Thermal alteration of asteroids: evidence from meteorites

Planetary and Space Science 48 (2000) 887–903

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Thermal alteration of asteroids: evidence from meteorites  Klaus Keil Hawai’i Institute of Geophysics and Planetology, School of Ocean and Earth Science and Technology, University of Hawai’i at Manoa, Honolulu, Hawai’i 96822, USA Received 23 December 1999; accepted 3 March 2000

Abstract The world’s meteorite collections contain meteorites from at least ∼27 primitive, chondritic and ∼108 partially or totally melted asteroids. Detailed studies of these meteorites have shown that all of their asteroidal parent bodies have been thermally altered by internal ◦ heating to some degree. These alterations range from low temperature, aqueous processes (∼0 to ¡ ∼300 C), to thermal metamorphism ◦ ◦ ◦ (∼400 to ∼950 C), to partial melting and formation of unfractionated (∼980–1050 C) and fractionated (∼1000 to ¿ 1250 C) residues, ◦ to partial and complete melting, dierentiation and fractional crystallization of asteroids (∼1150 to 1250 C). The most likely heat source was the decay of short-lived radionuclides, notably 26 Al. These thermal alterations took place penecontemporaneously on all asteroids of which we have samples, and in the rst few Ma of solar system history. The asteroid 4 Vesta, the likely parent body of the HED meteorites, is a highly dierentiated object, may have a metal core, and can be viewed as the smallest of the terrestrial planets. It accreted, was heated, was partially to completely melted, and formed an extrusive basaltic crust, all within a few Ma of c 2000 Elsevier Science Ltd. All rights reserved. formation of CAIs and the dawn of the solar system.

1. Introduction Recently, Meibom and Clark (1999) reviewed what is known about the number of asteroids that are represented in the world’s meteorite collections. They conclude, in agreement with earlier studies (e.g. Scott, 1977; Wasson, 1995) that the collections contain meteorite groups (¿ 5 members), grouplets (¡ 5 members) and unique meteorites representing at least ∼135 asteroidal parent bodies (Table 1; note that this is possibly a lower limit, as not all meteorites in one group necessarily come from only one parent body). This is based on a combination of chemical, mineralogical and isotopic compositions and textures of meteorites that appear to be incompatible with an origin on one and the same parent body. This is an astonishingly large number, and detailed laboratory studies of these meteorites should allow important conclusions to be drawn about the thermal histories and evolutions of a large number of these asteroids. It is noteworthy that amongst these ∼135 parent bodies, only 27 represent primitive, chondritic asteroids, whereas the remaining ∼108 were partially or totally melted and dierentiated. However, in spite of the remarkably large number of parent bodies that are represented by meteorites, there are  Modi ed invited talk to the “Asteroids, Comets, Meteors” Conference, July 26 –30, 1999, Cornell University, Ithaca, NY, USA.

still a number of problems and uncertainties associated with meteorite-asteroid comparisons, and these must be kept in mind when meteorite properties are used to decipher the geological evolution of their parent bodies. A few examples are listed here: (a) Only the howardite–eucrite–diogenite meteorites (HEDs) can be related with reasonable certainty to a speci c asteroid, 4 Vesta (McCord et al., 1970; Binzel and Xu, 1993), whereas for all other meteorites, the speci c source asteroids are unknown. It should be noted that Gaey and Gilbert (1998) recently argued that the S(IV)type asteroid 6 Hebe is the parent body of the H-group ordinary chondrites and the IIE iron meteorites. This suggestion is fraught with considerable uncertainties, because it requires complex impact processes such as the formation of metal melt pools by partial impact melting to have altered the surface of the asteroid. However, detailed observations on meteorites and terrestrial impact craters as well as experimental and theoretical studies have shown that partial melting by impact and formation of large partial melt pools on small asteroids is impossible (Keil et al., 1997). (b) Considerable uncertainties exist in relating speci c meteorite types to asteroidal spectral types. The classic controversy of the S asteroids is a good example: Are the surfaces of the abundant S asteroids altered (space

c 2000 Elsevier Science Ltd. All rights reserved. 0032-0633/00/$ - see front matter PII: S 0 0 3 2 - 0 6 3 3 ( 0 0 ) 0 0 0 5 4 - 4

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K. Keil / Planetary and Space Science 48 (2000) 887–903

Table 1 Number of asteroidal meteorite parent bodies (∼135) represented in the world’s meteorite collections, based on known meteorite groups, grouplets, and unique meteorites (modi ed after Meibom and Clark, 1999) 13 chondrite groups [enstatite chondrites of types H and L (EH, EL); ordinary chondrites of types H, L, and LL; carbonaceous chondrites of the Ivuna type (CI), the Mighei type (CM), the Renazzo type (CR), the Ornans type (CO), the Vigarano type (CV), the Karoonda type (CK), and the type (CH); and the Rumuruti type (R)]. 14 chondrite grouplets=unique chondrites [e.g., Kagangari type (K); low-FeO ordinary chondrites]. 11 dierentiated groups [howardites-eucrites-diogenites (HED); Main Group pallasites; Eagle Station Trio pallasites; mesosiderites; ureilites; aubrites; brachinites; winonaites-IAB-IIICD irons; acapulcoites-lodranites; angrites; IIE irons]. 12 dierentiated ungrouped=unique meteorites (e.g., pyroxene pallasites; Divnoe). 10 dierentiated magmatic iron meteorite groups (e.g., IIAB; IIIAB; IVA). ∼ 75 ungrouped=unique irons (dierentiated).

weathered) and are some of these asteroids the parent bodies of the ordinary chondrites (the most common meteorites falling on Earth)? Alternatively, are they mostly dierentiated asteroids and the parents of the relatively rare stony-iron meteorites? (e.g. Bell et al., 1989; Chapman, 1996 and references therein). (c) Although there are meteorites from ∼135 dierent parent bodies in the world’s collections, the sampling of the asteroidal belt by meteorites is nevertheless incomplete. For example, there are no meteorites in the collections with the spectral properties of the D and P asteroids (e.g. Gaey et al., 1989). (Note that there are also meteorites in our collections such as the angrites with distinct spectral properties that have not been observed in the asteroid belt.) (d) Many of the ∼135 parent bodies of which there are samples in the collections are also incompletely sampled. For example, although there are numerous meteorites from the metallic cores of 10 dierent parent bodies (in the form of the 10 magmatic iron meteorite groups; e.g. Scott and Wasson, 1975), there are no meteorites representing the silicate mantles and crusts of these differentiated asteroids. Furthermore, the aubrites, ureilites, acapulcoites-lodranites, and brachinites are ultrama c, partial melt residues from at least four partially melted parent bodies, but there are no equivalent basaltic meteorites in the collections that should have crystallized from the rst partial silicate melts that formed by melting and dierentiation of these bodies. While the basaltic melts could have been lost into space by explosive pyroclastic volcanism (e.g. Wilson and Keil, 1991), it is not clear if the parent

bodies of these meteorites experienced suciently high degrees of partial melting to form metallic cores (as is possibly the case for the aubrite parent body; Casanova et al., 1993), or whether their partial melting was arrested at relatively low temperatures. Therefore, the assignment of meteorites and their parent bodies to categories of thermal alteration throughout this paper suers from these uncertainties in sampling. Here, the thermal alterations experienced by asteroids are summarized, based on evidence derived from studies of selected examples of major meteorite groups and a few important grouplets. Note that the focus of this paper is on alterations that can be ascribed to internal heating processes on asteroids and excludes alteration effects due to shock impact heating (e.g. Scott et al., 1989; Stoer et al., 1991; Keil et al., 1997; Rubin et al., 1997). In spite of the uncertainties mentioned above, it is evident from decades of studies of meteorites that virtually all parent bodies of which we have samples have been thermally altered to some degree. These alterations range from low-temperature, (“aqueous”) alteration (∼0 to ◦ ◦ ¡ ∼300 C) through metamorphism (¿ ∼400 to ∼950 C ◦ for parent body metamorphism; up to 1500 C if metasomatic replacement reactions took place in the solar nebula), to partial melting and formation of unfractionated ◦ ◦ (∼980 to 1050 C) and fractionated (∼1000 to 1250 C) partial melt residues, to partial to complete melting, dierentiation and fractional crystallization (∼1150 to ◦ 1250 C) of asteroids. Major results and speculations regarding the igneous=volcanic processes on 4 Vesta, a dierentiated asteroid that can be viewed as the smallest of the terrestrial planets, will be brie y summarized. The heat source, and the chronological history of asteroid heating, thermal alteration, and igneous processing will also be discussed. 2. Low-temperature (“aqueous”) alteration of asteroids ◦ (∼0 to ¡ ∼300 C) That some primitive chondritic meteorites have been aqueously altered on their parent bodies was suggested a long time ago (e.g. DuFresne and Anders, 1962). In the past decade, numerous studies have shown that aqueous alteration was indeed a widespread process that affected a number of meteorite types, including carbonaceous (e.g. CI, CM, CO, CV, CR) and ordinary chondrites (e.g. Zolensky and McSween, 1988; Brearley and Jones, 1998, and references therein). This aqueous alteration resulted in the formation of hydrated phyllosilicates (e.g. saponite, smectite, phlogopite, serpentine, vermiculite) and biopyriboles, amphibole, talc, tochilinite and magnetite (e.g. Zolensky and McSween, 1988; Tomeoka and Buseck, 1990; Brearley, 1997a, b; Krot et al., 1998a) and a host of carbonates and hydrated sulfates which sometimes occur in veins (e.g. Richardson 1978; Bunch and

K. Keil / Planetary and Space Science 48 (2000) 887–903

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Table 2 ◦ Temperature estimates (in C) for low-temperature (“aqueous”) alteration of the CI, CM, CO, CV3, CR, and ordinary chondrite (H, L, LL) meteorites. Suggested locations of alterations on the asteroidal parent bodies (APB) or in the solar nebula (S) are indicated Temperature

Location

References

APB APB APB APB APB

DuFresne and Anders (1962) Zolensky et al. (1993) Hayatsu and Anders (1981) Bunch and Chang (1980) Clayton and Mayeda (1984, 1999)

APB APB APB S APB APB APB APB

Clayton (1993); Clayton and Mayeda (1999) McSween (1999) DuFresne and Anders (1962) Ikeda and Prinz (1993) Zolensky et al. (1993) Hayatsu and Anders (1981) Bunch and Chang (1980) Zolensky (1984)

APB

Zolensky et al. (1993)

S APB S APB APB

Hashimoto and Grossman (1987) Keller and Buseck (1990); Keller et al. (1994) Keller and Buseck (1991) Lee et al. (1996) Krot et al. (1998a, b)

S APB

Ichikawa and Ikeda (1995) Zolensky et al. (1993)

APB

Alexander et al. (1989)

◦

CI (20–150 ) ◦ 20 ◦ 50 –150 ◦ 85 ◦ ¡ 125 ◦ 150 ◦

CM (∼0 to ¡ 170 ) ◦ ∼0–25 ◦ ¡ 20 ◦ 20 ◦ ∼25 to ∼225 ◦ ¡ 50 ◦ 105 –125 ◦ ¡ 125 ◦ ¡ 170 ◦

CO (¡ 50 ) ◦ ¡ 50 ◦

CV 3 (¡ 100 to ¡ ∼300 ) ◦ ∼50 to ∼200 ◦ ¡ 100 ◦ ¡ ∼125 ◦ ¡ 150 ◦ ¡ ∼300 ◦

CR (∼25 to ∼225 ) ◦ ∼25 to ∼225 ◦ ¡ 150 ◦

H; L and LL ordinary chondrites (¡ ∼260 C) ◦ ¡ ∼260

Chang, 1980; Zolensky and McSween, 1988; Lee, 1993; Endreand Bischo, 1996). There is widespread agreement that aqueous alteration of CIs took place on their parent body (e.g. Bischo, 1998). However, for other meteorites, there is no universal agreement on where this alteration took place (solar nebula or on the parent body) and, for some meteorite types such as the CV3 chondrites, models for their alteration and the formation of secondary alteration products (many of which are anhydrous) are very complex. For example, for the Allende-like oxidized CV3 chondrites, it has been suggested that evaporation in the solar nebula of earlier condensed materials in regions with enhanced ◦ dust=gas ratios produced hot (1400–1500 C), oxidized gas that reacted with chondrules and calcium–aluminum-rich inclusions (CAIs) to form secondary phases by metasomatic alteration (e.g. Kurat et al., 1989; Palme and Fegley, 1990; Ikeda and Kimura, 1995; Weisberg and Prinz, 1998). The asteroidal scenario suggests that the oxidized CV3s were aqueously altered to various degrees on their parent body. Subsequent metamorphism caused dehydra-

tion of water-bearing phases and formation of secondary, water-free phases, possibly in the presence of aqueous solutions (e.g. Krot et al., 1995, 1998a, b, 1999, and references therein). Selected temperature estimates for the solar nebula replacement reactions have been included in Table 3 with those for meteorites that were metamorphosed on their parent bodies. These estimates are much higher than those for aqueous alteration (Table 2) and are higher than those for metamorphic (Table 3) alteration of CV3s in a parent body setting. In view of these controversies, it is not surprising that the nature and physical settings of these alteration processes have been the subject of intense debate. Speci cally, these alterations, including aqueous and replacement alterations, for some meteorite types and certain of their constituents, have been interpreted by some to have taken place prior to accretion, and in many studies it is suggested that this was in the solar nebula (e.g. Grossman and Larimer, 1974; Barshey and Lewis, 1976; MacPherson et al., 1985; Peck and Wood, 1987; Hua et al., 1988; Kurat et al., 1989; Ikeda and Kimura, 1995; Kimura and

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Table 3 ◦ Temperature estimates (in C) for thermal metamorphism of enstatite, ordinary and carbonaceous chondrites. Suggested locations of alterations on the asteroidal parent bodies (APB) or in the solar nebula (S) are indicated Petrologic type

Temperature

Location

References

APB APB APB APB APB APB APB APB

Dodd (1981); Olsen and Bunch (1984); Clayton (1993); McSween (1999)

APB APB Unspeci ed

Zolensky et al. (1993) Lipschutz et al. (1999) Akai (1992)

∼300–900 b ◦ ¿ 500 ◦ ∼500–600 ◦ 500 –700 ◦ ¡ 500–600 ◦ 600 –700

Unspeci ed APB APB APB Unspeci ed APB

Akai (1992) Tomeoka et al. (1989) Lipschutz et al. (1999) Zolensky et al. (1993) Akai (1988, 1990) Paul and Lipschutz (1990)

◦

APB APB Unspeci ed APB S S S S

Weinbruch et al. (1994) Krot et al. (1995) Fuchs (1971) Lee et al. (1996) Ikeda and Kimura (1995) Kimura and Ikeda (1995) Hua and Buseck (1995) Palme and Fegley (1990)

◦

EH; EL; H; L; LL (∼400 to ∼950 ) Type 3 Type 4 Type 5 Type 6 H6 L6 LL6 LL6

◦

∼400–600 ◦ 600 –700 ◦ 700 –750 ◦ 750 –950 ◦ 725 –742 ◦ 808–820 ◦ 800 ◦ 800 –960

Nakamuta and Motomura (1989)

McSween and Patchen (1989)

◦

CI (∼400 ) “Metamorphosed” “Metamorphosed” “Metamorphosed”

◦

∼40 ◦ 600 –700 ◦ 600 –800 a ◦

CM (∼500 to ∼700 ) “Metamorphosed” “Metamorphosed” “Metamorphosed” “Metamorphosed” “Metamorphosed” “Metamorphosed”

◦

◦

CV 3 (350 to ∼1500 ) “Metamorphosed” “Metamorphosed” “Metamorphosed” Vigarano Allende Allende Kaba, Mokoia Allende

∼350 ◦ 400 –500 ◦ ¿ 400–450 ◦ 400 –500 ◦ 600 –800 ◦ ¡ 830 ◦ ¿ 800 ◦ ∼1400–1500

◦

CK (590–630 ) ◦ APB Type 4 590 –630 a Lower temperature for long duration and higher temperature for shorter duration metamorphism. b Depending on speci c meteorite and duration of metamorphism.

Ikeda, 1995; Hua and Buseck, 1995; Weisberg and Prinz, 1998, and references therein). Others have suggested that these alterations took place on precursor planetesimals to the parent bodies and that the altered materials were later mixed by impact into the parent bodies (e.g. Metzler et al., 1992; Bischo, 1998). It should be noted that Wilson et al. (1999) proposed an alternative process to impact mixing. They showed that aqueous alteration in bodies originally accreting from mixtures of silicate and ice grains was commonly accompanied by substantial gas production. Internal pressurization of asteroids by these gases and their sudden release could have let to partial or complete disaggregation of the body and reaccretion into second or third generation bodies, where further alteration could have been repeated until the heat source (probably 26 Al) was exhausted. Still others have suggested that these alterations took place prior to accretion into the speci c parent body, but the location(s)

Clayton and Mayeda (1999)

where these prior alterations are to have taken place are not speci ed (e.g. Kimura and Ikeda, 1998, and references therein). Finally, extensive evidence has been presented that these alterations actually took place on the parent bodies themselves (e.g. McSween, 1977; Richardson, 1978; Kerridge et al., 1979; Bunch and Chang, 1980; Housley and Cirlin, 1983; Tomeoka and Buseck, 1985, 1990; Keller and Buseck, 1990; Kojima et al., 1993; Keller et al., 1994; Krot et al., 1995, 1997a, 1998a, b, 1999; Kojima and Tomeoka, 1996; Browning et al., 1996; Lee et al., 1996; Nomura and Miyamoto, 1998) (also see abstracts in “Workshop on Parent-Body and Nebular Modi cation of Chondritic Materials, 1997”). Powerful evidence for parent body alteration comes from a comparison of the lifetime of the solar nebula (estimated to be ¡ 10 Ma with large uncertainties; Strom et al., 1993; Podosek and Cassen, 1994; Cameron, 1995), with that of the ages of secondary phases

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Fig. 1. 53 Mn–53 Cr evolution diagram for four fayalite (Fe2 SiO4 ) grains in three porphyritic olivine–pyroxene chondrules #s 6, 8 and 9 from the oxidized CV3 chondrite Mokoia. The slope of the line yields an initial 53 Mn abundance at the time of fayalite formation of (53 Mn= 55 Mn)0 = 2:32 (±0:18) × 10−6 : 53 Cr= per mil deviation from the 53 Cr= 52 Cr ratio measured in a terrestrial fayalite standard. Errors are 2 standard deviations of the mean. Reprinted with permission from Hutcheon et al. (1998). Copyright 1998, American Association for the Advancement of Science.

in carbonaceous chondrites, relative to primitive calcium– aluminum-rich inclusions (CAIs). If these ages were to exceed the solar nebula lifetime, then these phases most likely did not form in the nebula but after accretion on the parent body. For example, these ages have been determined by short-lived radionuclides such as 53 Mn, which decays with a half-life of 3.7 Ma to 53 Cr. This method has been applied to dating of fayalite grains in the carbonaceous chondrite Mokoia (a member of the Bali-like oxidized sub-group of the CV3 chondrites). These fayalite grains are thought by some to have formed in the solar ◦ nebula by high-temperature (¿ 800 C) reactions between SiO gas, released by decomposition of pyroxene, and magnetite (Hua and Buseck, 1995). Alternatively, Krot et al. (1998b) proposed that this fayalite formed by aqueous ◦ alteration–dehydration reactions at temperatures ¡ 300 C on the Mokoia parent body. Hutcheon et al. (1998) have shown that this fayalite contains up to ∼1300 per mil excess 53 Cr from the decay of 53 Mn (Fig. 1). The slope of the “isochron” line in Fig. 1 indicates an initial 53 Mn abundance corresponding to a ( 53 Mn= 55 Mn)0 = 2:32 (±0:18) × 10−6 at the time the fayalite formed. The ratio determined for the Mokoia fayalite gives an age interval after CAI formation of ∼7–14 (or 17; Birck and Allegre, 1988) Ma; this range re ects uncertainties in the solar system initial 53 Mn= 55 Mn ratio which is not well de ned (e.g. Birck and Allegre, 1988; Lugmair and Shukolyukov,

891

1998; Birck et al., 1999). This age is equal to, or exceeds, the estimated lifetime of the solar nebula and therefore strongly suggests that the fayalite and, by implication, other associated secondary phases formed during aqueous alteration. It should be noted, however, that 53 Mn–53 Cr ion microprobe ages of secondary carbonates in the complex Kaidun breccia yield ages within 1 Ma of CAIs (Hutcheon et al., 1999). This suggests an early stage of aqueous alteration and, because of the variety of lithic clasts from dierent bodies in Kaidun, that recognizable parent bodies had formed and been fragmented within 1 Ma of formation of CAIs. In Table 2 are summarized temperature estimates for low-temperature (“aqueous”) alteration of meteorites as well as suggested alteration locations. Temperature estimates for some meteorite groups such as the CM and CV3 chondrites vary considerably. This results from the use of dierent methods (e.g. oxygen isotopes; mineral stabilities) that are used to estimate these temperatures. The range of temperature estimates also results from the fact that dierent authors have used minerals that are stable in dierent temperature regimes to place limits on temperatures of alteration. As is indicated by the data in Table 2, parent body aqueous alteration of all chondritic meteorite types was a low-temperature process that apparently took place be◦ tween ∼0 and ¡ 300 C. However, ordinary chondrites are much less aqueously altered than many carbonaceous chondrites, and convincing evidence that mild aqueous alteration has taken place on the parent bodies of all ordinary chondrites (H, L, LL) has only recently been presented (Grossman et al., 2000). Early workers (Kurat, 1969; Christophe Michel-Levy, 1976) had suggested that the white matrix in the highly unequilibrated Tieschitz ordinary chondrite formed by aqueous alteration on the parent body. McSween and Labotka (1993) proposed that oxidation during metamorphism of ordinary chondrites was promoted by interaction of minerals with small amounts of water vapor as the oxidizing phase. Later studies have shown that aqueous alteration is recorded not only in Tieschitz (e.g. Hutchison et al., 1998), but also in several other highly unequilibrated ordinary chondrites such as Semarkona and Bishunpur (e.g. Krot et al., 1997b; Alexander et al., 1989; Grossman et al., 1997). Recently, “bleached” chondrules that were aected by aqueous uids during low-temperature alteration on the parent bodies have been found in all unequilibrated H, L and LL chondrites studied and, in lower abundances, in some thermally metamorphosed petrologic types 4 – 6 ordinary chondrites as well (Grossman et al., 2000). Thus, mild aqueous alteration played a role in the evolution of all of the ordinary chondrite parent bodies. The recent discovery of ancient (extraterrestrial) liquid water in inclusions within purple NaCl in the Monahans, Texas, H5 ordinary chondrite (Zolensky et al., 1998) supports these conclusions. (Since Monahans is a regolith breccia, it cannot be ruled out that the water

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was introduced into its regolith by impacts of comets or water-bearing asteroids). 3. Thermal metamorphism of asteroids (∼400 to ◦ ∼950 C) Although some researchers have suggested that metamorphism took place prior to aqueous alteration and accretion (e.g. Kimura and Ikeda, 1998), recent evidence suggests that the parent bodies accreted cold and were heated after accretion, presumably by short-lived radionuclides such as 26 Al (e.g. McSween et al., 1988; Wilson et al., 1999, and references therein). Reactions with water, presumably formed by the melting of ice, caused mild (e.g. ordinary chondrites) to strong (e.g. CIs) aqueous alteration. The only chondrites that appear to have escaped aqueous alteration are the enstatite chondrites. [El Goresy et al. (1988) suggested that rare hydrated (?) Na–Cr sul des in the EHs Qingzhen and Y-691 are of extraterrestrial origin, implying that some mild aqueous alteration had occurred on the EH parent body. This assertion is based on the tenuous evidence of the low analytical totals of the Na–Cr sul des which they attributed to the presence of water.] As temperatures increased in parent bodies, some volatiles degassed and aqueous alteration was replaced by metamorphism. At the low end of the temperature scale, this metamorphism may have occurred in the presence of aqueous uids that originated by dehydration of hydrated minerals, as suggested, for example, by Krot et al. (1995, 1998a, b, 1999) for the CV3s. In other meteorite types and their parent bodies, or at higher temperatures after water had been driven o, metamorphism was dry (e.g. OCs). This metamorphism resulted in formation of new minerals by replacement of other minerals and glass, homogenization of originally zoned minerals, re-equilibration of elements between coexisting phases, and coarsening and recrystallization of the chondritic textures, eventually resulting in the loss of recognizable chondrules. Metamorphic series were rst recognized in case of the H, L and LL ordinary chondrite parent bodies. These chondrites were divided into petrologic types 3– 6, with type 3 being the least and type 6 the most highly metamorphosed rocks (Van Schmus and Wood, 1967). Later, the least metamorphosed chondrites of type 3 were subdivided into subtypes ranging from type 3.0 to 3.9 (e.g. Huss et al., 1981; Sears and Hasan, 1987). Peak temperature estimates ◦ ◦ ◦ ◦ range from 400 –600 C for type 3 to 750 –950 C for type 6 chondrites, based on mineral compositional (e.g. Dodd, 1981; Olsen and Bunch, 1984; McSween and Patchen, 1989; Nakamuta and Motomura, 1999) and oxygen isotopic data (Clayton, 1993) (Table 3). That enstatite chondrites were thermally metamorphosed was also suggested early (e.g. Anders, 1964; Van Schmus and Wood, 1967; Keil, 1968). However, only discovery

in Antarctica of several dozens of new enstatite chondrites has provided the data base for con rmation of petrologic types 3–5 among EH and types 3– 6 among EL chondrites (e.g. Zhang et al., 1995; Zhang and Sears, 1996, Rubin, 1997). Clearly, the EH and EL parent bodies experienced metamorphism under very dry and reducing conditions (e.g. Keil, 1989; Fogel et al., 1989), but temperature estimates are less certain than those for ordinary chondrites. This is due to the lack of suitable geothermometers (e.g. Dodd, 1981); variable re-equilibration of dierent minerals during cooling (e.g. metal, silicates, sul des), used in calculations of metamorphic temperatures with the kamacite– quartz–enstatite–oldhamite–troilite geothermometer (e.g. Fogel et al., 1989; Zhang and Sears, 1996); uncertainties in FeO contents of nearly FeO-free enstatite due to secondary uorescence of Fe X-rays from near-by metal (Wasson et al., 1994); and the fact that multiple thermal events (nebular heating, parent body metamorphism, shock) may be recorded in single meteorites (Zhang et al., 1995). In fact, many enstatite chondrites such as Abee and RKPA80259 have been altered by impact melting and may owe their petrologic type 4 – 6 characteristics to this process rather than internally driven metamorphism (Rubin and Scott, 1997). Some authors list similar temperatures for the petrologic types of EH and EL chondrites as is suggested for the ordinary chondrites (e.g. McSween, 1999), although they realize that this is a simpli cation and that these estimates are uncertain (McSween, personal communication, 1999). Recent temperature estimates for EHs using the alabandite-niningerite, metal-phosphide and sphalerite geothermometers increase from 300 – 400 to 600 ◦ –800 C with petrographic indications of increasing metamorphism, not drastically dierent from estimates for ordinary chondrites (Zhang and Sears, 1996, and references therein). For simplicity, and with awareness of the uncertainties in the estimates of enstatite chondrite metamorphic temperatures, Table 3 follows McSween (1999) and lists the same temperatures for enstatite as for ordinary chondrites. Table 3 also lists selected temperature estimates for metamorphism of the CI, CM, CV3 and CK meteorites. The solar nebular setting proposed by a number of authors (e.g. Weinbruch et al., 1990; Palme and Fegley, 1990; Ikeda and Kimura, 1995) for the origin of replacement minerals in the Allende CV3 chondrite requires very high ◦ temperatures, ranging from 600 to 1500 C. However, a scenario in which metamorphism took place on the Allende parent body would require much lower tempera◦ tures of ∼350–700 C. It should be noted that there are some uncertainties in classifying some of these metamorphosed carbonaceous chondrites. For example, on petrographic grounds, Akai (1988) and Zolensky et al. (1993) classi ed B7904 and Y86720 (for which they estimate ◦ metamorphic temperatures of 500 –700 C), as having been derived from CM precursors. On the other hand, Clayton and Mayeda (1999), based on oxygen isotopes, classify

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these meteorites as having been derived from CI precursors. 4. Partial melting of asteroids and formation of ◦ unfractionated (∼980–1050 C) and fractionated ◦ (∼1000 to ¿ 1250 C) partial melt residues Here are summarized the properties of meteorites from asteroids that experienced partial melting, but the meteoritic record is incomplete and only partial melt residues (unfractionated and fractionated) are preserved. No samples exist of the rocks that should have crystallized from the partial melts. In this category belong meteorites from three parent bodies, the acapulcoites–lodranites, brachinites, and ureilites. The acapulcoites–lodranites have roughly chondritic composition but highly recrystallized, achondritic textures. They were therefore referred to as primitive (i.e. in composition) achondrites (i.e. in texture) (Prinz et al., 1983). Detailed petrologic, chemical and isotopic studies show that the acapulcoites are partial melt residues having resulted from low degrees (∼1–3 vol%) of partial melting of a chondritic precursor material in the tempera◦ ture range of ∼980–1050 C (McCoy et al., 1996). In this range, only partial melting of Fe,Ni–FeS but not of silicates took place and, due to the low degree of partial melting, Fe,Ni–FeS melts moved only over very short distances (mm=cm scale). Thus, acapulcoites have retained their primitive, chondritic compositions in spite of temperatures at or slightly above the Fe,Ni–FeS cotectic. Similar oxygen isotopic compositions between lodranites and acapulcoites are consistent with an origin from the same parent body (McCoy et al., 1996, 1997b). Lodranites are also partial melt residues, but are the products of higher ◦ (∼5–15 vol%) partial melting at higher (∼1150–1200 C) temperatures (McCoy et al., 1997b). This resulted in the production of Fe,Ni–FeS and silicate (plagioclase-pyroxene, i.e. basaltic) partial melts and, because of the higher degree of partial melting, these partial melts migrated more extensively through the parent body. Thus, lodranites are fractionated partial melt residues with fractionated compositions when compared to primitive, unfractionated chondritic compositions. The rocks that should have crystallized from the silicate partial melts (i.e. basalts) are unknown, probably because these melts were ejected o the parent body by explosive pyroclastic volcanism early in the history of the solar system and were lost into space (Wilson and Keil, 1991; McCoy et al., 1997a). Brachinites are unfractionated or slightly fractionated partial melt residues, possibly derived from carbonaceous chondrite precursors, that resulted from relatively low degrees (1–10 vol%) of partial melting (Nehru et al., 1996; Goodrich, 1998). The minimum temperatures of partial melting, based on closure temperatures of a number of ◦ mineral geothermometers, vary between 800 and 1246 C

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(Nehru et al., 1996). Since some brachinites lost their basaltic (plagioclase-pyroxene) component, they must have ◦ been heated to at least 1050–1200 C. Again, basaltic rocks that should have crystallized from these partial melts are not known, and these melts were probably lost into space by explosive volcanism (Nehru et al., 1996). Ureilites are ultrama c rocks and formed either as partial melt residues or as cumulates from a highly evolved parent body (e.g. Goodrich, 1992, and references therein). It has been suggested that these rocks were derived from a chondritic precursor by 10 –30 vol% partial melting at ◦ temperatures ¿ ∼1250 C (references in Goodrich, 1992). This is a minimum temperature estimate based on the equilibrium closure temperature of pyroxene geothermometers. As with all parent bodies and their meteorites described in this chapter, there are no basaltic rocks known that should have crystallized from the partial melts, prompting the suggestion that these melts were lost into space by explosive volcanism (Scott et al., 1993). ◦

5. Partial to complete melting (∼1150 to 1250 C), dierentiation and fractional crystallization of asteroids Here, examples are discussed of meteorites from asteroids that generally experienced relatively high degrees of partial melting, including the angrites, mesosiderites, aubrites, winonaites=IAB-IIICD irons, pallasites, and the large number of iron meteorite groups, grouplets and ungrouped, unique irons. The HED meteorites that include the most abundant asteroidal basalts (eucrites) are discussed in the following section. The angrites are a grouplet of four members that are roughly basaltic in composition but, unlike terrestrial and lunar basalts, are depleted in alkalis and are high in Ca and Al (e.g. Mittlefehldt et al., 1998, and references therein). Melting experiments by Jurewicz et al. (1993) suggest that angrites may have crystallized from partial melts that ◦ formed at ∼1200 C by ∼20 vol% partial melting of CM or CV precursor materials. Note, however, that dierences in the 54 Cr= 52 Cr ratios of carbonaceous chondrites and angrites suggest that the latter did not form by partial melting of the former (Shukolyukov and Lugmair, 1998; Shukolyukov et al., 1999). No other rocks exist that can be attributed to the angrite parent body. The mesosiderites are stony-iron breccias of variable metamorphic grade that consist of clasts or veins of metallic Fe,Ni and clasts of basaltic, gabbroic, pyroxenitic and minor dunitic rocks. Metallographic cooling rates for the metal phase are amongst the slowest known and, for the ◦ temperature interval of ∼500–350 C, vary between ∼0:01 ◦ and 1 C=Ma (references in Mittlefehldt et al., 1998). Cooling rate estimates for the silicate lithologies vary depending upon the method and applicable temperature inter◦ ◦ ◦ val and are ¿ ∼100 C=Ka at 1000 C and 6 1 C=Ma ◦ at 250 C. Thus, mesosiderites appear to be mixtures of

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core (metal) and crustal=surface igneous lithologies (silicate clasts). However, the silicate clasts which, based on their compositions and textures, must be of crustal=surface origin, were also deeply buried to account for the very slow cooling at low temperatures. Various models have been proposed for the origin of these enigmatic rocks. All models involve major impact mixing (and possibly collisional break-up followed by gravitational reassembly) of one or two highly dierentiated parent bodies that had experienced high degrees of partial melting and formation of a metal core (possibly still liquid when impact mixing occurred) (references in Mittlefehldt et al., 1998). The aubrites are highly reduced, brecciated enstatite pyroxenites (e.g. Keil, 1989) and consist mostly of pyroxenitic clasts with irregular intergrown grain boundaries, suggestive of plutonic co-crystallization (Okada et al., 1988). They are clearly related to some enstatite chondritelike precursor material and formed on a parent body that experienced high degrees of partial melting at high tem◦ peratures (up to ∼1500 C; McCoy et al., 1999), which may have resulted in the formation of a metal core (e.g. Hewins and Newsom, 1988). Aubrites contain Si-bearing metallic Fe,Ni nodules similar in composition to the Horse Creek iron, a meteorite that may be from the same parent body as the aubrites (Casanova et al., 1993). Chondritic abundances of trace elements in the metal nodules and Horse Creek suggest that these objects were not part of a central, fractionally crystallized core of the parent body but, rather, are metal nodules that never segregated into a central core (Casanova et al., 1993). Thus, neither basaltic rocks (Wilson and Keil, 1991) nor fragments of the core of the highly dierentiated aubrite parent body are represented among known meteorites. The only known unbrecciated aubrite (Shallowater) appears to be a sample from a second, highly dierentiated aubrite parent body and is thought to have formed by the impact of a solid, enstatite–chondrite-like asteroid with a totally to partly molten aubrite asteroid (Keil et al., 1989). The winonaites are unfractionated to slightly fractionated partial melt residues of a chondritic precursor ma◦ terial and resulted from heating to ¡ ∼1200 C and low degrees of partial melting (e.g. Benedix et al., 1998). Similarities in oxygen isotopes and other properties of winonaites to silicate inclusions in IAB and IIICD irons suggest that the IAB irons and, with less certainty, the IIICD irons, may be from the same parent body as the winonaites (e.g. Benedix et al., 1998, 2000). The association of the winonaites=IAB-IIICD irons suggest that their parent body must have experienced high degrees of partial melting at ◦ temperatures ¿ 1200 C. A host of models have been proposed for the origin of these meteorites and the history of their parent body, most of which require extensive partial melting (e.g. Middlefehldt et al., 1998, and references therein). The three pallasite types (Main Group, Eagle Station, and pyroxene pallasites) probably represent three highly

dierentiated parent bodies (Meibom and Clark, 1999), although the pyroxene pallasite grouplet is populated by only two meteorites which have not been studied in great detail [note that Mittlefehldt (1999), based on minor element compositions of olivine, suggested that the Springwater Main Group pallasite may possibly represent a fourth pallasite group and a fourth pallasite parent body]. The Main Group pallasites and IIIAB irons apparently formed on the same parent body (Scott, 1977), and high degrees ◦ of partial melting of 45 –90 vol% at ¿ 1200 C are required for their formation (Takahashi, 1983; Taylor, 1992; Taylor et al., 1993). In fact, the required high degree of partial melting has prompted Taylor et al. (1993) to suggest that the pallasite parent bodies had magma oceans. There are 11 major iron meteorite groups known besides the IAB-IIICD irons. The IIE irons contain complex silicate inclusions, and suggestions for their origin are as diverse as those for the IAB-IIICD irons, but major partial melting of their parent body was clearly involved (e.g. Mittlefehldt et al., 1998). The remaining 10 iron meteorite groups (Table 1) show siderophile trace element trends compatible with fractional crystallization in the cores of 10 distinct parent bodies. Depending upon the assumed composition of the precursor material (especially sulfur concentration), core formation in these 10 asteroidal-sized bodies requires high degrees of partial melting of ∼40– ◦ 60 vol% at temperatures of ∼1350–1500 C (Taylor, 1992; Taylor et al., 1993). In addition, there are about 75 iron meteorite grouplets, ungrouped, and unique iron meteorites known that do not fall in the groups mentioned above (Wasson, 1995). These irons are interpreted to each represent a separate parent body, and most show evidence for fractional crystallization (Scott, 1979). Therefore, they must once have been part of fractionally crystallizing asteroidal cores and originated on parent bodies that must have experienced high temperatures that resulted in high degrees of partial melting and fractionation. 6. Asteroid 4 Vesta, the “smallest terrestrial planet” Asteroid 4 Vesta is a fascinating asteroid: it is a dierentiated object with a basaltic crust and exposed (impactexcavated) ma c (pyroxenitic; olivine-bearing) mantle material (Binzel et al., 1997; Gaey, 1997) that survived essentially intact the 4.57 Ga of solar system history. It most likely also has a metallic core (e.g. Righter and Drake, 1997). The asteroid can therefore be thought of as the smallest of the terrestrial planets [an ellipsoid of semi-axes of 289, 280, and 229 (±5) km (Thomas et al., 1997)]. Asteroid 4 Vesta is also unique because similarities in its re ectance spectra to laboratory spectra of the numerous HED meteorites (e.g. McCord et al., 1970) suggest that the HEDs are actually samples of this asteroid (Binzel and Xu, 1993). Study of the HEDs and of 4 Vesta

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have shed light on the complex melting, dierentiation and fractionation processes that took place on this and other dierentiated asteroids. It also has relevance to the evolution of dierentiated, asteroidal-sized bodies that are thought to have been the building blocks of the terrestrial planets (e.g. Taylor and Norman, 1990). The well-studied HEDs represent a large suite of crustal rocks, with the eucrites being basalts and cumulate gabbros. The diogenites are (usually) brecciated, coarse-grained orthopyroxenites, and the howardites are regolith breccias composed of mixtures of mostly eucritic and diogenitic materials. Regolith breccias contain solar wind-implanted noble gases and, thus, some of their constituents must have once resided at the very surface of the asteroid. It should be noted that some eucrites and diogenites are also polymict breccias (e.g. Mittlefehldt et al., 1998, and references therein). Extensive laboratory studies of HEDs by a host of analytical techniques, and modeling studies of igneous=volcanic and core formation processes, combined with ground and space-based telescopic studies of the asteroid, have allowed development of detailed hypotheses of various aspects of the geological evolution of 4 Vesta (e.g. papers and references in Meteorit. Planet. Sci. 32, 813–980, 1997). In addition, although there are similarities in the compositions and textures of the rocks that resulted from igneous dierentiation on small asteroidal-sized bodies to those of the terrestrial planets and the Moon, there are signi cant dierences in the physical processes of magmatism-volcanism. Below are summarized selected important results and hypotheses regarding the igneous= volcanic processes on 4 Vesta and small asteroidal-sized bodies in general. 6.1. Petrogenesis of HEDs Numerous hypotheses have been suggested for the petrogenesis of the HEDs and, thus, for the thermal history of their presumed parent body, 4 Vesta (see references in Mittlefehldt et al., 1998). For example, Stolper (1977) carried out melting experiments with basaltic eucrites and concluded that these rocks represent primary partial melts formed by ∼4–15 vol% partial melting in the temperature ◦ range of ∼1150–1190 C of a chondritic (CM) source material (e.g. Jurewicz et al., 1993). The cumulate eucrites are thought to have crystallized from fractionally crystallizing ma c melts, and the diogenites are pyroxenite cumulates from a fractionally crystallizing magma. The model of Stolper (1977) requires relatively low degrees of partial melting of the parent body. Other models suggest high degrees of partial melting of the HED parent body and formation of a magma ocean (e.g. Ikeda and Takeda, 1985; Taylor et al., 1993; Ruzicka et al., 1997), including formation of a metal core (e.g. Righter and Drake, 1997). In many of these models, basaltic eucrites are not

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considered to be primary partial melts, but the products of fractionally crystallized residual melts (e.g. Ruzicka et al., 1997; Righter and Drake, 1997). 6.2. No tectonically controlled magmatism On small parent bodies, the complex processes of heating, melting, dierentiation, and crystallization=solidi cation of igneous rocks took place in the rst few million years of solar system history [e.g. Lugmair and Shukolyukov, 1998; and discussion in Section 7 below]. As a result, tectonically controlled magmatism, which is prevalent on Earth and requires long time periods of melting and dierentiation, is absent on asteroidal parent bodies (e.g. Scott et al., 1989). 6.3. Basalts present only on parent bodies ¿ ∼100 km in radius As mentioned above, partial melting of chondritic precursor materials of dierentiated parent bodies will pro◦ duce Fe,Ni–FeS partial melts at ∼980–1050 C and pyroxene-plagioclase (basaltic) partial melts at ∼1000– ◦ 1250 C, depending on the exact composition of the starting material. On parent bodies ¡ 100 km in radius, these low-density basaltic partial melts, upon extrusion to the asteroid surfaces and exposure to the vacuum of space, would be accelerated by explosive pyroclastic volcanism to velocities greater than the escape velocities of these small parent bodies, provided the melts contain a few hundred ppm of volatiles (Wilson and Keil, 1991). As a result, the small pyroclasts, estimated to range in size from ∼30 to 4000 m (Wilson and Keil, 1996a), would be accelerated into space and lost by spiraling into the sun early in the history of the solar system. Thus, on such small parent bodies, basalts are not expected to be present (Fig. 2). This is the case for the parent bodies of the aubrites (Wilson and Keil, 1991; Muenow et al., 1992, 1995), the ureilites (Warren and Kallemeyn 1992; Scott et al., 1993), the acapulcoites-lodranites (McCoy et al., 1996, 1997a, b), and brachinites. In fact, it has also been suggested that, under favorable conditions, small parent bodies could lose part of their high-density Fe,Ni–FeS partial melts, thus explaining the magmatic iron meteorite groups whose magmas appear to have been depleted in sulfur (Keil and Wilson, 1993). However, on larger parent bodies, unrealistically high volatile contents of the melts would be required to accelerate the pyroclast-gas mixtures into space. In case of 4 Vesta, a magma gas content ncrit of ¿ 3:8 wt% would be required if eruption velocities of pyroclasts are to exceed the escape velocity of ∼380 m=s (Fig. 2). Such high volatile contents are unknown from the volatile-rich terrestrial magmas, let alone from the “dry” magmas of 4 Vesta (e.g. Mittlefehldt, 1987). As a result, 4 Vesta is covered by basalts and the eucrites are thought

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5 –20 m. Deep intrusions are estimated to have had lateral extents of ¡ 30 km and thicknesses of ¡ 3 m, whereas shallow intrusions are estimated to have had vertical extents of ¡ 10 km and widths of ∼1 m. Dikes that reached the surface would have fed eruptions that lasted ∼8– 60 h, with eusion rates of between ∼0:05 and ¿ 3m3 =s=m from ssure vents (this range is similar to that of current basaltic eruptions on Earth). 6.5. No pyroclastic deposits

Fig. 2. Values of the escape speed, V , for asteroids of radii, R, assuming a density of 3500 kg=m3 . Also shown as a function of R are the critical values of magma gas content, ncrit , at which the eruption velocity, U , equals the escape speed, V . For the approximate radius of 265 km for asteroid 4 Vesta, an unrealistically high magma gas content ncrit of ¿ 3:8 wt% would be required to overcome the escape velocity of ∼380 m=s to eject into space basaltic pyroclast=gas mixtures. Thus, 4 Vesta is suciently large to retain basalts on its surface. Modi ed from Wilson and Keil (1991). Reprinted from Earth Planet. Sci. Letters, copyright 1991, with permission of Elsevier Science Publishers.

to be impact-ejected meteoritic fragments of these basalts (Binzel and Xu, 1993). There are two other asteroidal parent bodies that have retained basalts, namely those of the mesosiderites and angrites. Based on the work of Wilson and Keil (1991), one would predict that these bodies also must have had radii ¿ 100 km in order to prevent loss of basaltic melts into space due to explosive pyroclastic volcanism (alternatively, basaltic melts would also have been retained if the magmas were essentially free of volatiles, which is not likely in view of the compositions of the mesosiderites and angrites). It is interesting to note that Haack et al. (1996), based on metallographic cooling rates and thermal models, concluded that the mesosiderite parent body must have been ∼200– 400 km in radius. 6.4. Nature of intrusions and extrusions Wilson and Keil (1996b) modeled intrusive and eruptive volcanic processes as a result of large-scale partial melting on the low-gravity asteroid 4 Vesta. These studies have allowed estimation of the dimensions of some volcanic features ( ows, dikes, etc.) and allow comparison of these processes and features to those on a large, high-gravity planet like Earth. These results indicate, for example, that basaltic surface lava ows on 4 Vesta have had widths of a few hundred m to several km, lengths between a few km to several tens of km, and thicknesses of

Eruptions of volatile-bearing magmas to the surfaces of Earth and Moon have resulted in the formation of widespread pyroclastic deposits (e.g. Wilson and Head, 1981, 1983). If ascending basaltic magmas on 4 Vesta contained any volatiles at all, pyroclastic deposits should be present on the asteroid and, in the form of glass spherules and shards, be recognizable in HED meteorites. However, extensive thin section searches in HEDs for such pyroclasts by Yamaguchi (personal communication, 1995) did not discover any pyroclasts. Thus, either magmas on 4 Vesta were totally devoid of dissolved volatiles, or the physics of pyroclastic eruptions on a body with the physical properties of 4 Vesta (e.g. gravitational and rock mechanical forces; lack of atmosphere) are suciently different to severely aect the nature of pyroclastic eruptions. In fact, Wilson and Keil (1997) have shown that lava fountains on 4 Vesta are extremely optically dense and, hence, pyroclast droplets do not “see the sky” and do not cool o during ight (Fig. 3). As a result, pyroclastic droplets will land while molten to form lava lakes that will feed lava ows, which will crystallize basalts that are indistinguishable from those that formed by extrusion of magmas. For example, for a magma discharge rate from a ssure vent of 0:3 m3 =s=m and a released magma gas content of, say, 100 ppm (Table 4), the eruption speed would be 20:1 m=s; the maximum droplet range would be 1550 m; the thickness of the outer layer of the lava fountain from within which pyroclastic droplets would cool before landing would be 8.3 m; and the fractional area of the fountain within which the droplets would not cool before landing would be 99.5%. Thus, very few pyroclasts would be deposited and preserved, unlike on Earth and Moon, and in agreement with evidence from HED meteorites. 6.6. Global metamorphism of eucrites A number of models have been proposed to explain why most eucrite meteorites have been annealed, an important topic because it has relevance to the thermal history of 4 Vesta. Takeda (1997), for example, proposed a layered crustal model, where the degree of annealing is related to the depth of origin of the rocks in the crust. To explain some thermally annealed eucrites that do not t into this sequence, Takeda and Graham (1991) suggested that

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Fig. 3. Lava fountain from the August, 1984 eruption of Pu’u ‘O’o on the island of Hawaii, as a qualitative example of lava fountaining on the asteroid 4 Vesta (courtesy of Peter Mouginis-Mark). The fountain is approximately 200 m high. The interior, red-hot pyroclast/gas mixtures remain hot during the entire ight and, upon landing, these will form lava ponds that will feed lava ows. The basalts crystallizing from these ows are indistinguishable from those that formed by extrusion of magma. Only the outer, dark, cool pyroclasts will be deposited as solids to form pyroclastic deposits. Unlike on Earth and Moon, the fractional area on 4 Vesta within which landing droplets did not cool, is ¿ 99% for reasonable magma gas contents (see Table 4). This explains the lack of pyroclastic materials in HED meteorites.

Table 4 For a magma discharge rate from a ssure vent of 0:3 m3 =s=m and released magma gas contents n, in ppm, are given the gas-droplet eruption speed ug , in m=s; the maximum droplet range D, in m; the thickness of the outer layer , in m, of the fountain from within which pyroclastic droplets cool; and the fractional area F, in %, within which landing droplets did not cool and will coalesce to form lava ponds that will feed lava ows indistinguishable from those that formed by extrusion (modi ed after Wilson and Keil, 1997) n (ppm) 10 30 100 300 1000

ug (m=s)

D (m)

 (m)

F (%)

10.0 12.5 20.1 33.9 61.6

380 600 1550 4400 15000

1.0 2.0 8.3 40.0 240

99.7 99.7 99.5 99.1 98.4

heating by impact or burial by later ows could be responsible for this thermal metamorphism. However, Keil et al. (1997) showed that impact heating is not a viable process for global metamorphism of the eucrites, and Yamaguchi et al. (1996, 1997) argued that the omnipresent thermal metamorphism of the eucrites could not be the result of contact metamorphism by burial below younger

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Fig. 4. Schematic diagram of the model by Yamaguchi et al. (1997) for the development of the basaltic crust of 4 Vesta, resulting in abundant metamorphosed eucrites. Plotted are the depth of basalts as a function of time. Assuming, for example, that volcanism lasted for 1 Ma and that the rate of eruption was constant, then a basalt erupted at the beginning of volcanism (curve A) is buried progressively deeper, ending up at the base of the crust (at 10 –25 km depth) after 1 Ma. A basalt erupted 0.5 Ma later, midway through the magmatic episode, ends up buried about half way down in the crust (curve B). Basalt A reaches a higher metamorphic temperature and cools slower than basalt B, but both would be metamorphosed. A basalt erupted near the end of volcanism would remain at or near the surface and not signi cantly heated by burial. This model predicts a wide variety of metamorphic temperatures and duration of heating (and abundant metamorphosed basalts), consistent with observations of eucrites. Reprinted with permission from Yamaguchi et al. (1997), copyright by the American Geophysical Union.

ows. Instead, these authors suggested that high eruption rates of eucritic magmas and quick burial of earlier by later ows would allow heat to be conducted upwards from the mantle through the crust that would metamorphose these basalts on a global scale (Fig. 4). The data and hypotheses summarized above illustrate the signi cance of asteroid 4 Vesta for understanding the earliest igneous processes on small, dierentiated bodies. Much is yet to be learned about this asteroid, and exploration by spacecraft of this smallest of the terrestrial planets is therefore of great scienti c importance. 7. Heat sources for thermal alteration of asteroids, and chronology of thermal events 7.1. Heat source One of the great mysteries of meteoritics has been the heat source(s) responsible for the internally-driven thermal processing of the parent bodies of meteorites. A number of processes have been proposed (see summary in Scott et al., 1989). However, in recent years discussions have concentrated mostly on electrical conduction heating by the T-Tauri solar wind from the pre-main sequence Sun (e.g. Sonett et al., 1970), or the decay of short-lived 26 Al, rst proposed by Urey (1955). This isotope decays into 26 Mg with a half-life of 0.73 Ma, and the resulting excess

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26 Mg was rst discovered in a CAI from the CV3 chondrite Allende (Lee et al., 1976). Since then, such excesses have been found in other CAIs and in chondrules from carbonaceous and ordinary chondrites (e.g. Hutcheon and Hutchison, 1989; MacPherson et al. 1995, and references therein; Srinivasan et al., 1996a, b; Russell et al., 1996, 1998; Kita et al., 1998; Mostefaoui et al., 1999). Particularly noteworthy is that excess 26 Mg has recently been detected in plagioclase of the eucrite Piplia Kalan (Srinivasan et al., 1998). These observations suggest that 26 Al was widespread in the early solar system and was most likely the heat source for thermal alteration of asteroids. By now, the decay products of about a dozen short-lived radionuclides have been discovered in meteorites (e.g. Podosek and Nichols 1997). However, if 26 Al were present in the bulk material that accreted to form asteroids in canonical abundances (26 Al= 27 Al = 5 × 10−5 , as is the case for some CAIs), all asteroidal-sized bodies would totally melt, due to the high abundance of aluminum in meteorites (e.g. Lee et al., 1976). However, since the world’s collections of meteorites contains samples of ∼27 parent bodies that did not melt (Table 1), clearly not all asteroids accreted with the canonical 26 Al= 27 Al in their bulk material (or they accreted after some of the 26 Al had already decayed).

7.2. Penecontemporaneous onset of thermal alteration of asteroids Studies of the decay products of short-lived radionuclides in meteorites have also contributed signi cantly to understanding the chronology of geological activities on asteroids. For example, it was pointed out above that 53 Mn–53 Cr systematics indicate that secondary fayalite in the parent body of the CV3 chondrite Mokoia formed in the presence of live 53 Mn (half-life 3.7 Ma) when the initial 53 Mn abundance was (53 Mn= 55 Mn)0 = 2:32 (±0:18) × 10−6 (Hutcheon et al., 1998). Similar ratios were measured for two other aqueously altered carbonaceous chondrite parent bodies, namely 2:0 × 10−6 for carbonates in the CI chondrites Ivuna and Orgueil (Endre and Bischo, 1996; Hutcheon and Phinney, 1996), and 2:5×10−6 in carbonates from the CM chondrite Nogoya (Browning and Hutcheon, preliminary, unpublished data). Thus, aqueous alteration took place penecontemporaneously on the CI, CM and CV parent bodies, roughly 7–17 Ma after the formation of CAIs. Similarly, the signatures of live 53 Mn were also found in other chondrite parent bodies (enstatite and ordinary chondrites) and CAIs (e.g. Lugmair and Shukolyukov, 1998). It was also found in meteorites whose parent bodies experienced partial or complete melting, such as the acapulcoite, brachinite, angrite, mesosiderite, Main Group pallasite, Eagle Station trio pallasite, IIIAB iron, eucrite, and the ungrouped (unique) Divnoe meteorite parent bodies (references in Lugmair and

Shukolyukov, 1998). Thus, these studies indicate that the onset of thermal alteration of asteroids, ranging from lowtemperature aqueous alteration at temperatures of ∼0 to ◦ ¡ ∼300 C to partial and complete melting at temperatures ◦ of ∼980 to 1200 C, was penecontemporaneous (within a few Ma) on many dierent asteroidal parent bodies. It should be noted, however, that the duration of thermal alteration (e.g. thermal metamorphism) often extended over a hundred Ma or longer, as is indicated by younger absolute ages of some meteorites based on long-lived radionuclides (e.g. Turner, 1988; Bogard and Garrison, 1995) and resulted from the relatively slow cooling of the more interior portions of even small asteroids (e.g. Haack et al., 1990). 7.3. Melting and dierentiation took place in the ÿrst ∼10 Ma of solar system history The mere presence of the decay products of short-lived radionuclides in meteorites is very signi cant, because it indicates that these nuclides were live at the time of formation of the carrier phases=meteorites. It also indicates that the time interval between the introduction into the solar nebula of freshly synthesized short-lived radionuclides and the formation of the carrier phases=meteorites was within a few times the half-lives of these nuclides. What is so astonishing is that the signatures of short-lived radionuclides are not only apparent in mm-sized chondrules and mm–cm-sized CAIs (including excess 41 K from the decay of 41 Ca with a half-life of only 0.10 Ma; Srinivasan et al., 1996c), but in the minerals and rocks of highly altered and even highly dierentiated asteroids as well. For example, the 53 Mn= 53 Cr systematics discussed earlier indicates that aqueous alteration of carbonaceous chondrite parent bodies took place within 7–17 Ma of formation of CAIs. Similarly, the (53 Mn= 55 Mn)0 ratios measured for meteorites from partially melted=dierentiated parent bodies of the acapulcoite, brachinite, angrite, mesosiderite, Main Group pallasite, Eagle Station trio pallasite, IIIAB iron, eucrite, and the ungrouped (unique) Divnoe meteorites, suggest that these parent bodies must have accreted and been partially=completely melted in the same time frame. The detection in eucrites of the decay products of shortlived radionuclides with shorter half-lives than 53 Mn suggest that these time intervals were even shorter in case of the HED parent body. For example, excess 60 Ni, which is the decay product of 60 Fe (half-life 1.5 Ma), has been detected in the Chervony Kut and Juvinas basaltic eucrites, suggesting that these particular extrusive basalts formed within only a few Ma of accretion and dierentiation of their parent body, presumably 4 Vesta (Shukolyukov and Lugmair, 1993a, b). Finally, evidence has recently been presented that plagioclase in the Piplia Kalan eucrite has excess 26 Mg from the decay of 26 Al (half-life 0.73 Ma),

K. Keil / Planetary and Space Science 48 (2000) 887–903

con rming that the time interval between formation of CAIs and solidi cation of this basalt was on the order of ∼5 Ma (Srinivasan et al., 1998). These data lead to the conclusion that the processes of formation of solid materials in the early solar system, accretion into asteroids, heating and partial to complete melting, dierentiation, fractionation, extrusion of partial melts and their crystallization, all took place within the rst ∼10 Ma of solar system history (relative to CAIs). If the HED meteorites indeed are impact ejecta o 4 Vesta, this smallest of the terrestrial planets accreted, was heated up, partially to completely melted, and formed extrusive basalts all within a very short time interval of a few Ma after CAI formation! Acknowledgements Discussions with H.Y. McSween, Jr. and D.W.G. Sears, and numerous helpful suggestions for improvements of the manuscript by T. Fagan, A. Meibom, A. Krot, E.R.D. Scott and G.J. Taylor and by reviewers B.E. Clark and A. Shukolyukov and an anonymous reviewer are gratefully acknowledged. I thank Peter Mouginis-Mark for providing the original photograph for Fig. 3 and Anders Meibom for help with its reproduction. This work was supported in part by NASA grant NAG 5-4212. This is Hawaii Institute of Geophysics and Planetology Publication no. 1097 and School of Ocean and Earth Science and Technology Publication no. 5078. References Anders, E., 1964. Origin, age and composition of meteorites. Space Sci. Rev. 3, 583–714. Akai, J., 1988. Incompletely-transformed serpentine-type phyllosilicates in the matrix of Antarctic CM chondrites. Geochim. Cosmochim. Acta 52, 1593–1599. Akai, J., 1990. Mineralogical evidence of heating events in Antarctic carbonaceous chondrites, Y-86720 and Y-82162. Proc. NIPR Symp. Antarct. Meteorites 3, 55–68. Akai, J., 1992. T–T–T diagram of serpentine and saponite, and estimation of metamorphic heating degree of Antarctic carbonaceous chondrites. Proc. NIPR Symp. Antarct. Meteorites 5, 120–135. Alexander, C.M.O’D., Barber, D.J., Hutchison, R., 1989. The microstructure of Semarkona and Bishunpur. Geochim. Cosmochim. Acta 53, 3045–3057. Barshey, S.S., Lewis, J.S., 1976. Chemistry of primitive solar material. Ann. Rev. Astron. Astrophys. 14, 81–94. Bell, J.F., Davis, D.R., Hartmann, W.K., Gaey, M.J., 1989. Asteroids: The big picture, In: Binzel, R.P., Gehrels, T., Matthews, M.S. (Eds.), Asteroids II. University of Arizona Press, Arizona, pp. 921–945. Benedix, G.K., McCoy, T.J., Keil, K., Bogard, D.D., Garrison, D.H., 1998. A petrologic and isotopic study of winonaites: evidence for early partial melting, brecciation, and metamorphism. Geochim. Cosmochim. Acta 62, 2535–2553. Benedix, G.K., McCoy, T.J., Keil, K., Love, S.G., 2000. A petrologic study of the IAB iron meteorites: constraints on the formation of the IAB-winonaite parent body. Meteorit. Planet. Sci. 35, in press.

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