are too fine-scale for accurate analysis,10% magnetite, 7% pyrrhotite, 5% poorlycrystalline ferrihydrite (5Fe2O3 � 9H2O), andsmaller amounts of pentlandite and otherphases (Bland et al., 2002; B&J, 1998, p. 192).Carbonates constitute B5% of the meteorites:the largest have been dated by 53Mn–53Crisotope systematics at B6–20Ma after CAIformation (Endress et al., 1996).

CI1 chondrites also contain a small fractionof isolated olivine and pyroxene grains up to400mm in size with chemical, oxygen-isotopiccompositions, and rounded inclusions of metal-lic Fe,Ni and trapped melt indicating that theywere derived from chondrules (Leshin et al.,1997). Brearley and Jones (1998) estimatethe abundance of olivine and pyroxene ato1 vol.%, but X-ray diffraction studies byBland et al. (2002) indicate 7wt.% olivine. Thedifference probably reflects a high abundance ofolivine crystallites embedded in phyllosilicates(P. A. Bland, private communication). CI1chondrites also contain rare refractory grainswith oxygen-isotopic compositions comparableto those in CAIs (see Scott and Krot, 2001).

Why do the most highly altered chondritesprovide the best match with the composition ofthe solar photosphere? We have yet to answerthis question satisfactorily, but many factorsare probably responsible. First, chondrules thatare deficient in volatiles relative to matrix ma-terial were scarce in the CI precursor material.Second, the rocks appear to have been affectedby closed-system alteration. The prominentsulfate veins in CI1 chondrites appear to haveformed in fractures on Earth (Gounelle andZolensky, 2001) and not in the parent body,so it is likely that aqueous solutions did notpercolate through the rock. Third, CI1 chond-rites are unique fine-grained breccias as theyare entirely composed of diverse, altered lithicfragments less than a few hundred micrometersin size (Morlok et al., 2002). They appear to bevery representative samples of one or more


asteroids that were thoroughly scrambled.Fourth, they probably formed further fromthe Sun than other chondrite groups and there-fore acquired higher concentrations of carbon-aceous, volatile-rich material.

1.07.5.8.2 CM2 chondrite matrices

In CM2 chondrites, the interstitial matrixmaterial and the matrix rims around chond-rules are very similar mineralogically (Zolenskyet al., 1993). Matrices are dominated by phyllo-silicates, which are also present in chondrulesand CAIs, and formed as a result of aqueousalteration, but the location is controversial(see Chapter 1.09). Phyllosilicates are mostly10–100 nm in size and are largely serpentineswith diverse morphologies, degrees of crystal-linity, and compositions (Barber, 1981; Zolen-sky et al., 1993). Crystals of the Fe3þ -richserpentine, cronstedite, reach up to 1� 10 mmin size and may be intergrown with tochilinite,(Fe,Mg)(OH)2 � (Fe,Ni)S (see B&J, 1998,p. 202). Metzler et al. (1992), Bischoff (1998),and Lauretta et al. (2000) argue that alterationoccurred mainly in small precursor planet-esimals prior to accretion of the CM body.Hanowski and Brearley (2000, 2001) disagreepointing to the occurrence of cross-cutting veinsof alteration minerals, loss of calcium fromchondrule mesostases and deposition as carbon-ate in the surrounding matrix, and otherfeatures favoring in situ alteration on asteroids.

CM2 matrices also contain small olivinegrains down to 0.1 mm in size, which are closeto Fo100 in composition (Barber, 1981). Theseforsterites lack defects, are sometimes euhedral,and appear to have formed before the phyllo-silicates (Barber, 1981; B&J, 1998, p. 216).Euhedral Fa40–50 olivines are also embedded inphyllosilicate (Lauretta et al., 2000). Pyroxenesare rarer and are dominantly magnesium-rich,Fso2 (Brearley, 1995). Some olivines andpyroxenes are especially enriched in mangan-ese (1–2wt.% MnO) and are comparable incomposition to the low iron, high manganeseolivines and pyroxenes present in IDPs (Klocket al., 1989). These were called LIME silicates(low-iron, manganese enriched) by Klock et al.,who argued that they are nebular condensates.

As much as 15% of CM matrices may beamorphous silicate, which is associated withcomparable volumes of very poorly crystallinephyllosilicates (Barber, 1981). In ALH81002,5 vol.% of the matrix rims are composed ofamorphous regions a few micrometers acrossthat are rich in silicon, aluminum, magnesium,iron, and presumably oxygen, and contain5–40 nm grains of iron-rich olivines and

metallic Fe,Ni (Lauretta et al., 2000). Chizma-dia and Brearley (2003) found more abundantamorphous material in the CM2 chondrite,Y 791198. (Origins of matrix components arereviewed in Section 1.07.5.8.11.)

Rim thickness appears to be correlated withobject size (Metzler et al., 1992). However, rimsaround grains of calcite, tochilinite, and otheralteration products may be artifacts of altera-tion rather than preaccretionary features.Trigo-Rodriguez et al. (2006) infer that the‘‘primary accretionary rocks’’ of Metzler et al.(1992) did not form in the nebula but rather inasteroidal regoliths. Hua et al. (2002) comparedthe concentrations of trace and minor elementsin matrix rims around different types of chond-rules and other objects in two CM chondritesand found no relationship between bulk com-positions of rims and enclosed objects. Abun-dances of moderately volatile elements in CMmatrices are depleted relative to CI chondrites,but enriched relative to bulk CM chondrites(Bland et al., 2005; Figure 29).

1.07.5.8.3 CR2 chondrite matrices

CR2 chondrite matrices are predomi-nantly composed of olivine with grain sizesof 0.2–0.3 mm accompanied by intergrowths ofserpentine and saponite with dimensions of20–300 nm and minor sulfide and magnetitegrains 0.1–25 mm in size (Zolensky et al., 1993;Endress et al., 1994). CR2 chondrites also con-tain B3 vol.% of more altered matrix lumpsmostly 0.1–1mm in size, which are similar inbulk composition to the other matrix material(Endress et al., 1994). However, phyllosilicates,magnetite, and chondrule fragments are moreabundant in the matrix lumps than in othermatrix material, and sulfides are richer in nickelsuggesting that the lumps are chondritic claststhat formed in a totally different environment.

1.07.5.8.4 CO3 chondrite matrices

The matrix of the type 3.0 CO chondrite,ALHA77307, is very different mineralogic-ally from matrices in CO3.1–3.6 chondrites(Table 5). It is largely composed of amorphoussilicate material (Figure 30), which forms dis-crete regions 1–5 mm in size, and 30–40 vol.% ofheterogeneously distributed tiny olivine crystalsand crystal aggregates (Brearley, 1993). Matrixrims around chondrules and refractory inclu-sions are fairly homogeneous (Figure 31), min-eralogically and chemically indistinguishable,and nearly isochemical with interstitial material(Brearley et al., 1995; Chizmadia et al., 2005).Rims are slightly lower in manganese and

2.0

1.0

0.1

CM chondrites

Bulk (Kallemeyn and Wasson, 1981)Matrix (Bland et al., 2005)

Abu

ndan

ce/C

I/Yb

Lu Ni Cr Mn Li As Au Cu K Ag Sb Ga Ge Rb Cs Bi Pb Zn Se In

2.0

1.0

0.1

Abu

ndan

ce/C

I/Yb

Lu Ni Cr Mn Li As Au Cu K Ag Sb Ga Ge Rb Cs Bi Pb Zn Se In2.0

1.0

0.1

Abu

ndan

ce/C

I/Yb

Lu Ni Cr Mn Li As Au Cu K Ag Sb Ga Ge Rb Cs Bi Pb Zn Se In2.0

1.0

0.1

Abu

ndan

ce/C

I/Yb

Lu Ni Cr Mn Li As Au Cu K Ag Sb Ga Ge Rb Cs Bi Pb Zn Se In


CV chondrites


CO chondrites

Bulk (Kallemeyn et al., 1994)

CR chondrites

Renazzo matrixAl Rais matrix

Elements in order of volatility

Figure 29 Elemental abundances in matrices of CM, CV, CR, and CO chondrites normalized to bulk CIchondrites using the refractory element Yb; data from Bland et al. (2005); figure from Huss et al. (2005).Matrices, like bulk chondrites, show increasing depletions with increasing elemental volatility. However,matrices are less depleted than bulk chondrites. Siderophile and chalcophile elements in CM, CO, and CVchondrites tend to have higher abundances than adjacent lithophiles, which are marked by gray bands.

Reproduced by permission of Astronomical Society of the Pacific from Huss et al. (2005).


OL

Pyx OL

0.2 µm

Kam

FeS

Figure 30 Transmission electron micrograph of thematrix in ALHA77307 (CO3.0) showing that it islargely composed of amorphous material with em-bedded crystallites of kamacite (Kam), pyrrhotite(FeS), olivine (OL), and low-calcium pyroxene(Pyx). Matrices of all carbonaceous chondritesprobably resembled this matrix material prior to as-teroidal alteration and metamorphism. Reproducedby permission of Elsevier from Brearley (1993).

10.00

1.00

0.10

0.01Al Ca Ti Mg Si Cr Mn Na

Ele

men

t/CO

cho

ndrit

e

Figure 31 Elemental abundances in matrix material infor rims on chondrules in ALHA77307 (type 3.0) obtaanalysis (Brearley et al., 1995) and for bulk matrix in Oactivation (data from Rubin and Wasson, 1988). Chondrbut deviate significantly from the bulk chondrite compo

Reproduced by permission of Elsev


calcium than interstitial matrix (Table 6). As inother chondrites, matrix material is broadlychondritic in composition but aluminum, nick-el, and potassium are enriched relative to sil-icon and CI chondrites, and calcium andtitanium are depleted. Moderately volatile ele-ments are typically enriched relative to bulkCO chondrite by factors of 2–3 (Figure 31; seealso Bland et al., 2005).

The amorphous phase in the ALHA77307matrix is heterogeneous on the submicrometerscale but homogeneous on a 10 mm scale (Table6). It is mostly composed of silicon, iron,magnesium, and oxygen with relatively highconcentrations of aluminum, nickel, sulfur, andphosphorus. Its Ca/Al ratio is much lower thanthose in chondrule mesostases, which are closerto chondritic values, precluding an origin fromchondrule glass (Brearley, 1993). Brearley’schemical and textural studies also suggest thatthe amorphous component was not formed byshock melting of matrix material, as Bunchet al. (1991) proposed for the matrix rims. Insome areas, the amorphous phase appears toshow incipient formation of phyllosilicates.

Three types of olivine occur in the amorphousphase of ALHA77307 (Brearley, 1993): (1)abundant irregularly shaped and poorly crys-talline olivines o300 nm in size with Fa10–70;(2) well-crystallized, subhedral, larger forsteritecrystals 200 nm to 44mm in size (Fo95–98); and

K Ni Fe Cu Ga Ge Se Zn S

ALHA77307 rims

Rim 1

Rim 2

Rim 3

Rim 4

Rim 5

Ornans matrix

CO3 chondrites normalized to bulk chondrite: datained by synchrotron X-ray fluorescence microprobernans (type 3.3) analyzed by instrumental neutronule rims in ALHA77307 are uniform in compositionsition. Matrix in Ornans is relatively unfractionated.ier from Brearley et al. (1995).

Table 6 Composition of interstitial matrix and matrix rims on chondrules in the CO3.0 chondrite,ALHA77307, and amorphous matrix phases in ALHA77307 and Acfer 094.

Weight% Interstitial matrix 3areas ALHA77307

Matrix rims 8chondrules

ALHA77307

Amorphous phaseALHA77307

matrix (CO3.0)

Amorphous phaseAcfer 094 matrix(ungrouped)

No. anal. 36 106 8 6SiO2 33.6 (0.9) 33.3 (0.7) 53 (9) 47 (4)TiO2 0.07 (0.02) 0.06 (0.01) o0.04 NAAl2O3 4.4 (0.3) 4.7 (0.3) 5.2 (1.6) 3.8 (0.9)Cr2O3 0.43 (0.04) 0.36 (0.02) 0.44 (0.18) 0.7 (0.2)FeO 36.2 (2.5) 37.3 (2.1) 20 (6) 26 (3)MnO 0.27 (0.04) 0.16 (0.05) 0.24 (0.20) 0.6 (B0.2)MgO 16.7 (1.5) 15.8 (0.8) 8.3 (1.7) 10.1 (3)CaO 1.1 (0.2) 0.76 (0.12) 0.41 (0.38) 1.4 (0.4)Na2O 0.32 (0.06) 0.25 (0.06) NA B0.7K2O 0.11 (0.02) 0.14 (0.05) NA NANiO 3.4 (0.4) 3.9 (0.7) 4.7 (2.6) 5.2 (2.2)P2O5 0.32 (0.09) 0.25 (0.09) 3.6 (2.1) NASO3 3.4 (0.8) 3.0 (0.6) 4.5 (2.8) 4.0 (2.0)Total 100 100 100 100

Sources: Brearley (1993) and Greshake (1997).Note: Matrix analyzed by electron microprobe: amorphous phase by analytical transmission electron microscopy. Totals are normalizedto 100%. Figures in parentheses are 1s of means of analyzed regions for matrix and 1s of analyses for amorphous phase.NA, not analyzed.


(3) micrometer-sized, isolated crystals andsimilar-sized aggregates of 100 nm grains ofwell-crystallized manganese-rich forsterites,with up to 2wt.% MnO. Pyroxene, which isnot as abundant as olivine, is mostly poor incalcium and iron and also occurs in severalvarieties. Orthopyroxene forms angular crystals1–1.5mm in size and is intergrown with clinopyr-oxene in aggregates of 0.2–0.3mm crystals.High-resolution TEM images of an intergrow-th indicate cooling from the protopyroxenestability field above 1,000 1C at B1,000–5,000 1Ch�1 (Jones and Brearley, 1988).Manganese-rich enstatite crystals are B0.1mmin size, largely clinoenstatite, indicative of rapidcooling from above 1,000 1C at 41,000 1Ch�1,and are associated with manganese-rich olivinesin micrometer-sized aggregates.

Some amorphous regions in ALHA77307contain rounded grains of kamacite with4–5% Ni and nickel-bearing sulfides, whichare botho200 nm in size. Elsewhere, magnetiteoccurs as crystals as small as 50 nm andaggregates o10 mm in size. The close proxim-ity of magnetite-rich and metal-sulfide-richregions suggests that they formed separatelyunder very different conditions in the nebula(Brearley, 1993). From a weak correlationbetween the Mg/(MgþFe) ratios of the oli-vines and the adjacent amorphous material,Brearley (1993) proposed that the olivines hadcrystallized from the amorphous phase duringlow-temperature annealing. However, the verydifferent properties of the magnesian and fer-roan olivines suggest they formed under differ-ent conditions. Presumably, the larger, highly

crystalline forsterites crystallized at highertemperatures. Brearley (1993) argued againsta condensation origin for the iron-poor,manganese-rich silicates as proposed by Klocket al. (1989), and suggested that they formed byannealing of amorphous material.

Matrices in the more metamorphosed CO3.1–3.6 chondrites are unlike the ALHA77307matrix as they are dominated by submicrometergrains of FeO-rich olivine Fo40–60 with minorpyroxene, spinels and metallic Fe,Ni (see B&J,1998, pp. 217–220). In Lance and Ornans, thematrix, like the mesostasis in chondrules, alsocontains minor phyllosilicates and other alter-ation products. Lance matrix contains magne-sium-rich serpentine and poorly crystallinematerial between FeO-rich olivines (Kellerand Buseck, 1990). Hydrous ferric oxides arealso present on grain boundaries in Kainsazand Warrenton indicating incipient alteration.Matrix olivines in type Z3.4 chondrites havemore equilibrated compositions than those intypes 3.1–3.3 (B&J, 1998, p. 220). Ornansmatrix is close to bulk CO composition (Rubinand Wasson, 1988; Figure 31).

The evidence that CO3 chondrites form ametamorphic sequence (Scott and Jones, 1990;B&J, 1998, p. 38) suggests that they all origin-ally contained matrix material like that inALHA77307. The amorphous material inALHA77307, which shows only incipient for-mation of fayalitic olivine and phyllosilicates,could have been readily replaced by fayaliticolivine by further fluid-assisted metamorphismto produce the type 3.1–3.6 matrices. Additionalstudies are needed to see if the minor feldspathic

KABA

fa

fa

Ca−Fe px

(a)

(b)

50 µm

ALHA81258

50 µm

Ca−Fe px

Figure 32 Backscattered electron images showingmatrix material in the Bali-like oxidized CV3 chond-rite, Kaba (a) and the Allende-like, oxidized CV3chondrite, ALHA81258 (b). Kaba matrix containsfayalite (fa), Ca–Fe pyroxene, fine-grained, iron-richolivine, pentlandite, magnetite, and phyllosilicates.The ALHA81258 matrix lacks phyllosilicates andcontains coarse-grained laths of fayalitic olivine(light-gray) and patches of Ca–Fe pyroxene plus ne-pheline and magnetite. Both matrices lack the amor-phous silicate present in ALHA77307 and appear to

be highly altered.


component between FeO-rich olivines in types3.1–3.6 (B&J, 1998, p. 217) is related to theamorphous component in ALHA77307.

1.07.5.8.5 CK3 chondrite matrices

Matrices of CK4–6 chondrites are composedlargely of olivine, low-calcium pyroxene andplagioclase grains that are mostly 20–100 mm insize with minor magnetite and sulfide (B&J,1998, p. 239). CK3 chondrite matrices appearto resemble those in CV3 chondrites, which arediscussed below. Matrix textures in the KobeCK4 chondrite were interpreted by Tomeokaet al. (2001) to result from partial meltingduring shock metamorphism, although Kobeitself is only classified as a shock stage S2.

1.07.5.8.6 CV3 chondrite matrices

Matrix material in CV3 chondrites occurs asrims on chondrules and other components, aswell as interstitial material, which is coarsergrained. The volume of a matrix rim isapproximately equal to the volume of theenclosed chondrule (Paque and Cuzzi, 1997).Matrices are mineralogically diverse but aredominated by iron-rich olivine, Fa30–60, whichis coarser than in other chondrite groups. Itforms elongate grains up to 20 mm in length andsubhedral finer grains o100 nm to 3 mm in theinterstices (B&J, 1998, p. 220; Watt et al.,2006). Olivines contain planar defects andsmall inclusions of sulfide, metal or magnetite,and poorly graphitized carbon (Tomeoka andBuseck, 1990; Keller et al., 1994; Brearley,1999). Reduced CV3 chondrites appear to bethe least altered as they have abundant metallicFe,Ni rather than magnetite: minor matrixminerals include low-calcium pyroxene, metal-lic, Fe,Ni, and FeS. In the Bali-like oxidizedCV3 chondrites, matrices lack metal andcontain abundant phyllosilicates, fayalite(Fa495), calcium-rich pyroxene, pentlandite,and magnetite in addition to iron-rich olivine(Figure 32a). In the Allende-like oxidized CV3s,hydrated silicates and fayalite are rare or absent,but Ca–Fe-rich pyroxene, nepheline, and pent-landite are abundant, and magnetite is present(Krot et al., 1995, 1998a, b, c, 2001f; Brearley,1999; B&J, 1998, pp. 220–225) (Figure 32b).

Proponents of a nebular origin for the sec-ondary minerals in oxidized CV3 chondritesargue that fayalitic olivine condensed in thesolar nebula at high fO2

following evaporationof chondritic dust, and coated preexistingchondrules (Nagahara et al., 1988; Palme andFegley, 1990; Weinbruch et al., 1990). Kimuraand Ikeda (1998) infer that anhydrous

alteration minerals like nepheline in Allendeformed at B600–800 1C whereas estimatesfrom zoning in olivine point to maximumtemperatures of B400–550 1C. They, therefore,conclude that the anhydrous alterationoccurred prior to accretion of the chondriticcomponents. Brenker et al. (2000) who studiedaggregates of andradite and hedenbergite20–50 mm in size in the Allende matrix, inferthat the Ca–Fe-rich pyroxene cooled rapidlyfrom 41,050 1C at 410 1Ch�1. They arguethat a localized nebular environment is morecompatible with this thermal history than anasteroidal environment.

Evidence favoring progressive alteration inasteroids includes fayalite veins that cross-cutchondrules and matrices, replacement ofchondrules by fayalite-rich matrix material in


dark inclusions, and comparable alterationminerals in chondrules and matrices (Krotet al., 1995, 1998a, b, c, 2001f; Kimura andIkeda, 1998). In Bali-like CV3s, saponite re-places calcium-rich minerals in chondrules andCAIs; magnetite is replaced by fayalite andcalcium-rich pyroxene. In Allende-like CV3s,fayalitic olivine replaces low-calcium pyroxene;magnetite–sulfide grains in matrices andchondrules are replaced by fayalitic olivineand calcium-rich pyroxenes; nepheline and cal-cium-rich pyroxene form veins in and aroundCAIs and dark inclusions. Isotopic data alsofavor asteroidal sites for alteration. Coexistingmagnetite and fayalite in Mokoia show largemass fractionation in oxygen isotopes, consist-ent with aqueous alteration (Choi et al., 2000).Nepheline lacks evidence for 26Mg excesses in-dicating formation 43–5Myr after CAI for-mation (MacPherson et al., 1995). 53Mn–53Crisotope systematics suggest that fayalite formedB7–16Myr after CAIs (Hutcheon et al., 1998).

The mineralogy of CV3 chondrite matricesprior to alteration is not clear. The two reducedCV3 chondrites that have been studied by TEMare not ideal samples. In Leoville, which is mod-erately shocked, Nakamura et al. (1992) foundaggregates of rounded, 10–100nm olivine grainswith interstitial amorphous material, but attrib-uted this to shock. In Vigarano, which is a brec-cia of altered and unaltered material (Lee et al.,1996), oxygen-isotopic mapping of the matrixrevealed 16O-rich forsterite grains 2mm in sizethat may be fragments of AOA or related ob-jects, as well as similar-sized 16O-poor fayaliteand enstatite grains (Kunihiro et al., 2005b).Clearly this technique is essential for unravelingthe origins of matrix components.

1.07.5.8.7 Matrix of ungrouped C chondrites,Acfer 094, and Adelaide

These two chondrites are among the leastaltered and metamorphosed chondrites and likeALHA77307, their matrices contain presolarsilicates (Nagashima et al., 2004; Kobayashiet al., 2005). Acfer 094, which is closely linked tothe CO and CM groups, has abundant matrixmaterial and higher concentrations of presolarSiC than any other chondrite (Newton et al.,1995). Its matrix is composed of 30 vol.%forsterites 10–300 nm in size and Fa0–2 incomposition, 40% amorphous material, and20% enstatite with Fs0–3 and grain sizes of200–300 nm (Greshake, 1997). The enstatitesshow intergrowths of ortho- and clinoenstatiteindicating rapid cooling at B1,000 1Ch�1. Minorcomponents included 5% sulfides, o1%phyllosilicates, and o2% ferrihydrite.

The highly unusual combination of amor-phous material and forsterite in Acfer 094matches that in the CO3.0 chondrite,ALH77307. The amorphous phases in thetwo chondrites are compositionally similar(Table 6) and both contain submicrometerenstatites and forsterites with high concentra-tions of MnO (0.6–2wt.%; Greshake, 1997).These concentrations are higher than inchondrules and comparable to those in high-manganese, low-iron silicates reported inIDPs by Klock et al. (1989). Acfer 094, likeALH77307, contains o1% phyllosilicate inthe amorphous material and 100–300 nmsulfide grains (pyrrhotite and minor pent-landite) as in ALHA77307, but it lacksmatrix metal grains. Since ferrihydrite ispresent in the amorphous material in Acfer094, metal may once have been present.Greshake (1997) proposed that the amorphousmaterial was a presolar or solar condensateand that the FeO-poor silicates formed in thesolar nebula by condensation or annealing ofamorphous material.

Adelaide, which also has links to CO andCM groups, has a matrix that lacks phyllosil-icates and is dominated by fayalitic olivineFa31–84 occurring as aggregates of 100–200 nmcrystals and platy and corroded single crystals(Brearley, 1991). Amorphous material, whichis rich in iron and sulfur, is present but isless abundant than in ALHA77307. High-manganese enstatite forms 1–1.5 mm clustersof 100–500 nm crystals but forsterites are absent.Pentlandite forms aggregates of 10–30 nmcrystals and magnetite occurs as 400–6,500 nmclusters of 100 nm grains.

1.07.5.8.8 H–L–LL3 chondrite matrices

Matrix material in type 3 ordinary chond-rites forms rims on chondrules, metal grains,and other components and occurs as lumps(B&J, 1998, pp. 231–237). In the LL3.0 chond-rite, Semarkona, matrix consists largely ofphyllosilicates formed by aqueous alterationwith fine-grained olivine Fa5–38 and low-calciumpyroxene, calcite, and magnetite (Alexanderet al., 1989a). Magnetite grains show large massfractionation of oxygen isotopes suggesting thata limited supply of aqueous fluid was largelyconsumed during asteroidal alteration (Choiet al., 1998). Klock et al. (1989) identified anadditional population of forsterites Fao1 andenstatites Fso2 having MnO concentrations of0–1.4wt.% that are correlated with FeO. Thehigh-manganese silicates are comparable incomposition to the iron-poor, manganese-richsilicates identified in IDPs (Klock et al., 1989).


Matrices in type 3.1–3.6 chondrites arelargely composed of FeO-rich olivines, whichmay have skeletal or elongate morphologies,and an amorphous feldspathic phase (Nagahara,1984; B&J, 1998, pp. 231–237). The origins ofthese phases are uncertain. Alexander et al.(1989b) argued that interchondrule matrix andmatrix rims in ordinary chondrites were derivedfrom broken chondrules, contrary to Brearleyet al. (1989), who suggested an origin by anneal-ing of amorphous material. Ikeda et al. (1981)identified two types of matrix in ordinary chond-rites with abundant amorphous material, onecontained Fe,Ni droplets 10–100nm in size.They speculated that the amorphous materialformed by shock. Fayalitic olivine (Fa54–94)occurs in the matrices of Krymka (LL3.1) andBishunpur where it rims metal grains (Weisberget al., 1997; Lauretta et al., 2001). These authorsargued that it formed in the solar nebula, notduring aqueous or hydrothermal alteration.Strong evidence for asteroidal alteration ispresent in Tieschitz (H3.6), which contains ne-pheline and albite in veins between chondrulesand in chondrule mesostases (Hutchison et al.,1998). This so-called white matrix formed whenaqueous fluids attacked chondrule mesostasesforming voids.

1.07.5.8.9 K3 chondrite matrix (Kakangari)

Kakangari has many unusual properties aschondrules and matrix both contain enstatiteand forsterite, and matrix silicates are moremagnesian than those in chondrules (Weisberget al., 1996; B&J, 1998, p. 82). Matrix andchondrules have similar bulk chemical compo-sitions though matrix is richer in 16O (Weisberget al., 1996).

Kakangari matrix contains mostly euhedralto subhedral crystals with grain sizes of 200 nmto 1.5mm, which occur as individual crystals andpolycrystalline units or clusters 2–8mm in size,some of which have well-annealed textures. Itis composed of 50 vol.% enstatite, Fs2–5, B20vol.% forsterite, Fa0–3, 10 vol.% albite crystals250–1,000 nm in size, Oro3, and o1 vol.%anorthite up to 500 nm in size. Minor phasesinclude chromium-spinel, troilite grains 1–5mmin size and B10 vol.% ferrihydrite, which formsacicular aggregates up to 2mm in length inpores. The lower FeO concentrations in olivinerelative to coexisting pyroxene in matrix andchondrules as in EH3 chondrites, suggest thatreduction occurred during cooling as olivinediffusion rates are faster than those in pyroxene(Scott and Taylor, 1983; B&J, 1998, pp. 72–74).

Olivine occurs mainly as inclusions in ensta-tite and is also found in albite, spinel, and

troilite (Brearley, 1989). At least three kinds ofpolycrystalline clusters appear to have formedseparately: (1) enstatite-rich units with forsteriteinclusions, (2) coarser versions with albite andolivine inclusions, and (3) forsterite–anorthiteunits with no enstatite. Enstatite crystals areintergrowths of ortho- and clinopyroxene andtheir microstructures indicate cooling from41,000 1C at B1,000 1Ch�1. Brearley (1989)suggests that the Kakangari matrix formed fromamorphous or partly crystalline particleso10mm in size that were annealed at 1,100–1,200 1C or possibly higher, and then rapidlycooled in an hour. The chondrules in Kakangaricould have formed from similar material thatwas heated to higher temperatures, partlymelted, and quenched at comparable rates,provided that the chondrules acquired lower16O concentrations when molten.

The bulk composition of Kakangari suggeststhat the hypothesized precursor material forthe chondrules and matrix had near-CI levels ofmoderately volatile elements that were not lostduring high-temperature processing. Reductionprobably occurred during this processing.

1.07.5.8.10 Heavily altered, matrix-rich lithicclasts

Heavily altered, matrix-rich chondritic clastsare very different from the most common lithicclasts in chondrites, which are fragments ofchondrules and CAIs or lithic fragmentsresembling the host rock. They are also unlikerare foreign chondritic clasts that are shockedand appear to have impacted the parent bodyat high speed after accretion, e.g., an LLchondrite clast in an H chondrite (Lipschutzet al., 1989).

Altered, matrix-rich chondritic clasts aresimilar in size to chondrules and CAIs in thehost chondrite, unshocked, and appear to havebeen gently incorporated into the host materialat an early stage. Prior to incorporation, theywere aqueously altered in a different location.They occupy a few volume% of CR2 (seeSection 1.07.5.7.3), CV3, CO3, CH, and CBchondrites. In CV3s they tend to have fewerand smaller chondrules than the host chond-rites and some are rimmed by matrix material(Kracher et al., 1985; Johnson et al., 1990).Although Palme et al. (1989) and Weisberg andPrinz (1998) argued that some are unalteredaggregates of nebular condensates, there is nowabundant evidence favoring formation byasteroidal alteration (Kojima et al., 1993;Kojima and Tomeoka, 1996; Krot et al.,1997, 1998a, b, c, 2001e). I–Xe ages of someCV3 clasts confirm that they were altered


before their hosts (Pravdivtseva et al., 2003).Dark clasts in CO3 chondrites are smaller thanin CV3 chondrites (50–200 mm; cf. millimeter tocentimeter in CV3s) (Itoh and Tomeoka, 2003).In CH3 chondrites and CB3 chondrites, matrix-rich chondritic clasts are the only matrix mate-rial and they are heavily altered to type 1 or2 levels, unlike the chondrules and metal grains,which are pristine (Greshake et al., 2002). In CHchondrites these clasts are mostly 5–200mm insize and have a size distribution comparable tothat of the chondrules (Scott, 1988), suggestingthat sizes of dark clasts are correlated withchondrule size.

Some authors infer that matrix-rich clasts aremerely pieces of the body that supplied thehost chondrite (e.g., Itoh and Tomeoka, 2003).However, their unusual mineralogy, limitedsize ranges, which are correlated with chond-rule size, and matrix rims on some clasts favora separate origin. The rims suggest they werelofted after alteration into the dusty environ-ment that formed rims on chondrules and othercomponents. Most matrix-rich clasts may befragments of bodies that formed before theparent body of the host chondrite, and werethen accreted with another batch of chondrulesand CAIs to form a later generation ofchondritic asteroids. Such planetesimals havebeen invoked by Metzler et al. (1992) andBischoff (1998). Early formed, volatile-richplanetesimals may have been disrupted by exp-losive volatile release (Wilson et al., 1999) orimpact fragmentation and dispersal.

The unique chondrite, Kaidun, appears tobe a breccia composed almost entirely of milli-meter- and submillimeter-sized fragments ofdiverse kinds of chondrites (Zolensky andIvanov, 2003). Kaidun may be a clast-rich,chondrule-poor chondrite that formed in anebula region where chondrules were muchrarer than clasts (Scott, 2002). CI chondritesand the Tagish Lake chondrite also appear tohave formed from submillimeter particles ofaltered chondritic material (Nakamura et al.,2003). Thus disagregation of early-formedplanetesimals may have been widespread.

1.07.5.8.11 Origins of matrix phases

Unaltered carbonaceous chondrite matricesare composed largely of amorphous silicate andsubmicrometer forsterites and enstatites andmicrometer-sized aggregates of such crystals(Scott and Krot, 2005a; Nuth et al., 2005). Inaddition, they contain up to a few volume per-cent of carbonaceous material, sulfides, partsper million levels of presolar silicate and oxidegrains (see Chapter 1.02), and solar refractory

grains. Micrometer-sized corundum, spinel,and hibonite grains of solar-system origin aremuch more common than presolar grains (e. g.,Virag et al., 1991; Strebel et al., 2000; Nittleret al., 1994) and probably formed as dust in theCAI-forming region (Scott and Krot, 2005a),or elsewhere when CAIs were forming.

Not enough is known about unalteredmatrices in ordinary and enstatite chondritesto infer their detailed mineralogy. However,enstatite chondrite matrices probably resemblethe matrix of Kakangari in having enstatite,metallic Fe,Ni, and sulfide as important con-stituents in addition to parts per million levelsof presolar grains (Ebata et al., 2006).

Amorphous ferromagnesian silicate is bestpreserved in the matrices of ALHA77307,Acfer 094, and Adelaide, though it is absentin Kakangari. It is more fractionated than bulkmatrix with higher Al/Ca, Ni/Si, and P/Si thanCI chondrites. Metallic Fe,Ni is abundant onlyin the ALHA77307 matrix but it may havebeen altered in the other chondrites. Relatedamorphous material may be present inCM, CO, and ordinary chondrites. The alumi-num-rich nature of amorphous material mayaccount for the high Al/Ca ratio of somematrix materials (Figure 27), the depletion ofrefractory elements with respect to aluminumin matrix materials in ordinary chondrites(Alexander, 1995), and the low Ca/Al ratio ofcertain chondrites like Adelaide and Kakangari(Brearley, 1991). The cause of the enrichmentof aluminum, nickel, and phosphorus in amor-phous matrix is uncertain as these elementshave different geochemical properties. Thisamorphous material did not form fromquenched melt from chondrule mesostases, byimpact in situ or elsewhere, or by dehydrationof phyllosilicates. It appears to have accretedinto chondrule rims as small particles (1–5 mmin size in ALHA77307).

Forsterites Fao5 and enstatites Fao5, whichare 10–300 nm in size, are found in the matricesof CM chondrites, ALHA77307, Acfer 094,and probably also in Semarkona, for whichgrain sizes were not reported. Enstatites andforsterite grains with higher FeO concentra-tions tend to be MnO-rich (0.6–2wt.%).Intergrowths of ortho- and clinoenstatite andforsterite–enstatite associations indicate thatall magnesian silicates cooled from 41,300Kat B103 to 104 1Ch�1 as micrometer- and sub-micrometer-sized grains prior to accretion.Oxygen-isotopic mapping of micrometer-sizedgrains shows that primitive matrices containB10–100 ppm of presolar silicates (Nagashimaet al., 2004). Therefore, the vast majority ofiron-poor matrix silicates must have formed inthe solar nebula.


Primitive chondrite matrices are surprisinglysimilar to porous chondritic aggregate IDPsin that both contain amorphous material,FeO-poor olivines, and pyroxenes, some ofwhich are MnO-rich (Klock et al., 1989), re-fractory oxides, and presolar silicate and oxidegrains (Scott and Krot, 2005a; Nuth et al., 2005).Oort cloud and Jupiter-family comets alsocontain amorphous and crystalline FeO-poorolivines and pyroxenes (Hanner, 1999; Woodenet al., 2005; Harker et al., 2005), and the sameminerals are observed in disks around proto-stars (see Nuth et al., 2005). 16O-rich refractorymicrograins are present in micrometeoritesand chondritic aggregate IDPs suggesting ad-ditional links between chondrite matrices andparticles from the outer solar system (Greshakeet al., 1996; Engrand and Maurette, 1998;Engrand et al., 1999). The amorphous aggre-gates in ALHA77307 that are rich in metal andsulfide may be related to the glass with embed-ded metal and sulfide (GEMS) grains that are amajor constituent of porous chondritic IDPsand resemble amorphous interstellar grains(Bradley et al., 1999; Rietmeijer, 1998, 2002).However GEMS are richer in SiO2 and muchsmaller: 100–500nm in size with kamacite andsulfide grains that are o5nm in size, cf.o200 nm in ALH77307. GEMS may haveformed by amorphization of interstellar grainsby cosmic-ray irradiation (like amorphous rimson some lunar grains) or by shocks (Bradley,1994; Bradley et al., 1999; Chapter 1.26).However, nearly all GEMS have solar-likeoxygen-isotopic compositions and are morelikely to be solar-system products (Messengeret al., 2003). Amorphous silicate material isknown to condense around evolved stars (Hillet al., 2001), and the amorphous matrix mate-rial may be a solar-system analog.

Two high-temperature processes have beenproposed for magnesian silicates in matrices:disequilibrium condensation (Klock et al., 1989;Nagahara et al., 1988; Greshake, 1997) andannealing of amorphous condensate particles(Brearley, 1993; Greshake, 1997). Annealingexperiments suggest that magnesium-rich sili-cates can crystallize from amorphous conden-sates in a few hours at 1,050K, or a year ormore below 1,000K (Hallenbeck et al., 2000;Rietmeijer et al., 2002). If the magnesian sili-cates formed close to the protosun by annealingor condensation, they could have been loftedacross the disk by outflows powered by recon-necting magnetic field lines, turbulence, or radi-ation pressure (Liffman and Brown, 1996b; Shuet al., 1996, 2001; Nuth et al., 2002). Alter-natively, Harker and Desch (2002) suggest thatmagnesian silicates in comets formed at5–10AU by annealing caused by nebular

shocks. We infer that magnesian silicates wereubiquitous in the solar nebula at 2–10AU as aresult of high-temperature processing andquenching in hours or days (Scott and Krot,2005a). The similarity of the cooling rates in-ferred for the magnesian silicates in the matrixand for chondrules suggests that matrix silicatesand chondrules may have been formed in relatedevents.

Submicrometer fayalitic olivines are hetero-geneously distributed in the matrices of prim-itive chondrites. In Acfer 094 they are absentbut the ALHA77307 matrix contains 30–40vol.%, which probably formed in situ by low-temperature annealing of amorphous material(Brearley, 1993). Submicrometer-sized fayaliticolivines are also present in matrices of CO3.1–3.6, type 3 ordinary chondrites and the inter-stices of larger olivines in CV3 chondrites. Anorigin for these olivines by annealing of amor-phous material also appears plausible but theevidence is weak. In some instances (e.g., inALH77307) annealing and crystallization ofolivines may have started prior to accretion, asKakangari matrix grains were extensivelyannealed before chondrules and matrix accreted.However, for the coarse, abundant fayaliticgrains and pure fayalite in CV3 chondrites, anasteroidal origin is indicated as their formationrequired an aqueous fluid (Krot et al., 2004f ).Although fayalitic olivine has been inferred tohave been a major constituent of the solar neb-ula (e.g., Nagahara et al., 1994; Ozawa andNagahara, 2001), we find no firm evidence fromchondrite matrices that any ‘‘crystalline,’’ FeO-rich silicate formed in the solar nebula.

Chemical differences between matrix andchondrules suggest that nearly all chondrulescould not have formed from matrix material.For example, matrix material is typically richerin FeO than chondrules, and mean refractoryelemental abundances deviate further from CIlevels than for chondrules. Some authors arguefrom chemical differences that chondrules andmatrices are chemically complementary andtherefore formed in the same location. Forexample, Fe/Si ratios in chondrules and matrixare lower and higher, respectively, than CIratios whereas the bulk compositions of CM,CO, CV, and other chondrites are close to solarimplying that chondrules and matrix werechemically fractionated from one another(Wood, 1996a). The case for Fe/Si is not veryconvincing, however, as there is a wide range ofCI-normalized Fe/Si ratios from 0.4 in LLchondrites to Z2 in CH and CB chondrites.However, for other elements, a stronger case canbe made. CI-normalized, mean Ti/Al ratios formatrix and chondrules in the Renazzo CR2chondrite are 0.5 and 1.5, respectively, whereas

Heating Mechanisms in the Early Solar System 55

bulk Ti/Al ratios for CR and other chondritesare within 10% of CI levels (Klerner and Palme,2000). Palme and Klerner (2000) infer thatchondrules and matrix formed from a singlereservoir that behaved as a closed system duringcondensation and processing of refractoryelements. Bland et al. (2005) infer that matricesand chondrules have complementary enrich-ments and depletions of volatile siderophileand chacolphile elements due to exchange inthe nebular region where chondrules formed.

1.07.6 FORMATION AND ACCRETIONOF CHONDRITIC COMPONENTS

Several common themes have emerged fromthis review of chondritic components. Type 3chondrites, for example, are more affected byalteration than generally realized. Only a hand-ful of chondrites meet the minimal requirementsfor pristine chondrites, and none of these es-caped the imprint of asteroidal modification.Whether this alteration occurred in parent bod-ies, as we infer, or in the nebula, the alterationenvironment was quite different from that at thelocation where each component formed.

Mean sizes of chondrules, CAIs, amoeboidolivine inclusions, dark clasts, and possiblymetallic Fe,Ni grains tend to be correlatedamong the chondrite groups. Thus, in groupswith relatively small chondrules, all compo-nents tend to be relatively small. This sortingprocess occurred after all components formed,possibly during turbulent accretion triggeredby chondrule formation (Cuzzi et al., 2001).Aerodynamic sorting should cause metal parti-cles to accrete with larger chondrules, as particlesare sorted according to the product of theirradius and density (Kuebler et al., 1999).However, Akridge and Sears (1999) have alsoargued for aerodynamic sorting but they pro-posed that components were sorted in anasteroidal regolith during degassing.

Many different processes were involved inmaking each chondritic component. Unalteredchondrite matrices may contain at least five dif-ferent types of micrometer-to-nanometer-sizedcomponents, which formed in diverse environm-ents: amorphous FeO-rich silicate, forsterite andenstatite crystals, refractory grains, presolargrains, and carbonaceous material. Chondrulesformed by several nebular processes (closed-system melting, condensation, and possiblyevaporation) and at least one asteroidal proc-ess (impact melting in regoliths). CAIs arelargely fine-grained condensates or coarse-grained igneous objects that probably formedfrom fine-grained precursors. An exception to

this preference for complexity is provided bythe amoeboid olivine inclusions: all AOAscould have formed by the same basic process:nebular condensation with only minor subse-quent thermal processing. Aluminum-richchondrules may provide a second exception,at least within carbonaceous chondrites.

Each chondrite component appears to havebeen manufactured predominantly by briefhigh-temperature processes. Chondrules andigneous CAIs have textures indicating coolingin minutes or hours. Many enstatite grains inmatrices contain ortho–clino intergrowths like-those in chondrules, and cooled in less than afew hours. Metal grains in CH chondritescooled from high temperatures in days, at most.However, the generation of group II REEpattern characteristic of many fine-grained CAIsis one process that required longer timescales ofmany years (Palme and Boynton, 1993).

The high-temperature processes that formedchondritic components typically involved gas–solid or gas–liquid exchange over part of thetemperature range 1,200–2,000K. Rapid cool-ing in the formation environment or removal tocooler localities ensured an extraordinary diver-sity of mineralogical and chemical compositionsamong chondritic components. Chemicalfractionations during brief high-temperatureprocessing were so ubiquitous and extensivethat they may be largely responsible for at leastfour of the major chemical variations shown bychondrites: variations in refractory and moder-ately volatile elements, metal–silicate ratios, andforsterite–enstatite variations.

The search for nebular condensates amongchondrites has generated several interestingbut seemingly false leads including wollasto-nite needles (e.g., Kerridge, 1993) and blue-luminescing enstatite rims (Weisberg et al.,1994; Hsu and Crozaz, 1998). FeO-rich matrixsilicates are also unlikely to be nebular con-densates. They appear to have formed predom-inantly by asteroidal processes, thoughsome may have formed by nebular annealingof FeO-rich amorphous material.

In the final section, we review the heat sourcesthat may have operated in the early solar systemto produce such diverse chondritic components.

1.07.7 HEATING MECHANISMS IN THEEARLY SOLAR SYSTEM

A wide variety of components in chondritesand chondritic IDPs were processed at hightemperatures in the solar nebula over a periodof several million years, and several heatingmechanisms were probably involved. Possible


heating mechanisms for making chondruleshave been reviewed by Cassen (1996), Boss(1996), Rubin (2000), and Ciesla (2005), andfor CAIs and chondrules by Jones et al.(2000a). Most heating models have focusedon chondrule origins, but Shu et al. (1996,2001) address both CAI and chondrule forma-tion close to the protosun; Hood (1998) andDesch and Connolly (2002) consider mecha-nisms for separately melting ferromagnesianand refractory aggregates to make chondrulesand igneous CAIs. Richter et al. (2006) inferthat igneous CAIs may have experiencednebula conditions like those calculated forchondrule-forming shocks. Russell et al.(2005) conclude that refractory inclusionsformed at the inner edge of the accretion diskwhereas chondrules formed close to the sitewhere chondrites accreted.

Three kinds of processes have been invokedfor forming chondrules: melting of dust aggre-gates (generally considered to be the standardmodel), condensation of melts, or melts andcrystals (Ebel and Grossman, 2000; Krot et al.,2001d; Varela et al., 2005), and impactsinto solid or partly molten bodies (e.g., Symeset al., 1998; Sanders, 1996). As discussed above,we believe that all three processes formedchondrules, and that the standard modelcannot account for olivine- and pyroxene-richchondrules.

Rubin (2000) has reviewed models forforming chondrules and their consistencywith various petrologic constraints (Table 7):condensation, exothermic chemical reactions,jetting during nebular collisions between parti-cles, ablation in protoplanetary atmospheres,nebular lightning (Desch and Cuzzi, 2000), sup-ernova shock waves, aerodynamic drag heatingat the accretion edge of the disk, magnetic re-connection flares, gas dynamic shock waves(Section 1.07.7.1), bow shocks from earlyformed planetesimals (Hood, 1998; Weidens-chilling et al., 1998; Ciesla et al., 2004a), FUOrionis outbursts, bipolar outflows (Section1.07.7.2), gamma-ray bursts (Duggan et al.,2003), radiative heating (Eisenhour and Buseck,1995), and impacts into asteroidal regolithsand molten planetesimals (Section 1.07.7.3).Additional mechanisms include shocks in theenvelopes of giant protoplanets (Nelson andRuffert, 2005), current sheets in magneticallyactive nebula regions (Joung et al., 2004), andablation of 1–500m-sized planetesimals byshock (Genge, 2000). Note that the constraintslisted in Table 7 apply to what we have called,‘‘archetypal chondrules,’’ which are most closelyrepresented by type II chondrules. Chronologi-cal and magnetic constraints are not included inTable 7.

Below we review three models that are cur-rently popular: nebular shocks, jets and outflowsnear the protosun, and impacts on planetesimals.

1.07.7.1 Nebular Shocks

Nebular shock waves are discontinuitiesbetween hot compressed gas moving fasterthan the local sound speed ¼ 1:2 � 105

�ffiffiffiffiffiffiffiffiffiffiffiffiffiffi300=T

pcm s�1Þ and cooler, less dense, slowly

moving gas (Cassen, 1996). Particles overtakenby shocks are suddenly enveloped in a blast ofwind moving at several kilometers per second,which subjects the particles to frictional heating.Particles are also heated by radiation from theshock front before they are overtaken and byconduction from hot gas. Desch and Connolly(2002) find that cooling rates in the range 10–1,000 1Ch�1 can be produced without invokingshocks of unreasonable speeds. They studied arange of shock speeds (5–10 km s�1) and gasdensities and found conditions under whichchondrules would partially melt, conditions thatwould cause dust to evaporate, and at the highershock speeds, conditions that would evaporatechondrules. The shock model appears to be ca-pable of forming droplets of the correctsize (Connolly and Love, 1998; Susa andNakamoto, 2002; Miura and Nakamoto,2005), it seems consistent with many petrologicconstraints from archetypal chondrules (Table7), and predicts a correlation between coolingrate and the concentration of chondrules, whichappears consistent with compound chondrulestatistics (Desch and Connolly, 2002). Shockheating models have also been investigated byIida et al. (2001) and Ciesla and Hood (2002)and are reviewed by Desch et al. (2005). Con-straints on shock models from the presence ofiron–sufide inclusions in chondrules are dis-cussed by Uesugi et al. (2005).

A major problem for the shock model hasbeen to find a source of powerful, pervasive, andrepeatable shocks. Several have been proposed:clumpy material falling into the nebula (Tanakaet al., 1998), bow shocks from planetesimalsscattered by Jupiter (Hood, 1998; Weidenschil-ling et al., 1998), spiral-arm instabilities in thesolar nebula (Wood, 1996b), X-ray flares (Naka-moto et al., 2005), and close encounters withsibling protostars (Bally et al., 2005). Boss andDurisen (2005) infer that clumps and spiral armscould generate B10km s�1 shocks in the aster-oid belt (Chapter 1.04).

1.07.7.2 Jets and Outflows

Two types of jet or outflow models have beenproposed. Liffman and Brown (1996a, b)

Table 7 Review by Rubin (2000) of chondrule-formation models and their consistency with petrologic and geochemical constraints.

Oxygenisotopes

Incompletemelting

Evaporationduring long-term heating

Relictgrains andmoderatevolatiles

Rapidcooling

Chondruleabundance

Chondrulesizea

Multipleheating

Micro-chondrulesin rims

Primarytroilite

Condensation � ? / ? � þ � � � þExothermic chemical reactions � þ þ þ � � � � � �Jetting ? þ þ þ þ � � � � þMeteor ablation ? þ þ þ þ � � � � þNebular lightning þ þ þ þ þ þ þ þ þ þSupernova shock waves þ þ þ þ þ � � � � þAerodynamic drag heating � þ þ þ þ ? � � � þMagnetic reconnection flares þ þ þ þ þ þ þ þ þ þGas dynamic shock waves þ þ þ þ þ þ þ þ þ þPlanetesimal bow shocks ? þ þ þ þ þ ? þ ? þFU Orionis outbursts ? � � � � � � � � �Bipolar outflows � � � � � þ � � � �Gamma-ray bursts þ þ � þ þ þ ? � � þRadiative heating þ þ þ þ þ þ þ þ þ þPlanetesimal collision � � � � � þ � � � �Impact-melting in parent-bodyregolith

� þ þ þ � � � ? � þ

aChondrule size distribution within a group is much narrower than the total variation shown by all chondrites.Note: þ , consistent; �, inconsistent; ?, possibly consistent after ad hoc modifications; /, not applicable.


suggested that chondrules formed in a bipolaroutflow by ablation of planetesimals and werethen injected into the asteroid belt. Size sortingof metal grains and chondrules can occur in jet-flow models (Liffman, 2005), and the observedcorrelation between rim thickness and chond-rule size may have arisen when chondrulesreentered the dusty nebula at hypersonic speeds(Liffman and Toscano, 2000).

Shu et al. (1996, 2001) proposed that chond-rules and CAIs formed at the inner edge of theprotostellar disk (see Chapter 1.04). About0.05AU from the protosun, the rotation rate ofthe disk matches that of the protostar and thedisk is terminated. Inflowing gas is ionized andeither whipped into the protostar by the star’smagnetic field, or ejected in the bipolar out-flows, while rocks spiral inward until theyevaporate or are launched by the winds. CAIsare thought to form during quiescent condi-tions, whereas chondrules form when theposition of the inner edge of the disk fluctu-ates due to intense flares from the protosun.Shu and colleagues suggest that CAIs andchondrules are hurled outward by the so-calledX-wind, which they infer powers the bipolaroutflows. CAIs and chondrules fall back ontothe disk at several astronomical units or beyondand are available for accretion into asteroidsand planets or recycling back to the inner edgeof the disk. Smaller particles fall back at greaterdistances from the protosun. Detailed thermalhistories for chondrules and CAIs have not beencalculated for this model.

1.07.7.3 Impacts on Planetesimals

Four kinds of impacts have been proposedto form chondrules: hypervelocity impactsbetween solid bodies that melt regolith mate-rial (e.g., Symes et al., 1998; Sears, 2005;Bridges et al., 1998), hypervelocity impacts onVesta-sized and larger, partly molten bodies(Hutchison et al., 2005), impacts onto alreadymolten planetesimals at lower speeds thatpuncture the planet spewing forth molten dropsof silicate and metal (Sanders, 1996; Sandersand Taylor, 2005), and specifically for CBchondrites, an impact between undifferentiatedMoon-to-Mars-sized bodies that generated aplume of melt and vapor (Rubin et al., 2003;Krot et al., 2005a).

Many factors suggest that the vast majorityof chondrules could not form by impact inregoliths (Taylor et al., 1983). Molten dropletsare rare on the regolith of Vesta, the presumedparent of the eucrites, and meteorite evidencesuggests that chondrule-free asteroids did notform in the inner part of the asteroid belt

(Scott, 2002). (See Sears, 1998 for a contraryview.) However, the chromite-rich chondrulespresent in brecciated type 4–6 ordinary chond-rites (Krot and Rubin, 1993) and perhaps a fewpercent of other chondrules in H, L, and LLchondrites probably did form this way after thechondritic bodies had accreted. Since impactmelt will be ejected at speeds of several kilo-meters per second, much faster than the escapevelocity of the asteroid (less than a few hundredmeters per second), most impact melt will notre-accrete to the target asteroid. Impact meltshould be rarer on the surface of asteroids thanon the Moon (Taylor et al., 1983).

Any asteroid more than B15 km in radiusthat accreted o1–1.5Myr after CAIs formedwould have melted o3Myr later, given theevidence that 26Al was relatively homogeneou-sly distributed with an 26Al/27Al ratio ofB5� 10�5 when CAIs formed (e.g., Yoshinoet al., 2003; Hevey and Sanders, 2006). Theparent asteroids of differentiated asteroidsappear to have formed at this stage before thechondrite parent bodies, and several sufferedcatastrophic impacts during or soon after igne-ous differentiation (Keil et al., 1994). Sanders(1996) argues that chondrules formed in majorimpacts on molten planetesimals with thincrusts, and that convection was vigorousenough to prevent gravitational segregationof melted Fe,Ni droplets to the core (see alsoYoshino et al., 2003; Sanders and Taylor,2005). This model envisages that moltenmetal–silicate droplets were dispersed into thesolar nebula and that they accreted long afterthey had cooled and mixed with impact drop-lets from other asteroidal impacts, nebula dust,CAIs, and material from the cool unmeltedcrust of the target asteroids. The advantages ofthis model are that abundant chondrules aregenerated, the model readily accommodatesthe several Ma interval between CAI andchondrule formation, and, provided gas lossoccurred, it could explain why many chond-rules lost volatiles (Lugmair and Shukolyukov,2001). Issues that require further study arediscussed by Sanders and Taylor (2005).

ACKNOWLEDGMENTS

This research has made use of NASA’sAstrophysics Data System and was supportedin part by NASA grants NAG5-11591 andNAG5-11591 (K. Keil, P. I.) and NAG5-10610and NNG04GH02G (A. Krot, P. I.). We thankA. Davis and an anonymous reviewer for manyhelpful comments on the original chapter. Thisis SOEST publication #6882 and HIGPpublication #1460.

References 59

REFERENCES

Akaki T. and Nakamura T. (2005) Formation processes ofcompound chondrules in CV3 carbonaceous chondrites:constraints from oxygen isotope ratios and major ele-ment concentrations. Geochim. Cosmochim. Acta 69,2907–2929.

Akridge G., Benoit P. H., and Sears D. W. G. (1998)Regolith and megaregolith formation of H-chondrites:thermal constraints on the parent body. Icarus 132,185–195.

Akridge D. G. and Sears D. W. G. (1999) The gravitationaland aerodynamic sorting of meteoritic chondrules andmetal: experimental results with implications for chondri-tic meteorites. J. Geophys. Res. 104, 11853–11864.

Aleon J., Krot A. N., and McKeegan K. D. (2002a) Cal-cium–aluminum-rich inclusions and amoeboid olivineaggregates from the CR carbonaceous chondrites. Me-teorit. Planet. Sci. 37, 1729–1755.

Aleon J., Krot A. N., McKeegan K. D., MacPherson G. J.,and Ulyanov A. A. (2002b) Oxygen isotopic compositionof fine-grained Ca–Al-rich inclusions in the reduced CV3chondrite Efremovka. In Lunar Planet. Sci. XXXIII,#1426. The Lunar and Planetary Institute, Houston (CD-ROM).

Alexander C. M. O’D. (1995) Trace element contents ofchondrule rims and interchondrule matrix in ordinarychondrites. Geochim. Cosmochim. Acta 59, 3247–3266.

Alexander C. M. O’D. (2005) From supernovae to planets:the view from meteorites and interplanetary dust parti-cles. In Chondrites and the Protoplanetary Disk (eds. A.N. Krot, E. R. D. Scott, and B. Reipurth). Astron. Soc.Pacific Conf. Ser., vol. 341, pp. 972–1002.

Alexander C. M. O’D., Barber D. J., and Hutchison R.(1989a) The microstructure of Semarkona and Bishun-pur. Geochim. Cosmochim. Acta 53, 3045–3057.

Alexander C. M. O’D., Boss A. P., and Carlson R. W.(2001) The early evolution of the inner solar system: ameteoritic perspective. Science 293, 64–68.

Alexander C. M. O’D. and Grossman J. N. (2005)Alkali elemental and potassium isotopic compositionsof Semarkona chondrules. Meteorit. Planet. Sci. 40,541–556.

Alexander C. M. O’D., Grossman J. N., Wang J., ZandaB., Bourot-Denise M., and Hewins R. H. (2000) The lackof potassium-isotopic fractionation in Bishunpur chond-rules. Meteorit. Planet. Sci. 35, 859–868.

Alexander C. M. O’D., Hutchison R., and Barber D. J.(1989b) Origin of chondrule rims and interchondrulematrices in unequilibrated ordinary chondrites. EarthPlanet. Sci. Lett. 95, 187–207.

Alexander C. M. O’D. and Wang J. (2001) Iron isotopes inchondrules: implications for the role of evaporationduring chondrule formation. Meteorit. Planet. Sci. 36,419–428.

Amelin Y., Ghosh A., and Rotenberg E. (2005) Unravelingthe evolution of chondrite parent asteroids byprecise U–Pb dating and thermal modeling. Geochim.Cosmochim. 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.

Anders E. and Kerridge J. F. (1988) Future directions inmeteorite research. In Meteorites and the Early SolarSystem (eds. J. F. Kerridge and M. S. Matthews).University of Arizona Press, Tucson, AZ, pp. 1155–1186.

Artioli G. and Davoli G. (1994) Low-Ca pyroxenes fromLL group chondritic meteorites: crystal structural studiesand implications for their thermal histories. Earth Planet.Sci. Lett. 128, 469–478.

Bally J., Moeckel N., and Throop H. (2005) Evolution ofUV-irradiated protoplanetary disks. In Chondrites andthe Protoplanetary Disk (eds. A. N. Krot, E. R. D. Scott,

and B. Reipurth). Astron. Soc. Pacific Conf. Ser., vol.341, pp. 81–106.

Barber D. J. (1981) Matrix phyllosilicates and associatedminerals in C2M carbonaceous chondrites. Geochim.Cosmochim. Acta 45, 945–970.

Begemann F. (1980) Isotopic anomalies in meteorites. Rep.Prog. Phys. 43, 1309–1356.

Bell J. F., Davis D. R., Hartmann W. K., and Gaffey M. J.(1989) Asteroids: the big picture. In Asteroids II (eds. R. P.Binzel, T. Gehrels, and M. S. Matthews). University ofArizona Press, Tucson, AZ, pp. 921–945.

Bennett M. E. and McSween H. Y., Jr. (1996) Revisedmodel calculations for the thermal histories of ordinarychondrite parent bodies. Meteorit. Planet. Sci. 31,783–792.

Binzel R. P., Lupishko D. F., Martino M. D., Whiteley R.J., and Hahn G. J. (2002) Physical properties of near-Earth objects. In Asteroids III (ed. R. P. Binzel). Univer-sity of Arizona Press, Tucson, AZ, pp. 255–271.

Bischoff A. (1998) Aqueous alteration of carbonaceouschondrites—a review. Meteorit. Planet. Sci. 33,1113–1122.

Bischoff A. (2000) Mineralogical characterization of prim-itive, type 3 lithologies in Rumuruti chondrites. Meteorit.Planet. Sci. 35, 699–706.

Bischoff A. (2001) Fantastic new chondrites, achondrites,and lunar meteorites as the result of recent meteoritesearch expeditions in hot and cold deserts. Earth MoonPlanet. 85, 87–97.

Bischoff A. and Keil K. (1984) Al-rich objects in ordinarychondrites: related origin of carbonaceous and ordinarychondrites and their constituents. Geochim. Cosmochim.Acta 48, 693–709.

Bischoff A., Keil K., and Stoffler D. (1985) Perovskite–hibonite–spinel-bearing inclusions and Al-rich chond-rules and fragments in enstatite chondrites. Chem. Erde44, 97–106.

Bischoff A., Palme H., Schultz L., Weber D., Weber H. W.,and Spettel B. (1993) Acfer 182 and paired samples, aniron-rich carbonaceous chondrite: similarities withALH85085 and relationship to CR chondrites. Geochim.Cosmochim. Acta 57, 2631–2648.

Bizzarro M., Baker J. A., Haack H., and Lundgaard K. L.(2005) Rapid timescales for accretion and melting of dif-ferentiated planetesimals inferred from 26Al–26Mg chro-nometry. Astrophys. J. 632, L41–L44.

Bland P. A., Alard O., Benedix G. K., Kearsley A. T.,Menzies O. N., Watt L. E., and Rogers N. W. (2005)Volatile fractionation in the early solar system andchondrule/matrix complementarity. Proc. Natl. Acad.Sci. 102, 13755–13760.

Bland P. A., Cressey G., Alard O., Rogers N. W., Forder S.D., and Gounelle M. (2002) Modal mineralogy of car-bonaceous chondrites and chemical variation in chond-rite matrix. In Lunar Planet. Sci. XXXIII, #1754. TheLunar and Planetary Institute, Houston (CD-ROM).

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

Bonal L., Quirico E., Bourot-Denise M., and MontagnacG. (2006) Determination of the petrologic type of CV3chondrites by Raman spectroscopy of included organicmatter. Geochim. Cosmochim. Acta 70, 1849–1863.

Boss A. P. (1996) A concise guide to chondrule formationmodels. In Chondrules and the Protoplanetary Disk (eds.R. H. Hewins, R. H. Jones, and E. R. D. Scott). Cam-bridge University Press, Cambridge, pp. 257–263.

Boss A. P. and Durisen R. H. (2005) Chondrule-formingshock fronts in the Solar Nebula: a possible unified sce-nario for planet and chondrite formation. Astrophys. J.621, L137–L140.

Bradley J. P. (1994) Chemically anomalous, preaccretion-ary irradiated grains in interplanetary dust particles fromcomets. Science 265, 925–929.


Bradley J. P., Keller L. P., Snow T. P., Hanner M. S., FlynnG. J., Gezo J. C., Clemette S. J., Brownlee D. E., andBowey J. E. (1999) An infrared spectral match betweenGEMS and interstellar grains. Science 285, 1716–1718.

Brearley A. J. (1989) Nature and origin of matrix in theunique type 3 chondrite, Kakangari. Geochim. Cosmo-chim. Acta 53, 2395–2411.

Brearley A. J. (1991) Mineralogical and chemical studies ofmatrix in the Adelaide meteorite: a unique carbonaceouschondrite with affinities to ALHA77307 (CO3). LunarPlanet. Sci. 22, 133–134.

Brearley A. J. (1993) Matrix and fine-grained rims in theunequilibrated CO3 chondrite, ALHA77307: origins andevidence for diverse, primitive nebular dust components.Geochim. Cosmochim. Acta 57, 1521–1550.

Brearley A. J. (1995) Aqueous alteration and brecciation inBells, an unusual, saponite-bearing, CM chondrite. Geo-chim. Cosmochim. Acta 59, 2291–2317.

Brearley A. J. (1996) Nature of matrix in unequilibratedchondrites and its possible relationship to chondrules. InChondrules and the Protoplanetary Disk (eds. R. H. He-wins, R. H. Jones, and E. R. D. Scott). CambridgeUniversity Press, Cambridge, pp. 137–151.

Brearley A. J. (1999) Origin of graphitic carbon and pent-landite inclusions in matrix olivines in the Allende me-teorite. Science 285, 1380–1382.

Brearley A. J., Bajt S., and Sutton S. (1995) Distribution ofmoderately volatile trace elements in fine-grained chond-rule rims in the unequilibrated CO3 chondrite, ALHA77307. Geochim. Cosmochim. Acta 59, 4307–4316.

Brearley A. J., Hutcheon I. D., and Browning L. (2001)Compositional zoning and Mn–Cr systematics in car-bonates from the Y791198 CM2 carbonaceous chondrite.In Lunar Planet. Sci. XXXII, #1458. The Lunar andPlanetary Institute, Houston (CD-ROM).

Brearley A. J. and Jones R. H. (1998) Chondritic meteor-ites. In Planetary Materials (ed. J. J. Papike). Rev.Mineral., vol. 36, pp. 3-1–3-398.

Brearley A. J., Scott E. R. D., Keil K., Clayton R. N.,Mayeda T. K., Boynton W. V., and Hill D. H. (1989)Chemical, isotopic and mineralogical evidence for theorigin of matrix in ordinary chondrites. Geochim. Cosmo-chim. Acta 53, 2081–2093.

Brenker F. E., Palme H., and Klerner S. (2000) Evidencefor solar nebula signatures in the matrix of the Allendemeteorite. Earth Planet. Sci. Lett. 178, 185–194.

Bridges J. C., Franchi I. A., Hutchison R., Sexton A. S.,and Pillinger C. T. (1998) Correlated mineralogy, chem-ical compositions, oxygen isotopic compositions and sizeof chondrules. Earth Planet. Sci. Lett. 155, 183–196.

Bunch T. E., Schultz P., Cassen P., Brownlee D., PodolakM., Lissauer J., Reynolds R., and Chang S. (1991) Aresome chondrule rims formed by impact processes?Observations and experiments. Icarus 91, 76–92.

Burbine T. H., McCoy T. J., Meibom A., Gladman B., andKeil K. (2002) Meteoritic parent bodies: their numberand identification. In Asteroids III (ed. R. P. Binzel).University of Arizona Press, Tucson, AZ, pp. 653–666.

Buseck P. R. and Hua X. (1993) Matrices of carbonaceouschondrite meteorites. Annu. Rev. Earth Planet. Sci. 21,255–305.

Busemann H., Bauer H., and Wieler R. (2002) Phase Q—acarrier for subsolar gases. In Lunar Planet. Sci. XXXIII,#1462. The Lunar and Planetary Institute, Houston.

Campbell A. J., Humayun M., Meibom A., Krot A. N.,and Keil K. (2001) Origin of zoned metal grains in theQUE94411 chondrite. Geochim. Cosmochim. Acta 65,163–180.

Campbell A. J., Humayun M., and Weisberg M. K. (2002)Siderophile element constraints on the formation ofmetal in the metal-rich chondrites Bencubbin, Weather-ford, and Gujba. Geochim. Cosmochim. Acta 66,647–660.

Campbell A. J., Zanda B., Perron C., Meibom A., andPetaev M. I. (2005) Origin and thermal history of Fe–NiMetal in primitive chondrites. In Chondrites and theProtoplanetary Disk (eds. A. N. Krot, E. R. D. Scott, andB. Reipurth). Astron. Soc. Pacific Conf. Ser., vol. 341,pp. 407–431.

Campins H. and Swindle T. D. (1998) Expected character-istics of cometary meteorites. Meteorit. Planet. Sci. 33,1201–1211.

Cassen P. (1996) Overview of models of the solar nebula:potential chondrule-forming events. In Chondrules andthe Protoplanetary Disk (eds. R. H. Hewins, R. H. Jones,and E. R. D. Scott). Cambridge University Press, Cam-bridge, pp. 21–28.

Cassen P. (2001a) Unresolved questions regarding theorigins of Solar System solids. Phil. Trans. Roy. Soc.Lond. A 359, 1935–1946.

Cassen P. (2001b) Nebular thermal evolution and theproperties of primitive planetary materials. Meteorit.Planet. Sci. 36, 671–700.

Chaussidon M., Libourel G., and Krot A. N. (2006) Originof olivines in type I chondrules: constraints from oxygenand magnesium isotopic compositions. In Lunar Planet.Sci. XXXVII, #1335. The Lunar and Planetary Institute,Houston (CD-ROM).

Chapman C. R. (2004) Space weathering of asteroidsurfaces. Annu. Rev. Earth Space Sci. 32, 539–567.

Chizmadia L. J. and Brearley A. J. (2003) Mineralogy andtextural characteristics of fine-grained rims in Yamato791198 CM2 carbonaceous chondrite: constraints on thelocation of aqueous alteration. In Lunar Planet. Sci.XXXIV, #1419. The Lunar and Planetary Institute, Hou-ston (CD-ROM).

Chizmadia L. J., Rubin A. E., and Wasson J. T. (2002)Mineralogy and petrology of amoeboid olivine inclusionsin CO3 chondrites: relationship to parent-body aqueousalteration. Meteorit. Planet. Sci. 37, 1781–1796.

Chizmadia L. J., Scott E. R. D., and Krot A. N. (2005)Fine-grained rims around chondrules and refractory in-clusions in ALHA77307 are compositionally similar.Meteorit. Planet. Sci. 40, A29.

Choi B. G., Krot A. N., and Wasson J. T. (2000) Oxygenisotopes in magnetite and fayalite in CV chondrites Kabaand Mokoia. Meteorit. Planet. Sci. 35, 1239–1248.

Choi B. G., McKeegan K. D., Krot A. N., and Wasson J.T. (1998) Extreme oxygen-isotope compositions inmagnetite from unequilibrated ordinary chondrites. Na-ture 392, 577–579.

Ciesla F. J. (2005) Chondrule-forming processes—an over-view. In Chondrites and the Protoplanetary Disk (eds. A.N. Krot, E. R. D. Scott, and B. Reipurth). Astron. Soc.Pacific Conf. Ser., vol. 341, pp. 811–820.

Ciesla F. J. and Hood L. L. (2002) The nebular shock wavemodel for chondrule formation: shock processing in aparticle-gas suspension. Icarus 158, 281–293.

Ciesla F. J., Hood L. L., and Weidenschilling S. J. (2004a)Evaluating planetesimal bow shocks as sites for chond-rule formation. Meteorit. Planet. Sci. 39, 1809–1821.

Ciesla F. J., Lauretta D. S., Cohen B. A., and Hood L. L.(2003) A nebular origin for chondritic fine-grainedphyllosilicates. Science 299, 549–552.

Ciesla F. J., Lauretta D. S., and Hood L. L. (2004b) Thefrequency of compound chondrules and implications forchondrule formation. Meteorit. Planet. Sci. 39, 531–544.

Clayton R. N. (1993) Oxygen isotopes in meteorites. Annu.Rev. Earth Planet. Sci. 21, 115–149.

Clayton R. N. (2002) Self-shielding in the solar nebula.Nature 415, 860–861.

Clayton R. N., Mayeda T. K., and Molini-Velsko C. A.(1985) Isotopic variations in solar system material: evap-oration and condensation of silicates. In Protostars andPlanets II (eds. D. C. Black and M. S. Matthews).University of Arizona Press, Tucson, AZ, pp. 755–771.

References 61

Cohen B. A. and Coker R. F. (2000) Modeling of liquidwater on CM meteorite parent bodies and implicationsfor amino acid racemization. Icarus 145, 369–381.

Connolly H. C., Jr. and Desch S. J. (2004) On the origin ofthe ‘‘kleine Kugelchen’’ called chondrules. Chem. Erde64, 95–125.

Connolly H. C., Jr. and Hewins R. H. (1995) Chondrules asproducts of dust collisions with totally molten dropletswithin a dust-rich nebular environment: an experimentalinvestigation. Geochim. Cosmochim. Acta 59, 3131–3246.

Connolly H. C., Jr., Huss G. R., and Wasserburg G. J.(2001) On the formation of Fe–Ni metal in Renazzo-likecarbonaceous chondrites. Geochim. Cosmochim. Acta 65,4567–4588.

Connolly H. C., Jr., Jones B. D., and Hewins R. H. (1998)The flash melting of chondrules: an experimentalinvestigation into the melting history and physical na-ture of chondrule precursors. Geochim. Cosmochim. Acta62, 2725–2735.

Connolly H. C., Jr. and Love S. G. (1998) The formation ofchondrules: petrologic tests of the shock wave model.Science 280, 62–67.

Cosarinsky M., Leshin L. A., MacPherson G. J. Krot A.N., and Guan Y. (2003) Oxygen isotopic composition ofCa–Fe-rich silicates in and around an Allende Ca–Al-rich inclusion. In Lunar Planet. Sci. XXXIV, #1043. TheLunar and Planetary Institute, Houston (CD-ROM).

Cuzzi J. N. (2004) Blowing in the wind. Part III: Accretionof dust rims by chondrule-sized particles in a turbulentprotoplanetary nebula. Icarus 168, 484–497.

Cuzzi J. N., Hogan R. C., Paque J. M., and DobrovolskisA. R. (2001) Size-selective concentration of chondrulesand other small particles in protoplanetary nebula tur-bulence. Astrophys. J. 546, 496–508.

Desch S. J., Ciesla F. J., Hood L. L., and Nakamoto T.(2005) Heating of chondritic materials in solar nebulashocks. In Chondrites and the Protoplanetary Disk (eds.A. N. Krot, E. R. D. Scott, and B. Reipurth). Astron.Soc. Pacific Conf. Ser., vol. 341, pp. 849–872.

Desch S. J. and Connolly H. C., Jr. (2002) A model of thethermal processing of particles in solar nebula shocks:application to the cooling rates of chondrules. Meteorit.Planet. Sci. 37, 183–207.

Desch S. J. and Cuzzi J. N. (2000) The generation oflightning in the solar nebula. Icarus 143, 87–105.

Duggan P., McBreen B., Carr A. J., Winston E., VaughanG., Hanlon L., McBreen S., Metcalfe L., Kvick A., andTerry A. E. (2003) Gamma-ray bursts and X-ray meltingof material to form chondrules and planets. Astron. As-trophys. 409, L9–L12.

Ebata S., Nagashima K., Itoh S., Kobayashi S., SakamotoN., Fagan T. J., and Yurimoto H. (2006) Presolar silicategrains in enstatite chondrites. In Lunar Planet. Sci.XXXVII, #1619. The Lunar and Planetary Institute,Houston (CD-ROM).

Ebel D. S. and Grossman L. (2000) Condensation in dustenriched systems. Geochim. Cosmochim. Acta 64, 339–366.

Eisenhour D. D. and Buseck P. R. (1995) Chondruleformation by radiative heating: a numerical model.Icarus 117, 197–211.

Endress M., Keil K., Bischoff A., Spettel B., Clayton R. N.,and Mayeda T. K. (1994) Origin of the dark clasts in theAcfer 059/El Djouf 001 CR2 chondrite. Meteoritics 29,26–40.

Endress M., Zinner E., and Bischoff A. (1996) Early aque-ous activity on primitive meteorite bodies. Nature 379,701–703.

Engrand C. and Maurette M. (1998) Carbonaceous micro-meteorites from Antarctica. Meteorit. Planet. Sci. 33,565–580.

Engrand C., McKeegan K. D., and Leshin L. A. (1999)Oxygen isotopic compositions of individual minerals in

Antarctic micrometeorites: further links to carbonaceouschondrites. Geochim. Cosmochim. Acta 63, 2623–2636.

Fagan T. J., Krot A. N., and Keil K. (2000) Calcium,aluminum-rich inclusions in enstatite chondrites (I): min-eralogy and textures. Meteorit. Planet. Sci. 35, 771–783.

Fagan T. J., Krot A. N., Keil K., and Yurimoto H. (2004)Oxygen isotopic evolution of amoeboid olivine aggre-gates in the reduced CV3 chondrites Efremovka,Vigarano, and Leoville. Geochim. Cosmochim. Acta 68,2591–2611.

Fagan T. J., Yurimoto H., Krot A. N., and Keil K. (2002)Contrasting oxygen isotopic evolution of fine and coarserefractory inclusions from the CV3 Efremovka. NIPRAntarct. Symp. Meteorit. 27, 15–17.

Folco L., Mellini M., and Pillinger C. T. (1997) Equili-brated ordinary chondrites: constraints for thermal his-tory from iron–magnesium ordering in orthopyroxene.Meteorit. Planet. Sci. 32, 567–575.

Franchi I. A., Baker L., Bridges J. C., Wright I. P., andPillinger C. T. (2001) Oxygen isotopes and the early solarsystem. Phil. Trans. Roy. Soc. Lond. A 359, 2019–2035.

Gaffey M. J. and Gilbert S. L. (1998) Asteroid 6 Hebe: theprobable parent body of the H-type ordinary chondritesand the IIE irons. Meteorit. Planet. Sci. 33, 1281–1295.

Ganguly J. and Tirone M. (2001) Relationship betweencooling rate and cooling age of a mineral: theory andapplications to meteorites. Meteorit. Planet. Sci. 36, 167–176.

Genge M. J. (2000) Chondrule formation by the ablationof small planetesimals. Meteorit. Planet. Sci. 35,1143–1150.

Gilmour J. D. (2002) The solar system’s first clocks. Science297, 1658–1659.

Gilmour J. D., Pravdivtseva O. V., Busfield A., andHohenberg C. M. (2006) The I-Xe chronometerand the early solar system. Meteorit. Planet. Sci. 41,19–31.

Gooding J. L., Keil K., Fukuoka T., and Schmitt R. A.(1980) Elemental abundances in chondrules fromunequilibrated chondrites: evidence for chondrule originby melting of preexisting materials. Earth Planet. Sci.Lett. 50, 171–180.

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.

Goswami J. N., Marhas K. K., and Sahijpal S. (2001) Didsolar energetic particles produce the short-lived nuclidespresent in the early solar system? Astrophys. J. 549, 1151–1159.

GounelleM., SpurnyP., andBlandP.A. (2006)Theorbit andatmospheric trajectory of the Orgueil meteorite from his-torical records. Meteorit. Planet. Sci. 41, 135–150.

Gounelle M. and Zolensky M. E. (2001) A terrestrial originfor the sulfate veins in CI1 chondrites. Meteorit. Planet.Sci. 36, 1321–1329.

Grady M. M. (2000) Catalogue of Meteorites. CambridgeUniversity Press, Cambridge, UK, 689pp.

Graf T. and Marti K. (1994) Collisional records inLL-chondrites. Meteoritics 29, 643–648.

Graf T. and Marti K. (1995) Collisional history of Hchondrites. J. Geophys. Res. 100, 21247–21264.

Greshake A. (1997) The primitive matrix components ofthe unique carbonaceous chondrite Acfer 094: a TEMstudy. Geochim. Cosmochim. Acta 61, 437–452.

Greshake A., Hoppe P., and Bischoff A. (1996) Miner-alogy, chemistry, and oxygen isotopes of refractory in-clusions from stratospheric interplanetary dust particlesand micrometeorites. Meteorit. Planet. Sci. 31, 739–748.

Greshake A., Krot A. N., Meibom A., Weisberg M. K.,Zolensky M. E., and Keil K. (2002) Heavily-hydratedlithic clasts in CH chondrites and the related, metal-richchondrites Queen Alexandra Range 94411 and Hamma-dah al Hamra 237. Meteorit. Planet. Sci. 37, 281–293.


Grimm R. E. (1985) Penecontemporaneous metamorphism,fragmentation, and reassembly of ordinary chondrite as-teroids. J. Geophys. Res. 90, 2022–2028.

Grossman J. N. (1996) Chemical fractionations of chond-rites. In Chondrules and the Protoplanetary Disk (eds.R. H. Hewins, R. H. Jones, and E. R. D. Scott). Cam-bridge University Press, Cambridge, pp. 243–253.

Grossman J. N., Alexander C. M. O. D., Wang J. H.,and Brearley A. J. (2000) Bleached chondrules: evidencefor widespread aqueous processes on the parent asteroidsof ordinary chondrites. Meteorit. Planet. Sci. 35,467–486.

Grossman J. N., Alexander C. M. O. D., Wang J. H., andBrearley A. J. (2002) Zoned chondrules in Semarkona:evidence for high- and low-temperature processing. Me-teorit. Planet. Sci. 37, 49–73.

Grossman J. N. and Brearley A. J. (2005) The onset ofmetamorphism in ordinary and carbonaceous chondrites.Meteorit. Planet. Sci. 40, 87–122.

Grossman L., Ganapathy R., Methot R. L., and Davis A.M. (1979) Trace elements in the Allende meteorite am-oeboid olivine aggregates. Geochim. Cosmochim. Acta 43,817–829.

Grossman L. and Larimer J. W. (1974) Early chemicalhistory of the solar system. Rev. Geophys. Space Phys. 12,71–101.

Grossman L. and Olsen E. (1974) Origin of high-temper-ature fraction of C2 chondrites. Geochim. Cosmochim.Acta 38, 178–187.

Grossman J. N., Rubin A. E., Nagahara H., and King E.A. (1988) Properties of chondrules. In Meteorites and theEarly Solar System (eds. J. F. Kerridge and M. S. Mat-thews). University of Arizona Press, Tucson, AZ,pp. 619–659.

Grossman L. and Steele I. M. (1976) Amoeboid olivineaggregates in the Allende meteorite. Geochim. Cosmo-chim. Acta 40, 149–155.

Guan Y., Huss G. R., Leshin L. A., MacPherson G. J., andMcKeegan K. D. (2006) Oxygen isotope and 26Al–26Mgsystematics of aluminum-rich chondrules from unequil-ibrated enstatite chondrites. Meteorit. Planet. Sci. 41,33–47.

Guan Y., McKeegan K. D., and MacPherson G. J. (2000)Oxygen isotopes in calcium–aluminum-rich inclusionsfrom enstatite chondrites: new evidence for a single CAIsource in the solar nebula. Earth Planet. Sci. Lett. 181,271–277.

Guimon R. K., Symes S. J. K., Sears D. W. G., and BenoitP. H. (1995) Chemical and physical studies of type 3chondrites XII: the metamorphic history of CV chond-rites and their components. Meteoritics 30, 704–714.

Haack H., Taylor G. J., Scott E. R. D., and Keil K.(1992) Thermal history of chondrites: hot accretion vs.metamorphic reheating. Geophys. Res. Lett. 19, 2235–2238.

Haack H., Warren P. H., and Rasmussen K. L. (1990) Ef-fects of regolith/megaregolith insulation on the coolinghistories of differentiated asteroids. J. Geophys. Res. 95,5111–5124.

Hallenbeck S. L., Nuth J. A., and Nelson R. N. (2000)Evolving optical properties of annealing silicate grains:from amorphous condensate to crystalline material. As-trophys. J. 535, 247–255.

Hanner M. S. (1999) The silicate material in comets. SpaceSci. Rev. 90, 99–108.

Hanowski N. P. and Brearley A. J. (2000) Iron-rich aureolesin the CM carbonaceous chondrites, Murray, Murchison,and Allan Hills 81002: evidence for in situ aqueous alter-ation. Meteorit. Planet. Sci. 35, 1291–1308.

Hanowski N. P. and Brearley A. J. (2001) Aqueous alter-ation of chondrules in the CM carbonaceous chondrite,Allan Hills 81002: implications for parent body altera-tion. Geochim. Cosmochim. Acta 65, 495–518.

Harker D. E. and Desch S. L. (2002) Annealing of silicatedust by nebular shocks at 10 AU. Astrophys. J. 565,L109–L112.

Harker D. E., Woodward C. E., and Wooden D. H. (2005)The dust grains from 9P/Tempel 1 before and after theencounter with Deep Impact. Science 310, 278–280.

Hashimoto A. and Grossman L. (1987) Alteration ofAl-rich inclusions inside amoeboid olivine aggregates inthe Allende meteorite. Geochim. Cosmochim. Acta 51,1685–1704.

Hevey P. J. and Sanders I. S. (2006) A model for planet-esimal meltdown by 26Al, and its implications for mete-orite parent bodies. Meteorit. Planet. Sci. 41, 95–106.

Hewins R. H. (1997) Chondrules. Annu. Rev. Earth Planet.Sci. 25, 61–83.

Hewins R. H. and Connolly H. C., Jr. (1996) Peak tem-peratures of flash-melted chondrules. In Chondrules andthe Protoplanetary Disk (eds. R. H. Hewins, R. H. Jones,and E. R. D. Scott). Cambridge University Press, Camb-ridge, pp. 197–204.

Hewins R. H., Connolly H. C., Jr., Lofgren G. E., andLibourel G. (2005) Experimental constraints on chond-rule formation. In Chondrites and the ProtoplanetaryDisk (eds. A. N. Krot, E. R. D. Scott, and B. Reipurth).Astron. Soc. Pacific Conf. Ser., vol. 341, pp. 286–316.

Hewins R. H. and Herzberg C. T. (1996) Nebular turbu-lence, chondrule formation, and the composition of theEarth. Earth Planet. Sci. Lett. 144, 1–7.

Hewins R. H., Jones R. H., and Scott E. R. D. (1996)Chondrules and the Protoplanetary Disk. CambridgeUniversity Press, Cambridge, 346pp.

Hewins R. H., Yu Y., Zanda B., and Bourot-Denise M.(1997) Do nebular fractionations, evaporative losses, orboth, influence chondrule compositions? Antarct. Mete-orit. Res. 10, 275–298.

Hill H. G. M., Grady C. A., Nuth J. A., Hallenbeck S. L.,and Sitko M. L. (2001) Constraints on nebular dynamicsand chemistry based on observations of annealedmagnesium silicate grains in comets and disks surround-ing Herbig Ae/Be stars. Proc. Natl. Acad. Sci. 98,2182–2187.

Hiroi T., Zolensky M. E., and Pieters C. M. (2001) TheTagish Lake meteorite: a possible sample from a D-typeasteroid. Science 293, 2234–2236.

Hiyagon H. (1998) Distribution of oxygen isotopes in andaround some refractory inclusions from Allende.In Lunar Planet. Sci. XXIX, #1582. The Lunar andPlanetary Institute, Houston (CD-ROM).

Hiyagon H. (2000) An ion microprobe study of oxygenisotopes in some inclusions in Kainsaz and Y-81020 CO3chondrites. Symp. Antarct. Meteorit. 25, 19–21.

Hiyagon H. and Hashimoto A. (1999a) 16O excesses inolivine inclusions in Yamato-86009 and Murchisonchondrites and their relation to CAIs. Science 283,828–831.

Hiyagon H. and Hashimoto A. (1999b) An ion microprobestudy of oxygen isotopes in various types of inclusions inYamato-791717 (CO3) chondrite: oxygen isotopes vs.fayalite content in olivine. Symp. Antarct. Meteorit. 24,37–39.

Hong Y. and Fegley B., Jr. (1998) Experimental studies ofmagnetite formation in the solar nebula. Meteorit.Planet. Sci. 33, 1101–1112.

Hood L. (1998) Thermal processing of chondrule precur-sors in planetesimal bow shocks. Meteorit. Planet. Sci.33, 97–107.

Hsieh H. H. and Jewitt D. (2006) A population of comets inthe main asteroid belt. Science 312, 561–563.

Hsu W. and Crozaz G. (1998) Mineral chemistry and theorigin of enstatite in unequilibrated enstatite chondrites.Geochim. Cosmochim. Acta 62, 1993–2004.

Hsu W., Wasserburg G. J., and Huss G. R. (2000) Hightime resolution by use of the 26Al chronometer in the

References 63

multi-stage formation of a CAI. Earth Planet. Sci. Lett.182, 15–29.

Hua X., Wang J., and Buseck P. R. (2002) Fine-grainedrims in the Allan Hills 81002 and Lewis Cliff 90500 CM2meteorites: their origin and modification. Meteorit.Planet. Sci. 37, 229–244.

Huss G. R., Alexander C. M. O’D., Palme H., Bland P. A.,and Wasson J. T. (2005) Genetic relationships betweenchondrules, fine-grained rims, and interchondrule matrix.In Chondrites and the Protoplanetary Disk (eds. A. N.Krot, E. R. D. Scott, and B. Reipurth). Astron. Soc.Pacific Conf. Ser., vol. 341, pp. 701–731.

Huss G. R., Keil K., and Taylor G. J. (1981) The matricesof unequilibrated ordinary chondrites: implications forthe origin and history of chondrites. Geochim. Cosmo-chim. Acta 45, 33–51.

Huss G. R., MacPherson G. J., Wasserburg G. J., RussellS. S., and Srinivasan G. (2001) Aluminum-26 in calcium–aluminum-rich inclusions and chondrules from unequil-ibrated ordinary chondrites. Meteorit. Planet. Sci. 36,975–997.

Hutcheon I. D., Huss G. R., and Wasserburg G. J. (1994) Asearch for 26Al in chondrites: chondrule formation timescales. In Lunar Planet. Sci., The Lunar and PlanetaryInstitute, Houston, pp. 587–588.

Hutcheon I. D. and Jones R. H. (1995) The 26Al–26Mgrecord of chondrules: clues to nebular chronology. InLunar Planet. Sci. XXVI, The Lunar and Planetary In-stitute, Houston, pp. 647–648.

Hutcheon I. D., Krot A. N., Keil K., Phinney D. L., andScott E. R. D. (1998) 53Mn–53Cr dating of fayalite for-mation in the CV3 chondrite Mokoia: evidence for as-teroidal alteration. Science 282, 1865–1867.

Hutcheon I. D., Krot A. N., and Ulyanov A. A. (2000) 26Alin anorthite-rich chondrules in primitive carbonaceouschondrites: evidence chondrules post-date CAI. In LunarPlanet. Sci. XXXI, #1869. The Lunar and Planetary In-stitute, Houston (CD-ROM).

Hutchison R., Alexander C. M. O. D., and Bridges J. C.(1998) Elemental redistribution in Tieschitz andthe origin of white matrix. Meteorit. Planet. Sci. 33,1169–1179.

Hutchison R., Bevan A. W. R., Agrell S. O., and AshworthJ. R. (1979) Accretion temperature of the Tieschitz, H3,chondritic material. Nature 280, 116–119.

Hutchison R., Bridges J. C., and Gilmour J. D. (2005)Chondrules: chemical, petrographic and chronologicclues to their origin by impact. In Chondrites and theProtoplanetary Disk (eds. A. N. Krot, E. R. D. Scott, andB. Reipurth). Astron. Soc. Pacific Conf. Ser., vol. 341,pp. 933–950.

Hutchison R., Williams I. P., and Russell S. S. (2001)Theories of planetary formation: constraints from the studyof meteorites. Phil. Trans. Roy. Soc. Lond. A 359, 2077–2091.

Iida A., Nakamoto T., Susa H., and Nakagawa Y. (2001) Ashock heating model for chondrule formation in a proto-planetary disk. Icarus 153, 430–450.

Ikeda Y., Kimura M., Mori H., and Takeda H. (1981)Chemical compositions of matrices of unequilibrated or-dinary chondrites. Mem. Nat. Inst. Polar Res. Spec. Issue20, 124–144.

Imai H. and Yurimoto H. (2000) Oxygen and magnesiumisotopic distributions in a Type C CAI from the Allendemeteorite. In Lunar Planet. Sci. XXXI, #1510. The Lunarand Planetary Institute, Houston (CD-ROM).

Imai H. and Yurimoto H. (2001) Two generations of olivinegrowth in an amoeboid olivine aggregate from the Allendemeteorite. In Lunar Planet. Sci. XXXII, #1580. The Lunarand Planetary Institute, Houston (CD-ROM).

Ireland T. R. and Fegley B., Jr. (2000) The solar system’searliest chemistry: systematics of refractory inclusions.Int. Geol. Rev. 42, 865–894.

Ito M., Yurimoto H., and Nagasawa H. (1999) Oxygenisotope microdistribution vs. composition in melilite/fa-ssaite in the Allende CAI 7R-19-1: a new evidence formultiple heating. In Lunar Planet. Sci. XXX, #1538. TheLunar and Planetary Institute, Houston (CD-ROM).

Itoh S., Kojima H., and Yurimoto H. (2004b) Petrographyand oxygen isotopic compositions in refractory inclu-sions from CO chondrites. Geochim. Cosmochim. Acta68, 183–194.

Itoh S., Rubin A. E., Kojima H., Wasson J. T., and Yuri-moto H. (2002) Amoeboid olivine aggregates and AOA-bearing chondrule from Y-81020 CO 3.0 chondrite: dis-tribution of oxygen and magnesium isotopes. In LunarPlanet. Sci. XXXIII, #1490. The Lunar and PlanetaryInstitute, Houston (CD-ROM).

Itoh S., Russell S. S., and Yurimoto H. (2004a) Amoeboidolivine aggregates from the Semarkona LL3.0 chondrite.Workshop on Chondrites and the Protoplanetary Disk,#9019. The Lunar and Planetary Institute, Houston(CD-ROM).

Itoh D. and Tomeoka K. (2003) Dark inclusions in CO3chondrites: new indicators of parent-body processes.Geochim. Cosmochim. Acta 67, 153–169.

Itoh S. and Yurimoto H. (2003) Contemporaneous forma-tion of chondrules and refractory inclusions in the earlySolar System. Nature 423, 728–731.

Jarosewich E. (1990) Chemical analyses of meteorites: acompilation of stony and iron meteorite analyses. Mete-oritics 25, 323–337.

Johnson C. A., Prinz M., Weisberg M. K., Clayton R. N.,and Mayeda T. K. (1990) Dark inclusions in Allende,Leoville, and Vigarano: evidence for nebular oxidation ofCV3 constituents. Geochim. Cosmochim. Acta 54, 819–830.

Jones R. H. (1990) Petrology and mineralogy of type II,FeO-rich chondrules in Semarkona (LL3.0): origin byclosed system fractional crystallization, with evidence forsupercooling. Geochim. Cosmochim. Acta 54, 1785–1802.

Jones R. H. (1994) Petrology of FeO-poor, por-phyritic pyroxene chondrules in the Semarkona chond-rite. Geochim. Cosmochim. Acta 58, 5325–5340.

Jones R. H. (1996a) FeO-rich, porphyritic pyroxene chond-rules in unequilibrated ordinary chondrites. Geochim.Cosmochim. Acta 60, 3115–3138.

Jones R. H. (1996b) Relict grains in chondrules: evidencefor chondrules recycling. In Chondrules and the Proto-planetary Disk (eds. R. H. Hewins, R. H. Jones, and E.R. D. Scott). Cambridge University Press, Cambridge,pp. 163–172.

Jones R. H. and Brearley A. J. (1988) Kinetics of theclinopyroxene–orthopyroxene transition: constraints onthe thermal histories of chondrules and type 3–6 chond-rites. Meteoritics 23, 277–278.

Jones R. H., Grossman J. N., and Rubin A. E. (2005)Chemical, mineralogical and isotopic properties ofchondrules: clues to their origin. In Chondrites and theProtoplanetary Disk (eds. A. N. Krot, E. R. D. Scott, andB. Reipurth). Astron. Soc. Pacific Conf. Ser., vol. 341,pp. 251–285.

Jones T. D., Lebovsky L. A., Lewis J. S., and Marley M. S.(1990) The composition and origin of the C, P, and Dasteroids: water as a tracer of thermal evolution in theouter belt. Icarus 88, 172–192.

Jones R. H., Lee T., Connolly H. C., Jr., Love S. G., andShang H. (2000a) Formation of chondrules and CAIs:theory vs. observation. In Protostars and Planets IV (eds.V. Mannings, A. P. Boss, and S. S. Russell). Universityof Arizona Press, Tucson, AZ, pp. 927–962.

Jones R. H., Leshin L. A., Guan Y., Sharp Z. D.,Durakiewicz T., and Schilk A. J. (2004) Oxygen isotopeheterogeneity in chondrules from the Mokoia CV3 car-bonaceous chondrite. Geochim. Cosmochim. Acta 68,3423–3438.


Jones R. H. and Lofgren G. E. (1993) A comparison ofFeO-rich, porphyritic olivine chondrules in unequilibrat-ed chondrites and experimental analogues. Meteoritics28, 213–221.

Jones R. H., Saxton J. M., Lyon I. C., and Turner G.(2000b) Oxygen isotopes in chondrule olivine and iso-lated olivine grains. Meteorit. Planet. Sci. 35, 849–857.

Joung M. K. R., Mac Low M.-M., and Ebel D. S. (2004)Chondrule formation and protoplanetary disk heating bycurrent sheets in nonideal magnetohydrodynamic turbu-lence. Astrophys. J. 606, 532–541.

Kallemeyn G. W., Rubin A. E., and Wasson J. T. (1991)The compositional classification of chondrites. Part V:The Karoonda (CK) group of carbonaceous chondrites.Geochim. Cosmochim. Acta 55, 881–892.

Kallemeyn G. W., Rubin A. E., and Wasson J. T. (1994)The compositional classification of chondrites. Part VI:The CR carbonaceous chondrite group. Geochim. Cosmo-chim. Acta 58, 2873–2888.

Kallemeyn G. W., Rubin A. E., and Wasson J. T. (1996)The compositional classification of chondrites. Part VII:The R chondrite group. Geochim. Cosmochim. Acta 60,2243–2256.

Kallemeyn G. W. and Wasson J. T. (1981) The compos-itional classification of chondrites. Part I: The carbon-aceous chondrite groups. Geochim. Cosmochim. Acta 45,1217–1230.

Keil K. (1989) Enstatite meteorites and their parent bodies.Meteoritics 24, 195–208.

Keil K. (2000) Thermal alteration of asteroids: evidencefrom meteorites. Planet. Space Sci. 48, 887–903.

Keil K., Haack H., and Scott E. R. D. (1994) Catastrophicfragmentation of asteroids: evidence from meteorites.Planet. Space Sci. 42, 1109–1122.

Keil K., Stoffler D., Love S. G., and Scott E. R. D. (1997)Constraints on the role of impact heating and melting inasteroids. Meteorit. Planet. Sci. 32, 349–363.

Keller L. P. and Buseck P. R. (1990) Matrix mineralogy ofthe Lance CO3 carbonaceous chondrite. Geochim. Cosmo-chim. Acta 54, 1155–1163.

Keller L. P., Clark J. C., Lewis C. F., and Moore C. B.(1992) Maralinga, a metamorphosed carbonaceouschondrite found in Australia. Meteoritics 27, 87–91.

Keller L. P., Thomas K. L., Clayton R. N., Mayeda T. K.,DeHart J. M., and McKay D. S. (1994) Aqueous alterationof the Bali CV3 chondrite: evidence from mineralogy, min-eral chemistry, and oxygen isotopic compositions. Geochim.Cosmochim. Acta 58, 5589–5598.

Kerr R. A. (2001) A meteoriticist speaks out, his rocksremain mute. Science 293, 1581–1584.

Kerridge J. F. (1993) What can meteorites tell us aboutnebular conditions and processes during planetesimalaccretion? Icarus 106, 135–150.

Kerridge J. F. and Matthews M. S. (1988) Meteorites andthe Early Solar System. Arizona State University,Tucson, 1269pp.

Kim G. L., Yurimoto H., and Sueno S. (1998) 16O-enrichedmelilite and anorthite coexisting with 16O-depleted me-lilite in a CAI. In Lunar Planet. Sci. XXIX, #1344. TheLunar and Planetary Institute, Houston (CD-ROM).

Kimura M., El Goresy A., Palme H., and Zinner E. (1993)Ca-, Al-rich inclusions in the unique chondrite ALH85085:petrology, chemistry, and isotopic compositions. Geochim.Cosmochim. Acta 57, 2329–2359.

Kimura M., Hiyagon H., Palme H., Spettel B., Wolf D.,Clayton R. N., Mayeda T. K., Sato T., Suzuki A., andKojima H. (2002) Yamato 792947, 793408 and 82038: themost primitive H chondrites, with abundant refractoryinclusions. Meteorit. Planet. Sci. 37, 1417–1434.

Kimura M. and Ikeda Y. (1998) Hydrous and anhydrousalteration in Kaba and Mokoia CV chondrites. Meteorit.Planet. Sci. 33, 1139–1146.

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).Astron. Soc. Pacific Conf. Ser., vol. 341, pp. 558–587.

Kita N. T., Nagahara H., Togashi S., and Morishita Y.(2000) A short formation period of chondrules in thesolar nebula. Geochim. Cosmochim. Acta 48, 693–709.

Kleine T., Mezger K., Palme H., Scherer E., and Munker C.(2005) Early core formation in asteroids and late accretionof chondrite parent bodies: evidence from 182Hf–182W inCAIs, metal-rich chondrites, and iron meteorites. Geo-chim. Cosmochim. Acta 69, 5805–5818.

Klerner S. and Palme H. (2000) Large titanium/aluminumfractionation between chondrules and matrix in Renazzoand other carbonaceous chondrites. Meteorit. Planet.Sci. 35, A89.

Klock W., Thomas K. L., McKay D. S., and Palme H.(1989) Unusual olivine and pyroxene composition in in-terplanetary dust and unequilibrated ordinary chond-rites. Nature 339, 126–128.

Kobayashi S., Tonotani A., Sakamoto N., Nagashima K.,Krot A. N., and Yurimoto H. (2005) Presolar silicategrains from primitive carbonaceous chondrites Y-81025,ALHA 77307, Adelaide and Acfer 094. In Lunar Planet.Sci. XXXVI, #1931. The Lunar and Planetary Institute,Houston (CD-ROM).

Kojima T. and Tomeoka K. (1996) Indicators of aqueousalteration and thermal metamorphism on the CV parentbody: microtextures of a dark inclusion. Geochim. Cosmo-chim. Acta 60, 2651–2666.

Kojima T., Tomeoka K., and Takeda H. (1993) Unusualdark clasts in the Vigarano CV3 carbonaceous chondrite:record of the parent body process. Meteoritics 28, 649–658.

Komatsu M., Krot A. N., Petaev M. I., Ulyanov A. A., KeilK., and Miyamoto M. (2001) Mineralogy and petrograp-hy of amoeboid olivine aggregates from the reduced CV3chondrites Efremovka, Leoville, and Vigarano: productsof nebular condensation, accretion, and annealing. Mete-orit. Planet. Sci. 36, 629–641.

Kong P., Deloule E., and Palme H. (2000) REE-bearingsulfide in Bishunpur (LL3.1)—a highly unequilibratedordinary chondrite. Earth Planet. Sci. Lett. 177, 1–7.

Kong P., Mori T., and Ebihara M. (1997) Compositionalcontinuity of enstatite chondrites and implications forheterogeneous accretion of the enstatite chondrite parentbody. Geochim. Cosmochim. Acta 61, 4895–4914.

Kong P. and Palme H. (1999) Compositional and geneticrelationship between chondrules, chondrule rims, metal,and matrix in the Renazzo chondrite. Geochim. Cosmo-chim. Acta 63, 3673–3682.

Kracher A., Keil K., Kallemeyn G. W., Wasson J. T., Clay-ton R. N., and Huss G. I. (1985) The Leoville acccretionarybreccia. Proc. 16th Lunar Planet. Sci. Conf.: J. Geophys.Res. 90, D123–D135.

Kring D. A. and Boynton W. V. (1990) Trace-elementcompositions of Ca-rich chondrules from Allende:relationships between refractory inclusions andferromagnesian chondrules. Meteoritics 25, 377.

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

Krot A. N., Brearley A. J., Ulyanov A. J., Biryukov V. V.,Swindle T. D., Keil K., Mittlefehldt D. W., Scott E. R.D., and Nakamura K. (1998a) Mineralogy, petrography,bulk chemical, iodine–xenon, and oxygen isotopic com-positions of dark inclusions in the reduced CV3 chond-rite Efremovka. Meteorit. Planet. Sci. 34, 67–89.

Krot A. N., Fagan T. J., Keil K., McKeegan K. D., Sa-hijpal S., Hutcheon I. D., Petaev M. I., and Yurimoto H.(2004a) Ca, Al-rich inclusions, amoeboid olivine

References 65

aggregates, and Al-rich chondrules from the unique car-bonaceous chondrite Acfer 094. Part I: Mineralogy andpetrology. Geochim. Cosmochim. Acta 68, 2167–2184.

Krot A. N., Fagan T. J., Yurimoto H., and Petaev M. I.(2005b) Origin of low-Ca pyroxene in amoeboid olivineaggregates: evidence from oxygen isotopic compositions.Geochim. Cosmochim. Acta 69, 1873–1881.

Krot A. N., Huss G. R., and Hutcheon I. D. (2001a)Corundum–hibonite refractory inclusions from Adelaide:condensation or crystallization from melt? Meteorit.Planet. Sci. 36, A105.

Krot A. N., Hutcheon I. D., and Huss G. R. (2001b)Aluminum-rich chondrules and associated refractory in-clusions in the unique carbonaceous chondrite Adelaide.Meteorit. Planet. Sci. 36, A105–A106.

Krot A. N., Hutcheon I. D., and Keil K. (2002a) Anorthite-rich chondrules in the reduced CV chondrites: evidencefor complex formation history and genetic links betweenCAIs and ferromagnesian chondrules. Meteorit. Planet.Sci. 37, 155–182.

Krot A., Ivanova M. A., and Wasson J. T. (1993) Theorigin of chromitic chondrules and the volatility of Crunder a range of nebular conditions. Earth Planet. Sci.Lett. 119, 569–584.

Krot A. N. and Keil K. (2002) Anorthite-rich chondrules inCR and CH carbonaceous chondrites: genetic linkbetween Ca, Al-rich inclusions and ferromagnesian chond-rules. Meteorit. Planet. Sci. 37, 91–111.

Krot A. N., Libourel G., and Chaussidon M. (2006a)Oxygen isotope compositions of chondrules in CRchondrites. Geochim. Cosmochim. Acta 70, 767–779.

Krot A. N., Libourel G., Goodrich C. A., and Petaev M. I.(2004b) Silica-rich igneous rims around magnesianchondrules in CR carbonaceous chondrites: evidencefor fractional condensation during chondrule formation.Meteorit. Planet. Sci. 39, 1931–1955.

Krot A. N., MacPherson G. J., Ulyanov A. A., and PetaevM. I. (2004c) Fine-grained, spinel-rich inclusions fromthe reduced CV chondrites Efremovka and Leoville. PartI: Mineralogy, petrology, and bulk chemistry. Meteorit.Planet. Sci. 39, 1517–1553.

Krot A. N., McKeegan K. D., Huss G. R., Liffman K.,Sahijpal S., Hutcheon I. D., Srinivasan G., Bischoff A.,and Keil K. (2006b) Aluminum–magnesium and oxygenisotope study of relict Ca-Al-rich inclusions in chond-rules. Astrophys. J. 639, 1227–1237.

Krot A. N., McKeegan K. D., Leshin L. A., MacPhersonG. J., and Scott E. R. D. (2002b) Existence of an 16O-richgaseous reservoir in the solar nebula. Science 295,1051–1054.

Krot A. N., McKeegan K. D., Russell S. S., Meibom A.,Weisberg M. K., Zipfel J., Krot T. V., Fagan T. J., andKeil K. (2001c) Refractory Ca, Al-rich inclusions and Al-diopside-rich chondrules in the metal-rich chondritesHammadah al Hamra 237 and QUE 94411. Meteorit.Planet. Sci. 36, 1189–1217.

Krot A. N., Meibom A., Russell S. S., Alexander C. M. O.,Jeffries T. E., and Keil K. (2001d) A new astrophysicalsetting for chondrule formation. Science 291, 1776–1779.

Krot A. N., Meibom A., Weisberg M. K., and Keil K.(2002c) The CR chondrite clan: implications for early so-lar system processes. Meteorit. Planet. Sci. 37, 1451–1490.

Krot A. N., Petaev M. I., and Bland P. A. (2004f) Multipleformation mechanisms of ferrous olivine in CV carbo-naceous chondrites during fluid-assisted metamorphism.Antarct. Meteorite Res. 17, 154–172.

Krot A. N., Petaev M. I., Meibom A., and Keil K. (2001e)In situ growth of Ca-rich rims around Allende dark in-clusions. Geochem. Int. 36, 351–368.

Krot A. N., Petaev M. I., Russell S. S., Itoh S., Fagan T. J.,Yurimoto H., Chizmadia L., Weisberg M. K., KomatsuM., Ulyanov A. A., and Keil K. (2004d) Amoeboid olivine

aggregates in carbonaceous chondrites: records of nebularand asteroidal processes. Chem. Erde 64, 185–239.

Krot A. N., Petaev M. I., Scott E. R. D., Choi B.-G., Zolen-sky M. E., and Keil K. (1998b) Progressive alteration inCV3 chondrites: more evidence for asteroidal alteration.Meteorit. Planet. Sci. 33, 1065–1085.

Krot A. N., Petaev M. I., and Yurimoto H. (2004e) Am-oeboid olivine aggregates with low-Ca pyroxenes: agenetic link between refractory inclusions and chond-rules? Geochim. Cosmochim. Acta 68, 1923–1941.

Krot A. N., Petaev M. I., Zolensky M. E., Keil K., Scott E.R. D., and Nakamura K. (1998c) Secondary calcium–iron-rich minerals in the Bali-like and Allende-like oxi-dized CV3 chondrites and Allende dark inclusions. Me-teorit. Planet. Sci. 33, 623–645.

Krot A. N. and Rubin A. E. (1993) Chromite-rich maficsilicate chondrules in ordinary chondrites: formation byimpact melting. In Lunar Planet. Sci. XXXIII, The Lunarand Planetary Institute, Houston, pp. 827–828.

Krot A. N., Scott E. R. D., and Reipurth B. (eds.) (2005c)Chondrites and the Protoplanetary Disk. Astron. Soc.Pacific Conf. Ser., vol. 341, 1029pp.

Krot A. N., Scott E. R. D., and Zolensky M. E. (1995)Mineralogical and chemical modification of componentsin CV3 chondrites: nebular or asteroidal processing?Meteoritics 30, 748–775.

Krot A. N., Scott E. R. D., and Zolensky M. E. (1997)Origin of fayalitic olivine rims and lath-shaped matrixolivine in the CV3 chondrite Allende and its dark inclu-sions. Meteorit. Planet. Sci. 32, 31–49.

Krot A. N., Ulyanov A. A., Ivanova M. A., and Russell S. S.(2006c) Origins of chondrules in the metal-rich carbon-aceous chondrites. In Lunar Planet. Sci. XXXVII, #1224.The Lunar and Planetary Institute, Houston (CD-ROM).

Krot A. N., Ulyanov A. A., Meibom A., and Keil K.(2001f) Forsterite-rich accretionary rims around Ca, Al-rich inclusions from the reduced CV3 chondrite Efre-movka. Meteorit. Planet. Sci. 36, 611–628.

Krot A. N. and Wasson J. T. (1995) Igneous rims on low-FeO and high-FeO chondrules in ordinary chondrites.Geochim. Cosmochim. Acta 59, 4951–4966.

Krot A. N., Yurimoto H., Hutcheon I. D., and MacPher-son G. J. (2005d) Relative chronology of CAI andchondrule formation: evidence from chondrule-bearingigneous CAIs. Nature 434, 998–1001.

Kuebler K. E., McSween H. Y., Jr., Carlson W. D., andHirsch D. (1999) Sizes and masses of chondrulesand metal-troilite grains in ordinary chondrites: possibleimplications for nebular sorting. Icarus 141, 96–106.

Kunihiro T., Nagashima K., and Yurimoto H. (2005b)Microscopic oxygen isotopic homogeneity/heterogeneityin the matrix of the Vigarano CV3 chondrite. Geochim.Cosmochim. Acta 69, 763–773.

Kunihiro T., Rubin A. E., McKeegan K. D., and Wasson J.T. (2004a) Initial 26Al/27Al in carbonaceous-chondritechondrules: too little 26Al to melt asteroids. Geochim.Cosmochim. Acta 68, 2947–2957.

Kunihiro T., Rubin A. E., McKeegan K. D., and Wasson J.T. (2004b) Oxygen-isotopic compositions of relict andhost grains in chondrules in the Yamato 81020 CO3.0chondrite. Geochim. Cosmochim. Acta 68, 3599–3606.

Kunihiro T., Rubin A. E., and Wasson J. T. (2005a)Oxygen-isotopic compositions of low-FeO relicts in high-FeO host chondrules in Acfer 094, a type 3.0 carbon-aceous chondrite closely related to CM. Geochim.Cosmochim. Acta 69, 3831–3840.

Kurat G. (1988) Primitive meteorites: an attempt towardsunification. Proc Trans. Roy. Soc. Lond. A 325, 459–482.

Langenhorst F., Poirier J.-P., Deutsch A., and HornemannU. (2002) Experimental approach to generate shock veinsin single-crystal olivine by shear melting. Meteorit.Planet. Sci. 37, 1541–1553.


Larimer J. W. (1967) Chemical fractionations in meteorites:I. Condensation of the elements. Geochim. Cosmochim.Acta 31, 1215–1238.

Larimer J. W. and Anders E. (1967) Chemical fraction-ations in meteorites. Part II: Abundance patterns andtheir interpretation. Geochim. Cosmochim. Acta 31,1239–1270.

Larimer J. W. and Wasson J. T. (1988) Refractory litho-phile elements. In Meteorites and the Early Solar Sys-tem (eds. J. F. Kerridge and M. S. Matthews).University of Arizona Press, Tucson, AZ, pp. 394–415.

Lauretta D. S., Buseck P. R., and Zega T. J. (2001) Opaqueminerals in the matrix of the Bishunpur (LL3.1) chondrite:constraints on the chondrule formation environment.Geochim. Cosmochim. Acta 65, 1337–1353.

Lauretta D. S., Hua X., and Buseck P. R. (2000) Miner-alogy of fine-grained rims in the ALH 81002 CM chond-rite. Geochim. Cosmochim. Acta 64, 3263–3273.

Lauretta D. S., Lodders K., and Fegley B. (1998) Kamacitesulfurization in the solar nebula. Meteorit. Planet. Sci.33, 821–833.

Lauretta D. S., Lodders K., Fegley B., and Kremser D. T.(1997) The origin of sulfide-rimmed metal grains in or-dinary chondrites. Earth Planet. Sci. Lett. 151, 289–301.

Lauretta D. S. and McSween H. Y., Jr. (2006) Meteoritesand the Early Solar System II. University of ArizonaPress, Tucson, AZ, 914pp.

Lee M. R., Hutchison R., and Graham A. L. (1996) Aque-ous alteration in the matrix of the Vigarano (CV3) car-bonaceous chondrite. Meteorit. Planet. Sci. 31, 477–483.

Leshin L. A., Rubin A. E., and McKeegan K. D. (1997) Theoxygen isotopic composition of olivine and pyroxene fromCI chondrites. Geochim. Cosmochim. Acta 61, 835–845.

Libourel G. and Krot A. N. (2006) Origin of olivines intype I chondrules: petrologic and chemical constraints. InLunar Planet. Sci. XXXVII, #1334. The Lunar and Plan-etary Institute, Houston (CD-ROM).

Liffman K. (2005) Chondrule and metal grain size sortingfrom jet flows. Meteorit. Planet. Sci. 40, 123–138.

Liffman K. and Brown M. (1996a) The motion and sizesorting of particles ejected from a protostellar accretiondisk. Icarus 116, 275–290.

Liffman K. and Brown M. J. I. (1996b) The protostellar jetmodel of chondrule formation. In Chondrules and theProtoplanetary Disk (eds. R. H. Hewins, R. H. Jones,and E. R. D. Scott). Cambridge University Press, Camb-ridge, pp. 285–302.

Liffman K. and Toscano M. (2000) Chondrule fine-grainedmantle formation by hypervelocity impact of chondruleswith a dusty gas. Icarus 143, 106–125.

Lin Y. and Kimura M. (1998) Petrographic and miner-alogical study of new EH melt rocks and a new enstatitechondrite grouplet. Meteorit. Planet. Sci. 33, 501–511.

Lin Y., Kimura M., Miao B., Dai D., and Monoi A. (2006)Petrographic comparison of refractory inclusions fromdifferent chemical groups of chondrites. Meteorit. Planet.Sci. 41, 67–81.

Lipschutz M. E., Gaffey M. J., and Pellas P. (1989)Meteoritic parent bodies: nature, number, size and rela-tionship to present-day asteroids. In Asteroids II (eds. R.P. Binzel, T. Gehrels, and M. S. Matthews). University ofArizona Press, Tucson, AZ, pp. 740–777.

Lodders K. and Fegley B., Jr. (1998) The Planetary Scien-tist’s Companion. Oxford University Press, 371pp.

Lodders K. and Osborne R. (1999) Perspectives on thecomet-asteroid-meteorite link. Space Sci. Rev. 90, 289–297.

Lofgren G. (1996) A dynamic crystallization model forchondrules melts. In Chondrules and the ProtoplanetaryDisk (eds. R. H. Hewins, R. H. Jones, and E. R. D. Scott).Cambridge University Press, Cambridge, pp. 187–196.

Luck J. M., Ben-Othman D., Barrat J. A., and Albarede F.(2003) Coupled 63Cu and 16O excesses in chondrites.Geochim. Cosmochim. Acta 67, 143–151.

Lugmair G. W. and Shukolyukov A. (2001) Early solarsystem events and timescales. Meteorit. Planet. Sci. 36,1017–1026.

Lyons J. R. and Young E. D. (2005) CO self-shielding asthe origin of oxygen isotope anomalies in the early solarsystem. Nature 435, 317–320.

MacPherson G. J., Davis A. M., and Grossman J. N.(1989) Refractory inclusions in the unique chondriteALH85085. Meteoritics 24, 297.

MacPherson G. J., Davis A. M., and Zinner E. K. (1995)The distribution of aluminum-26 in the early solar sys-tem. Meteoritics 30, 365–386.

MacPherson G. J. and Delaney J. S. (1985) A fassaite-twoolivine–pleonaste-bearing refractory inclusion from Kar-oonda. In Lunar Planet. Sci. XVI, The Lunar and Plan-etary Institute, Houston, pp. 515–526.

MacPherson G. J., Grossman L., Hashimoto A., Bar-Matthews M., and Tanaka T. (1984) Petrographic studiesof refractory inclusions from the Murchison meteorite.Proc. 15th Lunar Planet. Sci. Conf.: J. Geophys. Res.Suppl. 89, C299–C312.

MacPherson G. J., Hashimoto A., and Grossman L. (1985)Accretionary rims on inclusions in the Allende meteorite.Geochim. Cosmochim. Acta 49, 2267–2279.

MacPherson G. J. and Huss G. R. (2005) Petrogenesis of Al-rich chondrules: evidence from bulk compositions andphase equilibria. Geochim. Cosmochim. Acta 69, 3099–3127.

MacPherson G. J. and Russell S. S. (1997) Origin ofaluminum-rich chondrules: constraints from major-element chemistry. Meteorit. Planet. Sci. 32, A83.

MacPherson G. J., Wark D. A., and Armstrong J. T. (1988)Primitive material surviving in chondrites: refractory in-clusions. In Meteorites and the Early Solar System (eds.J. F. Kerridge and M. S. Matthews). University ofArizona Press, Tucson, AZ, pp. 746–807.

Marhas K. K., Goswami J. N., and Davis A. M. (2002)Short-lived nuclides in hibonite: evidence for solar-sys-tem evolution. Science 298, 2182–2185.

Matsunami S., Ninagawa K., Nishimura S., Kubono N.,Yamamoto I., Kohata M., Wada T., Yamashita Y., LuJ., Sears D. W. G., and Nishimura H. (1993) Thermo-luminescence and compositional zoning in the mesostasisof a Semarkona group A1 chondrule and new insightsinto the chondrule-forming process. Meteoritics 57,2101–2110.

May C., Russell S. S., and Grady M. M. (1999) Analysis ofchondrule and CAI size and abundance in CO3 and CV3chondrites: a preliminary study. In Lunar Planet. Sci.XXX, #1688. The Lunar and Planetary Institute, Hou-ston (CD-ROM).

McCoy T. J., Scott E. R. D., Jones R. H., Keil K., andTaylor G. J. (1991) Composition of chondrule silicates inLL3-5 chondrites and implications for their nebular his-tory and parent body metamorphism. Geochim. Cosmo-chim. Acta 55, 601–619.

McKeegan K. D., Chaussidon M., and Robert F. (2000a)Incorporation of short-lived 10Be in a calcium–alumi-num-rich inclusion from the Allende meteorite. Science289, 1334–1337.

McKeegan K. D., Greenwood J. P., Leshin L., and Cosa-rinsky M. (2000b) Abundance of 26Al in ferromagnesianchondrules of unequilibrated ordinary chondrites. In Lu-nar Planet. Sci. XXX, #2009. The Lunar and PlanetaryInstitute, Houston (CD-ROM).

McKeegan K. D., Leshin L. A., Russell S. S., and MacP-herson G. J. (1998) Oxygen isotopic abundances in cal-cium–aluminum-rich inclusions from ordinarychondrites: implications for nebular heterogeneity. Sci-ence 280, 414–418.

McKeegan K. D., Sahijpal S., Krot A. N., Weber D., andUlyanov A. A. (2006) Preservation of primary oxygenisotopic compositions in Ca, Al-rich inclusions from CHchondrites. Earth Planet. Sci. Lett. (submitted).

References 67

McSween H. Y., Jr. (1977a) Petrographic variations amongcarbonaceous chondrites of the Vigarano type. Geochim.Cosmochim. Acta 41, 1777–1790.

McSween H. Y., Jr. (1977b) On the nature and origin ofisolated olivine grains in carbonaceous chondrites. Geo-chim. Cosmochim. Acta 41, 411–418.

McSween H. Y., Jr. (1979) Are carbonaceous chondritesprimitive or processed—a review. Rev. Geophys. SpacePhys. 17, 1059–1078.

McSween H.Y., Jr., Ghosh A., Grimm R. E., Wilson L., andYoung E. D. (2002) Thermal evolution models ofasteroids. In Asteroids III (eds. W.F. Bottke Jr., A.Cellino, P. Paolicchi, and R. P. Binzel). University ofArizona Press, Tucson, AZ, pp. 559–571.

McSween H. Y. and Labotka T. C. (1993) Oxidationduring metamorphism of the ordinary chondrites. Geo-chim. Cosmochim. Acta 57, 1105–1114.

McSween H. Y., Jr. and Patchen A. D. (1989) Pyroxenethermobarometry in LL chondrites and implications forparent body metamorphism. Meteoritics 24, 219–226.

McSween H. Y., Jr. and Richardson S. M. (1977) The com-position of carbonaceous chondrite matrix. Geochim.Cosmochim. Acta 41, 1145–1161.

Meibom A. and Clark B. E. (1999) Evidence for theinsignificance of ordinary chondrite material in the as-teroid belt. Meteorit. Planet. Sci. 34, 7–24.

Meibom A., Desch S. J., Krot A. N., Cuzzi J. N., Petaev M.I., Wilson L., and Keil K. (2000) Large-scale thermalevents in the solar nebula: evidence from Fe, Nimetal grains in primitive meteorites. Science 288, 839–841.

Meibom A., Petaev M. I., Krot A. N., Keil K., and WoodJ. A. (2001) Growth mechanism and additionalconstraints on FeNi condensation in the solar nebula.J. Geophys. Res. 106, 32797–32801.

Meibom A., Petaev M. I., Krot A. N., Wood J. A., and KeilK. (1999) Primitive FeNi metal grains in CH carbon-aceous chondrites formed by condensation from a gas ofsolar composition. J. Geophys. Res. 104, 22053–22059.

Messenger S., Keller L. P., Stadermann F. J., Walker R.M., and Zinner E. (2003) Samples of stars beyond thesolar system: silicate grains in interplanetary dust. Sci-ence 300, 105–108.

Messenger S., Sandford S., and Brownlee D. (2006) Thepopulation of starting materials available for solar-sys-tem construction. In Meteorites and the Early Solar Sys-tem II (eds. D. S. Lauretta and H. Y. Jr. McSween).University of Arizona Press, Tucson, AZ, pp. 187–208.

Metzler K., Bischoff A., and Stoffler D. (1992) Accretionarydust mantles in CM chondrites: evidence for solar nebulaprocesses. Geochim. Cosmochim. Acta 56, 2873–2897.

Misawa K. and Fujita T. (1994) A relict refractory inclu-sion in a ferromagnesian chondrule from the Allendemeteorite. Nature 368, 723–726.

Misawa K. and Fujita T. (2000) Magnesium isotopicfractionations in barred olivine chondrules from the Al-lende meteorite. Meteorit. Planet. Sci. 35, 85–94.

Misawa K. and Nakamura N. (1988) Highly fractionatedrare-earth elements in ferromagnesian chondrules fromthe Felix (CO3) meteorite. Nature 334, 47–49.

Misawa K. and Nakamura N. (1996) Origin of refractoryprecursor components of chondrules from carbonaceouschondrites. In Chondrules and the Protoplanetary Disk(eds. R. H. Hewins, R. H. Jones, and E. R. D. Scott), pp.99–105.

Miura H. and Nakamoto T. (2005) A shock-wave heatingmodel for chondrule formation. Part II: Minimum size ofchondrule precursors. Icarus 175, 289–304.

Morfill G. E., Durisen R. H., and Turner G. W. (1998) Anaccretion rim constraint on chondrule formationtheories. Icarus 134, 180–184.

Morlok A., Floss C., Zinner E., Bischoff A., Henkel T.,Rost D., Stephan T., and Jessberger E. K. (2002) Traceelements in CI chondrites: a heterogeneous distribution.

In Lunar Planet. Sci. XXXIII, #1269. The Lunar andPlanetary Institute, Houston (CD-ROM).

Mostefaoui S., Kita N. T., Tachibana S., Nagahara H.,Togashi S., and Morishita Y. (2000) Correlated initialaluminum-26/aluminum-27 with olivine and pyroxeneproportions in chondrules from highly unequilibratedordinary chondrites. Meteorit. Planet. Sci. 35, A114.

Mostefaoui S., Kita N. T., Togashi S., Tachibana S.,Nagahara H., and Morishita Y. (2002) The relative for-mation ages of ferromagnesian chondrules inferred fromtheir initial aluminum-26/aluminum-27 ratios. Meteorit.Planet. Sci. 37, 421–438.

Nagahara H. (1984) Matrices of type 3 ordinary chond-rites—primitive nebular records. Geochim. Cosmo-chim. Acta 48, 2581–2595.

Nagahara H., Kita N. T., Ozawa K., and Morishita Y.(1999) Condensation during chondrule formation:elemental and Mg isotopic evidence. In Lunar Planet.Sci. XXX, #1342. The Lunar and Planetary Institute,Houston (CD-ROM).

Nagahara H., Kushiro I., and Mysen B. (1994) Evapora-tion of olivine: low pressure phase relations of the olivinesystem and its implication for the origin of chondriticcomponents in the solar nebula. Geochim. Cosmochim.Acta 58, 1951–1963.

Nagahara H., Kushiro I., Mysen B., and Mori H. (1988)Experimental vaporization and condensation of olivinesolid solution. Nature 331, 516–518.

Nagashima K., Krot A. N., and Yurimoto H. (2004) Stardustsilicates from primitive meteorites. Nature 428, 921–924.

Nagashima K., Takeda Y., Itoh S., and Yurimoto H.(2003) 16O-rich olivine in igneous rim of type I chondrulefrom CR2 chondrite. Meteorit. Planet. Sci. 38, A77.

Nakamoto T., Hayashi M. R., Kita N. T., and TachibanaS. (2005) Chondrule-forming shock waves in the solarnebula by X-ray flares. In Chondrites and the Proto-planetary Disk (eds. A. N. Krot, E. R. D. Scott, andB. Reipurth). Astron. Soc. Pacific Conf. Ser., vol. 341,pp. 883–902.

Nakamura T., Noguchi T., Zolensky M. E., and TanakaM. (2003) Mineralogy and noble-gas signatures of thecarbonate-rich lithology of the Tagish Lake carbon-aceous chondrite: evidence for an accretionary breccia.Earth Planet. Sci. Lett. 207, 83–101.

Nakamura T., Tomeoka K., Takaoka N., Sekine T., andTakeda H. (2000) Impact-induced textural changes ofCV carbonaceous chondrites: experimental reproduction.Icarus 146, 289–300.

Nakamura T., Tomeoka K., and Takeda H. (1992) Shockeffects of the Leoville CV carbonaceous chondrite: atransmission electron microscope study. Earth Planet.Sci. Lett. 114, 159–170.

Nelson A. F. and Ruffert M. (2005) A proposed origin forchondrule-forming shocks in the solar nebula. In Chond-rites and the Protoplanetary Disk (eds. A. N. Krot, E. R.D. Scott, and B. Reipurth). Astron. Soc. Pacific Conf.Ser., vol. 341, pp. 903–912.

Newsom H. E. and Drake M. J. (1979) The origin of metalclasts in the Bencubbin breccia. Geochim. Cosmochim.Acta 43, 689–707.

Newton J., Bischoff A., Arden J. W., Franchi I. A., GeigerT., Greshake A., and Pillinger C. T. (1995) Acfer 094, auniquely primitive carbonaceous chondrite from the Sa-hara. Meteorit. Planet. Sci. 30, 47–56.

Newton J., Franchi I. A., and Pillinger C. T. (2000) Theoxygen isotopic record in enstatite meteorites. Meteorit.Planet. Sci. 35, 689–698.

Niemeyer S. (1988) Isotopic diversity in nebular dust: thedistribution of Ti isotopic anomalies. Geochim. Cosmo-chim. Acta 52, 2941–2954.

Nittler L. R., Alexander C. M. O. D., Gao X., Walker R.M., and Zinner E. K. (1994) Interstellar grains from theTieschitz ordinary chondrite. Nature 370, 443–446.


Noguchi T. (1993) Petrology and mineralogy of CK chond-rites: implications for the metamorphism of the CKchondrite parent body. Proc. NIPR Symp. Antarct.Meteorit. 6, 204–233.

Nuth J. A., III., Brearley A. J., and Scott E. R. D. (2005)Microcrystals and amorphous material in comets andprimitive meteorites: keys to understanding processes inthe early solar system. In Chondrites and the Protoplan-etary Disk (eds. A. N. Krot, E. R. D. Scott, and B. Re-ipurth). Astron. Soc. Pacific Conf. Ser., vol. 341, pp. 675–700.

Nuth J. A., Reitmeijer F. J. M., and Hill H. G. M. (2002)Condensation processes in astrophysical environ-ments: the composition and structure of cometary grains.Meteorit. Planet. Sci. 37, 1579–1590.

Okazaki R., Takaoka N., Nagao K., Sekiya M., andNakamura T. (2001) Noble-gas-rich chondrules in anenstatite meteorite. Nature 412, 795–798.

Okazaki R., Takaoka N., Nakamura T., and Nagao K.(2000) Cosmic-ray exposure ages of enstatite chondrites.Antarct. Meteorit. Res. 13, 153–169.

Ott U. (1993) Interstellar grains in meteorites. Nature 364,25–33.

Ott U. (2001) Presolar grains in meteorites: an overviewand some implications. Planet. Space Sci. 49, 763–767.

Ozawa K. and Nagahara H. (2001) Chemical and isotopicfractionations by evaporation and their cosmochemicalimplications. Geochim. Cosmochim. Acta 65,2171–2199.

Pack A., Palme H., and Shelley J. M. G. (2005) Origin ofchondritic forsterite grains. Geochim. Cosmochim. Acta69, 3159–3182.

Pack A., Yurimoto H., and Palme H. (2004) Petrographicand oxygen-isotopic study of refractory forsterites fromR-chondrite Dar al Gani 013 (R3. 5-6), unequilibratedordinary and carbonaceous chondrites. Geochim. Cosmo-chim. Acta 68, 1135–1157.

Palme H. (2000) Are there chemical gradients in the innersolar system? Space Sci. Rev. 92, 237–262.

Palme H. (2001) Chemical and isotopic heterogeneity inprotosolar matter. Phil. Trans. Roy. Soc. Lond. A 359,2061–2075.

Palme H. and Boynton W. V. (1993) Meteoritic constraintson conditions in the solar nebula. In Protostars andPlanets III (eds. E. H. Levy and J. I. Lunine). Universityof Arizona Press, Tucson, AZ, pp. 979–1004.

Palme H. and Fegley B. (1990) High temperature conden-sation of iron-rich olivine in the solar nebula. EarthPlanet. Sci. Lett. 101, 180–195.

Palme H., Hutcheon I. D., and Spettel B. (1994) Compo-sition and origin of refractory-metal-rich assemblages ina Ca,Al-rich Allende inclusion. Geochim. Cosmochim.Acta 58, 495–513.

Palme H. and Klerner S. (2000) Formation of chondrulesand matrix in carbonaceous chondrites. Meteorit. Planet.Sci. 35, A124.

Palme H., Kurat G., Spettel B., and Burghele A. (1989)Chemical composition of an unusual xenolith of the Al-lende meteorite. Z. Naturforsch. 44a, 1005–1014.

Palme H., Larimer J. W., and Lipschutz M. E. (1988)Moderately volatile elements. In Meteorites and the EarlySolar System (eds. J. F. Kerridge and M. S. Matthews).University of Arizona Press, Tucson, AZ, pp. 436–461.

Palme H. and Wlotzka F. (1976) A metal particle from aCa, Al-rich refractory inclusion from the meteorite,Allende, and the condensation of refractory siderophileelements. Earth Planet. Sci. Lett. 33, 45–60.

Paque J. M. and Cuzzi J. N. (1997) Physical characteristicsof chondrules and rims, and aerodynamic sorting in thesolar nebula. Lunar Planet. Sci. 28, 1071–1072.

Patzer A. and Schultz L. (2002a) Noble gases in enstatitechondrites. Part I: exposure ages, pairing, and weather-ing effects. Meteorit. Planet. Sci. 36, 947–961.

Patzer A. and Schultz L. (2002b) Noble gases in enstatitechondrites. Part II: Meteorit. Planet. Sci. 37, 601–612.

Pauls A. and Gladman B. (2005) Decoherence time scalesfor ‘‘meteoroid streams.’’ Meteorit. Planet. Sci. 40,1141–1256.

Petaev M. I., Meibom A., Krot A. N., Wood J. A., and KeilK. (2001) The condensation origin of zoned metal grainsin Queen Alexandra Range 94411: implications for theformation of the Bencubbin-like chondrites. Meteorit.Planet. Sci. 36, 96–106.

Petaev M. I. and Wood J. A. (2000) The condensation withpartial isolation (CWPI) model of condensation in thesolar nebula. Meteorit. Planet. Sci. 33, 1123–1137.

Podosek F. A., Ott U., Brannon J. C., Neal C. R., Bern-atowicz T. J., Swan P., andMahan S. E. (1997) Thoroughlyanomalous chromium in Orgueil. Meteorit. Planet. Sci. 32,617–627.

Pravdivtseva O. V., Krot A. N., Hohenberg C. M., MeshikA. P., Weisberg M. K., and Keil K. (2003) The I–Xerecord of alteration in the Allende CV chondrite. Geo-chim. Cosmochim. Acta 67, 5011–5026.

Reipurth B. (2005) The early Sun: evolution and dynamicenvironment. In Chondrites and the Protoplanetary Disk(eds. A. N. Krot, E. R. D. Scott, and B. Reipurth). As-tron. Soc. Pacific Conf. Ser., vol. 341, pp. 54–77.

Reisener R., Meibom A., Krot A. N., Goldstein J. I., andKeil K. (2000) Microstructure of condensate Fe–Nimetal particles in the CH chondrite PAT91546. In LunarPlanet. Sci. XXXI, #1445. The Lunar and Planetary In-stitute, Houston (CD-ROM).

Richter F. M., Mendybaev R. A., and Davis A. M. (2006)Conditions in the protoplanetary disk as seen by type BCAIs. Meteorit. Planet. Sci. 41, 83–93.

Rietmeijer F. J. M. (1998) Interplanetary dust particles. InPlanetary Materials (ed. J. J. Papike). Rev. Mineral., vol.36, pp. 2-1–2-95.

Rietmeijer F. J. M. (2002) The earliest chemical dustevolution in the solar nebula. Chem. Erde 62, 1–45.

Rietmeijer F. J. M., Hallenbeck S. L., Nuth J. A., andKarner J. M. (2002) Amorphous magnesiosilicate smokesannealed in vacuum: the evolution of magnesium silicatesin circumstellar and cometary dust. Icarus 156, 269–286.

Roedder E. (1981) Significance of Ca–Al-rich silicate meltinclusions in olivine crystals from the Murchison type IIcarbonaceous chondrite. Bull. Mineral. 104, 339–353.

Rubin A. E. (1990) Kamacite and olivine in ordinarychondrites: intergroup and intragroup relationships.Geochim. Cosmochim. Acta 54, 1217–1232.

Rubin A. E. (1995) Petrologic evidence for collisional hea-ting of chondritic asteroids. Icarus 113, 156–167.

Rubin A. E. (2000) Petrologic, geochemical and experi-mental constraints on models of chondrule formation.Earth Sci. Rev. 50, 3–27.

Rubin A. E. (2002) Post-shock annealing of Miller Range99301 (LL6): implications for impact heating of ordinarychondrites. Geochim. Cosmochim. Acta 66, 3327–3337.

Rubin A. E. (2003) Chromite-plagioclase assemblages as anew shock indicator: implications for the shock andthermal histories of ordinary chondrites. Geochim. Cosmo-chim. Acta 67, 2695–2709.

Rubin A. E. (2004) Postshock annealing and postannealingshock in equilibrated ordinary chondrites: implicationsfor the thermal and shock histories of chondritic aster-oids. Geochim. Cosmochim. Acta 68, 673–689.

Rubin A. E. (2005) What heated the asteroids? Sci. Am.292, 80–87.

Rubin A. E. and Brearley A. J. (1996) A critical evaluationof the evidence for hot accretion. Icarus 124, 86–96.

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 and sil-icate globules in Gujba: a new Bencubbin-like meteoritefall. Geochim. Cosmochim. Acta 67, 3283–3298.

References 69

Rubin A. E., Sailer A. L., and Wasson J. T. (1999) Troilitein the chondrules of type 3 ordinary chondrites:implications for chondrule formation. Geochim.Cosmochim. Acta 63, 2281–2298.

Rubin A. E., Scott E. R. D., and Keil K. (1997) Shockmetamorphism of enstatite chondrites. Geochim. Cosmo-chim. Acta 61, 847–858.

Rubin A. E., Ulff-Møller F., Wasson J. T., and Carlson W.D. (2001) The Portales Valley meteorite breccia: evidencefor impact-induced melting and metamorphism of anordinary chondrite. Geochim. Cosmochim. Acta 65, 323–342.

Rubin A. E. and Wasson J. T. (1988) Chondrules in theOrnans CO3 meteorite and the timing of chondrule for-mation relative to nebular fractionation events. Geochim.Cosmochim. Acta 52, 425–432.

Rubin A. E. and Wasson J. T. (2005) Non-spherical lobatechondrules in CO3.0 Y-81020: general implications forthe formation of low-FeO porphyritic chondrules in COchondrites. Geochim. Cosmochim. Acta 69, 211–220.

Rubin A. E., Zolensky M. E., and Bodnar R. J. (2002) Thehalite-bearing Zag and Monahans (1998) meteorite brec-cias: shock metamorphism, thermal metamorphism andaqueous alteration on the H-chondrite parent body.Meteorit. Planet. Sci. 37, 124–141.

Russell S. S., Huss G. R., Fahey A. J., Greenwood R. C.,Hutchison R., and Wasserburg G. J. (1998) An isotopicand petrologic study of calcium–aluminum-rich inclu-sions from CO3 meteorites. Geochim. Cosmochim. Acta62, 689–714.

Russell S. S., Krot A. N., Huss G. R., Keil K., Itoh S.,Yurimoto H., and MacPherson G. J. (2005) The geneticrelationship between refractory inclusions and chond-rules. In Chondrites and the Protoplanetary Disk (eds.A. N. Krot, E. R. D. Scott, and B. Reipurth). Astron.Soc. Pacific Conf. Ser., vol. 341, pp. 317–350.

Russell S. S., MacPherson G. J., Leshin L. A., andMcKeegan K. D. (2000) 16O enrichments in aluminum-rich chondrules from ordinary chondrites. Earth Planet.Sci. Lett. 184, 57–74.

Russell S. S., Srinivasan G., Huss G. R., Wasserburg G. J.,and MacPherson G. J. (1996) Evidence for widespread26Al in the solar nebula and new constraints for nebulatimescales. Science 273, 757–762.

Ruzicka A., Killgore M., Mittlefehldt D. W., and Fries M.D. (2005) Portales Valley: petrology of a metallic-meltmeteorite breccia. Meteorit. Planet. Sci. 40, 261–295.

Ruzicka A., Snyder G. A., and Taylor L. A. (1998) Mega-chondrules and large, igneous-textured clasts in Julesberg(L3) and other ordinary chondrites: vapor-fractionation,shock-melting, and chondrule formation. Geochim. Cosmo-chim. Acta 62, 1419–1442.

Ryerson F. J. and McKeegan K. D. (1994) Determinationof oxygen self-diffusion in akermanite, anorthite,diopside, and spinel: implications for oxygen isotopicanomalies and the thermal histories of Ca–Al-richinclusions. Geochim. Cosmochim. Acta 58, 3713–3734.

Sahijpal S., McKeegan K. D., Goswami J. N., and Davis A.M. (2000) Oxygen isotopic compositions of Murchisonhibonites with wide-ranging radiogenic and neutron-richstable isotope anomalies. In Lunar Planet. Sci. XXXI,#1502. The Lunar and Planetary Institute, Houston(CD-ROM).

Sanders I. S. (1996) A chondrule-forming scenario involvingmolten planetesimals. In Chondrules and the ProtoplanetaryDisk (eds. R. H. Hewins, R. H. Jones, and E. R. D. Scott).Cambridge University Press, pp. 327–334.

Sanders I. S. and Taylor G. J. (2005) Implications of 26Al innebular dust: formation of chondrules by disruption ofmolten planetesimals. In Chondrites and the Protoplane-tary Disk (eds. A. N. Krot, E. R. D. Scott, and B.Reipurth). Astron. Soc. Pacific Conf. Ser., vol. 341, pp.915–932.

Schersten A., Elliott T., Hawkesworth C., Russell S., andMasarik J. (2006) Hf–W evidence for rapid differentia-tion of iron meteorite parent bodies. Earth Planet. Sci.Lett. 241, 530–542.

Schmitt R. T. (2000) Shock experiments with the H6chondrite Kernouve: pressure calibration of microscopicshock effects. Meteorit. Planet. Sci. 35, 545–560.

Schneider D. M., Benoit P. H., Kracher A., and Sears D.W. G. (2003) Metal size distribution in EH andEL chondrites. Geophys. Res. Lett. 30(No. 8), 1420doi:10.1029/2002GL016672.

Scott E. R. D. (1988) A new kind of primitive chondrite,Allan Hills 85085. Earth Planet. Sci. Lett. 91, 1–18.

Scott E. R. D. (2002) Meteoritic evidence for the accretionand collisional evolution of asteroids. In Asteroids III(eds. W. F. Jr. Bottke, A. Cellino, P. Paolicchi, and R. P.Binzel). University of Arizona Press, Tucson, AZ, pp.697–709.

Scott E. R. D., Barber D. J., Alexander C. M., HutchisonR., and Peck J. A. (1988) Primitive material surviving inchondrites: matrix. In Meteorites and the Early SolarSystem (eds. J. F. Kerridge and M. S. Matthews).University of Arizona Press, Tucson, AZ, pp. 718–745.

Scott E. R. D. and Jones R. H. (1990) Disentangling neb-ular and asteroidal features of CO3 carbonaceous chond-rites. Geochim. Cosmochim. Acta 54, 2485–2502.

Scott E. R. D., Jones R. H., and Rubin A. E. (1994) Clas-sification, metamorphic history, and pre-metamorphiccomposition of chondrules. Geochim. Cosmochim. Acta58, 1203–1209.

Scott E. R. D., Keil K., and Stoffler D. (1992) Shock met-amorphism of carbonaceous chondrites. Geochim. Cos-mochim. Acta 56, 4281–4293.

Scott E. R. D. and Krot A. N. (2001) Oxygen isotopiccompositions and origins of calcium–aluminum-richinclusions and chondrules. Meteorit. Planet. Sci. 36,1307–1319.

Scott E. R. D. and Krot A. N. (2005a) Thermal processingof silicate dust in the solar nebula: Clues from primitivechondrite matrices. Astrophys. J. 623, 571–578.

Scott E. R. D. and Krot A. N. (2005b) Chondritic mete-orites and the high-temperature nebular origins of theircomponents. In Chondrites and the Protoplanetary Disk(eds. A. N. Krot, E. R. D. Scott, and B. Reipurth). As-tron. Soc. Pacific Conf. Ser., vol. 341, pp. 15–53.

Scott E. R. D., Love S. G., and Krot A. N. (1996) For-mation of chondrules and chondrites in the protoplan-etary nebula. In Chondrules and the Protoplanetary Disk(eds. R. H. Hewins, R. H. Jones, and E. R. D. Scott).Cambridge University Press, Cambridge, pp. 87–96.

Scott E. R. D. and Newsom H. (1989) Planetary compo-sitions—clues from meteorites and asteroids. Z. Nat-urforsch. 44a, 924–934.

Scott E. R. D. and Taylor G. J. (1983) Chondrules andother components in C, O, and E chondrites: similaritiesin their properties and origins. J. Geophys. Res. 88(sup-pl.), B275–B286.

Scott E. R. D., Taylor G. J., Newsom H., Herbert F.,Zolensky M., and Kerridge J. F. (1989) Chemical, ther-mal and impact processing of asteroids. In Asteroids II(eds. R. P. Binzel, T. Gehrels, and M. S. Matthews).University of Arizona Press, Tucson, AZ, pp. 701–739.

Sears D. W. G. (1998) The case for rarity of chondrules andcalcium–aluminum-rich inclusions in the early solar sys-tem and some implications for astrophysical models.Astrophys. J. 498, 773–778.

Sears D. W. G. (2005) The Origin of Chondrules andChondrites. Cambridge University Press, Cambridge,222pp.

Sears D. W. G., Batchelor D. J., Lu J., and Keck B. D.(1991) Metamorphism of CO and CO-like chondrites andcomparisons with type 3 ordinary chondrites. Proc.NIPR Symp. Antarct. Meteorit. 4, 319–343.


Sears D. W. G. and Hasan F. A. (1987) The type three or-dinary chondrites: a review. Surv. Geophys. 9, 43–97.

Sears D. W. G., Huang S., and Benoit P. H. (1995b) Chond-rule formation, metamorphism, brecciation, an importantnew primary chondrule group, and the classification ofchondrules. Earth Planet. Sci. Lett. 131, 27–39.

Sears D. W. G., Huang S., and Benoit P. H. (1996) Open-system behavior during chondrule formation. In Chond-rules and the Protoplanetary Disk (eds. R. H. Hewins, R.H. Jones, and E. R. D. Scott). Cambridge UniversityPress, Cambridge, pp. 221–231.

Sears D. W. G., Morse A. D., Hutchison R., Guimon R.K., Jie L., Alexander C. M. O. D., Benoit P. H., WrightI., Pillinger C., Xie T., and Lipschutz M. E. (1995a)Metamorphism and aqueous alteration in low petrogra-phic type ordinary chondrites. Meteoritics 30, 169–181.

Sheng Y. J., Hutcheon I. D., and Wasserburg G. J. (1991)Origin of plagioclase–olivine inclusions in carbonaceouschondrites. Geochim. Cosmochim. Acta 55, 581–599.

Shu F. H., Shang H., Gounelle M., and Glassgold A. E.(2001) The origin of chondrules and refractory inclusionsin chondritic meteorites. Astrophys. J. 548, 1029–1050.

Shu F. H., Shang H., and Lee T. (1996) Toward anastrophysical theory of chondrites. Science 271, 1545–1552.

Simon S. B., Davis A. M., Grossman L., and McKeegan K.D. (2002) A hibonite–corundum inclusion from Murchi-son: a first generation condensate from the solar nebula.Meteorit. Planet. Sci. 37, 533–548.

Simon S. B., Yoneda S., Grossman L., and Davis A. M.(1994) A CaAl4O7-bearing refractory spherule from Mu-rchison: evidence for very high-temperature melting inthe solar nebula. Geochim. Cosmochim. Acta 58, 1937–1949.

Skinner W. R. and Leenhouts J. M. (1993) Size distribu-tions and aerodynamic equivalence of metal chondrulesand silicate chondrules in Acfer 059. In Lunar Planet. Sci.XXIV, The Lunar and Planetary Institute, Houston, pp.1315–1316.

Spurny P., Oberst J., and Heinlein D. (2003) Photographicobservations of Neuschwanstein, a second meteoritefrom the orbit of the Pribram chondrite. Nature 423,151–153.

Srinivasan G., Huss G. R., and Wasserburg G. J. (2000a) Apetrographic, chemical and isotopic study of calcium–aluminum-rich inclusions and aluminum-rich chondrulesfrom the Axtell (CV3) chondrite. Meteorit. Planet. Sci.35, 1333–1354.

Srinivasan G., Krot A. N., and Ulyanov A. A. (2000b)Aluminum–magnesium systematics in anorthite-richchondrules and calcium–aluminum-rich inclusions fromthe reduced CV chondrite Efremovka. Meteorit. Planet.Sci. 35, A151–A152.

Steele I. M. (1986) Compositions and textures of relic fors-terite in carbonaceous and unequilibrated ordinarychondrites. Geochim. Cosmochim. Acta 50, 1379–1395.

Steele I. M. (1989) Compositions of isolated forsterites in Or-nans (C3O). Geochim. Cosmochim. Acta 53, 2069–2079.

Stoffler D. (1997) Minerals in the deep Earth: a messagefrom the asteroid belt. Science 278, 1576–1577.

Stoffler D., Keil K., and Scott E. R. D. (1991) Shock met-amorphism of ordinary chondrites. Geochim. Cosmo-chim. Acta 55, 3845–3867.

Strebel R., Huth J., and Hoppe P. (2000) In situ location bycathodoluminescence and SIMS isotopic analysesof small corundum grains in the Krymka meteorite. InLunar Planet. Sci. XXXI. The Lunar and PlanetaryInstitute, Houston (CD-ROM).

Sugiura N., Miyazaki A., Hiyagon H., Kimura K., andPetaev M. I. (2006) Nebular history of amoeboid olivineaggregates. In Lunar Planet. Sci. XXXVII, #1266. TheLunar and Planetary Institute, Houston (CD-ROM).

Sugiura N., Miyazaki A., Kimura M., and Hiyagon H. (2004)Cathodoluminescence study of amoeboid olivine aggregatesin Yamato 81020. Meteorit. Planet. Sci. 39, A101.

Susa H. and Nakamoto T. (2002) On the maximal size ofchondrules in shock wave heating model. Astrophys. J.564, L57–L60.

Symes S. J. K., Sears D. W. G., Huang S. X., and Benoit P.H. (1998) The crystalline lunar spherules: their formationand implications for the origin of meteoritic chondrules.Meteorit. Planet. Sci. 33, 13–29.

Tanaka K. K., Tanaka H., Nakazawa K., and NakagawaY. (1998) Shock heating due to accretion of a clumpycloud onto a protoplanetary disk. Icarus 134, 137–154.

Taylor S. R. (2001) Solar System Evolution: A New Per-spective. Cambridge University Press, Cambridge, 460pp.

Taylor G. J., Maggiore P., Scott E. R. D., Rubin A. E., andKeil K. (1987) Original structures and fragmentationhistories of asteroids: evidence from meteorites. Icarus69, 1–13.

Taylor G. J., Scott E. R. D., and Keil K. (1983) Cosmicsetting for chondrule formation. In Chondrules and theirOrigin (ed. E. A. King). The Lunar and Planetary Insti-tute, Houston, pp. 262–278.

Thiemens M. H. (1999) Mass-independent isotopic effectsin planetary atmospheres and the early solar system.Science 283, 341–345.

Tissandier L., Libourel G., and Robert F. (2002) Gas–metal interactions and their bearing on chondrule for-mation. Meteorit. Planet. Sci. 37, 1377–1389.

Tomeoka K. and Buseck P. R. (1990) Phyllosilicates in theMokoia CV carbonaceous chondrite: evidence for aque-ous alteration in an oxidizing environment. Geochim.Cosmochim. Acta 54, 1745–1754.

Tomeoka K. and Itoh D. (2004) Sodium-metasomatism inchondrules in CO3 chondrites: relationship to parentbody thermal metamorphism. Meteorit. Planet. Sci. 39,1359–1373.

Tomeoka K., Ohnishi I., and Nakamura N. (2001) Silicatedarkening in the Kobe CK chondrite: evidence for shockmetamorphism at high temperatures. Meteorit. Planet.Sci. 36, 1535–1545.

Tomeoka K. and Tanimura I. (2000) Phyllosilicate-richchondrule rims in the Vigarano CV3 chondrite: evidencefor parent-body processes. Geochim. Cosmochim. Acta64, 1971–1988.

Tomeoka K., Yamahana Y., and Sekine T. (1999) Exper-imental shock metamorphism of the Murchison CM car-bonaceous chondrite. Geochim. Cosmochim. Acta 63,3683–3703.

Trieloff M., Jessberger E. K., Herrwerth I., Hopp J., FieniC., Ghelis M., Bourot-Denise M., and Pellas P. (2003)Structure and thermal history of the H-chondriteasteroid revealed by thermochronometry. Nature 422,502–506.

Trigo-Rodriguez J. M., Rubin A. E., and Wasson J. T.(2006) Non-nebular origin of dark mantles aroundchondrules and inclusions in CM chondrites. Geochim.Cosmochim. Acta 70, 1271–1290.

Uesugi M., Akaki T., Sekiya M., and Nakamura T. (2005)Motion of iron sulfide inclusions inside a shock-meltedchondrule. Meteorit. Planet. Sci. 40, 1103–1114.

Van Schmus W. R. and Wood J. A. (1967) A chemical-petrologic classification for the chondritic meteorites.Geochim. Cosmochim. Acta 31, 747–765.

Varela M. E., Kurat G., and Zinner E. (2005) A liquid-supported condensation of major minerals in the solarnebula: evidence from glasses in the Kaba (CV3) chond-rite. Icarus 178, 553–569.

Virag A., Zinner E., Amari S., and Anders E. (1991) An ionmicroprobe study of corundum in the Murchison mete-orite—implications for 26Al and 16O in the early solarsystem. Geochim. Cosmochim. Acta 55, 2045–2062.

References 71

Wark D. and Boynton W. V. (2001) The formation of rimson calcium–aluminum-rich inclusions. Step I: Flash heat-ing. Meteorit. Planet. Sci. 36, 1135–1166.

Wasson J. T. (1985) Meteorites: Their Record of Early SolarSystem History. W. H. Freeman, New York.

Wasson J. T. (2000) Oxygen-isotopic evolution of the solarnebula. Rev. Geophys. Space Phys. 38, 491–512.

Wasson J. T. and Kallemeyn G. W. (1988) Compositions ofchondrites. Phil. Trans. Roy. Soc. Lond. A 325, 535–544.

Wasson J. T., Krot A. N., Lee M. S., and Rubin A. E.(1995) Compound chondrules. Geochim. Cosmochim.Acta 59, 1847–1869.

Wasson J. T. and Rubin A. E. (2003) Ubiquitious low-FeOrelict grains in type II chondrules and limited overgrow-ths on phenocrysts following the final melting events.Geochim. Cosmochim. Acta 67, 2239–2250.

Wasson J. T., Yurimoto H., and Russell S. S. (2001)16O-rich melilite in CO3.0 chondrites: possible formationof common, 16O-poor melilite by aqueous altera-tion. Geochim. Cosmochim. Acta 65, 4539–4549.

Watt L. E., Bland P. A., Prior D. J., and Russell S. S.(2006) Fabric analysis of Allende matrix using EBSD.Meteorit. Planet. Sci. 41, 989–1001.

Weber D. and Bischoff A. (1997) Refractory inclusions inthe CR chondrite Acfer 059-El Djouf 001: petrology,chemical composition, and relationship to inclusion pop-ulations in other types of carbonaceous chondrites.Chem. Erde 57, 1–24.

Weber D., Zinner E., and Bischoff A. (1995) Trace elementabundances and magnesium, calcium, and titanium iso-topic compositions of grossite-containing inclusions fromthe carbonaceous chondrite Acfer 182. Geochim. Cosmo-chim. Acta 59, 803–823.

Weidenschilling S. J., Marzari F., and Hood L. L. (1998)The origin of chondrules at Jovian resonances. Science279, 681–684.

Weinbruch S., Muller W. F., and Hewins R. H. (2001) Atransmission electron microscope study of exsolutionand coarsening in iron-bearing clinopyroxene from syn-thetic analogues of chondrules. Meteorit. Planet. Sci. 36,1237–1248.

Weinbruch S., Palme H., Muller W. F., and El Goresy A.(1990) FeO-rich rims and veins in Allende forsterite:evidence for high-temperature condensation at oxidizingconditions. Meteoritics 25, 115–125.

Weinbruch S., Palme H., and Spettel B. (2000) Refractoryforsterite in primitive meteorites: condensates from thesolar nebula? Meteorit. Planet. Sci. 35, 161–171.

Weisberg M. K., Connolly H. C., Jr., and Ebel D. S. (2004)Petrology andorigin of amoeboidolivine aggregates inCRchondrites. Meteorit. Planet. Sci. 39, 1741–1753.

Weisberg M. K. and Prinz M. (1998) Fayalitic olivine inCV3 chondrite matrix and dark inclusions: a nebularorigin. Meteorit. Planet. Sci. 33, 1087–1099.

Weisberg M. K. and Prinz M. (2000) The GrosvenorMountains 95577 CR1 chondrite and hydration of theCR chondrites. Meteorit. Planet. Sci. 35, A168.

Weisberg M. K., Prinz M., Clayton R. N., and Mayeda T.K. (1993) The CR (Renazzo-type) carbonaceous chond-rite group and its implications. Geochim. Cosmochim.Acta 57, 1567–1586.

Weisberg M. K., Prinz M., Clayton R. N., Mayeda T. K.,Grady M. M., Franchi I., Pillinger C. T., and KallemeynG. W. (1996) The K chondrite grouplet. Geochim. Cosmo-chim. Acta 60, 4253–4263.

Weisberg M. K., Prinz M., Clayton R. N., Mayeda T. K.,Sugiura N., Zashu S., and Ebihara M. (2001) A newmetal-rich chondrite grouplet. Meteorit. Planet. Sci. 36,401–418.

Weisberg M. K., Prinz M., and Fogel R. A. (1994) Theevolution of enstatite and chondrules in unequilibratedenstatite chondrites. Meteoritics 29, 362–373.

Weisberg M. K., Prinz M., Kennedy A., and Hutcheon I.D. (1991) Trace elements in refractory-rich inclusions inCR2 chondrites. Meteoritics 26, 407–408.

Weisberg M. K., Prinz M., and Nehru C. E. (1990) TheBencubbin chondrite breccia and its relationship to CRchondrites and the ALH85085 chondrite. Meteoritics 25,269–279.

Weisberg M. K., Zolensky M. E., and Prinz M. (1997)Fayalitic olivine in matrix of the Krymka LL3.1 chond-rite: vapor–solid growth in the solar nebula. Meteorit.Planet. Sci. 32, 791–801.

Wiens R. C., Huss G. R., and Burnett D. S. (1999) Thesolar oxygen isotopic composition: predictions and im-plications for solar nebula processes. Meteorit. Planet.Sci. 34, 99–107.

Wieler R. (2002) Cosmic-ray-produced noble gases in mete-orites. In Noble Gases in Geochemistry and Cosmochem-istry (eds. D. Porcelli, C. J. Ballentine, and R. Wieler).Rev. Mineral. Geochem., pp. 125–170.

Williams C. V., Keil K., Taylor G. J., and Scott E. R. D.(1999) Cooling rates of equilibrated clasts in ordinarychondrite regolith breccias: implications for parent bodyhistories. Chem. Erde 59, 287–305.

Wilson L., Keil K., Browning L. B., Krot A. N., andBourcier W. (1999) Early aqueous alteration, explosivedisruption, and reprocessing of asteroids. Meteorit.Planet. Sci. 34, 541–557.

Wlotzka F. (2005) Cr spinel and chromite as petrogeneticindicators in ordinary chondrites: equilibration temper-atures of petrologic types 3.7 to 6. Meteorit. Planet. Sci.40, 1673–1702.

Wlotzka F., Palme H., Spettel B., Wanke H., FredrikssonK., and Noonan A. F. (1983) Alkali differentiation in LLchondrites. Geochim. Cosmochim. Acta 47, 743–757.

Wolf D. and Palme H. (2001) The solar system abundanceof phosphorus and titanium and the nebular volatility ofphosphorus. Meteorit. Planet. Sci. 36, 559–571.

Wood J. A. (1963) On the origin of chondrules and chond-rites. Icarus 2, 152–180.

Wood J. A. (1967) Chondrites: their metallic minerals,thermal histories, and parent planets. Icarus 6, 1–49.

Wood J. A. (1988) Chondritic meteorites and the solarnebula. Annu. Rev. Earth Planet. Sci. 16, 53–72.

Wood J. A. (1996a) Unresolved issues in the formation ofchondrules and chondrites. In Chondrules and the Proto-planetary Disk (eds. R. H. Hewins, R. H. Jones, and E.R. D. Scott). Cambridge University Press, Cambridge,pp. 55–69.

Wood J. A. (1996b) Processing of chondritic and planetarymaterial in spiral density waves in the nebula. Meteorit.Planet. Sci. 31, 641–645.

Wood J. A. (2000a) Pressure and temperature profiles in thesolar nebula. Space Sci. Rev. 92, 87–93.

Wood J. A. (2000b) The beginning: swift and violent. SpaceSci. Rev. 92, 97–112.

Wood J. A. (2004) Formation of chondritic refractory in-clusions: the astrophysical setting. Geochim. Cosmochim.Acta 68, 4007–4021.

Wood J. A. and Hashimoto A. (1993) Mineral equilibriumin fractionated nebula systems. Geochim. Cosmochim.Acta 57, 2377–2388.

Wooden D. H., Harker D. E., and Brearley A. J. (2005)Thermal processing and radial mixing of dust: evidencefrom comets and primitive chondrites. In Chondrites andthe Protoplanetary Disk (eds. A. N. Krot, E. R. D. Scott,and B. Reipurth). Astron. Soc. Pacific Conf. Ser., vol.341, pp. 774–808.

Woolum D. S. and Cassen P. (1999) Astronomical con-straints on nebular temperatures: implications for plan-etesimal formation. Meteorit. Planet. Sci. 34, 897–907.

Yoshino T., Walter M. J., and Katsura T. (2003) Core for-mation triggered by permeable flow. Nature 422, 154–157.


Yoshitake M., Koide Y., and Yurimoto H. (2005)Correlations between oxygen-isotopic composition andpetrologic setting in a coarse-grained Ca, Al-rich inclu-sion. Geochim. Cosmochim. Acta 69, 2663–2674.

Young E. D. (2000) Assessing the implications of K isotopecosmochemistry for evaporation in the preplanetary so-lar nebula. Earth Planet. Sci. Lett. 183, 321–333.

Young E. D. (2001) The hydrology of carbonaceous chond-rite parent bodies and the evolution of planet progenitors.Phil. Trans. Roy. Soc. Lond. A 359, 2095–2110.

Young E. D., Ash R. D., England P., and Rumble D. (1999)Fluid flow in chondritic parent bodies: deciphering thecompositions of planetesimals. Science 286, 1331–1335.

Young E. D., Zhang K. K., and Schubert G. (2003) Condi-tions for pore water convection within carbonaceouschondrite parent bodies—implications for planetesimal sizeand heat production. Earth Planet. Sci. Lett. 213, 249–259.

Yu Y. and Hewins R. H. (1998) Transient heating andchondrule formation: evidence from sodium loss in flashheating simulation experiments. Geochim. Cosmochim.Acta 62, 159–172.

Yurimoto H., Ito M., and Nagasawa H. (1998) Oxygenisotope exchange between refractory inclusion in Allendeand solar nebula gas. Science 282, 1874–1877.

Yurimoto H., Koike O., Nagahara H., Morioka M., andNagasawa H. (2000) Heterogeneous distribution of Mgisotopes in anorthite single crystal from Type B CAIs inAllende meteorite. In Lunar Planet. Sci. XXXI, #1593.The Lunar and Planetary Institute, Houston (CD-ROM).

Yurimoto H. and Kuramoto K. (2002) A possible scenariointroducing heterogeneous oxygen isotopic distributioninto protoplanetary disks. Meteorit. Planet. Sci. 37,A153.

Yurimoto H., Rubin A. E., Itoh S., andWasson J. T. (2001)Non-stoichiometric Al-rich spinel in an ultrarefractoryinclusion from CO chondrite. In Lunar Planet. Sci.

XXXIII, #1557. The Lunar and Planetary Institute,Houston (CD-ROM).

Yurimoto H. and Wasson J. T. (2002) Extremely rapidcooling of a carbonaceous-chondrite chondrule contain-ing very 16O-rich olivine and a 26Mg excess. Geochim.Cosmochim. Acta 66, 4355–4363.

Zanda B. (2004) Chondrules. Earth Planet. Sci. Lett. 224,1–17.

Zanda B., Bourot-Denise M., Perron C., and Hewins R. H.(1994) Origin and metamorphic redistribution of silicon,chromium, and phosphorus in the metal of chondrites.Science 265, 1846–1849.

Zhang Y., Huang S., Schneider D., Benoit P. H., DeHart J.M., Lofgren G. E., and Sears D. W. G. (1996) Pyroxenestructures, cathodoluminescence and the thermal history ofthe enstatite chondrites. Meteorit. Planet. Sci. 31, 87–96.

Zinner E. and Gopel C. (2002) Aluminum-26 in H4 chond-rites: implications for its production and its usefulness asa fine-scale chronometer for early solar system events.Meteorit. Planet. Sci. 37, 1001–1013.

Zolensky M., Barrett R., and Browning L. (1993) Miner-alogy and composition of matrix and chondrule rims incarbonaceous chondrites. Geochim. Cosmochim. Acta 57,3123–3148.

Zolensky M. and Ivanov A. (2003) The Kaidun microbrec-cia meteorite: a harvest from the inner and outer asteroidbelt. Chem. Erde 63, 185–246.

Zolensky M. E. 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.

Zolensky M. E., Mittlefehldt D. W., Lipschutz M. E.,Wang M.-S., Clayton R. N., Mayeda T. K., Grady M.M., Pillinger C., and Barber D. (1997) CM chondritesexhibit the complete petrologic range from type 2 to 1.Geochim. Cosmochim. Acta 61, 5099–5115.

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