A Critical Evaluation of the Evidence for Hot Accretion

ICARUS

124, 86–96 (1996) 0191

ARTICLE NO.

A Critical Evaluation of the Evidence for Hot Accretion ALAN E. RUBIN Institute of Geophysics and Planetary Physics, University of California, Los Angeles, California 90095-1567 E-mail: [email protected] AND

ADRIAN J. BREARLEY Institute of Meteoritics, Department of Earth and Planetary Sciences, University of New Mexico, Albuquerque, New Mexico 87131 Received October 12, 1995; revised May 13, 1996

bility of hot-accretion models in a nebular environment: The low cooling rates required of these models are inconsistent with abundant evidence which mandates rapid cooling rates for chondrules. Fine-grained, FeO-rich silicate matrix material in type-3 chondrites is relatively rich in volatile elements and could not have equilibrated along with magnesian chondrules at $1100 K in the same parent-body environment. In addition, in order to achieve sufficient heat retention on chondritic asteroids residing in an appreciably cooler nebular environment, hot-accretion scenarios require accretion rates four orders of magnitude greater than those inferred theoretically for planetary bodies. Nevertheless, a surprisingly large number of arguments has been framed that attempt to show that hot accretion of ordinary chondrites (OC) has in fact occurred. These fall into eight categories: (1) lack of correlation between petrologic type and Ca in low-Ca pyroxene, (2) pyroxene structure and composition, (3) polycrystalline taenite in equilibrated OC, (4) compositionally and texturally heterogeneous metallic Fe-Ni grains in equilibrated OC, (5) misshapen chondrules in type-3 OC, (6) the apparent igneous nature of matrix materials in type-3 OC, (7) depletion of volatile trace elements with petrologic type, and (8) natural remanent magnetization. In this manuscript we present a comprehensive evaluation of the evidence for hot accretion. As discussed below, an appraisal of these arguments finds them to be flawed; prograde thermal metamorphism is more plausible.

Although the type-3 to -6 ordinary chondrite recrystallization sequence is attributed to prograde thermal metamorphism by most researchers, various studies have inferred support for the hot-accretion alternative on the basis of the composition and fine structure of pyroxene, the composition and texture of metallic Fe–Ni, misshapen chondrules, the texture of fine-grained silicate matrix material, the depletion of volatile trace elements in chondrite whole rocks, and natural remanent magnetization. Critical assessment of these studies shows that the arguments favoring hot accretion are based largely on inadequate data, flawed experiments, or improbable interpretations. The properties of ordinary chondrites are more readily understood in terms of prograde metamorphism of initially cold, unequilibrated type-3 material and episodes of brecciation, shock heating, and, in some cases, aqueous alteration.  1996 Academic Press, Inc. INTRODUCTION

Anhydrous chondrite groups are classified into petrologic types 3–6 reflecting increasing degrees of textural recrystallization, mineral chemical equilibration, and bulk loss of volatile trace elements (Van Schmus and Wood 1967; Tandon and Wasson 1968; Lipschutz and Woolum 1988). Two models have been developed to explain this sequence: (1) prograde thermal metamorphism (e.g., Sorby 1877; Merrill 1921; Wood 1962; Van Schmus and Koffman 1967; Dodd 1969; Lux et al. 1980; McSween et al. 1988) wherein initially cold, unequilibrated material resembling a type-3 chondrite is progressively heated on a parent body, and (2) autometamorphism or hot accretion (e.g., Anders 1964; Larimer and Anders 1967; Hutchison et al. 1980; Hutchison and Bevan 1983; Watanabe et al. 1985; Smith and Launspach 1991; Hutchison 1996) wherein batches of hot (1000–1200 K) chondritic material accreted to a parent body and cooled at different rates. Haack et al. (1992) demonstrated the physical implausi-

EVALUATION OF THE ARGUMENTS FOR HOT ACCRETION

Lack of Correlation between Petrologic Type and Ca in Low-Ca Pyroxene Low-Ca pyroxene in H-L-LL3 chondrites shows variable Ca contents, presumably reflecting rapid crystallization 86

0019-1035/96 $18.00 Copyright  1996 by Academic Press, Inc. All rights of reproduction in any form reserved.

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during chondrule formation. Similarly, the compositions of Ca-rich clinopyroxenes in type-3 OC are dependent on igneous crystallization within their parent chondrules. In type-4 and -5 chondrites, there is incomplete metamorphic overprinting of clinopyroxenes, leading to a decrease in the range of clinopyroxene compositions with increasing petrologic type (McSween and Patchen 1989). Low-Ca pyroxene and Ca-rich clinopyroxene are incompletely equilibrated in chondrites of petrologic type ,6 (McSween and Patchen 1989). Models of prograde metamorphism assume that, with increasing temperature, orthopyroxene acquires Ca at the expense of augite. It is further assumed that the augite did not exsolve from orthopyroxene during postmetamorphic cooling (Dodd 1981). Hutchison et al. (1980) analyzed low-Ca pyroxene in 10 H3 to H6 chondrites and did not find a positive correlation between the Ca content of pyroxene cores and petrologic type. They concluded that Ca was not subjected to solid-state partitioning between orthopyroxene and diopside in a metamorphic environment. Instead, they proposed igneous partitioning of Ca in hot H6 chondrites via crystallization of substantial amounts of very-low-Ca clinopyroxene or protopyroxene from chondrule liquids; orthopyroxene formed by inversion during slow cooling. However, the Ca-abundance histograms in Fig. 2 of Hutchison et al. (1980) show that pyroxene in type-3 OC peaks at lower Ca contents than that in type-5 or -6 OC (i.e., p0.1 vs. p0.4–0.6 wt.% Ca). This suggests that there may indeed be a correlation between the Ca content of pyroxene cores and petrologic type. Other workers have analyzed low-Ca pyroxene in much larger numbers of OC and demonstrated this correlation. Dodd (1969; Figs. 9 and 10) compiled data for 80 OC and found a clear trend between petrologic type and the Ca content of low-Ca pyroxene in L4-6 and LL4-6 chondrites; the trend for H4-6 chondrites is similar but less clear mainly because of the scatter in the Ca content of pyroxene in H4 chondrites (which Dodd attributed to the greater modal abundance of low-Ca pyroxene in H4 than in L4 or LL4 chondrites). Scott et al. (1986; Fig. 5) reviewed the data for 51 OC (including H3-5, L3-6, and LL3-6 samples) and demonstrated a progressive increase in the Wo content of low-Ca pyroxene with petrologic type: Wo 0.4–1.2 in most type-3 and -4 OC; Wo 1.2–1.6 in most type-5 OC; and Wo 1.6–2.2 in nearly all type-6 OC. These data strongly support the assumption that orthopyroxene acquired Ca at the expense of augite in a metamorphic environment. Heyse (1978; Fig. 2) found that the Wo content of orthopyroxene was highly correlated (r 5 0.966) with the mean grain size of plagioclase in 20 LL4-6 chondrites. This correlation is consistent with prograde metamorphism because the disappearance of igneous glass from chondrule mesostases and the appearance and growth of plagioclase

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is the hallmark of the type-3 to -6 recrystallization sequence (Van Schmus and Wood 1967). Pyroxene Structure and Composition Several studies have attempted to show that the structure of pyroxene in equilibrated chondrites is incompatible with prograde metamorphism. However, pyroxene structural data should be interpreted with caution. For example, use of experimental data to constrain pyroxene cooling rates and thermal histories is problematic because good experimental calibrations of microstructures are rare and even small differences in composition between natural and experimental samples can result in considerable uncertainties. Furthermore, slight errors in cation site occupancy can cause very large uncertainties in calculated cooling rates. Striated orthopyroxene. Ashworth (1980) examined undeformed (i.e., dislocation-free) striated orthopyroxene in H4 Quenggouk and H5 Allegan by high-voltage transmission electron microscopy and found that these orthopyroxene grains had numerous evenly spaced, thin lamellae of clinopyroxene. He interpreted the striated orthopyroxene grains as having formed from inverted protopyroxene during slow cooling. Ashworth et al. (1984) annealed striated orthopyroxene from Quenggouk at 1100 K for 1 week and found that it had converted into normal orthopyroxene (as would be expected in a type-6 chondrite). Because prograde metamorphism is assumed to have lasted far longer than a week, they inferred that there should be no striated orthopyroxene remaining in type-4 or -5 OC. They concluded that the striated orthopyroxene did not form by prograde metamorphism, but instead resulted from primary cooling after hot accretion. R. Hutchison (personal communication, 1996) concluded that if pyroxene characteristic of a type-4 OC can be transformed into that characteristic of a type-6 OC by annealing at 1000 K for 1 week (i.e., if prograde thermal metamorphism is that rapid), then type-5 OC should be extremely rare or absent. The structure of low-Ca pyroxene is very similar in type4, -5, and -6 OC; many grains of orthopyroxene containing few defects and relatively few clinopyroxene lamellae occur in chondrites of each of these petrologic types. In fact, Ashworth (1980, 1981) found that the pyroxene in H4 Quenggouk and H5 Allegan are so similar that a single description sufficed. Type-5 and -6 chondrites are not characterized mainly by the structural state of low-Ca pyroxene (if they were, there would be little difference between the types), but by other features including the presence of coarse-grained plagioclase feldspar and diopside and extensive chondrule-matrix recrystallization (Van Schmus and Wood 1967). These features are unlikely to be reproduced in a type-3 or -4 OC after annealing at 1100 K for 1 week (e.g., Lipschutz and Ikramuddin 1977).

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Because the Quenggouk pyroxene sample that Ashworth et al. (1984) annealed had a substantial portion of orthopyroxene lamellae to begin with, it was not representative of the nearly pure low-Ca clinopyroxene found in highly unequilibrated type-3 chondrites. In order to determine how prograde metamorphism would affect the clinopyroxene typically found in chondrules, Jones and Brearley (1988) and Brearley and Jones (1988, 1993) experimentally annealed synthetic clinopyroxenes at temperatures up to 1318 K for periods of 1–3 weeks. They found that for grains of compositions Fs0 and Fs15 that consisted initially of 100% twinned clinopyroxene, only very minor inversion to orthopyroxene occurred even at the highest annealing temperatures. This failure to produce rapid inversion is consistent with the long periods of time thought necessary to produce orthopyroxene from clinopyroxene during prograde metamorphism. The rate of transformation appears to be strongly dependent on the initial abundance of orthopyroxene intergrown with the clinoenstatite (Brearley and Jones 1993). The fact that type-4 and -5 OC do contain striated orthopyroxene simply provides additional evidence that these rocks did not reach temperatures of 1100 K for extended periods; this is consistent with the metamorphic temperatures of #1020 K suggested by Dodd (1981) for type #5 OC and the very sluggish kinetics of the clino–orthopyroxene transformation at low temperatures. Pyroxene spinodal decomposition textures. In the process known as spinodal decomposition a homogeneous solid solution decomposes into regions of different composition by simultaneous creation of sinusoidal composition waves throughout the entire grain (e.g., Putnis and McConnell 1980). Watanabe et al. (1985) observed that spinodal decomposition textures in pyroxene (augite exsolving from pigeonite) in type-4 to -5 OC have the same p20-nm periodicities (i.e., the same size modulation patterns) as those in type-3 OC. [Pyroxene in L6 Satsuma was found to have abundant dislocations but no exsolution textures.] Watanabe et al. stated that during prograde metamorphism, Ca diffusion must have been limited to less than p10 nm in order for the equilibrated OC to have preserved the modulation patterns of the type-3 OC. On this basis, they rejected prograde metamorphism and concluded that the spinodal decomposition textures must have formed during cooling after hot accretion. On the other hand, Ashworth (1981) found spinodal decomposition textures in pyroxene in H3 but not H4–6 chondrites. Watanabe et al. (1985) interpreted these data to mean that H3 chondrites cooled faster than H4–6 chondrites, but they did not address the problem of why some type-4 to -5 OC should contain pyroxene with spinodal decomposition textures whereas others do not. The interpretation of pyroxene microstructures as an

indicator of cooling rate is complicated; caution should be exercised when applying experimental results to natural samples (e.g., McCallister and Yund 1977; Grove 1982; Buseck et al. 1982). One possible explanation for the presence of spinodal decomposition textures in type-4 to -5 OC is that they are relicts of chondrule formation which survived metamorphism. However, this possibility appears to be incompatible with the observation that Ca has obviously diffused over considerable distances in pyroxenes in type-4 and -5 chondrites as indicated by the change in CaO content as a function of petrologic type (e.g., Dodd 1969; Scott et al. 1986). It is possible that coarsening of the spinodal lamellae is inhibited during prograde metamorphism because of the preferential diffusion of Ca from pigeonite into orthopyroxene. The loss of Ca from the pigeonite effectively reduces the amount of Ca available for coarsening of the spinodally exsolved augite. Such a process is supported by the observation that the pyroxene rims in radial pyroxene chondrules extend to much more calcic compositions in unequilibrated chondrites than in those of higher petrologic type (Kitamura et al. 1983; Watanabe et al. 1985). An alternative explanation is that the textures actually develop during metamorphism. Spinodal decomposition textures form not only as a result of cooling, but may also be produced by isothermal annealing, as long as the compositions lie within the spinodal solvus. For example, the experimental delineation of the coherent spinodal solvus by McCallister and Yund (1977) involved annealing of homogeneous clinopyroxene between 1570 K and 1370 K. McCallister (1978) found that the wavelength of the modulated spinodal textures decreased as temperature decreased, but increased during annealing. Some of the spinodal decomposition textures observed in type-4 and -5 OC could have formed by annealing of a homogeneous pyroxene during prograde metamorphism. Chondrules which cooled too rapidly to develop spinodal textures during primary cooling could have developed these textures during whole-rock metamorphism. In fact, To¨pel-Schadt and Mu¨ller (1985) suggested that the fine scale of spinodal decomposition textures in chondrules from L/LL3.4 Chainpur, LL3.4 Mezo¨-Madaras, and H/L3.6 Tieschitz could have formed either by primary cooling or by verylow-temperature annealing. Additional detailed studies are required to address this problem. Cooling rates based on compositional zoning in low-Ca pyroxene. Watanabe et al. (1985) found that pyroxene in chondrules in L3.7–3.9 ALHA77252 and L5 Fukutomi (listed as L4–5 by Watanabe et al.) exhibits compositional zoning wherein Ca increases from core to rim up to a value of Wo18 accompanied by a significant increase in FeO; Cafree regions in these pyroxenes are orthopyroxene. These observations were contrasted with those of an earlier study of radial pyroxene chondrules in L3 Y-74191 (Kitamura

EVALUATION OF HOT ACCRETION

et al. 1983) which showed that Ca increases from core to rim to a value exceeding Wo30, but that FeO remains constant; Ca-free regions in these pyroxenes are clinopyroxene. Watanabe et al. interpreted these differences in compositional zoning and pyroxene fine structure to indicate that type-3 OC cooled more rapidly than type-4 to -5 OC. There are considerable problems with this interpretation because it is based on experimental data (Hewins et al. 1981) for the evolution of Fe zoning during the primary igneous crystallization of the chondrule. Because chondrules formed as isolated objects in free space, such zoning cannot be used to infer anything about the postaccretionary cooling rate of equilibrated or unequilibrated OC. Accretion must have occurred after chondrule formation and solidification, because all OC (except for the most recrystallized type-6 samples in which primary textures have been obliterated), contain chondrules of different textures which formed at different cooling rates. It is far more likely that the observed Fe enrichment at the margins of chondrule pyroxenes in the type-3.7 to -3.9 and type-4 to -5 chondrites studied by Watanabe et al. (1985) is due to postaccretionary metamorphism as reported by Tsuchiyama et al. (1988) and Jones (1993). Experimental data (Brearley and Jones 1993) show that rare thick orthopyroxene lamellae in H4 Conquista could not have formed during single-stage cooling from an accretion temperature of 1000–1200 K as expected in autometamorphism. Brearley and Jones (1993) found a two-stage cooling history (i.e., rapid cooling after chondrule formation to temperatures below p770 K followed by parentbody annealing) more plausible. The presence of Fe-rich orthopyroxene containing lamellae of low-Ca clinopyroxene in type-4 to -5 L chondrites as reported by Watanabe et al. (1985) is a very significant problem for the autometamorphic model. This type of intergrowth can only form as a result of rapid cooling from the protopyroxene field. The clinopyroxene represents the metastable quench product of protopyroxene. However, the low-Ca pyroxenes in a single radial pyroxene chondrule in L5 Fukutomi have core compositions (Fs28–38; Watanabe et al. 1985) which extend well beyond the stability field of protopyroxene (Fs24; Huebner 1982). This range is similar to that observed in type-IIB chondrules in LL3.7 Parnallee (R. H. Jones, personal communication, 1995). The explanation provided by Watanabe et al. (1985) that these Fe-rich, low-Ca pyroxenes formed by a single-stage process during cooling is inadequate. The formation of pyroxenes with this composition and microstructure requires a two-stage process: (1) initial crystallization of protoenstatite at high temperature followed by sufficiently rapid cooling to below p770 K (to prevent complete inversion to orthopyroxene and the quenching of protopyroxene to clinopyroxene), and (2) reheating of the resultant intergrowth of orthopyroxene

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and clinopyroxene (to equilibrate the pyroxene to a new, Fe-rich composition lying beyond the field of protopyroxene). The formation of this microstructure by autometamorphism is untenable. It would require chondrites to have accreted at p1100–1200 K and then to have cooled slowly enough to allow diffusion of Fe into the chondrules. However, this would have caused complete inversion of the protopyroxene to clinopyroxene and obliterated the spinodal-decomposition textures indicative of rapid cooling. Autometamorphism would require extensive diffusion over long distances (tens to hundreds of micrometers) while preserving a microstructure which requires diffusion over only a few nanometers. Cooling rates from Fe–Mg ordering. The ordering of Fe and Mg between the M1 and M2 sites in pyroxene is an extremely useful tool for measuring the low-temperature cooling history of meteorites and terrestrial rocks (Ganguly 1982). Some determinations of the cooling rates of type4 to -6 OC have been carried out which show relatively high cooling rates (degrees per day) down to p800–900 K with no evidence for metamorphic reequilibration (e.g., Molin et al. 1994; Artioli and Davoli 1994). These observations appear to support either an autometamorphic model or a rubble-pile accretionary model. Although the data suggest rapid cooling rates, care should be taken in interpreting the data. Any application of cooling-rate models using intercrystalline exchange is made on the assumption that the crystal has a constant composition throughout its cooling history. However, this is not the case for either prograde metamorphism or autometamorphism. The autometamorphic model requires that accreting chondrules with different compositions equilibrate in type-4 to -6 chondrites during cooling. The prograde metamorphic model allows equilibration of such chondrules to occur during both reheating and cooling. Thus, in both models, the composition of any given pyroxene grain will change during its cooling history. The effect of this change on Fe–Mg intercrystalline ordering is essentially unknown, and, thus, meaningful cooling rates cannot be determined. Given the observations that pyroxenes in chondrules in unequilibrated chondrites show highly complex zoning and intergrowths (Tsuchiyama et al. 1988; Jones 1993) and that the calculated cooling rate is extremely sensitive to the measured site occupancies, the errors associated with such determinations are commonly large (e.g., Ganguly 1982; Saxena 1983; Tribaudino and Talarico 1992). For example, in some cases, an error of 60.005 in the occupancy of Fe in the M2 site can cause an order of magnitude difference in the calculated cooling rate (Tribaudino and Talarico 1992). Cooling rates are also sensitive to differences in pyroxene composition and the presence of minor elements such

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as Al. Saxena (1983) recommended that several crystals from an individual rock be studied to minimize errors in cooling-rate determinations. However, in the studies reported for OC, typically only a single crystal from each meteorite has been examined. Aberrant grains introduced during brecciation (Scott et al. 1985; Rubin et al. 1988; Rubin 1990) must also be avoided. For example, Artioli and Davoli (1994) studied a clinoenstatite from the highly shocked Jolomba LL6 breccia; however, any clinoenstatite in a brecciated type-6 chondrite must be an aberrant grain because the clinoenstatite–orthoenstatite transition occurs at #900 K (Boyd and England 1965; Grover 1972), appreciably lower than the equilibration temperatures of type-6 OC (1090–1200 K, Olsen and Bunch 1984). The measured cooling rate of this grain cannot be considered to be equivalent to that of the host rock. Until these problems and complexities are taken into consideration, cooling rates for chondritic meteorites based on intercrystalline partitioning should be viewed with caution. Polycrystalline Taenite Bevan and Axon (1980, 1981) observed polycrystalline taenite (multizoned with respect to Ni) within metallic spherules in H/L3.6 Tieschitz. They interpreted this phase as a relict solidification structure wherein the Ni zoning was produced in part during solidification; they inferred that the polycrystalline taenite had failed to anneal into monocrystalline taenite because of rapid cooling (1700 to 1000 K within days to weeks). Hutchison et al. (1980) observed polycrystalline taenite in H/L3.9 Bremervo¨rde, H4 Menow, H4 Quenggouk, H6 Butsura, and rare regions of H6 Kernouve. By analogy with the rapid-cooling origin proposed for this phase in Tieschitz, Hutchison et al. proposed that all H3–6 chondrites containing polycrystalline taenite cooled rapidly from 1700 K, an idea inconsistent with prograde metamorphism. It seems clear that metallic spherules containing polycrystalline taenite in Tieschitz and other type-3 OC were once molten; they either were directly melted by the chondrule-formation process or were expelled from spinning chondrules by centrifugal force (e.g., Fig. 8 of Grossman and Wasson 1985). However, the textures of many other occurrences of polycrystalline taenite in OC do not indicate that the metal was once molten. Scott and Rajan (1981a,b) pointed out that the Ni zoning in the polycrystalline taenite in Tieschitz could not have been produced by solidification zoning; they aruged that cooling rates in equilibrated chondrites that were slow enough to permit significant growth of kamacite would erase prior solidification zoning in taenite by solid-state diffusion (as also shown in the computer models of Willis and Goldstein 1981). Although Willis and Goldstein (1981) found that solidification zoning would

not be erased during relatively rapid cooling (.1000 K Ma21), as might be expected for type-3 OC during autometamorphism, appreciable amounts of kamacite (such as occur in most type-3 OC including Tieschitz) would not grow at these rates. This invalidates the assumption of Hutchison et al. (1980) that equilibrated OC containing polycrystalline taenite cooled rapidly. Although the Ni zoning in polycrystalline taenite would have been obliterated by solid-state diffusion during prograde metamorphism (particularly at slow cooling rates such as the 2 K Myr 21 rate measured for metal in Tieschitz; Willis and Goldstein 1981), the polycrystallinity of the taenite need not have been affected. It seems likely that polycrystalline taenite in type-3 OC is a premetamorphic relict; given the low diffusion coefficients of Fe and Ni in taenite, metal grains in type-3 OC may not have been annealed sufficiently to have formed large single crystals of taenite (Scott and Rajan 1981a). Although Bevan and Axon (1981) found little deformation in Tieschitz, they concurred with Scott and Rajan (1981a) that, in some equilibrated OC, polycrystalline taenite may have formed by shock deformation and subsequent annealing. Heterogeneous Metal Grains Smith and Launspach (1991) observed compositionally and texturally heterogeneous metallic Fe–Ni grains in L6 Bruderheim. Some grains have very complex structures and consist of plessitic cores with marginal regions composed of kamacite, taenite, and tetrataenite. The authors concluded that such structures were unlikely to have survived solid-state diffusion during prograde metamorphism. They proposed that chondrites accreted at high temperatures, causing the silicates to equilibrate but leaving metallic Fe–Ni grains unaffected. They assumed that because the metal grains did not form a three-dimensional network, solid-state diffusion of Fe and Ni was rendered ineffective. The establishment of compositional and textural heterogeneities in the metal grains was ascribed to primary cooling, autometamorphism, and subsequent low-temperature annealing during burial. The authors are correct in stating that complex compositional and textural heterogeneities in metallic Fe–Ni grains are unlikely to have survived prograde metamorphism. Because diffusion of Ni and Fe in metal is much faster than diffusion of Mg and Fe in silicates, heterogeneous metal is unlikely to have survived temperatures high enough to have equilibrated the silicates. However, Bruderheim is a fragmental breccia of shock stage S4, indicating that it experienced an equilibration shock pressure of at least 15–25 GPa (Sto¨ffler et al. 1991; Schmitt and Sto¨ffler 1995). At this shock pressure, some metallic Fe–Ni grains are partly melted and opaque veins form. Shock heating and quenching are most likely respon-

EVALUATION OF HOT ACCRETION

sible for producing the heterogeneities in metallic Fe–Ni in Bruderheim. Smith et al. (1993) completed p3800 random analyses of metal grains in 11 Antarctic OC (including many type6 samples) and observed wide variations in Ni and Co compositions. Most distributions are bimodal: one cluster, representing kamacite, has low Ni and high Co; the other cluster, representing mainly taenite, has large variations in Ni and is negatively correlated with Co. The kamacite Co range in the individual L6 chondrites is considerably wider than the L-group range determined by Rubin (1990). Smith et al. (1993) concluded that the scatter in their analyses was inherited from the time of accretion and is contrary to what is expected in chondrites that underwent prograde metamorphism. A major problem with the data of Smith et al. (1993) is that their set of analyses includes many grain boundaries wherein the electron beam overlapped adjacent metallic phases (e.g., kamacite–taenite; kamacite–tetrataenite). For example, the negative trend between Co and Ni among their ‘‘taenite’’ analyses (i.e., analyses with p25–45 wt.% Ni) is probably due to overlap between taenite and kamacite. The rare analyses in LL5-6 Y74646 that show p45 wt.% Ni and p1.0 wt.% Co may represent overlap between tetrataenite (high Ni, low Co) and the low-Ni phase with 37 wt.% Co found in highly oxidized LL chondrites (e.g., Afiattalab and Wasson 1980; Rubin 1990). Additional scatter among the analyses of Smith et al. (1993) may be due to aberrant metal grains incorporated into the rocks during brecciation. Although Rubin (1990) typically analyzed #12 kamacite grains per meteorite, these data were sufficient to show that at least 33% of type4 to -6 OC contain aberrant kamacite grains differing from the mean kamacite Co composition by more than three standard deviations. Smith et al. (1993) presented hundreds of analyses of kamacite in many of their meteorites; dozens of aberrant grains would be expected. Thus, the wide scatter in their data probably results from analyses of grain boundaries and aberrant grains; it does not constitute evidence against prograde metamorphism. The occurrence of rare martensite grains in the unshocked (shock stage S1) H6 chondrite, Kernouve (R. Hutchison, personal communication 1995), was proffered as evidence inconsistent with prograde metamorphism. However, it seems likely that these rare grains are aberrant and were incorporated into the chondrite during postmetamorphic brecciation. This is not an ad hoc explanation. The data of Rubin (1990) suggest that far more than 33% of equilibrated OC are aberrant-grain-bearing, postmetamorphic, fragmental breccias. Hutson (1989) and Rubin (1992) reported a series of elongated metal masses in Kernouve that appear to be relicts of partially obliterated metal shock veins; this suggests that Kernouve was annealed after being shocked.

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Martensite grains are not consistent with hot accretion because this model requires type-6 chondrites to have cooled too slowly to cause the formation of martensite. If the parameters of the model were somehow significantly modified to allow much more rapid cooling of type-6 chondrites, then martensite should be common. However, this is not the case (Christophe Michel-Le´vy 1981). Without brecciation, martensite could not occur along with kamacite and taenite in an unshocked chondrite. Misshapen Chondrules Hutchison et al. (1979) reported that a small proportion of chondrules in Tieschitz are nonspherical (e.g., dumbbell shaped) and seem to have molded themselves around one another while they were at least partly molten. In some cases, plastic chondrules appear to have been deformed, causing the solidified rims to rupture as chondrule liquid oozed out. Such misshapen chondrules are not confined to Tieschitz, but have been observed in many type-3 chondrites. Hutchison et al. (1979) concluded that these chondrules had been partly molten when they accreted together with somewhat cooler chondrules, clasts, metallic Fe–Ni and sulfide to the surface of a hot (1100 K) asteroid. Some misshapen chondrules may merely be chondrule fragments. Many, however, are probably adhering or enveloping compound chondrules (or compound-chondrule fragments) that fused in the nebula (Wasson et al. 1995). Most of these conjoined objects are probably siblings that collided shortly after forming in the same heating event. Wasson et al. (1995) also described chondrules that appear to have had their surfaces ruptured with resultant expulsion of interior liquid and formation of a portion of a new chondrule surface. Such chondrules were labeled as having ‘‘ghost-rims’’ and it was suggested that they suffered a collision while they were still molten. Objects adjacent to compound chondrules in type-3 OC are separated by intervening matrix material. As addressed in the following section, matrix material has never been melted and thus does not represent liquid extruded from chondrules. Hence, any complementarity in shape between adjacent objects separated by matrix material (e.g., as reported in H/L3.6 Tieschitz and LL3.1 Krymka; Holme`n and Wood 1986) is either due to coincidence, to jostling and wedging during chondrite compaction, or to uniaxial deformation during shock. Sneyd et al. (1988) studied 15 OC and found a positive correlation between shock facies and percent flattening of chondrules. Scott et al. (1992) proposed that chondrules became elongated during shock events as they were squeezed into matrix pores in the plane of the shock front. Such shock-deformation of chondrules is probably

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responsible for some of the misshapen, apparently molded chondrules in type-3 chondrites. Apparent Igneous Nature of Matrix Materials Hutchison and Bevan (1983) reviewed the evidence that matrix material in type-3 OC formed from a melt. Matrix material surrounds misshapen, apparently ruptured chondrules in Teischitz (e.g., Hutchison et al. 1979). Metalsulfide spherules in Tieschitz occur amidst matrix material (Fig. 1b of Bevan and Axon 1980). Matrix material is very fine grained (0.1 em) and contains amorphous material (Ashworth 1977, 1981); some matrix material in LL3.1 Bishunpur contains subspherical vugs (J. R. Ashworth, personal communication 1982 to Hutchison and Bevan 1983). Bishunpur matrix material also appears to have acted like an intrusive liquid, filling embayments in metalsulfide aggregates; similar embayments in chondrules are also filled with matrix material. On the basis of such textural evidence, Hutchison and Bevan (1983) inferred that matrix material was probably hotter than 1200 K during accretion. However, ATEM (analytical transmission electron microscope) studies of six type-3 OC revealed the clastic (unmelted) nature of interchondrule matrix material (Alexander et al. 1989). Hutchison (1987) described an unusual isolated clast from Tieschitz consisting of a twinned clinoenstatite fragment rimmed with Cr- and Mn-rich augite and surrounded by fine-grained matrix material. Because augite of this composition has not been reported from chondrules, Hutchison suggested that, during crystallization of white matrix at high temperatures, the Cr- and Mn-rich augite formed by partial replacement of clinoenstatite. Although Tieschitz is cited by Hutchison et al. (1979) as a meteorite which exemplifies hot-accretionary textures, there is considerable evidence of secondary mobilization of material. Christophe Michel-Le´vy (1976) suggested that Tieschitz underwent dissolution of mesostasis glass which was then redistributed to form white matrix due to the activity of parent-body fluids. This suggestion differed from that of Hutchison et al. (1979) who argued that white matrix represented chondrule mesostasis squeezed out of hot, plastic chondrules. More recently, Alexander et al. (1994) and Hutchison et al. (1994) reported evidence which lends support to the original suggestion by Christophe MichelLe´vy (1976). Ion microprobe analyses indicate that white matrix and the mesostases of chondrules are enriched in the halogens F and Cl and the alkalis K, Rb, and Ba. Hutchison et al. (1994) concluded that because some Tieschitz chondrules have regions of mesostasis near the chondrule margins that have been altered to Na–Al–Clrich material, the fluids responsible for this alteration must have come from outside the chondrules. These observations imply that halogen- and alkali-rich fluids promoted

extensive remobilization and mass transfer of material between the matrix and the chondrules. Tiny (p4 em) rimmed barred olivine microchondrules in the matrix of H3.4 Sharps (Nagahara 1984) could not have survived melting and are inconsistent with an igneous origin for matrix material. Nagahara (1984) also described many highly unequilibrated assemblages in type-3 OC matrix material where no reaction rim occurs at the interface between adjacent phases. These include a Fs4Wo3.5 pyroxene grain from Semarkona forming an overgrowth on Fs2Wo40 pyroxene, Fa50 olivine overgrowing Fs5 pyroxene in Tieschitz, and Fa75 olivine surrounding a Fa1 olivine core in LL3.1 Krymka. However, it is possible that fayalitic olivine surrounding less-ferroan mafic silicate grains in these chondrites reflects late-stage parent-body aqueous alteration (Wasson and Krot 1994). As shown above, misshapen chondrules were not deformed while accreting to the parent body. Many are primitive chondrule fragments or compound chondrules that were surrounded by cold, fine-grained matrix material in the nebula prior to, or commensurate with, accretion. The occurrence of matrix material around the unusual fragment with Cr- and Mn-rich augite in Tieschitz (Hutchison 1987) suggests that this fragment is also a primitive object, possibly derived from a remelted chondrule. Chondrite compaction (e.g., Wasson 1995) is probably responsible for the filling of embayments by matrix material in chondrules and metal-sulfide aggregates. The subspherical vugs within Bishunpur matrix are more difficult to understand, but it is possible that they are pores in poorly compacted clumps of matrix material. The presence of high D/H ratios in the matrices of unequilibrated chondrites also provides evidence against hot accretion. Although the D/H enrichments have been widely ascribed to organic compounds (McNaughton et al. 1982; Yang and Epstein 1983; Alexander et al. 1990), it now appears that much of the D (at least in LL3.0 Semarkona) is in structurally bound water in phyllosilicate minerals (Deloule and Robert 1995). Within the uncertainties of their calculations, Deloule and Robert (1995) could not rule out the possibility that all of the D in Semarkona is in phyllosilicates rather than organic compounds. Both the organic compounds and the water in phyllosilicates probably formed by ion-molecule reactions at very low temperatures within the interstellar medium (e.g., Zinner 1988; Deloule and Robert 1995) or in low-pressure portions of the early nebula (Yung et al. 1988; Halbout et al. 1990). Studies of the distribution of D-rich materials in type-3 OC show that they are very sensitive to metamorphism; isotopically anomalous material seems to have been destroyed in petrologic types $3.6 (McNaughton et al. 1981, 1982; Yang and Epstein 1983; Alexander et al. 1990). Because the phyllosilicates in Semarkona probably formed

EVALUATION OF HOT ACCRETION

on a parent body (rather than in the solar nebula) it seems likely that the D-rich water within them would have occurred as ice and mixed with other chondritic components during agglomeration. It is untenable that ice or organic compounds could have been present in the kind of environment in which hot accretion occurs; it is not possible to accrete these materials into an OC at 1100–1200 K and preserve them even during rapid cooling. Similarly, the presence of (isotopically anomalous) interstellar diamond, SiC, and graphite in the matrices of OC of petrologic type #3.8 and their absence in higher petrologic types indicate that their survival is sensitive to ambient temperatures (Huss 1990; Huss and Lewis 1995); these grains could not have survived hot accretion. Even if type3 OC remained at high temperatures only briefly during accretion, the interstellar carbon compounds would still have resided in a hot nebula for an extended period prior to accretion. They are unlikely to have survived. To circumvent this problem, advocates of hot accretion appeal to a ‘‘patchy’’ distribution in the nebula and conclude that presolar grains survived in cooler regions (e.g., R. Hutchison, personal communication 1996). Of course, if there are many cooler regions in the nebula, then the meaning of ‘‘hot’’ accretion becomes unclear. Data from graphitic carbons constitute additional evidence against hot accretion. Graphitic material intimately mixed with the fine-grained matrix material of Tieschitz has been shown to be only partially ordered (i.e., to be poorly graphitized carbon) (Christophe Michel-Le´vy and Lautie 1981). This type of material could not have formed at high temperatures; under those conditions it would have been converted to crystalline, ordered graphite (e.g., Rietmeijer and Mackinnon 1985). In summary, matrix material in type-3 chondrites did not form at high temperatures. Depletion of Volatile Trace Elements with Petrologic Type Several studies (e.g., Larimer and Anders 1967; Binz et al. 1976) have found that the concentrations of volatile trace elements in OC are increasingly depleted with increasing petrologic type. For example, type-3 OC contain concentrations of Bi, In and Tl of 0.13 3 CI, 0.16 3 CI and 0.04 3 CI, respectively; the corresponding values for equilibrated OC are 0.02 3 CI, 0.004 3 CI and 0.007 3 CI, respectively (Binz et al. 1976). These data were interpreted as indicating that type-3 OC formed at lower nebular temperatures than more equilibrated OC and were thus able to acquire greater concentrations of volatile trace elements. This scenario is equivalent to the hot accretion model wherein equilibrated OC accreted at higher nebular temperatures than unequilibrated OC. However, experimental low-pressure annealing of samples of H/L3.6 Tieschitz demonstrate that volatile elements

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are lost progressively with increasing temperature (Ikramuddin et al. 1977). This annealing experiment mimics prograde thermal metamorphism and shows that hot accretion is not required by the correlation between volatiletrace-element depletion and petrologic type. Furthermore, the supercosmic concentrations of moderately volatile Rb (112 3 CI) and Cs (7 3 CI) in lightcolored equilibrated chondritic clasts in the Leighton H chondrite regolith breccia (McSween et al. 1981) appear to be inconsistent with a hot accretion origin because autometamorphism of equilibrated OC material is expected to cause depletions in volatile trace elements. The high concentrations of Rb and Cs in the equilibrated clasts may have instead arisen from inhomogeneous crystallization of feldspar (in which Rb and Cs are cited) during prograde thermal metamorphism (McSween et al. 1981). Natural Remanent Magnetization (NRM) Morden and Collinson (1992) analyzed the magnetic properties of 11 OC, repeatedly fragmenting, reorienting and re-measuring the samples until the samples were millimeter sized. They found that the orientations of the stable NRM of the OC were random at scales of p1 mm3. In addition, they determined that all of the samples have a magnetic fabric, analogous to a structural fabric in deformed rocks; in most of the samples the fabric is continuous. The dominant magnetic carrier was found to be finegrained taenite in type-3 OC and tetrataenite in type-4 to -6 OC; these phases have Curie points of ,900 K and p850 K, respectively (Morden and Collinson 1992). The existence of the continuous magnetic fabric was regarded as evidence that the OC are not fine-scale breccias. The authors ruled out prograde metamorphism because metamorphic heating above p850 K (as expected in type-4 to -6 OC; Dodd 1981) would erase the random orientations of the NRM. Morden and Collinson suggested that the metal grains condensed from a hot (possibly impact-generated) plasma cloud on a parent body; the grains cooled in the presence of a magnetic field and accreted along with other chondritic components to form a new hot body. Tetrataenite formed in the equilibrated OC after they were buried by additional layers of debris. The magnetic fabric formed during final chondrite lithification as loose fragments were welded together by an impact event. Most OC appear to be fragmental breccias that contain scattered metallic Fe–Ni and silicate grains of aberrant compositions that were incorporated into their hosts after metamorphic equilibration (Scott et al. 1985; Rubin et al. 1988; Rubin 1990). For example, among 20 olivine grains analyzed in H5 Anlong, 19 are equilibrated and have a mean composition of Fa19.260.4 ; one grain is Fa21.1 , more than four standard deviations away from the mean (Fig. 7.7.2 of Rubin et al. 1988). Among 11 kamacite grains

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analyzed in LL6 Dongtai, 10 are equilibrated and have a mean Co content of 17.6 6 1.0 mg/g; one kamacite grain contains 106 mg/g Co (Rubin 1990). Approximately half of the analyzed type-4 to -6 OC were found to contain aberrant olivine and/or kamacite grains; because relatively few grains were analyzed, this suggests that far more than half of the equilibrated OC are fragmental breccias. By analogy to some CM chondrites which contain millimetersize clasts that experienced different degrees of aqueous alteration (e.g., Murray; Rubin and Wasson 1986), it is plausible that OC are also brecciated on millimeter-size scales. In his study of the NRM of LL5 Olivenza, Collinson (1987) concluded that Olivenza could be a fine-grained breccia. Such fine-scale brecciation would be texturally unobservable in the (anhydrous) OC, but it could account for the random orientations of the stable NRM measured by Morden and Collinson (1992). We agree with Morden and Collinson (1992) that the magnetic fabric in the OC may have been created during the final lithification of these rocks by an impact event (e.g., Bischoff et al. 1983), but it seems likely that this took place after metamorphism, brecciation and the incorporation of aberrant grains. CONCLUSIONS

Although a large number of arguments favoring hot accretion has been made, none is compelling. The textural, compositional, isotopic, and magnetic properties of type3 to -6 ordinary chondrites and their constituent minerals are more readily understood in terms of prograde metamorphism of initially cold, unequilibrated type-3 material and episodes of brecciation, shock heating, and, in some cases, aqueous alteration. ACKNOWLEDGMENTS We thank E. R. D. Scott, R. H. Jones, and J. T. Wasson for valuable discussions. Helpful reviews were supplied by H. Y. McSween, T. J. McCoy, and R. Hutchison. This work was supported by NASA Grants NAGW-3535 (J. T. Wasson, UCLA) and NAGW-3347 (J. J. Papike, Institute of Meteoritics, University of New Mexico).

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