Chondritic Meteorites and the High-Temperature Nebular Origins of Their Components

Home , Chondrule, CI chondrite, CM chondrite, Enstatite chondrite, Grossite

Chondrites and the Protoplanetary Disk ASP Conference Series, Vol. 341, 2005 A. N. Krot, E. R. D. Scott, & B. Reipurth, eds.

Edward R. D. Scott and Alexander N. Krot Hawaii Institute of Geophysics and Planetology, University of Hawaii at Manoa, Honolulu, Hawaii 96822, USA

Abstract. We present an overview of the chondrite groups focusing on the major constraints that can be derived on thermal history of silicates in the solar nebula from the mineralogical, chemical, and isotopic properties of the components in the least-altered chondrites. Recent advances in developing a chronology for the formation of CAIs and chondrules and plausible models for their oxygen isotopic compositions provide the basis for understanding their origin. Evidence from short-lived and long-lived isotopes, oxygen isotopes, nuclear isotopic effects and petrologic studies all suggest that refractory inclusions and grains were the first solids to form in the protosolar disk, probably within a period of <0.3 Myr, when the protosun was accreting rapidly (possibly as a class 0 or I protostar). Refractory inclusions formed in an 16O-rich, reducing environment of near-solar composition, <10-4 bar pressure, and temperatures >1300 K. Most chondrules appear to have formed 1-3 Myr after refractory inclusions, when the protosun was accreting more slowly. Chondrules in a single chondrite group probably formed over a much shorter period. Sev- eral types of chondrules formed under diverse conditions that were generally more oxidizing with lower ambient temperatures and higher total pressures or high dust/gas ratios, so that liquids were stable for hours. Formation of type I chondrules involved melting, evaporation, condensation, and accretion of solid, partly melted and completely melted materials. Chondrite matrices are mixtures of materials that probably formed in diverse locations in the solar nebula and traces of presolar materials. Matrices in pristine carbonaceous chondrites are largely composed of crystalline Mg-rich silicates (forsterite and enstatite) and amorphous Fe-Mg silicate. The cooling rates, composition, and structure of the Mg-rich silicates suggest that they probably condensed during heating events that formed chondrules. The near-solar composition of the matrix implies that the amorphous silicates have similar origins and that gas and dust were not separated during condensation. Comets and chondritic, porous interplanetary dust particles have more pre-solar material than chondrites, but also contain abundant forsterite and enstatite crystals resembling those in matrices of primitive chondrite matrices. The associated amorphous silicate is Fe-rich suggesting that the Mg-rich silicates may have formed by nebular condensation rather than by annealing. Since forsterite and enstatite are abundant around many protostars, the processes that heated silicate dust in the solar nebula may be common to other protostellar disks.

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1. Introduction Chondrites are the meteorites that provide the best clues to the composition and origin of the solar system. They are derived from the subset of asteroids that did not melt and are capable of supplying tough rocks that can survive the journey to Earth and atmospheric entry. Additional clues come from the porous, anhydrous, interplanetary dust particles (IDPs), which have broadly chondritic compositions and probably come from comets. Chondrites and the chondritic, porous IDPs both contain a large fraction of crystalline silicates that formed at high temperatures. Since the crystalline silicates, with few exceptions, have solar isotopic compositions and 99.8±0.2% by mass of the interstellar silicates are amorphous (Kemper, Vriend, & Tielens, 2004), we can infer that nearly all of the crystalline silicates in asteroids and comets formed in the solar system. Our purpose here is to review with a minimum of jargon what can be inferred from chondrites, and to a much lesser extent, comets, about the high-temperature processes in the protoplanetary disk (or solar nebula) that converted amorphous silicate into crystals. Chondrules, which are the major constituents of most chondrites, are roughly millimeter-sized particles that were wholly or partly molten in the solar nebula and crystallized in minutes to hours between ~1800 and ~1300 K prior to accretion (Rubin 2000; Zanda 2004; Jones, Grossman, & Rubin, this volume). The two major minerals that crystallized in chondrules are olivine, (MgxFe1-x)2SiO4, and low-Ca pyroxene, MgxFe1-xSiO3, where x is the Mg/(Mg+Fe) ratio. The magnesian end- members are called forsterite, Mg2SiO4, and enstatite, MgSiO3. Olivine and low-Ca pyroxene are also major minerals in the chondrite matrix—the fine-grained silicate material that coats chondrules and other coarse chondritic ingredients and fills the interstices between them (Scott & Krot, 2003; Nuth et al. and Huss et al., this volume). The other important ingredient of chondrites are refractory inclusions, which are composed almost entirely of crystalline silicates and oxides that are rich in Ca, Al, and Ti and formed above 1300 K (MacPherson, 2003; MacPherson et al., this volume). Since Wood (1962, 1963) first proposed that chondrules formed in the solar nebula and were a major component of the material that accreted into the terrestrial planets, there have been three major workshops devoted to understanding chondrule origins. The workshop in Kauai in 2004 differed from the first two as a broader range of questions was addressed including the role of high-temperature processing in the formation of all the ingredients in chondrites, the detailed timescales for the formation of chondrules and refractory inclusions, constraints from grains that formed around other stars, and relationships between chondrites, comets and chondritic IDPs. The Kauai workshop also benefited considerably from vastly improved astronomical and astrophysical insights into star formation. Here, we review the major constraints from chondrites on the high temperature processing of silicates in the solar nebula prior to their accretion into planetesimals. We review the mineralogy of the refractory inclusions, chondrules, and matrices, the chondrite groups, the chemical and isotopic compositions of chondrites and their components, and attempt to develop a plausible scenario for their formation. We con- clude that thermally processed dust was ubiquitous in the solar nebula and was a major ingredient in asteroids and comets. Chondritic Meteorites and Their Components 17

2. Chondritic Components

The nature and abundances of the three major chondritic ingredients—refractory inclusions, chondrules and matrix material—vary widely among the different chondrite groups. Table 1 lists the proportions of these ingredients in each group, the mean

Table 1. Concentrations of chondritic components in the chondrite groups.

Group Type Refract. Chondr. Chondr. Fe,Ni Matrix Fall freq. Refract. Examples incls. (vol.%)+ mean metal (vol.%)§ (%)^ lith./Mg (vol. %) diam. (vol.%) rel. CI# (mm) Carbonceous CI 1 <0.01 <5 - <0.01 95 0.5 1.00 Orgueil CM 1-2 5 20 0.3 0.1 70 1.6 1.15 Murchison CO 3 13 40 0.15 1-5 30 0.5 1.13 Ornans CV 2-3 10 45 1.0 0-5 40 0.6 1.35 Vigarano, Allende CK 3-6 4 15 0.8 <0.01 75 0.2 1.21 Karoonda CR 1-2 0.5 50-60 0.7 5-8 30-50 0.3 1.03 Renazzo CH 3 0.1 ~70 0.05 20 5 0 1.00 ALH 85085

CBa 3 <0.1 40 ~5 60 <5 0 1.0 Bencubbin CBb 3 <0.1 30 ~0.5 70 <5 0 1.4 QUE 94411 Ordinary H 3-6 0.01-0.2 60-80 0.3 8 10-15 34.4 0.93 Dhajala L 3-6 <0.1 60-80 0.5 3 10-15 38.1 0.94 Khohar LL 3-6 <0.1 60-80 0.6 1.5 10-15 7.8 0.90 Semarkona Enstatite EH 3-6 <0.1 60-80 0.2 8 <0.1-10 0.9 0.87 Qingzhen EL 3-6 <0.1 60-80 0.6 15 <0.1-10 0.8 0.83 Hvittis Other K 3 <0.1 20-30 0.6 6-9 70 0.1 0.9 Kakangari R 3-6 <0.1 >40 0.4 <0.1 35 0.1 0.95 Rumuruti Sources of data: Scott & Krot (2003) and references listed therein. ALH = Allan Hills; QUE = Queen Alexandra Range. # Mean ratio of refractory lithophiles relative to Mg, normalized to CI chondrites. + Includes chondrule fragments and silicates inferred to be fragments of chondrites. ^ Fall frequencies based on 918 falls of differentiated meteorites and classified chondrites (Grady 2000). § Includes matrix-rich rock fragments, which account for all the matrix in CH and CB chondrites. chondrule sizes, and the abundances of Fe,Ni metal grains, which are located within chondrules or probably formed with them. All but two of the 15 chondrite groups fall into the ordinary, carbonaceous, or enstatite classes (Table 1). A critical step in understanding the origin of chondritic components was to identify the effects of metamorphism, aqueous alteration, shock, and brecciation in asteroids and to establish which chondrites could have been derived from a common source (Wood 1962; Zolensky & McSween 1988; Scott et al. 1989; Scott 2002). Van Schmus & Wood (1967) inferred that most chondrites had been heated in asteroids

18 Scott and Krot and devised various mineralogical and chemical criteria to divide the chondrite groups into six metamorphic (or petrologic) types (Keil, this volume). These criteria and the development of the electron microprobe quickly led to the establishment of a small group of type 3 chondrites as the least equilibrated or metamorphosed of the ordinary chondrites and the precursors to the strongly metamorphosed types 4-6 (Dodd et al. 1967). In part because CI chondrites are closest in composition to solar composition, they were classed as type 1 and thought to be the primary material from which types 2-6 were derived. However, types 1 and 2 are now considered to be products of aqueous alteration on asteroids, and not pristine nebular aggregates. Type 3 chondrites in the ordinary and CO groups were further subdivided into 10 subtypes: 3.0 (least metamorphosed) to 3.9 (Sears et al. 1980; Scott & Jones 1990). Most of the mineralogical and chemical characteristics of pristine chondrules were then established by studying chondrules in the rare type 3.0 chondrites (see Brearley & Jones 1998). However, CV and enstatite chondrites have generally re- sisted subdivision because of insufficient samples and confusion over effects due to shock, hydrothermal alteration, and brecciation. The Allende meteorite, which for many years was virtually the sole source of CAIs and the most studied chondrite, was once thought to be the most pristine chondrite. But it now appears to be type >3.6 on the basis of the degree of disorder in the graphitizable carbon (Bonal, Quirico, & Bourot-Denise 2004). In many groups, especially the ordinary and R groups, there are chondrites resembling type 3 chondrites that are actually breccias composed of fragments of material with diverse metamorphic histories in a chondrule-rich matrix. Thus, some chondrites may have formed millions or even billions of years after their ingredients were made. Below, we review the mineralogical and chemical properties of the three major chondritic components based largely on studies of the few type 2 and 3.0 chondrites that show the least metamorphism, alteration, and brecciation [Allan Hills (ALH) A77307 and Yamato 81020 (CO3.0), Semarkona (LL3.0), Acfer 094, Adelaide, and Lewis Cliff (LEW) 85332 (ungrouped carbonaceous chondrites), and the CR2 and CH3 chondrites]. The chondrite groups are discussed in section 3, and the isotopic properties of the chondritic components in section 4. Detailed accounts of the mineralogy of all chondritic components are given by Brearley & Jones (1998).

2.1. Refractory Inclusions The numerous varieties of refractory inclusions are divided into two basic types: Ca- Al-rich inclusions (CAIs), which are composed of refractory Ca-Al-Ti minerals, and amoeboid olivine aggregates (AOAs), which are composed of forsterite and Ca-Al-Ti mineral aggregates (MacPherson 2003; MacPherson et al., this volume). The two types of refractory inclusions are closely related, although CAIs formed at higher nebular temperatures. Each of the 15 chondrite groups contains roughly equal proportions of CAIs and amoeboid olivine aggregates, but the concentration of refractory inclusions varies enormously—between 0.01 and 10 vol.% (Table 1). The primary tool for understanding the mineralogy of the refractory inclusions (and many other chondritic ingredients) is the equilibrium mineral stability diagram for the solar nebula (Fig. 1), which is calculated from the thermodynamic properties of minerals and gaseous species (Grossman & Larimer 1974; Yoneda & Grossman Chondritic Meteorites and Their Components 19

1995; Petaev & Wood, this volume). Liquids only become stable if the total pressure is increased 10-100 × above the canonical pressure of 10-3 to 10-4 bar, or the dust-gas ratio is increased by a similar factor above the solar value (Wood 1963). Such dia- grams were first calculated on the assumption that chondritic components condensed from a hot and homogeneous nebula. However, the discovery of stardust and isotopic anomalies in chondritic components and the mineralogy and composition of the

Figure 1. Equilibrium diagram for a solar nebula at 10-3 bar showing mineral stability above 900 K (Davis & Richter 2003). At 900 K, half the atoms (0.55) in a CI chondrite are condensed; S and other volatile elements are in the gas. Minerals stable above 1400 K are present in refractory inclusions; minerals stable below 1400 K are found in chondrules. In a cooling nebula, only three minerals condense entirely from the gas: corundum (Al2O3), forsterite (Mg2SiO4), and Fe,Ni metal: the remainder form by reaction between solids and gas. Liquids are only stable if the total pressure or the dust/gas ratio is increased 10-100×. components show that thermal processing of the chondritic components was dominated by the effects of brief, localized heating in the nebula, rather than simple mono- tonic cooling of the nebula. Thus individual components only approached equilibrium over rather limited temperature intervals before being effectively isolated as a result of grain growth and aggregation, or loss of gas (Petaev & Wood 1998). All the minerals stable above 1400 K in Figure 1 are found in refractory inclusions, but the minerals in each inclusion typically reflect only a relatively narrow temperature interval. Some rare CAIs contain only corundum, hibonite, and

20 Scott and Krot perovskite, which are stable above ~1650 K (Fig. 2). More common are the CAIs containing spinel, melilite (solid solution of åkermanite, Ca2MgSi2O7, and gehlenite, Ca2Al2SiO7), Ca-Ti–pyroxene, and anorthite, which are stable at ~1400-1500 K. The amoeboid olivine aggregates contain nuggets of anorthite, Ca-pyroxene, and spinel, embedded in forsteritic olivine with Mg/(Fe+Mg) >0.99, which are all stable at 1350- 1450 K (Krot et al. 2004a). A few minerals that are found in CAIs, but not in Figure 1, such as grossite, CaAl4O7, may be stable under nebula conditions given the uncer- tainty in the thermodynamic data for these refractory minerals. The Ti4+/Ti3+ ratios in pyroxenes suggest that most CAIs formed in a reducing gas with O/H comparable to that of a nebula with solar composition; a few experienced more oxidizing conditions (Beckett et al. 1988).

Figure 2. Scanning electron microscope image using back-scattered electrons showing refractory inclusions in chondrites. (a) Ca-Al-rich inclusion in a CM chondrite made of corundum (cor), hibonite (hib), and perovskite (pv), which condensed in the solar nebula above 1650 K (Simon et al. 2002). The inferred sequence of mineral condensation matches the equilibrium diagram in Figure 1. Voids are black. (b) Core (lower right) and rim (upper left) of an amoeboid olivine aggregate in the ungrouped carbonaceous chondrite Acfer 094 composed of diopside (di, Ca-pyroxene) and anorthite (an) intergrowths enclosed by forsterite (fo) with occasional grains of metallic Fe,Ni. Forsterite is replaced by low-Ca pyroxene (px). Chondritic Meteorites and Their Components 21

Most refractory inclusions are irregularly shaped and fine-grained, and probably formed from minerals that condensed in the solar nebula (e.g., Fig. 2). The best studied, called “fluffy” Type A inclusions, are rich in melilite and have characteristic abundance patterns of the rare-earth elements (REE) indicative of condensation. The depletions of ultra-refractory, rare-earth elements in these Type A inclusions (called group II REE patterns) provide strong evidence that they condensed from the nebula after loss of an ultra-refractory component (Boynton 1975). Despite their name, the fluffy Type A inclusions are not loose clusters of grains, but dense aggregates of in- tergrown crystals indicating that they recrystallized at high temperatures (MacPher- son 2003). The other major variety of fine-grained refractory inclusions that appear to be products of nebular condensation are tamoeboid olivine aggregates. These range in texture from very porous aggregates of forsterite and CAI minerals to coarser, recrystallized, non-porous aggregates (Krot et al. 2004a). A few amoeboid olivine aggregates have group II rare-earth element patterns, but most have relatively uniformly high levels of refractory abundances (3-10 × CI levels), showing that they formed in a nebular region where ultrarefractory or even moderately refractory materials were not previously removed. By contrast, most coarse-grained CAIs have spheroidal shapes and igneous textures indicating crystallization from melts. They were initially called Ca-Al-rich chondrules (Christophe Michel-Lévy 1968), but the term, chondrule, is now reserved for objects that are rich in olivine and Ca-poor pyroxene. There are three major mineralogical types of igneous CAIs: compact Type A, which are rich in melilite; Type B, which have abundant Ca-Ti–pyroxene; and Type C, which are rich in anorthite (see MacPherson et al. 1988). The best studied CAIs are the cm-sized, Type B inclusions in CV chondrites, which have isotopic mass fraction patterns indicative of evaporation (Davis & Richter 2003). They probably formed from CAI material that condensed but was subsequently melted at ~1700 K, and cooled at ~0.5-50 K hr-1. Most igneously formed CAIs show rather uniform ~20-fold enrichments of rare-earth elements over CI chondrite concentrations, indicating that they represent the most refractory 5% by mass of condensable material (Boynton 1975; see MacPherson et al., this volume). Two kinds of rims composed entirely of CAI minerals are found on some CAIs (Fig. 3a). Melted and unmelted CAIs generally have multi-layered rims due to evaporation and condensation during thermal recycling, which are called Wark-Lovering rims. CAIs in CV chondrites and rarely in other groups may also have accretionary rims that form outside Wark-Lovering rims and are composed of forsterite-rich olivine with minor refractory minerals that closely match the mineralogy of the amoeboid olivine aggregates. The presence of these accretionary rims around CV CAIs may be related to their large size. The 15 chondrite groups have CAIs with diverse sizes and characteristic mineralogies: e.g., CM chondrites have spinel- and hibonite-rich, melilite-poor inclusions, CV chondrites have uniquely large CAIs rich in melilite and Ca-Ti- pyroxene, CH chondrites are unique in containing abundant, small grossite-rich CAIs (see Scott & Krot 2003; MacPherson, this volume). Interestingly, amoeboid olivine aggregates in different groups appear to be identical, except for changes caused by asteroidal alteration and metamorphism (Krot et al. 2004a). Although amoeboid olivine aggregates and Al-rich chondrules are chemically intermediate in composition between CAIs and ferromagnesian chondrules, their isotopic and chemical

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CAIs and ferromagnesian chondrules, their isotopic and chemical compositions suggest that chondrules and refractory inclusions formed in distinctly different regions (Russell et al., this volume). Ca-Al-rich inclusions and amoeboid olivine aggregates have unique O-isotopic compositions and formed at diverse temperatures above 1350 K in a nebula of solar composition at pressures of ~10-4 bar. They were then cooled quickly so that lower- temperature gas-solid reactions were kinetically inhibited.

Figure 3. Sketch showing diverse varieties of rims on Ca-Al–rich inclusions (a) and chondules (b). The outermost fine-grained rims of matrix material, which are nearly ubiquitous on chondrules and CAIs, were acquired after these objects cooled—probably as the chondritic components accreted together. The rims inside the matrix rims, which are less common (see text), contain minerals like those present in the enclosed CAI or chondrule and were acquired in the final stages of CAI or chondrule formation.

2.2. Chondrules Chondrules are round or irregularly shaped particles that were wholly or partly molten before they accreted (Fig. 4). They are the most abundant component in all chondrites except the CI, CM, and CK groups, which are dominated by matrix, and CB chondrites, which are dominated by metallic Fe,Ni (Table 1). Mean chondrule sizes in the groups vary between 0.05 and 1 mm, except for the CBa subgroup in which chondrules are typically 5 mm in size (Table 1). In addition to olivine and low-Ca pyroxene, which are the major minerals, chondrules commonly contain metallic Fe,Ni and troilite, FeS. Metallic Fe,Ni is stable above 1300 K in the canonical solar nebula with forsterite and enstatite (Fig. 1); troilite would form from Fe,Ni below 700 K (Grossman & Larimer 1974) or at higher temperatures in solid-gas fractionated systems (Wood & Hashimoto 1993). Some chondrules, called type I, which account for most chondrules in carbonaceous chondrites, have olivine and low-Ca pyroxene with Mg/(Mg+Fe) ratios >0.95 (>0.9 in ordinary chondrites). Other chondrules, Chondritic Meteorites and Their Components 23 called type II, have silicates with Mg/(Mg+Fe) ratios in the range 0.6-0.9 (Jones et al., this volume). Nearly all type I and II chondrules have porphyritic textures with crystals up to hundreds of micrometers in size called phenocrysts embedded in finer-grained material called mesostasis (Fig. 5). Experiments show that many of the features of these chondrules can be duplicated by partly or largely melting fine-grained material at 1700-2100 K and cooling at 10-1000 K hr-1, slowly enough to allow the phenocrysts to crystallize from the melt (Hewins et al., this volume). The non-porphyritic chondrules were heated more extensively, so that few seed crystals remained, and tend to have finer grained textures, which may be barred olivine, radial pyroxene, or cryptocrystalline (Fig. 4). They cooled at rates up to 1000 K hr-1 when they crystallized. Porphyritic chondrules are typically much more irregular in shape than non- porphyritic chondrules, which were melted sufficiently for surface tension to make them round. (Note that the porphyritic chondrule in Fig. 5a is atypical, illustrating the widespread tendency for authors to choose aesthetically pleasing rather than typical examples!) The most irregularly shaped, porphyritic chondrules, which are very abundant in CO chondrites, for example, may never have contained more melt than crystals (Hewins et al. 1997; Rubin & Wasson 2005). Since porphyritic chondrules dominate in all groups except CH and CB and account for >95% of chondrules in all other carbonaceous chondrites, we focus first on their origin.

Figure 4. Magnesium concentration map showing chondrules with diverse textures in the Tieschitz ordinary chondrite derived from Mg Kα X-rays generated in an electron microprobe. Most chondrules have porphyritic textures (large crystals in a fine-grained matrix) and are rich in olivine (PO) or pyroxene (PP), or less commonly both (POP). The Mg- rich ones are called type I (marked POI); type II chondrules have lower Mg/(Mg+Fe) ratios (POII). Chondrules without porphyritic textures include those with cryptocrystalline (C), radial pyroxene (RP) and barred-olivine (BO) textures.

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Figure 5. Back-scattered electron images of type I porphyritic chondrule in the Tieschitz chondrite containing numerous Mg-rich olivine crystals (ol) in a light-colored mesostasis (mes), which formed from residual silicate melt after the olivines crystallized. The white grains of metallic Fe,Ni concentrated around the rim were once molten also: some pro- trude from the chondrule rim, other grains may have escaped from the chondrule. The large speckled grains up to 300 µm in size are foreign Fe-rich olivine crystals from a type II chondrule, which probably collided with this type I chondrule, or its precursor aggregate, and are now dusted with metallic Fe. (b) Silicon-concentration map of a rounded, type I, porphyritic olivine chondrule in the CR chondrite, El Djouf 001, with two outer, coarse-grained, igneous rims: an inner pyroxene-rich rim (px), and an outer silica-rich rim (white). Black grains in the rims are Si-free, metallic Fe,Ni. The rims formed from partly melted material that accreted onto the chondrule (or solids that were partly melted after they accreted) and became progressively richer in Si and other volatile elements due to gas-melt interaction. Chondritic Meteorites and Their Components 25

2.2.1. Origin of Porphyritic Chondrules Porphyritic chondrules can be made in the laboratory by partly melting fine-grained dust of the appropriate composition and cooling at the appropriate rate to allow phenocrysts to form. However, just melting so-called “dust-balls” or “dusty clumps” without some evaporation and condensation cannot explain several observations. The most important is that the proportion of olivine to olivine plus pyroxene in porphyritic chondrules ranges from <1% to >99%. Since physical processes in the solar nebula cannot plausibly separate olivine and pyroxene dust and solid-state conversion of olivine to pyroxene is kinetically inhibited (Imae et al. 1993), it is rather unlikely that chondrules formed simply by melting aggregates of numerous dust grains. An- other problem is that dustballs or compacted dustballs have not been identified in chondrites, except for matrix lumps that are not compositionally appropriate precursors. All of the possible fine-grained precursors that have been identified in chondrules (e.g., aggregational chondrules) have igneous textures. Finally, it is not clear that millions of dry, cold, silicate dust grains could stick together long enough for an aggregate to reach the mass of a chondrule (Wood 1996). Many features of chondrules suggest that collisions between molten, partly molten, or hot, solid grains were an integral part of chondrule formation that allowed growth of mm-sized particles from dust. Half the chondrules in CV chondrites and 10% of those in ordinary chondrites have distinct rims with igneous textures and grain sizes of ~5-10 µm (Rubin 2000; Fig. 3b). These resulted from collisions of partly molten chondrules with dust or remelting of such accretionary rims (Krot & Wasson 1995; Rubin 2000; Hewins et al. and Jones et al., this volume). Other evidence for collisions during chondrule formation and multiple heating events comes from compound chondrules, which consist of two chondrules fused together (Wasson et al. 1995), and fragments of chondrules within chondrules (Fig. 6). Thus growth

Figure 6. Sketches showing how chondrules formed by repetitive processes involving heating and melting of dust plus collisions between solid particles and melted or partly melted objects. Such processes can account for the existence of fragments of chondrules and CAIs within chondrules and igneous rims and adhering chondrules on the exterior. Type I chondrule compositions were also modified by incorporation of refractory forsterites (Pack et al. 2004), gas-liquid equilibration (Krot et al. 2004b), loss of metallic Fe,Ni and possibly evaporation (Jones et al., this volume).

26 Scott and Krot

and consolidation of aggregates was an important part of chondrule formation: silicate melt allowed colliding particles to stick together, and high-temperatures allowed loosely consolidated grain clusters to weld together and coarsen. Type I porphyritic chondrules form a sequence from olivine-rich varieties that are rich in Ca and other refractory lithophiles (type IA) to pyroxene-rich chondrules that are volatile-rich and more ferroan (type IB) (see Fig. 1 of Hewins et al., this volume). An explanation for the diverse proportions of olivine and pyroxene (and the existence of low-Ca–pyroxene rims on olivine-rich type I chondrules) has been provided by Tissandier, Libourel, & Robert (2002), who allowed gaseous SiO to react with experimental melts, and Krot et al. (2004b), who observed Si-rich, igneous rims on type I chondrules in CR chondrites (Fig. 5b). In these chondrules, the composition of the mesostasis is chemically zoned with the concentrations of Si, Na, K, and Mn increasing from the olivine-rich core to the SiO2-rich rim, while Ca, Mg, and Al de- crease. The former elements are more volatile than the latter in the nebula (Fig. 1), suggesting that these trends probably result from condensation in the nebula. The pyroxene rims on type I chondrules in CR chondrites appear to result from formation of melts with high Mg/Si ratios at high temperatures followed by condensation of Si at lower temperatures, or by condensation into solids followed by remelting. The nebula gas may have been enriched in volatiles like Si as a result of frac- tional condensation—the continued removal of small fractions of condensate to prevent continued equilibration with the gas as a result, for example, of decreasing diffusion rates on cooling. Alternatively, the gas may have been enriched in volatiles by evaporation during chondrule formation. Clearly, the Si-rich rims formed by accretion of partly melted material, or accretion of solids followed by remelting, or by both mechanisms. Although many details remain to be understood, type I porphyritic chondrules are now inferred to be products of many solar nebula processes: accretion of melted, partly melted, and solid objects, melting, and evaporation-condensation processes. Refractory-rich forsterites and refractory-rich type I chondrules formed initially (Pack et al. 2004); pyroxene-rich type I chondrules formed subsequently from olivine-rich precursor chondrules and related olivine-rich aggregates. Studies of the Fe,Ni grains in type I chondrules complement the record in the chondrule silicates providing additional evidence for vaporization and condensation (Campbell et al., this volume). However, there are important differences between chondrite groups. In CR chondrites, the metal grains completely lack troilite, FeS, (Krot et al. 2002a), whereas those in ordinary chondrites appear to have formed under somewhat different conditions as some contained sulfur when molten (Rubin et al. 1999; Campbell et al., this volume). In addition, type I chondrules in Semarkona (LL3.0), appear to have been chemically modified after crystallization, probably during asteroidal alteration, and lack evidence for silica addition (Alexander & Grossman 2005). Type II porphyritic chondrules, unlike type Is, do not show such clear evidence for loss or gain of volatiles during thermal processing. Nevertheless, they were probably partly molten for an hour or more, as they cooled at ~10-1000 K hr-1 when crystallizing (Hewins et al., this volume). Such periods probably require dust- enriched nebular environments to stabilize FeO-rich melts and prevent loss of alka- lies and S (Ebel & Grossman 2000; Alexander 2004), which can be lost in minutes from synthetic analogs (Yu et al. 2003). However, Wasson & Rubin (2003) question Chondritic Meteorites and Their Components 27 these conclusions arguing that volatiles were retained because the chondrules crystallized as a result of numerous, much briefer melting events with cooling rates that were orders of magnitude higher than previously inferred.

2.2.2. Formation of Other Chondrules Some non-porphyritic chondrules, e.g., Mg-rich barred olivine chondrules, probably formed by similar processes to those described above except that they experienced more intense heating. However, non-porphyritic chondrules in CB and CH chondrites lack igneous rims and have unique compositions and skeletal olivine and cryptocrystalline textures (Krot et al. 2002a,b). These chondrules also lack relict grains and fine-grained matrix rims, and clearly formed in a very different way from the porphyritic chondrules. Their compositions and especially their 100-fold variations in their concentrations of refractory elements like Ca, Al, and Ti (e.g., Al/Mg ratios that are 0.02-3 × solar) favor an origin as condensates, and their textures suggest that they formed by condensation of melts. Metallic Fe,Ni grains associated with these chondrules appear to have condensed as solids above 1200 K following dust-gas enrichment and vaporization (Campbell et al., this volume). One variety of chondrules appears to have formed as a result of mixing of chondrule melts or precursor aggregates with refractory inclusions (Krot & Keil 2002; Krot, Hutcheon, & Keil 2002c). The Al-rich chondrules, which account for ~1% of chondrules in carbonaceous chondrites, commonly contain relict CAIs composed of anorthite, Ca-rich pyroxene and spinel and probably formed by accretion and melting of ferromagnesian silicate aggregates and fragments of CAIs (Fig. 6). The role of asteroid collisions in forming normal chondrules is controversial (Sanders & Taylor, this volume). However, a small fraction of chondrules in type 4-6 ordinary chondrites (<1%) are rich in chromite and plagioclase and appear to have formed by selective impact melting on the surface of an asteroid (Krot, Ivanova, & Wasson 1993; Rubin 2003). Chondrules of this composition are not found in type 3 chondrites, but impact melts formed largely from chromite and plagioclase are found in shocked chondrites of higher petrologic type. Most chondrules probably formed in the solar nebula as a result of repeated melting of solids, collisions between solid and melted or partly melted particles that allowed particle sizes to increase, and vapor condensation into melts. Thus, understanding chondrule formation is not just a question of identifying the heat source that melted dust, we must also understand the origin of the dust, how dust was enriched in the nebula so that liquids could exist for hours, and how repeated thermal processing allowed melts and solids to accrete. In addition, we need to know more about the relationship between chondrules, refractory inclusions and the fine-grained, matrix materials that coat chondrules and their igneous-textured rims.

2.3. Matrix Materials The matrix in chondrites is the volatile-rich material with grain sizes of 10 nm to ~5 µm that coats the chondrules and refractory inclusions, and may also fill the interstices between them (Scott & Krot 2003, 2005). It is a mixture of materials that probably formed in diverse locations in the solar system and contains small amounts of presolar materials, viz., interstellar organic material and circumstellar inorganic

28 Scott and Krot grains. Despite the heterogeneous nature of matrices, they are closer to solar composition than the bulk chondrites (see Huss et al., this volume). Studies of solar noble gases and irradiation tracks in regolith breccias show that the matrix rims on chondrules were not acquired in asteroidal regoliths (Nakamura et al. 1999; Metzler et al. 2004). The rims were probably formed when the chondrules and other ingredients accreted in a turbulent, dusty nebula (Cuzzi 2004). Because of its fine-grained nature, matrix was readily modified by heating and aqueous alteration on asteroids. Deci- phering how and where the matrix minerals formed—in asteroids or the nebula—has been highly controversial (Nuth et al. and Huss et al., this volume). Many matrices of C and O chondrites contain hydrated minerals that were once thought to have condensed at low temperatures from the solar nebula (e.g., Wright 2004). However, most authors now infer that ice accreted with matrix-rich chondrites and that hydrated minerals formed predominantly in asteroids by aqueous alteration (see also Ciesla et al. 2003). The matrices of many type 3 chondrites contain Fe-rich olivine that was once thought to have condensed at high temperatures in the solar nebula (e.g., Scott et al. 1988; Wright 2004). However, Fe-rich olivine, which is especially abundant in the Allende CV chondrite (Fig. 7a), has been shown to be a product of aqueous alteration in asteroids (Krot, Petaev, & Bland 2004c; Nuth et al., this volume). To understand what minerals were present in the fine-grained matrix materials that rimmed chondrules and refractory inclusions before they accreted into asteroids, we exclude matrices in those chondrites with chondrules and refractory inclusions containing hydrated minerals and other minerals that clearly formed after the components had accreted into asteroids as a result of alteration and metamorphism. This eliminates almost all of the ~4000 chondrites! Four remaining chondrites have matrices largely lacking hydrated minerals and Fe-rich olivine: these are Acfer 094 and ALHA77307, which are the least metamorphosed of the type 3.0 chondrites (Grossman & Brearley 2005), Adelaide, and Kakangari. Their matrices are largely composed of crystalline, Mg-rich silicates and amorphous, Fe-rich silicate with additional grains of metallic Fe,Ni, sulfides, refractory oxides, and presolar grains (Brear- ley 1989, 1993; Greshake 1997; see Scott & Krot 2005). We focus here on the silicates and oxides in matrices, but note that the organic materials in chondrite matrices and IDPs provide additional constraints on the thermal processing of dust in the nebula (Alexander, this volume).

2.3.1. Crystalline Mg-rich Silicates The crystalline silicates in the most pristine chondrites are grains of olivine and low- Ca pyroxene with Fe/(Fe+Mg) ratios of <0.05, which are mostly 100-1000 nm in size (Fig. 7b). They tend to occur as isolated crystals in the amorphous matrix, but clusters and aggregates are also present. Olivine tends to be more abundant than pyroxene, except in Kakangari. Both silicates may have high Mn concentrations (0.6-2 wt.% MnO), which are also observed in the Mg-rich olivine in amoeboid olivine aggregates, suggesting that the matrix silicates are also condensates (see Petaev & Wood, this volume). Many low-Ca pyroxenes have very fine-scale intergrowths of orthorhombic and monoclinic structures indicative of cooling at ≈1000 K hr-1 from 1300 K. The general similarity of this rate to the cooling rate of chondrules suggests Chondritic Meteorites and Their Components 29 that the crystalline matrix silicates may have formed during brief heating events that were related to those that produced chondrules.

Figure 7. (a) Back-scattered electron image of matrix in the Allende-like CV3 chondrite, ALHA81258, showing 10-20 µm laths of Fe-rich olivine (ol) and irregularly shaped grains of Ca-rich pyroxene (cpx). (b) Transmission electron microscope image of the matrix in ALHA77307 (CO3.0) showing vastly smaller grains of Mg-rich olivine and pyroxene (px), Fe,Ni, and FeS embedded in amorphous Fe-Mg-Si-O material (after Brear- ley 1993). The coarse-grained Allende-like matrix minerals in (a) probably formed from fine-grained material during asteroidal, hydrothermal alteration, whereas the silicates in (b) are probably nebula condensates.

2.3.2. Amorphous Silicate The amorphous silicate in chondrite matrices is composed largely of Si, Fe, Mg, and O with lesser amounts of the other elements present in chondrules (Al, Ni, Ca, S, Na, Cr, etc) and is Al-rich and Mg-poor relative to bulk matrix. Amorphous silicate is a major component of the matrices of three of the pristine chondrites (~20-80 vol.%),

30 Scott and Krot but is absent in Kakangari. In ALHA77307, the amorphous material forms 1-5 µm units, which are chemically very heterogenous and may contain micrometer-sized grains of Fe,Ni metal and sulfide. Since the bulk matrix composition is near solar, it is likely that the amorphous silicates, like the crystalline silicates, formed at high temperatures in the solar nebula. The matrix in Kakangari contains feldspar crystals (albite, NaAlSi3O8, and anorthite, CaAl2Si2O8) in place of amorphous material.

2.3.3. Presolar Grains The most remarkable ingredients in chondrites, which truly validate them as aggregates of materials that existed in the solar nebula, are the presolar grains that formed around diverse evolved stars (Zinner 2003; Clayton & Nittler 2004; Lodders & Amari 2005; Stroud, this volume). The first to be discovered were grains of graphite, diamond, and SiC, which are 0.1-3 µm in size, except for the diamonds, which are 2 nm in size. These grains are carriers of isotopically anomalous noble gas components and were found by acid dissolution of their meteoritic hosts. In situ detection of presolar grains by Alexander et al. (1990) using secondary ion mass spectrometry confirmed that presolar grains are located in the matrix. The list of presolar grains discovered using this technique now includes refractory oxides such as corundum (Al2O3), spinel (Mg2Al2O4), hibonite (CaAl12O19), metallic Fe,Ni, Si3N4, TiO2, and various carbides (Zinner 2003), which are all present at ppm levels or less. Presolar crystals of olivine and pyroxene are more abundant (~10-50 ppm levels). Amor- phous silicate grains with non-solar, isotopic compositions appear to be less common than crystalline silicates (Nguyen & Zinner 2004; Nagashima et al. 2004; Mostefaoui & Hoppe 2004). Presolar grains were discovered in CM chondrites, but acid dissolution of ordinary and enstatite chondrites led to a better appreciation that carbonaceous chondrites are not intrinsically more primitive than other classes. All chondrite classes appear to have sampled the same reservoir of presolar grains: differences in the nature and concentration of different presolar minerals probably reflect thermal processing in the nebula and asteroidal modification of the chondrites (Alexander 1993; Huss & Lewis 1995). Thus, even though the matrices of enstatite chondrites have scarcely been characterized, we infer that they contain ~2-15% of matrix that is genetically related to the matrices in other chondrites. Studies of presolar grains have also validated the inference from mineralogical studies that Acfer 094, ALHA77307 and Adelaide are probably the most pristine chondrites as they have the highest concentrations of presolar silicates (e.g., Kobayashi et al. 2005). (Chondritic, porous IDPs have higher concentrations.) The isotopic compositions of the presolar grains typically differ from terrestrial values by factors of 2-100. These isotopic effects are very much larger than the isotopic anomalies observed in CAIs, which are typically restricted to a few specific isotopes or inclusions, e.g., 10% anomalies in 50Ti in some CM group hibonites (Ire- land & Fegley 2000). Thus, there is a clear distinction, at least for crystalline materials, between the CAIs, chondrules, and matrix grains that formed in the solar system, and presolar grains (see Nittler, this volume). For amorphous silicates in IDPs, the distinction is less clearcut; see chapters by Dai & Bradley and Keller & Messenger (this volume). Chondritic Meteorites and Their Components 31

2.3.4. Refractory Solar Nebula Grains An important by-product of the search for presolar grains in the acid residues of chondrites was the discovery that chondrite matrices also contain ~100 ppm of micrometer-sized grains of the refractory oxides, corundum, spinel and hibonite that formed in the solar nebula (Choi et al. 1998, 1999; Zinner 2003). Some of the refractory grains with solar-system O- and Mg-isotopic compositions have morphologies like those of the circumstellar grains and probably condensed in the solar nebula, rather than in stellar outflows. Other micrometer-sized minerals including forsterite have been found in chondrite matrices with the characteristic 16O-rich signature of CAIs (Kunihiro et al. 2005). Thus the chondrite matrices truly contain the most ex- traordinary diversity of components from the most and the least thermally processed portions of the solar nebula.

2.4. Rock Fragments Chondrites are not simply aggregates of particles that formed in the nebula and traces of presolar material—they are cosmic garbage cans! Many contain so-called dark clasts, which are rock fragments ~100-1000 µm in size typically composed of aqueously altered material. Some rock fragments may be derived from near-surface regions of the parent asteroid or from hypervelocity asteroidal projectiles, as many chondrites are breccias of materials that formed in different locations long after chondrules formed (Bischoff et al. 2005). However, many dark inclusions in C chondrites are coated with matrix material like chondrules and have sizes that correlate with the mean size of chondrules in their host meteorites. This suggests that fragments of altered asteroids accreted in the nebula with chondrules to form planetesimals (Scott & Krot 2003). One remarkable chondrite, Kaidun, is composed entirely of millimeter-sized fragments from seven different groups of enstatite, ordinary, and carbonaceous chondrites and new kinds of C1 and C2 chondrites plus a few clasts of differentiated asteroids (Zolensky & Ivanov 2003; Mikouchi et al. 2005). Kaidun and its parent asteroid may have accreted when the disk lacked chondrules and matrix grains but contained abundant rock fragments (Scott 2002). The fragments appear to have accreted before the gas was dissipated in a region free of planetesimals by the same process that concentrated chondrules, viz., turbulent accretion (Cuzzi et al. 2001). The lack of rims composed of impact-generated dust may reflect the absence of organic glue. This explanation for the Kaidun breccia implies that a few planetesimals accreted after the inner part of the primordial disk evolved into a debris disk, but before all the gas had been lost. Chondrites also contain rare rock fragments that have been interpreted as pieces of igneous rock from melted asteroids (Kennedy et al. 1992; Hutchison, Bridges, & Gilmour, this volume). Given that the parent asteroids of the chondrites were probably tens of km or more in size, they probably accreted after 26Al had decayed sufficiently to prevent melting (see section 3.3). The low abundance of igneous rock fragments in chondrites may indicate that differentiated asteroids were largely formed inside 2 AU and were scattered into the belt after the parent asteroids of chondrites had accreted. However, Sanders (1996) suggests that asteroidal products may be ubiquitous in chondrites as he infers that most chondrules were derived from

32 Scott and Krot collisions between molten asteroids (see Sanders & Taylor, this volume). The origins of chondrules are discussed further in sections 6 and 7.

2.5. Alteration and Mixing of Chondritic Components No chondrites are entirely free from the effects of metamorphism, alteration, shock and other impact effects, and many of the best-studied type 2 and 3 chondrites, some of which are listed in Table 1, were severely affected. As a result, chemical, isotopic and mineralogical data for chondritic components need to be carefully assessed (see e.g., Jones et al., this volume). To illustrate the fundamental importance of understanding geological processes on asteroids in elucidating the nebula properties of the components in chondrites, we mention some of the minerals that were once thought to be nebular products. Miner- als such as nepheline (NaAlSiO4), sodalite (NaAlSiO4·NaCl), Fe-rich olivine, and fayalite (Fe2SiO4), which are all present in CAIs in Allende and some other CV3 chondrites and are not stable in a solar nebula at high temperatures (Fig. 1), are now recognized to be secondary, asteroidal products. The identification of chondrites like ALHA77307 and Acfer 094 as the best guides to the nebular properties of components in carbonaceous chondrites required an understanding of the origin of the abundant secondary minerals in Allende and the mechanism whereby Allende could form from chondritic components like those in ALHA77307. The timely fall in 1969 of the large Allende meteorite provided valuable clues that improved our understanding of the early solar system, but progress would have been more rapid if several tons of a less altered CV chondrite had fallen instead (MacPherson 2003)! Table 1 would contain far fewer groups but for the wealth of new meteorites recovered in the past 30 years from Antarctica and various deserts around the world (Sears 2004). Nearly every one of the most pristine chondrites was recovered by people hunting for meteorites. Continued recoveries are important for finding new chondrite groups and identifying unrecognized effects of asteroidal processing. One of the most surprising features of chondrites is their wide variety of components. One might expect that accretion processes and subsequent impacts over 4 Gyr in the asteroid belt would have eliminated heterogeneities in nebular solids, except for those due to ambient temperature. Although there was extensive mixing in the nebula before the chondritic ingredients accreted at low velocities, and numerous asteroidal impacts at high velocity after accretion, the asteroids have preserved an amazing variety of distinct kinds of chondrites. Below, we discuss the properties of the chondrite groups and the nature and location of their parent bodies.

3. Properties of Chondrite Groups The 15 chondrite groups are defined primarily by their chemical compositions, though secondary parameters such as O-isotopic compositions, chondrule size and mineralogy are also useful (Table 1). All but two groups (R and K) are divided into the ordinary, carbonaceous and enstatite (O, C, and E) classes. CI chondrites are used as the reference, because they best match the composition of the solar photosphere, neglecting H, He, C, N, O and the noble gases (Fig. 8). Element ratios in CI chondrites and solar photospheric values that have been corrected for settling of He and heavier elements provide the estimates for the bulk composition of the disk from Chondritic Meteorites and Their Components 33 which chondritic components and planets formed (Lodders 2003).

Figure 8. Element/Si ratios relative to mean CI chondrite values for 9 groups of chondrites arranged in order of decreasing O concentration, and the solar photosphere. CI chondrites are used as a standard as they are closest in composition to the surface of the Sun (based on Palme & Jones 2003).

The three groups of ordinary chondrites, H, L, and LL, probably formed close in time and space as their components are very similar in composition. Similarly, the EH and EL enstatite chondrites have highly reduced chondrules containing Si-rich metallic Fe,Ni and (Mg,Mn,Fe)S and probably formed in close proximity to each other. However, the carbonaceous chondrites are much more diverse. Three carbonaceous chondrite groups are rich in C, as their name implies (CR, CM, and CI chondrites contain 1.4-6 wt.% C, and up to 10% H2O), but others are not. Carbonaceous chondrites are actually defined by their abundances of refractory elements like Al and Ca, which exceed or equal those in CI chondrites (Table 1, Fig. 8). All but two carbonaceous groups have >30 vol.% matrix and chondrules that are very largely porphyritic. The matrix-poor groups, CH and CB, are unique as their chondrules are largely non-porphyritic and lack relict grains and matrix and igneous rims. These chondrites clearly formed in very different environments from the other chondrites. Note that the CBa chondrites were once classed as stony-iron, igneous meteorites because they are much richer in Fe,Ni metal than other chondrites and their chondrules are very large and have unusual textures. CI chondrites appear to be unique in that chondrules and CAIs have not been re- ported. However, they do contain a few vol.% of olivine and pyroxene grains with chemical and oxygen isotopic compositions and rounded inclusions of metallic Fe,Ni and trapped melt that show they probably come from heavily altered chondrules (Leshin et al. 1997). CI chondrites also contain rare refractory grains that may be

34 Scott and Krot

derived from CAIs (see Scott & Krot 2003). Several other kinds of ungrouped, matrix-rich carbonaceous chondrites have been described that are related to CI and CM chondrites including Tagish Lake, which fell in 2000 and is mineralogically less altered than CI chondrites (Zolensky et al. 2002; Krot et al. 2003).

3.1. Chemical Compositions of Chondrite Groups Chemical differences between groups are best understood in terms of their nebular volatility, which can be inferred from the equilibrium distribution of elements between minerals and solar nebula gas as a function of temperature (Fig. 1). Elements are divided into four classes: refractory elements like Ca, Al, and Ti that condense at 10-3 bar at temperatures above 1450 K, elements like Mg and Si that condense in the temperature range 1350-1450 K, moderately volatile elements like Na that condense in the range 650-1250 K, and highly volatile elements like S and the halogens that condense below 650 K. Elements may also be among the chondrite groups according to their preference for silicate (lithophile), Fe,Ni metal (siderophile), or sulfide (chalcophile) minerals (Palme 2000; Palme & Jones 2003). Figure 9, which shows element/Si ratios for CV chondrites as a function of con-

Figure 9. Concentrations of 40 elements in CV3 chondrites plotted as element/Si ratios normalized to CI chondrites as a function of solar nebula condensation temperature (temperature at which 50% of element is condensed at 10-4 bar). Highly volatile elements with condensation temperatures <750 K are uniformly depleted at ~0.25×CI levels without regard for their siderophile (metal-loving), lithophile (silicate-loving), or chalcophile (sulfide-loving) preferences. Refractory elements, which condense above 1550 K, are enriched above CI levels with lithophiles more enriched than most siderophiles. Moderately volatile elements that condense between 1250 and 750 K show a strong correlation between abundance and condensation temperature. Thus bulk chondrite compositions are controlled by nebular volatility and metal-silicate preferences. Data from Palme (2000) and Lodders & Fegley (1998). Chondritic Meteorites and Their Components 35 densation temperature, illustrates the strong dependence of elemental abundances in chondrites on their nebular volatility. Refractory elements are uniformly enriched relative to CI chondrites, lithophiles more so than siderophiles. Volatile elements show depletions relative to CI chondrites that correlate inversely with condensation temperature, irrespective of their preference for silicate or Fe,Ni. Highly volatile elements are uniformly depleted with Si-normalized concentrations at 0.25 times CI levels. In ordinary and enstatite chondrites, volatility and metal-silicate preference are both important. H chondrites are richer in siderophiles than L and LL because grains of metallic Fe,Ni accreted preferentially to silicates. Abundance patterns for other chondrite groups tend to resemble those for CV chondrites, except for CH and CB, which show larger enrichments of refractory siderophiles and larger depletions of moderately volatile elements than other chondrites. Enstatite chondrite abundance patterns are more erratic, perhaps reflecting highly reducing conditions during chondrule formation. In some cases, the processes that depleted or enriched elements according to their volatility are clearly related to the abundance and origin of the components. Thus, the bulk refractory abundance of chondrites was established in part by the concentration of refractory inclusions, but also by the concentration of refractories in the chondrules. However, the role of chondrule formation and its importance relative to nebular-wide thermal gradients in establishing the abundance of moderately volatile elements is especially controversial (e.g., Grossman 1996; Palme 2000; Cassen 2001; see also chapters by Yin and Alexander, this volume).

3.2. Sizes and Compositions of Components Each group of chondrites tends to have ingredients with a characteristic range of sizes. Mean chondrule sizes vary from 0.05-5 mm (Table 1) and the sizes of other components such as CAIs, amoeboid olivine aggregates, and dark clasts tend to correlate with chondrule size. Thus CR chondrites have chondrules and CAIs that are typically 0.3-2 mm in size, and metallic Fe,Ni grains and dark clasts that are 0.2-1 mm in size, whereas in CH chondrites all of these components have mean sizes of 20-100 µm (Figs. 10 a,b). Similarly, sizes of chondrules, CAIs, and amoeboid olivine aggregates in CV chondrites commonly exceed those in CO and CM chondrites. In the case of CB chondrites, two kinds of material evidently accreted into one body: CBa material with extremely large components up to several cm in size and CBb chondrites with sub-millimeter components (Figs. 10 c,d). Mean sizes of chondrules and Fe,Ni metal grains in ordinary and enstatite chondrites are not systematically related by their size and density in accordance with aerodynamic size sorting (Sears 2004). In addition, the cm-sized CAIs in CV chondrites, for example, are accompanied by a host of sub-millimeter CAIs that resemble those in other groups (MacPherson 2003). Nevertheless, some kind of aerodynamic sorting during accretion appears to have been responsible for the overall tendency for component sizes to be correlated (Cuzzi et al., this volume). Chemical compositions of chondrules in different groups are reviewed by Jones et al. (this volume). The chemical data, mineralogy and O-isotopic compositions (section 4.1) all indicate that there was little mixing between the chondrule popula- tions of C, O, and E chondrites. In addition, separate chondrule reservoirs are needed

36 Scott and Krot for the R and K chondrites, and the matrix-poor, carbonaceous chondrite groups, CH and CB.

Figure 10. Combined elemental maps with Mg (red), Ca (green), and Al (blue) for a) CR chondrite, b) CH chondrite, c) CBa chondrite, and d) CBb chondrite showing that chondrites in different groups can be distinguished by the size and nature of their major components. CR chondrites (a) have mm-sized, volatile-poor chondrules with coarse-grained igneous rims, grains of metallic Fe,Ni (black spherules) up to 1 mm in size, and abundant matrix (mx) that rims components. By contrast, CH chondrites (b) have smaller components that lack matrix rims. CB chondrites have very different properties from other chondrites and come in two varieties, the very coarse-grained CBa and the fine-grained CBb chondrites. (After Krot et al. 2003).

Bulk chemical and isotopic data for matrices of pristine chondrites are very limited (Huss et al., this volume); mineralogical studies of the few unaltered chondrites Chondritic Meteorites and Their Components 37 show that K and C chondrites have very different matrix minerals. We expect that E and O chondrite matrices also differ from those in C chondrites.

3.3. Parent Bodies of Chondrite Groups Studies of cosmic-ray exposure and shock ages and chondrite breccias suggest that each group of chondrites probably comes from one or a few parent asteroids that were probably many tens of kilometers or more in size (see Scott & Krot 2003). Es- tablishing where and when each chondrite group formed is a major goal of chondrite studies (Wood, this volume). Constraints on formation locations can be derived from a wide variety of meteorite and asteroidal properties as well as dynamical modeling of the early asteroid belt and the delivery of meteorites to Earth. However, chondrite groups cannot be arranged into a single, genetic sequence on the basis of their bulk chemical or O-isotopic properties, as these parameters are poorly correlated. This suggests that a second parameter, time of accretion, was as important as radial distance in determining bulk chondrite properties. The first step in establishing the formation locations of the chondrites is to identify their parent asteroids, or at least the likely class from which each chondrite group is derived. The most abundant asteroids, C-types, which dominate the outer belt, are clearly made of some kind of matrix-rich, carbonaceous chondrite (Burbine et al. 2002; Gradie, Chapman, & Tedesco 1989). Foreign rock fragments in primitive chondrites (dark clasts), regolith breccias, and the current micrometeorite population all suggest that CM-like material has always been dominant in the belt. Ordinary chondrites probably formed closer to the Sun than the matrix-rich, carbonaceous chondrites, judging from their close relationship to S-type asteroids (Chapman 2004), which dominate in the inner belt at ~2.2 AU and their higher peak metamorphic temperatures (see below). Enstatite chondrites are probably derived from closer to the Sun as they are closely related to the enstatite achondrites, whose supposed parents, the E type asteroids, are concentrated at 1.9-2.0 AU. Timescales for the accretion of planetesimals into larger bodies would have increased with increasing distance from the Sun because of decreasing surface density and increasing orbital periods. However, other factors may also have affected accretion. The absence of matrix-rich, chondrule-poor chondrites (except for CIs), which could have been sintered to form tough meteoroids, suggests that chondrules may have been required to initiate planetesimal accretion in the inner solar system where sticky organic materials were lacking. In addition, the surface density of solids would not have varied smoothly with heliocentric distance as density steps could have developed at locations corresponding to the decomposition temperature of interstellar organics (350-450 K; Nakano et al. 2003) and the sublimation temperature of ice (~180 K). These tar and snow lines would have advanced inwards as the nebula cooled and may also have helped to trigger accretion. Dynamical modeling of the early asteroid belt and the paucity of differentiated asteroids, except for Vesta and the parent bodies of the iron meteorites, suggest that most differentiated asteroids may have formed inside 2 AU and wandered into the belt as a result of scattering by planetary embryos (Bottke et al. 2005). This is consistent with inferences about the distribution of 26Al, the most plausible heat source (section 4.2). Bodies that accreted within 2 Myr of CAI formation would probably

38 Scott and Krot have melted unless thermally buffered by significant amounts of ice and other volatiles. The parent bodies of the chondrites may therefore have accreted in the asteroid belt 2-4 Myr after CAI formation when 26Al was no longer a potent heat source (Hevey & Sanders 2005; Sanders & Taylor, this volume). Tungsten isotope data for iron meteorites support earlier formation of differentiated asteroids (Kleine et al. 2005), but more radiometric age data are needed. A knowledge of the accretion locations and times of individual chondrite groups (and the interplanetary dust particles) will surely allow considerable insights into the origin and accretion of the chondritic components, asteroids and the terrestrial planets. However, this must await further insights from the isotopic properties of chondritic components, and formation ages for chondrules and CAIs.

4. Isotopic Compositions of Chondritic Components A wide variety of isotopic variations are found in chondrites, which provide key constraints on the origins of the chondritic components (e.g., Birck 2004; Podosek, this volume). These include (1) fractionation effects that are a linear function of isotope mass in CAIs (but not chondrules) that were produced during evaporation and condensation in the solar nebula, (2) nucleosynthetic effects preserved in presolar grains from diverse evolved stars, (3) isotopic variations due to decay of radioactive isotopes, (4) mass-independent effects in oxygen, which are ubiquitous in refractory inclusions and chondrules, and (5) spallation effects in rocks due to irradiation by energetic particles during transit to Earth and on asteroidal surfaces, and possible spallation effects in CAIs due to irradiation from the protosun. Here, we briefly discuss mass-related effects before focusing on the nucleosynthetic isotopic variations in presolar grains and the memories of these variations that are preserved in chondrites, isotopic variations due to short-lived isotopes, and the oxygen isotope effects.

4.1. Mass-dependent Isotopic Variations At equilibrium, mass-dependent isotopic differences between gas, solids, and melts are negligible at the high temperatures at which chondrules and CAIs formed: < 1-2 parts per thousand (‰). However, kinetic processes during condensation and evaporation can produce large mass-related isotopic effects, which provide important constraints on nebula conditions and timescales during high temperature processing (Davis & Richter 2003; Davis et al., this volume). For example, the range of mass- dependent variations for Ca isotopes is 20 × larger in CAIs than in terrestrial, lunar and bulk meteorite samples (Niederer & Papanastassiou 1984). Maximum fractionation occurs when the starting phase is continually homogenized and back reaction between the gas and the condensed phase is minimized. Because diffusion is much more rapid in liquids than in solids, maximum fractionation requires a liquid, but other factors such as gas composition and pressure are also critical. Mass fractionation effects are relatively minor for chondrules implying that conditions were close to equilibrium during gas-liquid interactions, whereas igneous CAIs typically show large effects (Davis et al., this volume). For example, igneous type B CAIs show enrichments of heavy isotopes of Si and Mg as a result of evaporation from melts during brief reheating. A rare subset of igneous CAIs called FUN inclusions shows larger mass fractionation effects for many elements including O Chondritic Meteorites and Their Components 39 than most other CAIs, e.g., 7-30 ‰ per atomic mass unit (amu) for Mg isotopes in FUN CAIs, cf. 0-11‰ amu-1 for normal CAIs (Davis & Richter 2003). The FUN inclusions also show large nucleosynthetic isotopic effects (Fractionation and Un- known Nuclear effects), but the reason for this coupling of mass and nuclear effects is not clear.

4.2. Nucleosynthetic Isotopic Variations As noted above, the isotopic deviations from terrestrial values in presolar grains due to nucleosynthetic effects are orders of magnitude larger than those preserved in CAIs, chondrules, and bulk matrix, which are measured in parts in 103 or 104 (δ or ε units, respectively). Thus, presolar particles and those that formed in the solar system can be readily distinguished by their mass-independent isotopic deviations (Fig. 11). The largest isotopic deviations in solar-system particles are found in the FUN-type

Figure 11. Deviations in the Nd isotopic ratios in a FUN-type Ca-Al–rich inclusion (EK1-4-1) and a presolar SiC grain normalized to 142Nd and a terrestrial standard. The isotopic variations in the presolar grain are 100 × larger than those in the CAI and of op- posite sign. The presolar grain contains predominantly s-process Nd, which is deficient in the CAI. (After Ott 1993).

CAIs and certain hibonite inclusions, which contain deviations in 50Ti, for example, of up to 10% (Ireland & Fegley 2000). The nucleosynthetic effects in normal CAIs are much smaller, e.g., variations of 5-30 ε units in 50Ti/46Ti (Niemeyer & Lugmair

40 Scott and Krot

1984), while chondrules have even smaller variations, e.g., a few ε-units or less in 50Ti/46Ti (Niemeyer 1988). Bulk chondrites and matrix samples also show isotopic variations of a few ε-units in 50Ti/46Ti (Niemeyer & Lugmair 1984), 54Cr (Podosek et al. 1999) and Mo isotopes (Yin, Jacobsen, & Yamashita 2002). The effects in CAIs, chondrules, matrices, and whole chondrites probably reflect the gradual isotopic homogenization by thermal processing in the solar nebula of diverse presolar grains that differ in their thermal stability, grain size, and other properties. This implies that FUN inclusions predate other CAIs (Wood 1998; Sahijpal & Goswami 1998), and that CAIs predate chondrules. The nucleosynthetic isotopic effects preserved in chondrules and chondrites have not yet provided firm constraints on their formation locations or mechanisms, though they clearly emphasize that the high-temperature processing experienced by chondrules was brief and/or localized.

4.3. Oxygen Isotopic Compositions The discovery by Clayton and colleagues of oxygen isotopic variations in CAIs and chondrules that could not be attributed to mass-dependent fractionation transformed the study of chondrites and cosmochemistry (see Clayton 1993, 2003). It galvanized the search for presolar grains (although ironically they are absent in CAIs), led to the demise of models for chondrites that relied solely on a hot, gaseous, and homogeneous nebula, provided a new way to classify meteorites, and generated important constraints on the origin of chondritic components, and insights into the history of minerals, moons, and planets. A key property of oxygen that was important in preserving isotopic heterogeneities in the protosolar disk is that oxygen is the only element with three or more isotopes that exists in several forms – both gaseous (CO and H2O) and solid (ice and rock) – over a wide range of nebula temperatures (Wood 1981). Oxygen isotopic data are plotted on a graph of 17O/16O vs. 18O/16O with units of deviations in parts in 103 from standard mean ocean water (Fig. 12). Most chemical reactions partition isotopes according to their masses along a single line of slope 0.52. Terrestrial samples lie on such a line showing that the Earth was initially isotopically homogeneous. Martian meteorites lie on a parallel line just above the terrestrial line. However, refractory inclusions, chondrules, and matrix samples plot on or near a slope-1 line with refractory inclusions furthest from the terrestrial line (Fig. 12). Vertical deviations from the terrestrial line are defined by the parameter ∆17O = δ17O - 0.52×δ18O (Clayton 1993). Mixtures of two components lie on a straight line connecting the end-members. Although the O-isotopic variations among CAIs were initially attributed to presolar grains, a local origin is now favored as O-bearing stardust in chondrites is largely 17O-rich and CAIs do not show isotopic abnormalities in other elements (Clayton & Nittler 2004). Laboratory experiments by Thiemens & Heidenreich (1983) showed that certain gas phase reactions could lead to products that differed in ∆17O from the reactants. Although various mechanisms have been proposed involving the symmetry of minor molecules such as O3, O2, and CO2 (Thiemens 1999), the most plausible explanation for the chondrite variations invokes isotopic self-shielding during UV photolysis of the abundant molecule CO in an initially 16O-rich protoplanetary disk or parental molecular cloud with ∆17O ~-25‰ (Clayton 2002; Lyons & Young 2005; Yurimoto & Kuramoto 2004; Yin 2004; Krot et al. 2005a). Astronomi- Chondritic Meteorites and Their Components 41 cal observations show that in molecular clouds, 12C17O and 12C18O can be preferentially dissociated, as UV photons that dissociate the vastly more abundant 12C16O cannot penetrate beyond the surface of the cloud (Federman et al. 2003). Atomic 17O and 18O can combine with hydrogen forming water ice, while the CO is enriched in 16O. In this model, the bulk solar system has a ∆17O value of ~-25‰, like most CAIs. If the inner nebula becomes enriched in H2O because of meter-sized, ice-rich bodies that drift rapidly towards the Sun and evaporate (Cuzzi & Zahlne 2004), then the inner nebula and chondrules formed therein will be enriched in 17O and 18O.

Figure 12. Oxygen isotopic plots of 17O/16O vs. 18O/16O with isotopic compositions plotted as deviations from standard ocean water in parts per 103 (δ-units). (a) Terrestrial samples show mass-dependent isotopic variations and plot on the line labeled terrestrial fractionation whereas refractory inclusions in most pristine chondrites plot in the lower left of the diagram. Chondrules and matrices plot closer to the terrestrial line within the box. CAIs in the Allende chondrite scatter along a slope-1 line. (b) Enlarged view of boxed region in (a) showing ranges of O isotopic compositions of large chondrules in carbonaceous chondrites, enstatite chondrites, ordinary chondrites (OCs), and R chondrites. There was little mixing between the chondrules in these four kinds of chondrites. (After Rubin 2000.)

Although many details of the 16O-fractionation process are not fully understood (Lyons & Young, this volume), one prediction of the CO self-shielding model has been confirmed. Ion probe analyses of the surfaces of lunar grains of metallic Fe (or Fe,Ni) revealed implanted solar wind with near-CAI oxygen isotopic compositions (Hashizume & Chaussidon 2005). What can be inferred about the origin of CAIs and chondrules from their O- isotopic compositions? First, most CAIs formed in a relatively 16O-rich environment that was probably solar in composition and quite distinct from the 16O-poor environment in which chondrules formed. Exceptions include some 16O-poor CAIs in CH chondrites, all CAIs in CB chondrites, and some igneously formed CAIs in CV and CR chondrites that appear to have partially exchanged oxygen with an 16O-poor gas

42 Scott and Krot when they were molten (Krot et al. 2005a). Minerals in these exceptional CAIs have oxygen isotopic compositions that plot in Figure 12 on the line marked “Allende CAIs”, which has a slope of ~1. Chondrules and matrix samples have oxygen isotopic compositions that lie close to the top of the Allende CAI line, near the terrestrial fractionation line (Fig. 12b). If the oxygen isotopic composition of the nebula became progressively more 16O-poor (Krot et al. 2005a), then chondrules formed after CAIs. Most individual chondrules in unaltered chondrites, like most CAIs, do not show mass-independent O-isotopic fractionations between, or within, minerals, except for 16O-rich relict grains. How- ever, chondrules from one chondrite (excluding those in E and CB chondrites) commonly scatter along or near the slope-1 line with a spread of ∆17O values of up to 5‰ (e.g., Clayton et al. 1985, 1991; Scott & Krot 2001; Jones et al., this volume). This shows that they cannot be derived from a single source such as a molten body. The wide range of O-isotopic compositions of chondrules in all chondrites shows that many distinctly different batches of chondrules were manufactured in the nebula (Fig. 12). Oxygen isotopes therefore confirm the mineralogical evidence that chondrules in enstatite, ordinary, carbonaceous, and R chondrites come from distinctly different sources. In addition, carbonaceous chondrules come from several sources, e.g., those in CR chondrites are mineralogically and isotopically distinct from chondrules in CM and CO chondrites. Other distinct sources are needed for the chondrules in CH, CB, and K chondrites. Al-rich chondrules, which are present in many groups, tend to have overlapping 16O-rich O-isotopic compositions, consistent with an origin as mixtures of chondrule and CAI material (Krot et al. 2005b). To illustrate that chondrites consistently provide exceptions for every rule, we note that chondrules in CH chondrites, which have the widest range of ∆17O values (+5 to –35‰), extend along the slope-1 line to below the CAI field (Kobayashi et al. 2003; Yoshitake & Yurimoto 2004). A few CH chondrules also have heterogeneous O isotopic compositions. The 16O-rich CH chondrules may have formed after other chondrules when 16O-rich CO began to dominate the nebula gas.

4.4. Short-lived Radioactive Isotopes Proof that short-lived isotopes decayed in chondritic components comes from excesses of the daughter isotopes that are correlated with the abundances of a stable isotope of the parent element in cogenetic minerals, e.g., 26Mg/24Mg ratios in diverse minerals of igneous CAIs are correlated with 27Al/24Mg ratios showing that 26Al decayed to 26Mg in situ (Table 2). The inferred initial concentrations of relatively long- lived isotopes like 53Mn and 182Hf appear to be compatible with steady-state interstellar abundances due to continuous galactic nucleosynthesis (Meyer, this volume). However, the inferred concentrations of 41Ca, 26Al, 60Fe and 36Cl are much higher than the steady-state abundances. These isotopes formed by nucleosynthesis in evolved stars or by spallation during energetic particle bombardment in the solar nebula or molecular cloud (Goswami et al., this volume). For 60Fe, a stellar origin is accepted, as energetic particle interactions are not a plausible source. However, 10Be, which was present in diverse kinds of CAIs (McKeegan, Chaussidon, & Robert 2000), is not made in stars, and inferred initial abundances of 10Be and 26Al are not correlated (Marhas et al. 2002). 10Be was probably acquired during spallation by pro- Chondritic Meteorites and Their Components 43 tosolar energetic particles close to the protosun, or possibly by trapping of galactic cosmic rays in the protosolar cloud (Desch et al. 2004).

Table 2. Short-lived nuclides present in chondrites.

Nuclide Half-life Daughter (Myr)

41Ca 0.10 41K 26Al 0.74 26Mg 10Be 1.5 10B 60Fe 1.5 60Ni 53Mn 3.7 53Cr 107Pd 6.5 107Ag 182Hf 9 182W See Goswami et al. (this volume) for complete list

Short-lived isotopes are critically important for understanding the early history of the solar system. Evidence for live 26Al in a CAI (Lee et al. 1976) led to the sug- gestion that solar system formation was triggered by a supernova (Cameron & Truran 1977). Discovery of evidence for 60Fe in an igneously formed meteorite (Shu- kolyukov & Lugmair 1993) and in chondrites (Tachibana & Huss 2003) has strength- ened the view that the Sun formed close to massive stars, not in isolation in a low- mass region like the Taurus-Auriga molecular cloud (Reipurth, this volume; Bally, Moeckel, & Throop, this volume). According to Hester & Desch (this volume), star formation may be triggered by the radiation and winds from massive stars that com- press the surrounding gas before the arrival of 60Fe in supernova ejecta. (For contrary views, see Gounelle et al. (2001) who disputed that massive stars are involved in forming Sun-like stars, and Hartmann (this volume) who infers that supernova ejecta were mixed into a large clump of molecular gas prior to disk formation.) The isotopes 26Al and 60Fe are also important, as they probably caused early- formed asteroids to melt: 26Al supplied more heat initially, but 60Fe may have become a significant heat source after a few Myr (Mostefaoui et al. 2005). Provided that these short-lived isotopes were homogeneously distributed, they can be used as chronometers for dating chondrule and CAI formation (Kita et al., this volume).

4.5. Radiometric Ages of CAIs and Chondrules After considerable efforts to identify and date pristine chondrules and CAIs that have not had their isotopic systems reset by asteroidal processing, a consistent chronology for the formation of chondrules and CAIs now appears to be developing (Kita et al., this volume). The most precise absolute ages of chondrules and CAIs come from the 207Pb-206Pb ages of Amelin et al. (2002, 2004): 4567.2±0.6 Myr for CAIs in CV chondrites, and 4566.7±1.0, 4564.7±0.7, and 4562.7±0.5 Myr for chondrules in CV, CR, and CB chondrites, respectively. Thus timescales for the total duration of CAI and chondrule formation are consistent with the lifetimes inferred for circumstellar disks around newborn stars (Reipurth and Hartmann, this volume).

44 Scott and Krot

Relative formation ages can be derived for chondrules and CAIs from initial 26Al/27Al ratios inferred from plots of 26Mg/24Mg against 27Al/24Mg (Kita et al., this volume). Detection of 26Mg excesses in bulk samples allows model ages to be inferred based on an assumed initial Mg isotopic ratio. The overall correlation between the absolute 207Pb-206Pb ages and the relative ages from inferred initial 26Al/27Al ratios for CAIs, chondrules and some rapidly cooled H4 chondrites has major implications (see Sanders & Taylor, this volume), viz., that 26Al was homogeneously distributed prior to chondrule and CAI formation and that it was not formed in the solar system. Homogeneity of 26Al is consistent with the uniformity of 26Mg/24Mg in various chondrite groups (Bizzarro, Baker, & Haack 2005). An important exception is provided by the rare FUN CAIs, which were deficient in 26Al, presumably because they formed prior to injection of 26Al into the disk. 53Mn-53Cr relative ages for chondritic components and old differentiated meteorites generally agree well with the Al-Mg and Pb- Pb ages within analytical uncertainties, supporting the homogeneous distribution of these nuclides (Wadhwa et al. 2005). Radiometric ages imply that CAI formation lasted for <0.3 Myr, or possibly <0.1 Myr (Bizzarro, Baker, & Haack 2004); chondrule formation started 0-2 Myr later and lasted for 1 to 4-5 Myr (Kita et al. and Russell et al., this volume). We note that chondrules in CB chondrites have been interpreted as products of impacts between planetary embryos after dust in the disk largely dissipated (Rubin et al. 2003; Krot et al. 2005c). If this is the case, the formation of normal chondrules may have lasted <3 Myr. Formation ages of amoeboid olivine aggregates are not well constrained, but limited Al-Mg data suggests they probably formed contemporaneously with CAIs or <1 Myr later (Itoh et al. 2002). Some CAIs were reheated 1-2 Myr after the majority of CAIs had formed, probably during chondrule formation. The period over which chondrules in a single group formed is poorly constrained. Al-Mg data for CO3 and LL3 chondrules suggest that chondrules in both groups may have formed from 1.5 to 2.5 Myr after CAIs (Kita et al., this volume). However, the analytical uncertainties do not preclude significantly shorter periods. Data for LL3.0-3.1 chondrules suggest that the olivine-rich chondrules are on average ~0.5 Myr older than pyroxene-rich chondrules (Mostefaoui et al. 2002; Kita et al., this volume).

5. Relationship Between Primitive Chondrite Matrices and Cometary Dust Because comets accreted at very low temperatures in the outer parts of the solar nebula, their silicates are commonly thought to have formed in molecular clouds prior to being coated with interstellar or solar nebula organics and ices (e.g., Brandt 1999; Hanner 1999). By contrast, asteroids, which by definition lack comas, are considered to have formed from silicates that were produced in the solar nebula and accreted at temperatures that were too high for water ice to be stable. Similarly, the chondritic, porous, anhydrous, interplanetary dust particles, which are thought to have come from comets (Rietmeijer 2002; Bradley 2003), are widely inferred to have little in common with chondrites. However, all of these views need revision. Astronomical studies of asteroids and comets (Weissman et al. 2002; Wooden 2002) and laboratory studies of matrices in the least altered chondrites (section 2.3) have blurred some of the distinctions between comets and asteroids and their off- spring. A major fraction of the silicates in the chondritic, porous IDPs and the comas Chondritic Meteorites and Their Components 45 of long-period comets are crystals that formed at high temperatures in the solar nebula (Hill et al. 2001; Messenger et al. 2003; Nuth et al., this volume). Moreover, the chemical and physical properties of these silicates suggest that cometary and chondritic crystals may have related origins (Scott & Krot 2005; Wooden et al., this volume). The silicate portions of the chondritic, porous IDPs, cometary silicates, and matrices of carbonaceous chondrites all contain submicrometer-sized forsterite and enstatite crystals with Fe/(Fe+Mg) <0.05 and amorphous Fe-Mg-Si-O particles (Brad- ley 2003; Wooden 2002; Wooden, Woodward, & Harker 2004; section 2.3). Chon- dritic, porous IDPs and chondrite matrices both contain submicrometer Mn-rich forsterite and enstatite (Klöck et al. 1989), and pyroxene crystals with ortho-clino intergrowths characteristic of rapid cooling at ~1300 K at ≈1000 K hr-1 (Bradley et al. 1983). Thus many crystalline silicates in comets and chondrite matrices probably formed in brief, localized thermal events. A tiny fraction of the crystalline silicates are presolar: ≈1000 ppm in IDPs and ≈10-100 ppm in chondrite matrices, the rest formed in the solar system (Messenger et al. 2003; section 2.3.3). Amorphous silicate particles in chondrite matrices, which are chemically heterogeneous and may contain grains of metallic Fe and sulfide (Brearley 1993), may be related to the particles in IDPs called GEMS, which are made of silicate Glass with Embedded Metal and Sulfides (Bradley et al. 1999; Chizmadia 2005). Small fractions of the amorphous silicates in chondritic, porous IDPs (~5%) are presolar as they have distinctive oxygen isotopic compositions (Messenger et al. 2003), and similar grains appear to be present in chondritic matrices (section 2.5.3). The origin of the large fraction of GEMS in IDPs that has solar-system oxygen isotopic compositions is controversial: they may be presolar (Bradley & Dai 2004; Westphal & Bradley 2004) or solar nebular in origin (Keller & Messenger, this volume). Thus some amorphous particles in IDPs and chondrites may also share common solar nebula origins. Additional similarities between chondritic, porous IDPs and chondrite matrices include the presence of refractory 16O-rich silicates and oxides like those in CAIs (section 2.3.4) and some apparently foreign material such as large sulfides up to 8 µm in size (Rietmeijer 2002; Bradley 2003). As in chondrites, the grain sizes of IDP components are non-random as similar-sized units in IDPs accreted together (Riet- meijer 2003). In both types of material, organic glue holds the components together. There is no sharp mineralogical distinction between chondritic IDPs with asteroidal and cometary orbits (Brownlee et al. 1993), and some chondritic, porous IDPs contain both hydrated and anhydrous silicates (Bradley 2003). Nevertheless, chondritic, porous IDPs may have concentrations of carbon and presolar grains that are considerably higher (≈10-100 × higher) than the matrices of the least altered chondrites (Messenger et al. 2003; Bradley 2003; section 2.3.3). Although comets clearly accreted at larger heliocentric distances and lower ambient temperatures, we suggest they are not fundamentally different from the parent bodies of many primitive carbonaceous chondrites. The latter also accreted with water ice, and at least one C type asteroid has exhibited cometary behavior (Hsieh, Jew- ett, & Fernandez 2004). Comets are certainly less affected by solar system processing, but nevertheless contain abundant crystalline silicates that resemble the crystalline silicates in primitive chondrite matrices.

46 Scott and Krot

6. Origins of Chondritic Components All but a tiny fraction of the components in chondrites formed at high temperatures in the solar nebula: CAIs, amoeboid olivine aggregates, chondrules and associated coarse grains of metallic Fe,Ni, plus matrix microcrystals of magnesian silicates and refractory silicates and oxides, and probably most of the amorphous silicate grains in matrices. Rare presolar survivors are found in matrices of primitive chondrites: crystalline silicate and oxide grains, amorphous silicate grains together with some carbo- nacous material. Recent advances in understanding the chronology of unaltered CAIs and chondrules and developing models for their oxygen isotopic compositions provide the basis for outlining a reasonably self-consistent scenario for the formation of chondritic components at high temperatures. Evidence from short-lived and long-lived isotopes, oxygen isotopes, nuclear isotopic effects and petrologic studies all suggest that refractory inclusions and grains were the first solids to form in the protosolar disk, probably within a short period of <0.3 Myr. Thus, CAIs probably formed when the protosun was accreting rapidly as a Class 0 or I protostar (Wood 2004; see Reiputh, this volume). The accretion of CAIs together with chondrules and matrix into chondritic asteroids and evidence for thermal processing of CAIs during chondrule formation (Russell et al., this volume) suggest that CAIs survived in the solar nebula for 1- 3 Myr. Nebula turbulence may have prevented all the refractory inclusions from drift- ing into the protosun during this period (Cuzzi et al. 2003). Refractory inclusions formed in what has been considered as a uniquely 16O-rich environment, but it now seems probable that the mean oxygen isotopic composition of this environment was solar. CAIs formed under reducing conditions at low total pressure under 10-4 bar and high temperatures above 1300 K. Some CAIs were subsequently melted and evaporated causing mass-dependent enrichment of heavier Mg and Si isotopes. Virtually all melted and unmelted CAIs acquired Wark-Lovering rims in the CAI-forming region by high-temperature gas-solid reactions. Amoeboid olivine aggregates and the closely related forsterite rims on CAIs were the last solids to form in the 16O-rich region. Rapid cooling or physical removal prevented reactions between refractory solids and gas at lower temperatures. The innermost edge of the disk near the proto-Sun appears to be the most plausible site for the formation of refractory inclusions (Shu et al. 2001; Wood 2004). Chondrules may have formed in the solar nebula soon after CAIs (Bizzarro, Baker, & Haack 2004), but chondrules in most chondrites appear to have formed 1-3 Myr after CAIs (Kita et al., this volume), when the protosun was accreting more slowly, possibly as a Class II object (Reipurth, this volume). Chondrules formed from dust under conditions that were generally more oxidizing than for CAIs, under lower ambient temperatures and higher dust/gas ratios, so that liquids were stable for hours. Several kinds of chondrules were formed in separate environments. Type I chondrule formation involved melting, evaporation, condensation, and accretion of solid and melted materials. Type II chondrules, which contain Fe-silicates, probably formed by a similar mechanism at higher dust/gas ratios, though they appear to have had a sim- pler history without evaporation and condensation. Some fraction of the refractory inclusions appears to have been melted in the chondrule-forming region 1-2 Myr after CAI formation (Russell et al., this volume). Chondritic Meteorites and Their Components 47

The major component of chondrite matrices—micrometer-sized crystalline magnesian silicates and amorphous Fe-Mg silicate particles—have many properties suggesting they formed in close proximity to chondrules at high temperatures. The cooling rates, composition, and structure of the magnesian silicates suggest that they probably condensed during heating events like those that formed chondrules (Scott & Krot 2003, 2005; Nuth et al., this volume). The near-solar composition of the matrix suggests that the amorphous silicates have similar origins. Matrix grains from various chondrule-forming regions were assembled at low temperatures with traces of refractory dust from the CAI-forming region and presolar dust and accreted to form rims on chondrules and refractory components (Fig. 13). (See Alexander and Huss et al., this volume for different conclusions about the origin of chondrite matrices.)

Figure 13. Schematic diagram showing how amorphous, presolar, silicate dust may have been thermally processed in the nebular disk before accretion into chondritic and cometary planetesimals (after Scott & Krot 2005). Dust that accreted at 2-3 AU into chondrite matrices contains traces of refractory dust (~10-5 to 10-4) that were dispersed from the inner edge of the disk by disk winds and turbulence, and presolar, crystalline silicates and refractory oxides (~10-5) and presolar, amorphous silicates and oxides (~10- 6), which escaped thermal processing. Most matrix silicates are crystalline, magnesian silicate and amorphous Fe-Mg silicate that condensed when dust aggregates were melted in the nebula to form chondrules, probably as a result of shocks. Comets accreted dust at >10 AU containing traces of refractory nebular dust (~10-6 to 10-5) and presolar, crystalline silicates and oxides (~0.1%). The remaining abundant, crystalline silicates in comets formed by condensation in the nebula; the proportions of presolar and solar, amorphous, silicate particles are less certain. Debris from planetesimal collisions at the midplane became increasingly important at later stages.

Given the unique chemical and isotopic properties of chondrules in many different chondrite groups and the evidence that CAIs and presolar grains were spread over the asteroid belt, we think it unlikely that chondrules in a single group could have formed over a period generally longer than ~0.3 Myr, i.e., a tenth of the total chon-

48 Scott and Krot drule formation period. Although the CAIs in each group have much in common, the mineralogical differences between groups are hard to explain. However, size sorting and alteration may account for some of the observed differences. If the abundance of refractory inclusions in the disk gradually declined, CV chondrites may have formed first, then the other matrix-rich C chondrites, followed later by O and E chondrites (Cuzzi et al. 2003). However, Al-Mg ages do not support later formation of ordinary chondrites: data for E chondrites are lacking (Kita et al., this volume). In addition to primitive nebular materials, chondrites also contain fragments of early-formed bodies that were aqueously altered or melted (section 2.4). Given a homogeneous 26Al distribution, it is almost inevitable that the parent asteroids of the chondrite groups accreted >2 Myr after CAIs formed, whereas the melted asteroids accreted <2 Myr after CAI formation. The abundance of altered chondritic rock fragments in chondrites suggests that carbonaceous asteroids that were subsequently altered hydrothermally accreted in the outer part of the asteroid belt while differentiated bodies at the inner edge of the main belt or inside 2 AU. We infer that both kinds of bodies formed from materials present in the chondrites, and, contrary to Sanders (1996), do not infer that most chondrules formed from molten asteroids. If any chondrites formed from colliding asteroids, it is likely to be the CB chondrites. As noted above, their properties are very different from those of other chondrites.

7. Nebular Heat Sources for Chondritic and Cometary Silicates Although the earliest refractory condensates in CAIs and AOAs appear to have formed from rapidly cooling nebular gas, probably close to the protosun (Petaev & Wood, this volume), the heating mechanism is poorly understood (Shu et al. 1996; Wood 2004). The most plausible mechanism for making chondrule and CAI melts is shock wave heating in the solar nebula. However, the sources of the shock waves remain controversial (Ciesla, this volume). Possible sources include bow shocks generated by planetesimals (Hood et al., this volume), spiral arms and clumps in a gravi- tationally unstable disk (Boss & Durisen, and Boley et al., this volume), close passes by protosolar companions (Reipurth, this volume; Bally et al., this volume), and X- ray flares (Nakamoto et al., this volume). The properties of chondrite matrices and anhydrous, porous chondritic IDPs suggest that thermal processing of interstellar dust was ubiquitous, brief, and localized, but limited in extent so that very small fractions of presolar materials survived. Annealing of presolar amorphous material either close to the protosun (Hill et al. 2001; Gail 2004) or by shocks (Harker & Desch 2002) is commonly thought to have generated cometary silicate crystals. However, the general absence of Mg-rich, Fe- poor, amorphous silicate from which the Mg-rich crystalline silicates could have formed by annealing favors an origin by condensation. If chondrules formed by shock heating, as seems plausible, it is likely that dust was vaporized by the shock (Desch & Connolly 2002) and rapidly condensed so that the bulk composition of the dust was not grossly disturbed. Since forsterite crystals and amorphous silicates are present around many T-Tauri and higher mass Herbig Ae/Be stars (Bouwman et al. 2003; Honda et al. 2003), the processes that heated and mixed silicate dust in the solar nebula (Fig. 13) may also have operated in other protostellar disks. Chondritic Meteorites and Their Components 49

Acknowledgments. We thank John Wood for his inspiring contributions to this field over many years; C. M. O’D. Alexander, A. E. Rubin, and B. Reipurth for their helpful reviews; and many other colleagues for helpful discussions. This work was supported by NASA grants 5-4212 (K. Keil, PI) and 5-10610 (A. Krot, PI).

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