Formation and alteration of CAIs in Cold Bokkeveld (CM2)

Geochimica

et Cosmochlmica

Acta. Vol.

58. No. 8. pp. 1913-1935.

Copyright

Pergamon

(ci 1994

Printed in thetJSA. All rights 00

1994

Elscvier Science Ltd reserved

I6-7037/94$6.00+ .OO

Formation and alteration of CAIs in Cold Bokkeveld (CM2) RICHARD C. GREENWOOD,’ MARTIN R. LEE,‘.* ROBERT HUTHISON,’

and DAVID J. BARBER’.’

‘The Natural History Museum, Cromwell Road, London SW7 5BD, UK ‘Department of Physics, University of Essex, Colchester CO4 3SQ, UK (Received Aug1rsf7, 1992; accepted in revisedji,rm Decmher

14. 1993)

Abstract-Calcium aluminium-rich inclusions (CAIs) from the carbonaceous chondrite Cold Bokkeveld (CM2) have been located and characterised in situ; the textural environment of CM2 CAIs has been studied for the first time. Based on their primary mineralogy, 345 CAls have been classified into four groups: ( 1) spine1 inclusions, (2) a hibonite inclusion, (3) spinel-pyroxene (sp-py) inclusions, and (4) spinel-pyroxene-olivine (sp-py-01) inclusions. In addition, refractory spherules and spinel-bearing isolated olivines are also recognised. CAIs in Cold Bokkeveld underwent extensive aqueous alteration and contain secondary minerals, including calcite and various phyllosilicates. These assemblages are similar to those in the chondrules and matrix of the meteorite, and it is concluded that all were altered more or less simultaneously within a parent body regolith. As is common in CM2s, most CAIs in Cold Bokkeveld have a rim sequence containing a prominent Fe-rich phyllosilicate layer. This formed early in the alteration sequence by the dissolution of melilite and subsequent precipitation of Fe-rich phyllosilicates from fluids with a high Fe/Mg ratio. CAIs in Cold Bokkeveld experienced two stages of deformation. An earlier nebula stage predated the formation of the dust mantles that surround most CAIs, and a later stage occurred in an asteroidal regolith. The CAIs in CM2s are smaller than those in CV3 chondrites because of the greater intensity of the earlier fragmentation episode that they experienced. Aqueous alteration of CAIs and regolith deformation appear to have been simultaneous. Phase relations and pyroxene compositional trends in sp-py and sp-py-01 inclusions indicate that they crystallised from liquids. Most spinel-bearing inclusions in Cold Bokkeveld have a normal Mg isotopic composition, so they could not have formed by the simultaneous melting and evaporation of primitive dust. One inclusion is enriched in the heavy isotopes of Mg and may have formed in this way. Two (or more) events are required to form most of the inclusions. The first produced the refractory compositions, either by vapour/solid condensation or slow evaporation of primitive dust. The second event melted the refractory precursors which later crystallised to form the observed textures and mineral assemblages. It has been suggested that refractory spherules with inwardly radiating hibonites formed from molten droplets at temperatures as high as 2135°C. However, the crystallisation sequences of these inclusions inferred from textural evidence are incompatible with those derived from phase relations and estimates of bulk composition. It is concluded that CM2 refractory spherules did not form by crystallisation of molten droplets, but were derived by fragmentation of larger inclusions in which they formed by partial melting.

INTRODUCTION

1985), and demonstrate the existence of distinct O-isotopic reservoirs in the early nebula (CLAYTON et al., 1973). A major obstacle to understanding the origin of CAIs is that all underwent some degree of secondary processing and alteration. Many are brecciated and the primary phases became altered to an often complex secondary mineral assemblage, in response to an influx of volatiles (Na, K, Fe. Cl, CO*, and HzO). ( MACPHERSON et al., 1988). Because of their small size (typically ~500 pm), the CAIs from CM2s (such as Murchison and Murray ) used in previous investigations were extracted and concentrated by mineral separation procedures ( MACPHERSON et al., 1983: FAHEY et al., 1987; IRELAND, 1988, 1990). One drawback of this approach is the loss of the textural relationship between a CA1 and its host. A second is that the freeze-thaw techniques that were frequently employed may have disaggregated friable CAIs; the populations studied may have been selectively biased. Our objective is to determine the mode of formation of CAIs in Cold Bokkeveld (CM2). Because the textural relationship between a CA1 and its host is critical in determining

CALCIUM ALUMINIUM-RICH inclusions (CAIs) are refractory objects found mainly in carbonaceous chondrites (excluding Cls) ( MACPHERSON et al., 1988). CAIs have been intensively studied since CHRISTOPHEMICHEL-LEVY (1968) compared a melilite-spine1 assemblage in Vigarano to the composition of early high-temperature condensate from a gas of solar composition ( LARIMER, 1967). Their importance in offering insight into the processes that operated in the early Solar System has been demonstrated by the results of this activity. CAIs, the oldest reliably dated Solar System materials ( TILTON, 1988), contain evidence for the existence of short-lived nuclides such as 26A1 (half-life 0.72 My) (WASSERBURG,

* Present address..Department of Geology and Geophysics, University of Edinburgh, Edinburgh EH9 3JW. UK. ‘Present address: The Hong Kong University of Science and Technology, Clear Water Bay. Kowloon. Hong Kong.

1913

1914

their

R. C. Greenwood

history,

we chose

to identify

and examine

the inclusions

et al.

Table 2. Modal analyses (Volume %)

in situ. Matrix Opaques

TECHNIQUES We used three demountable polished thin sections and one conventional polished section. An X-ray mapping technique was used to locate CAIs with the Al-rich index minerals spine1 and hibonite. Each inclusion was located on montages of backscattered electron images. Abundances of CAIs in each section are given in Table 1. Minerals were analysed by wavelength dispersive techniques, using either a Cambridge Instruments Microscan 9 electron microprobe (20 kV accelerating voltage, 25 nA current) or a Cameca SX50 electron microprobe (20 kV accelerating voltage, 20 nA current), and by energy dispersive techniques using an Hitachi S2500 SEM ( I5 kV accelerating voltage, 2 nA current) fitted with a Link Systems ANlO/ 55s analytical system. Selected CAIs were studied by analytical transmission electron microscopy (ATEM), using a JEOL 200CX TEM equipped with a Link Systems X-ray detector. CAIs were extracted from demountable polished thin sections by cementing Cu discs with central 50 pm-200 pm holes to the surface ofthe section with an epoxy resin, the middle of the hole being positioned over the CA1 of interest. After the resin had hardened, the mounting medium was melted and the Cu discs lifted off, complete with CAI. Each sample was briefly washed in acetone and then ion-thinned until the CA1 had perforated. Imaging and X-ray analysis were performed using an accelerating voltage of 200 kV. The X-ray spectra were quantified using the thin film approximation (CLIFF and LORIMER, 1975) and utilising a set of Kfactors experimentally determined from mineral standards. Magnesium isotopic abundances were measured on the Edinburgh University Cameca IMS-4f ion microprobe using a 14.5 keV Oprimary ion beam, focussed to a 20-40 Frn spot, with a primary beam current of 25 X 10e9 Amp. Magnesium was measured as positive secondary ions with m/dm of 3500-4000, which was sufficient to separate 24MgH+ from *‘Mg. Data were collected in an automated peak jumping mode with each analysis representing at least fifty cycles through the masses 24, 25, and 26. In each cycle mass 24 was counted for 1 s and masses 25 and 26 for 2 s each. Count rates on mass 24 were normally about 2.5 X IO’ cps. The Mg isotopic ratios measured by ion microprobe differ from true values because of fractionation produced during sputtering and secondary ion formation ( HUNEKE et al., 1983; MCKEEGAN et al., 1985). In addition. measured ratios vary depending on the composition of the material being analysed ( HUNEKE et al., 1983). To assess whether intrinsic mass-dependent fractionation is present in a sample, these effects must be compensated for using terrestrial standards of appropriate composition. Grains of Burma spine1 were mounted in each section in a position cofocal with the surface of the specimen. Measurements of the unknowns and the standard were interspersed and in the five days of data collection a total of sixty-eight individual analyses of Burma spine] were obtained. Day-to-day variation in the Mg isotopic ratios measured on the standard are given later. Petrography

and Mineral Chemistry of Cold Bokkeveld

To understand the post-formational history of the CAIs in a meteorite, their textural relationship to host matrix must be known.

Table

1. Abundance

of CAls

Section Number

No. of CAls

P4038” P4784’ P4504 P4507’

28 149 116 52

37 266 195 90

0.8 0.6 0.6 0.6

Total

345

588

0.6

*Demountable

Area of Section (mm2)

section

CAI (per mm2)

Single grains Mg-rich olivines Fe-bearing olivines Pyroxene Polymineralic fragments Chondrules and coarse-grained inclusions Fine-grained inclusions Sulphide-metal inclusions

82.7

(74.2)’ (I .O)‘

4.5 2.9 1.4 0.2

(6.8)’

12.8

(18.0)‘

9.1 3.2 0.5

Surface area counted = = 1 00mm2 Number of points = 1049 l Data in brackets McSween (I 979)

Although some details of the mineralogy of Cold Bokkeveld have been published (WOOD, 1967; CAILLERE and RAUTUREAU, 1974; KING and KING, 1978; MCSWEEN, 1979; BARBER, 1981; METZLER et al., 1992). no general description of its petrography or mineral chemistry is available. This deficiency is remedied here. Like other CM2s, Cold Bokkeveld is a breccia of various hightemperature macroscopic components (chondrules, single crystals and crystal fragments, refractory inclusions, lithic fragments) in a phyllosilicate-rich fine-grained matrix ( CAILLERE and RAUTUREAU, 1974; BARBER, 198 I ). Modal analyses of these constituents appear in Table 2, with data of MCSWEEN .( 1979). The different datasets presumably reflect real differences between the sections examined and indicate that the meteorite is heterogeneous. Cold Bokkeveld has a distinct planar fabric defined by the alignment of the long axes of chondrules, inclusions and crystal fragments (Fig. la,b). Compositional layering is developed locally and appears to be oriented parallel to this fabric. Similar fabrics occur in other carbonaceous and ordinary chondrites (MARTIN and MILLS, 1980; FUJIMURA et al., 1983; CAIN et al., 1986; SNEYD et al., 1988), and are ascribed to compaction in a regolith or impact-related processes. KING and KING ( 1978) noted that true chondrules, with unequivocal textural evidence of having crystallised from molten drops, are rare in Cold Bokkeveld. Even where chondrules have rounded margins there is often clear evidence that these have been abraded (Fig. la). Most microcrystalline, silicate-rich objects in the meteorite have angular margins and are probably fragments of disrupted chondrules (Fig. lb,c). As in other CM2 chondrites (GROSSMAN et al., 1988a). chondrules and chondrule fragments in Cold Bokkeveld are predominantly porphyritic olivine (PO) (Fig. la,b,c) and porphyritic olivinepyroxene (POP) varieties. However, barred olivine, radial pvroxene, and cryptocrystalline types also occur. Backscattered electron-imaging shows that among porphyritic chondrules, type-1 varieties (with MgOrich olivine) predominate over type-II varieties (with FeO-rich olivine). Chondrules and chondrule fragments range up to 1000 pm in diameter, but most are less than 400 pm. Statistical data for chondrule size in Cold Bokkeveld are given by KING and KING ( 1978). Fractured single crystals, mainly olivine, comprise 4.5 vol. % of the meteorite (Table 2) and range from < 1 pm to 500 pm in diameter. Single olivines have a compositional range of Fo~~.,_~~(Fig. 2), which is greater than any yet measured in a CM2 chondrite, but similar to that in Kaba (CV3) (SCOTT et al., 1988). Rounded, lithic fragments up to I500 pm in diameter are present in Cold Bokkeveld. They make sharp contacts against the enclosing matrix and are composed of CMZ-type material (Fig. Id). In addition to high temperature components, subhedral crystals of calcite, 20-60 pm dia., are abundant in the matrix (BARBER, 198 I ). Most are enclosed within sub-rounded bodies of serpentine and tochilinite (“poorly characterized phases” (PCP) of FUCHS et al., 1973). Semicircular sprays of coherently interstratified serpentine-tochilinite laths partially embay the margins of calcite crystals (LEE, 1993). indicating that at least one period of serpentine-tochilinite crystallisation followed calcite precipitation. Calcium sulphate is a widespread,

CAIs in Cold Bokkeveld

FIG. 1. (a) Type-11 ~~h}~~tic olivine chondrule enclosed by a thin dust mantle (DM). The long axis of the rhondrule is aligned parallel to a distinct planar fabric in the meteorite. Olivine phenocrysts along the lower right and middle left edges protrude from the chondrufe, suggesting that previously adhering mesostasis has been removed by abrasion (backscattered electron image). f b) PoIyminerali~ fragment comprising a large, strongly zoned olivine with adhering smaller oiivines and mesostasis. The fragment was presumably derived by disruption of a type-11 chondrule. It is encfosed by a dust mantle ( DM) and aligned parallel to the planar fabric of the meteorite (backs~ttered ejectron image). (ct Fragment of a disrupted type-11 chondrule partially enclosed by a dust mantle (DM). The dust mantle is layered and graded such that it coarsens outwards (backscattered electron image). (d) Rounded lithic fragment (diameter = 1000 GUI)composed of CMZ-type material. The contact between the fragment and the enclosing matrix is sharp (backscattered electron image). (e) Dust mantle (DM) fragment showing well-developed layering and grading. The fragment appears to have become detached from the macroscopic object it presumably once enclosed (backscattered electron image}. (f) Type II PCP grain enclosed by a thin, dark phyllosilicate-rich rim (backscattered electron image).

R. C. Greenwood

et al.

1) 2) 3) 4)

The number, percentage, and size range of the members of each type appear in Table 3a. The occurrence of accessory phases, secondary minerals, and dust mantles are summarised in Table 3b.

40N

Spinal

20-

Fo% FIG. 2. Histogram

of Cold Bokkeveld

matrix olivine Fo contents.

although minor, constituent of Cold Bokkeveld (LEE, 1993), where it fills matrix-cutting fractures. CAls, calcitised chondrules, and calcite crystals also contain calcium sulphate (LEE, 1993). In common with other CM2 meteorites ( MCSWEEN, 1979), matrix material is the major constituent of Cold Bokkeveld (82.7 vol 90, Table 2). Following SCOTTet al. ( 1988), we use the term matrix for the fine-grained, largely silicate material that is interstitial to chondrules, inclusions, and mineral grains. Also following SCOT et al. (1988) we place no arbitrary upper limit to the grain size of matrix but note that it generally comprises material finer than I pm. An exception to this are PCP grains (see below), which although coarser than other matrix components, are discussed here because they probably formed during parent body aqueous alteration (TOMEOKA and BUSECK, 1985; ZOLENSKY and MCSWEEN, 1988). The matrix of Cold Bokkeveld may be sub-divided into at least four components: 1) 2) 3) 4)

Spine1 inclusions, One hibonite inclusion, Spinel-pyroxene (sp-py) inclusions, and Spinel-pyroxene-olivine (sp-py-01) inclusions.

Dust mantles around CAIs, chondrules, etc.: Clasts of fine-grained dark material; Type-I and II PCP grains (TOMEOKA and BUSECK, 1985): Clastic matrix material ( METZLER et al.. 1992).

Each of these are discussed

in detail below:

Dust mantles ( MACPHERSON et al., 1985; METZLER et al., 1992) lo- 100 lrn thick are developed around most macroscopic objects (Fig. I a,b,c). They postdate the main phase of chondrule disaggregation (Fig. lb,c) are dark, and rich in phyllosilicates. They have a distinct fabric, concentric with the margin of the central object and are often layered, with the grain-size increasing outwards ( Fig. 1c,e ) Many mantles are partially disrupted (Fig. 1c ), and it is common to find lumps of this material which appear to have been detached from any central object (Fig. le). Angular to sub-rounded dark clasts, up to 1.2 cm dia., are an important matrix component in Cold Bokkeveld (GREENWOOD et al., 1993a). Their composition is similar to that of dust mantles, but the clasts lack a discernible structure. Sub-rounded grains, up to 100 Frn dia., of tochilinite (Type-I PCP), or intimate mixtures of tochilinite and serpentine (TypeII PCP) are present in the matrix (Fig. If) (TOMEOKA and BuSECK, 1985 ). PCP grains are often surrounded by a thin rim of Mg-rich phyllosilicate (Fig. If). Much of the matrix is an intimate mixture of mineral fragments and various types of phyllosilicate. METZLER et al. ( 1992) describe such areas as “elastic textured” and consider them to have been produced by late-stage regolith processes. In many areas of elastic matrix dust mantles have been partially (Fig. lc) or completely removed from the macroscopic objects. A planar fabric is present in some areas of elastic matrix, indicative of continuing compaction at a late stage in the development of the meteorite. CAls in Cold Bokkeveld On the basis of their primary silicate and oxide assemblages the CAIs located in this study have been classified into four principal types:

inclusions

These comprise blocky (Fig. 3a), fragmental (Fig. 3b), and banded varieties (Fig. 3~). Perovskite, as rounded grains < 5 pm dia., is a common accessory (Table 3b). Blocky varieties often contain large internal voids (Fig. 3a), with up to 30% porosity. Many spine1 inclusions are fragments of larger objects, with the main period of fragmentation predating the formation of the dust mantle. This is demonstrated by a fragmental spine1 inclusion (Fig. 3b) which is completely enclosed by a dust mantle. Most spine1 inclusions contain a 2-10 pm-wide layer of Fe-rich phyllosilicate (Table 3b), which may be complete (Fig. 3a), or discontinuous (Fig. 3b). In other inclusion types, this layer separates the diopside-rich portion of the rim sequence ( WARK and LOVERING, 1977) from the spine1 core (Fig. 4a,d). The outermost portion of the rim sequence, including the pyroxene layer, has been removed from many spine1 inclusions, apparently by abrasion. The Fe-rich phyllosilicate layer presumably formed by aqueous alteration of precursor material ( MACPHERSON et al., 1983. 1988), whose identity is discussed later. Banded spine1 inclusions are the least common variety. In the example shown (Fig. 3c), a core of calcite ( + perovskite) is enclosed by a band of spine1 (+ perovskite) up to 20 Frn wide, external to which is a 5-50 pm wide zone of tine-grained phyllosilicate of variable Mg/Fe ratio. The object has a sinuous morphology, suggesting that an apparently isolated area of spine1 and carbonate within the dust mantle may have been connected to the main body of the inclusion outside the plane of the section. This CA1 is unusual in retaining its complex morphology, although it has been pervasively altered. The abundance of secondary calcite in this object is reminiscent of the Murchison Blue Angel inclusion (ARMSTRONG et al., 1982). Hibonitr

inclusion

Among the CAIs we studied, hibonite is the major primary phase in only one. It occurs as anhedral grains, enclosing perovskite and rimmed by Fe-rich phyllosilicate (Fig. 3d). Minor spine1 occurs as anhedral grains interstitial to hibonite. Spindp_vroxenr

(sp-p_v) inclusions

These are morphologically more diverse than spine1 inclusions, and comprise fragmental (Fig. 4a), nodular (Fig. 4b), granular (Fig. 4c), and intricately lobate varieties (Fig. 4d). Unlike fragmental spine1 inclusions, fragmental sp-py inclusions have a selvedge of pyroxene, often only a few microns thick, outside the Fe-rich phyllosilicate layer (Fig. 4a). Both types were clearly derived from larger, disrupted Table 3a. Number, Bokkeveld

percentage

Inclusion Type

Number

Spine1 spherules Spine1 inclusions Hibonite inclusion Sp-py inclusions Sp-py-01 inclusions Spinel-bearing olivines

4 153 1 166 13 8

Total

345

Sp- spinel, py- pyroxene,

and size of CAls in Cold

ol- olivine

Percentage

1.2 44.3 0.3 48.1 3.8 2.3 100.0

Size Rangeurn) 50-108 1 O-378 77 28-510 66-438 70-400

1917

CAIs in Cold Bokkeveld

Table 3b. Accessory

and secondary

phases, accretionary

Hibonite Inclusion Type*

No.

%

No.

Spine1 spherules (4) Spine1 inclusions (153) Hibonite inclusion (I 1 Sp-py inclusions (166) Sp-py-01 inclusions (I 3) Spinel-bearing olivines (8)

2 17 1 5 0 0

50 11

2 86 1 67 3 0

3

rims

Perovskite

% 50 56 40 23

Fe-phyllosilicate No. %

No.

%

2 116 1 115 7 0

1 14 0 12 0 0

25 9

50 77 69 54

Calcite

7

Dust Mantles No. % 2 50 110 72 1 125 75 12 92 6 100

* Figures in brackets are the total number of CAls in each group (see Table 3a) Sp- spinel, py- pyroxene, ol- olivine % Figures refer to specific inclusion types, i.e. 23% of sp-py-01 inclusions contain perovskite

objects. Nodular sp-py inclusions do not show the development of an Fe-rich phyllosilicate layer, so pyroxene abuts spine1 (Fig. 4b). These inclusions frequently possess rounded cavities containing FeNi metal nuggets. The lack of Fe-rich phyllosilicate and the presence of Fe-Ni metal in nodular sp-py inclusions indicate that they have experienced less secondary alteration than other varieties. Granular sp-py inclusions are a previously unrecognised group of CAIs. They consist of loosely packed spine1 and pyroxene grains in a finely crystalline matrix of phyllosilicates. These CAIs are delicate objects and are unlikely to survive the freeze-thaw procedures used in previous studies. The boundary between granular sp-py inclusions and the surrounding matrix is generally indistinct. The spine1 to pyroxene ratio is highly variable, and in one example (Fig. 4c) spine1 forms less than 10% of the inclusion. Sp-py inclusions sometimes display complex morphologies (Fig. 4d). Here, only small remnants of the outer pyroxene layer are preserved, but the form of the adjacent Fe-rich phyllosilicate layer indicates that the intricately lobate shape of this inclusion is primary. Spinel-p~roxeneolivinr

(sp-p.v-01) inc~lusions

In Cold Bokkeveld sp-py-01 inclusions are either banded (Fig. 4e) or granular (Fig. 4f). In banded inclusions. spine1 occurs as ribbons, 5-15 pm wide, enclosed directly by pyroxene with no intervening Fe-rich phyllosilicate layer. Olivine of variable grain size abuts against pyroxene (Fig. 4e). Fe-rich phyllosilicate is present locally as anhedral masses interstitial to olivine. Fe-Ni-Pt metal nuggets, up to IO pm dia., partially fill cavities in some olivine-rich areas. Granular sp-py-01 inclusions are rounded objects = 400 pm dia. (Fig. 4f), composed of small spine1 and pyroxene grains enclosed in a matrix of finely crystalline phyllosilicate. predominantly Mg-rich serpentine. Forsterite crystals. also enclosed in Mg-rich phyllosilicate, are generally concentrated around the margins of the inclusions. In addition to the four principal types, spine1 spherules and spinelbearing isolated olivines are distinguished as two minor types of inclusions.

Spine1 and hibonite-bearing refractory spherules in CM2 chondrites ( MACDOUGALL 1981; MACPHERSON et al., 1983) are rare. Only four were encountered in this study. All are composed predominantly of spine1 (Fig. 5a). Hibonite laths, up to 20 pm long, form 5-10% of one, but the mineral is absent, or a trace constituent in the others. Void space in the spine1 spherules varies from 0 to 30% in area. One spherule has a continuous rim, 5-10 pm thick, comprising an outer zone of diopside separated from the spine1 core by a layer of Fe-rich phyllosilicate (Fig. 5a). Spin&earing

MINERALOGY

AND CHEMISTRY BOKKEVELD CAls

OF COLD

Spine1 Spine1 occurs in every inclusion studied. The Fe0 content of spine1 in all but sp-py inclusions is less than 2.0 wt% (Fig. 6a, Table 4). A minority of sp-py inclusions contain Fe-rich spinels with up to 1 I .8 wt% FeO, but in most it is less than 2.0 wt%. Spine1 grains within isolated olivines have CrZOs contents between 0.6-1.2 wt%. In sppy-01 inclusions, Cr,O, may be as high as 3.0 wt%, but is generally less than 1 wt%. In sp-py and spine1 inclusions, CrzO, is normally less than 1%. Vz03 is up to 0.8 wt% in some sp-py inclusions, but is more generally in the range 0.2-0.4 wt%. One drawback of investigating CAIs in situ is that the volume sampled is small, and possibly unrepresentative. The total area examined was 588 mm’ (Table 1). Assuming that the sections are 30 pm thick and using 2.75 g cme3 for the density of CM chondrites ( WASSON, 1974). the total mass examined was 0.05 g. To investigate spinels from a larger sample, we obtained a portion of a spinel-rich residue that had been prepared by acid dissolution of 68.5 g of Cold Bokkeveld. The procedures used are detailed by ASH et al. ( 1990). Fe0 contents of 106 spine1 grains from the residue are shown in Fig. 6b. Like those measured in situ, most (87%) spinels in the residue contain less than 2.0 wt% FeO. However. the remainder show a wider range of values (up to 24.3 wt% FeO) than those measured in situ. This presumably reflects the much larger volume of material used to produce the residue, and indicates that Cold Bokkeveld contains CAIs, as yet unidentified. with spinels of higher Fe0 content than those located in situ. Hibonite Hibonite [Ca(AI, Mg, Ti),,O,,J is a major component in only one inclusion (U98. Fig. 3d); in others it is an accessory (Table 3b) and generally occurs as laths enclosed by spinel. This paragenesis is reflected in its mineral chemistry. Hibonite in U98 (Table 4) has 0.80 wt% MgO and I .67 wt% TiOz, whereas accessory hibonite has 2.1-3.8 wt% MgO and 4.2-6.0 wt% TiOz. Similar variations, depending on whether hibonite is a major or minor phase, have been reported from Murchison ( MACPHERSON et al., 1983; IRELAND, 1988). MgO and TiOz variation in Cold Bokkeveld hibonite is consistent with the operation of the coupled substitution ( Ti4+ + Mg’+) for 2A13+ ( MACPHERSON et al., 1983). Fe0 contents range from 0.4-l .5 wt%; VzO, is generally less than 0.8 wt%. Perovskite The perovskite grains were too small for quantitative appear to be nearly pure CaTi03.

analysis,

but

isolated &vines

Some forsterite-rich isolated olivine grains contain euhedral, 5-35 pm dia., Mg-rich spinels (Fig. 5b), with a composition similar to that of spine1 in CAls and unlike that in ferromagnesian chondrules (see below).

Pyroxene Pyroxene rims on sp-py and sp-py-01 inclusions, although generally narrow. are usually strongly zoned, from Al- and Ti-poor diopside at the margins, to Al- and Ti-rich fassaite in the interior, close to the

1918

R. C. Greenwood

et al

FIG. 3. (a) Blocky spine] inclusion enclosed by dust mantle (DM), The spine1 (SP) core contains numerous large internal voids and small perovskite (Pv) grains. It is surrounded by a continuous bright rim of Fe-rich phyllosilicate ( Fe-Ph) (backscattered electron image). (b) Fragmental spine1 inclusion enclosed by a dark dust mantle (DM). The spine1 (Sp) core contains a single perovskite grain (Pv), and is enclosed by a discontinuous Fe-rich phyllosilicate rim (Fe-Ph). In the upper left corner of the inclusion the Fe-rich phyllosilicate rim terminates abruptly, indicating that it was fragmented prior to deposition of the enclosing dust mantle (backscattered electron image). (c) Banded spine] inclusion enclosed by a coarse-grained dust mantle (DM). The core of the inclusion is composed of large grains of blocky calcite (Cal) and smaller bright perovskite (Pv) grains. The core is enclosed by a multiple rim consisting of an inner layer of spine1 (Sp) and an outer layer of fine-grained phyllosilicate ( Mg-Ph) (backscattered electron image), (d) Hibonite inclusion partially enclosed by a dust mantle (DM). The inclusion is composed of a loosely packed aggregate of anhedral hibonite grains (Hib) and minor spine1 (Sp), both containing small, rounded perovskite (Pv) grains. A thin discontinuous rim of Fe-rich phyllosilicate ( Fe-Ph) surrounds the hibonite crystals (backscattered electron image 1.

contact either with spine1 or Fe-rich phyllosilicate. Representative analyses appear in Table 5. The AlzOx and Ti02 contents range up to 20.4 and 8.7 wt%, respectively, in sp-py inclusions and up to 22.8 and 8.6 wt%, respectively. in sp-py-01 inclusions. The pyroxenes from both types of inclusions display similar, overlapping trends on Al vs. Si and Ti vs. Al plots (Fig. 7a,b). Clinopyroxenes in CAIs are solid solutions of diopside (CaMgSi206). Ca-Tschermak’s molecule (CaAlzSi06), Ti-Tschermak’s molecule ( CaTi4+A1206), and a trivalent Ti-bearing component (expressed as CaTi “AISi06) ( STOLPER, 1982; SIMON et al., I99 1). The presence of these components can be identified by plotting Al vs. Si (Fig. 7a). Varying degrees of substitution of the other endmembers for diopside would result in points that plot along the tie lines joining the endmember compositions. The pyroxenes do not plot on a single tie-line. At high Si-contents diopside IS replaced by Ca-Tschermak’s molecule, but as Si decreases, substitution is by a mixture of CaTschermak’s and Ti-bearing pyroxene components.

Olivine Olivines

in

sp-py-01

inclusions

are

nearly

pure

forsterites

( Fo~~.~_,~). CaO contents lie in the range 0.1-0.5 wt%; Al,03 is generally less than 0. I wt%. Spinel-bearing isolated olivines are also nearly pure forsterites (Fog9 5_1oo). In comparison to olivines in sp-py-01 inclusions, these have significantly higher A1203 contents of up to 0.4 wt%, and slightly higher CaO contents of up to 0.6 wt%. Representative olivine analyses are given in Table 5. Phyllosilicates The phyllosilicates are mineralogically and chemically heterogeneous on a nanometre scale. Electron diffraction, lattice imaging, and X-ray analysis show that their mineralogy and morphology correlate well with chemical composition. Three mineralogical-chemical groups are distinguished:

CAIs in Cold Bokkeveld

RG .4. (a f Fragmental sp-py inclusion enclosed by dust mantle ( DM ) . A core of spinet( Sp) is enclosed by a nearcontinuous rim of Fe-rich phyllosilicate (Fe-Ph), exterior to which is a discontinuous pyroxene (Py) selvedge (backscattered electron image). (b) Nodular sp-py inclusion enclosed by dust mantle (DM). Porous areas of spine1 (Sp) are partially enclosed by pyroxene (Py). A number of rounded cavities are present containing nuggets of Fe-Ni metal (Fe-Ni) (backscattered electron image). (c) Granular sp-py inclusion composed of disseminated pyroxene (Py) and spine! (Sp) grams set in a phyllo~li~ate-huh mesostasis. In this example spinel forms only a minor component. Although there is evidence of a dust mantie (DM) in places, the contact between the inclusien and the su~ounding meteorite matrix is fairly indistinct (backscattered electron image). (d) Intricately lobate sp-py inciusion completely enclosed by a well-developed dust mantle (DM). This inclusion comprises a core area of calcite (Cal) and small anhedral pyroxene grains (Py) surrounded by complexly shaped areas of granular spine1 (Sp). A thin layer of Fe-rich phyllosilicate (Fe-Ph) is present along much of the inclusions margin. External to this locally is a selvedge of granutar pyroxene. Despite the preservation of a complex external morphology, this inclusion ctearty underwent a period ofdisaggregation prior to deposition of the dust mantle. Thus, the well-developed spine1 lobe in the upper left corner of the photo is sharply truncated along its contact with the dust mantle (backscattered electron image). (e) Banded sp-py-01 inclusion. Spinet (Sp) occurs as ribbons enclosed by pyroxene f Py). Anhedrat oiivine abuts the pyroxene bands and is enctosed locally by anhedrai masses of Fe-rich phytiosii~cate fFe-Ph) (backscattered electron image). (f) Granular sp-py-of inclusion. Large blocky olivines (01) and granular masses of pyroxene (Py) and spine1 (Spf are set in a Mg-phyllosii~cate rich ( Mg-Ph ) groundmass ( backscattered electron image f .

1919

1920

R. C. Greenwood et al.

Frc. 5. (a) Refmctory spherule composed of spine1 (Sp) and minor hibonite (Hibf. The spherufe has a high internal porosity with Fe-Ni metal (Fe-Ni) nuggets present locally within the voids. The spherule is enclosed by a rim sequence consisting ofan inner layer of Fe-rich phyllosilicate ( Fe-Ph) surrounded by an outer layer of pyroxene (Py) (backscattered electron image). (b) Spinel-bearing isolated olivine. Euhedral spine1 (Sp) crystal partially enclosed by large olivine (01) fragment. Remnants of a dust mantle (DM) are present locally (backscattered electron image). (c) Spine1 inclusion showing evidence of late-stage deformation. The inclusion is enclosed by a dust mantle (DM) and shows the development of a bright Fe-rich phyllosilicate rim (Fe-Ph) surrounding the spinel-bearing (Sp) areas. The long axis of the inclusion runs parallel to the prominent fabric which wraps around the large chondrule on the left-hand side of the field of view. Deformation of the inclusion took place after incorporation into the parent body and clearly postdates formation of the dust mantle (backscattered electron image). (d) Rotational deformation of a spine1 inclusion. In this example, a portion of the inclusion appears to have been sheared off and rotated parallel to the prominent north-south fabric seen on the right-hand side of the field of view. The inclusion is partially enclosed by a dust mantle (DM) and shows well-developed, Fe-rich phyllosilicate (Fe-Ph) rims around spine1 (Sp) (backscattered electron image). 1. Fe-serpentine, found in spine1 spherules, spine1 inclusions, and sp-py inclusions; 2. Fe- and Mg-serpentines. found in sp-py inclusions; 3. Mg-serpentine, found in sp-py-01 inclusions. F~wqwntine forms a well-defined layer in the rim sequences of many CAIs (Figs. 3a,b; 4a,d; 5a). Backscattered electron images show that the phyllosilicate layer is normally uniform, although a fibrous structure is sometimes discernible, with long axes normal to the contact between the phyllosilicate rim and spine1 core. TEM shows that the phyllosilicate layer is dominated by ~600-1000 nm by = 100-200 nm phyllosilicate laths (Fig. Sa). The long axes of adjacent laths commonly show parallel alignment, and in any one area share a similar crystallographic orientation. The phyllosilicates have a predominant ~0.71 nm lattice fringe spacing, although crystats with a a 1.40 nm spacing are occasionally seen, the basal row diffmction spots from which have a m~uIated intensity. Diffraction patterns

ofthe phyliosjlicate laths commonly show a continuous streaking of k # 3 reflections parallef to C *, indicative of stacking faults in the lattice (Fig. 8b). These phylIosili~tes are c~stailographi~ily similar to cronstedtites in CM2 matrices ( MOLLERet al.. 1979: BARBER, 1981). In Cold Bokkeveld, the Fe-serpentine group contains Fe and Si with minor Mg, AI, and occasionally Ca (Table 6), typical of meteoritic cronstedtite ( MOLLERet al.. 1979). Most rim phyllosilicate crystals are close to endmember cronstedtite (Fig. 9a,c). but more Mg-rich analyses trend towards the composition ofdust mantle phyllosilicates (Fig. 9a). Some phyllosilicates have minor concentrations of S and Ni. images of this phyllosilicate show cronstedtite laths (Fig. 8c.d) with intervening phyllosilicate crystals that have a Q 1.08 nm and x0.54 nm lattice fringe spacing, consistent with tochilinite (ZoLENSKY and MACKINNON~ 1984). Fe- & ~~~-,~~rpe~tj~e~ are abundant as inte~ranular phases within some granular sp-py inclusions. They comprise -0.7 nm serpentine

CA!s in Cold Bokkeveld

earing 01s I Sp-Py-01Inca R lncs Spine1lncs

0

10

5 Fe0

1921

calcite, calcium sulphate, and cronstedtite are also intergrown with the phyllosilicate, much of which is very finely crystalline (Fig. !Oe). and gives ring diffraction patterns with d-spacings consistent with lizard&e (Fig. IOf). However, in many areas phyllosilicate fthres. 215 nm in width. also occur (Fig. IOe). These produce diffuse spots superimposed on rings and have lattice fringes with u ‘= 0.73 nm spacing (Fig. IOe). The fibres may be straight. sinuous, or wavy and some have circular. chrysotile-like cross-sections. These crystals have similar morphology and d-spacings to some illustrated by Fig. Id of BARBER( I98 I ). from CM matrices. Compositionally. the Mg-rich serpentine group differs from others analysed during this study (Fig. 9a.c). Most analyses show minor concentrations of S ( z 3.0 wtX S03) and Ni ( c 1.5wt%NiO) (Table 6). which suggest the presence of tochilinite layers within the lizardite lattice. but no lattice discontinuities were imaged. However. rare crystals with a * I .78 nm lattice fringe spacing do occur, and have a structure consistent with coherently-interstratified serpentinc-tochilinite.

Wt% Calcite

90

Calcium carbonate occurs in 8% of the CAls (Table 3b). Electron diffraction indicates that it is predominantly calcite. not aragonite or vaterite. Electron microprobe analyses show the calcite to be pure CaCO,: minor concentrations of Fe and other cations are usually attributable to silicate inclusions, which is confirmed by ATEM.

80.

Calcium Sulphate

Calcium sulphate is an uncommon, finely crystalline phase within CAls, and it is mostly intimately intergrown with calcite. where it occludes intra-calcite dissolution and fracture pores ( LEE, 1993 ). The sulphate was identifi~ by TEM as an intergrowth of hemihydrate (CaS04*0.SH,0) and anhydrite (CaSO,). which was probably formed by dehydration ofongrnal gypsum. Gypsum occurs intimately associated with calcite in CAls from Murchison (ARMSTRONG et al.. 1982: MACPHERSONet al.. 1983) and Essebi (EL GOKESVet al.. 1984).

Fe0

Wt=%

Fta. 6. (a) Histogram of Fe0 content in spine! from Cold Bokkeveld CAls. (bf Histogram of Fe0 content in spine1 from a residue prepared by acid dissolution of 68.5 g of Cold Bokkeveld.

in a range of crystal habits, Most of this phyllosilicate forms straightsided laths = 700-800 nm by ~75-100 nm with a = 0.71 nm lattice fringe spacing. Between the laths lies a mass of porous, poorlycrystalline phyilosilicate (Fig. lOa). The laths give diffraction patterns typical of cronstedtite, including a well-developed streaking of k f 3 reflections in some. The poorly crystalline phyllosilicate comprises short, wavy crystallites a few tens of nanometres in width (Fig. 1Ob). intergrown with crystals that have circular cross-sections, 20-25 nm in diameter, typical of chrysotile (Fig. lOc,d). The poorly crystalline phyliosiiicate crystals have variable lattice fringe spacings, partly attributable to layers of tochi!inite/brucite inte~tratified with ~0.7 nm serpentine. The structure of this phyllosilicate is similar to that of”sponge” in the matrix of Nawapali, CM2 (BARBER,198 I ), which comprises plate-like crystals embedded within randomly oriented crystallites incorporating chrysotile fibres. Rare hemispherical Povlenlike sectored phyllosilicate crystals also occur within granular sp-py inclusions. Chemically, phyllosilicates from three granular sp-py inclusions vary in Mg-Fe from a cronstedtite-like composition to values more typtcal of dust mantles and matrix (Fig. 9b: Table 6). This is interpreted to reflect differences in the degree of admixture of an Fe-rich with a more Mg-rich phyllosilicate. In agreement with BARBER ( 1981). the coarse, lath-shaped crystals are inferred to be Fe-rich, whereas the poorly-crystalline sponge is more Mg-rich. f~~-.~~~r~~~?ti~es enclose high-tempe~ture mineraIs within granular sp-py-01 inclusions (Fig. 4f). In TEM, forsterite in particular is seen to have a highly irregular contact with phyllosilicate. Pentlandite,

~~agnes~urn Isotopes Spine! and olivine proved to be the only inclusion-free primary phases suitable for ion microprobe analysis. Their low Al/Mg ratios precluded the identification of radiogemc “Mg* produced by decay of extinct “Al, as is common in CAls ( WASSERBURT~, I985 ). The results of Mg isotopic measurements are given in Table 7. All data are from spinel. except for CBOLl and CBOL2. which are from otivine. An estimate of the mass-dependent fractionation intrinsic to the sample ( I;‘Mg)may be determined from the expression

where aZSMg = [(25Mg+/24Mgt)/0.12663 - I] X lOOO(%n/AMU); 0. I2663 = the accepted terrestrial value for 2sMg/2’Mg ( CATANZARO et al.. 1966): and -1Z5Mg.sm = the average of replicate analyses of the Burma spine1 standard calculated for each of the five days when data were collected (see Table 7 for values 1. The values of FMgfor each inclusion are plotted on Fig. I I. All of our individual measurements of Burma spine1 yielded FMg values between --4%0 and +4%0, delimited by the vertical lines on Fig. I 1. Most inclusions show no evidence of intrinsic mass-dependent fractionation, the F,, values falling within the +-4%0range (Fig. I I ). An exception is the spine1 inclusion COLDBl3. With &i, values of 5.6 + l.6%0 and 7.8 f I&&, it is clearly enriched in the heavier Mg isotopes. The isolated olivine CBOL has normal _FMpvalues (2.9 + 1.4%0and 3.8 f 1.9%0), whereas the spine1 it encloses has ~~~ = 10.3 ? 1.9%0,so is significantly enriched in the heavier Mg isotopes. Isotopic heterogeneity between coexisting spine1 and olivine, like that in CBOL. is a common feature of a group of objects in 0’3 and CM2 chondrites termed plagioclase-ofivine inclusions ( PO!s) (SHE~VC; et al.. 1991 ). Our isotopic data indicate that some isolated olivincs in Cold Bokkevetd may have come from disrupted POIs.

1922

R. C.

Greenwood et al.

Table 4. Spine1 and hibonite analyses. Phase

Spine1

Spine1

Spine1

Spine1

Spine1

Spine1

Spine1

Hib.

Hib.

Inc.#

Dll

D13

D44

DZ8

036”

U25*

Ull

U98

u57

MgO

26.88

26.73

27.98

24.01

20.07

28.70

27.63

0.80

2.98

A’& SiO,

77.29

70.54

70.54

67.02

64.96

70.30

70.55

88.29

81.77

0.19

0.90

-

0.70

1.10

0.40

0.18

0.14

0.27

0.20

-

8.48

0.40

0.34

1.67

5.96

0.09

0.80

cao TiO

0.26

a.17

0.30

0.55

V&23

0.45

0.26

0.34

0.70

n.d.

0.30

0.23

C’,Oo MnO

0.29

0.23

0.1t

0.20

0.60

-

0.64

0.08 8.34

0.07

0.12 0.31

0.67

0.23

5.40

11.35

0.50

99.67

99.52

99.50

98.84

98.35

100.60

M9 Al

0.96

0.96

1 .oo

0.89

0.77

1.01

0.98

2.00

2.00

1.98

1.96

1.97

1.96

1.97

Ca ii

0.00

0.00 0.00

0.00

0.00

0.01

0.01

0.00

1.02

1 .Ol

0.00

0.01

0.02

0.00

0.01

0.01

0.14

0.51

V

0.01

0.01

0.01

0.01

0.00

0.01

0.00

0.01

0.07

Cr

0.01

0.01

0.00

0.00

0.01

0.00

0.01

0.00

0.01

Fe0 Total

1.11 100.68

0.54

0.50

99.87

100.50

0.13 11.67

10.88

Number of cations per formula unit 0.50

Mn

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Fe

0.01

0.01

0.00

0.11

0.24

0.01

0.02

0.05

0.05

Total

2.99

2.99

3.00

2.99

3.00

3.01

2.99

13.02

13.03

0

4.00

4.00

4.00

4.00

4.00

4.00

4.00

19.00

19.00

indicated

(“1

All analyses

in wt%;

ali analyses

WDS

(-1 not detected,

n.d. not determined

Spine! inclusions

Dl 1, 013,

Sp-py-01 inclusion

U25;

044,

U57;

spinet-bearing

except

where

sp-py inclusions isolated

028,

D36

olivine Ul 1; hibonite

inclusion

U98.

Table 5. Pyroxene and olivine analyses. Phase

pyrox.

pyrox.

pyrox.

pyrox.

pyrox.

pyrox.

Inc.#

U80-lx

U36’

U80-4’

U81

u2

U38* 9.71

f&O

17.66

17.18

16.28

15.29

14.90

oliv.

oiiv.

oliv.

U81 56.16

U2 56.62

56.89

Ull

U65 56.33

2.90

4.32

6.03

8.94

15.42

16.75

0.10

0.22

SiO

53.84

52.84

50.84

49.41

42.47

40.10

41.76

42.43

41.84

41.65

CaO2

24.64

24.31

24.20

25.14

23.73

24.85

0.44

0.12

0.13

0.59

0.61

1.20

1.27

2.16

7.41

V&

0.10

0.19

0.42

Cr,O, MnO

0.11

0.08

AP, TiO

0.09 0.25 0.09

0.77

1.30

0.14

0.11

0.16

0.68

0.61

0.16

0.86

0.77

0.80

0.39

0.56

99.81

100.56

99.23

100.87

99.11

100.10

99.22

100.38

99.29

99.55

Mg

0.95

0.92

0.88

0.82

0.81

0.53

1.99

1.98

2.00

1.99

Al

0.12

0.18

0.26

0.38

0.67

0.73

0.00

0.00

0.00

0.01 0.99

Fe0 Total

Number of cations per formula unit

Si

1.94

1.90

f .85

1.77

I.56

1.47

0.99

0.95

0.94

0.95

0.97

0.93

0.98

0.99 0.01

0.99

Ca

0.00

0.00

0.01

Ti

0.00

0.02

0.03

0.03

0.06

0.20

0.00

0.00

0.00

0.00

V

0.00

0.00

0.00

0.00

0.01

0.01

0.00

0.00

0.00

0.00

Cr

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Mn

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Fe

0.03

0.04

0.02

0.02

0.00

0.03

0.02

0.02

0.01

0.01

Total

3.99

4.00

3.99

3.99

4.04

3.01

0

6.00

6.00

6.00

6.00

6.00

3.95 6.00

2.99 4.00

3.00 4.00

3.01 4.00

All analyses

in wt%;

all analyses

WDS

except

where indicated

4.00

1”)

(4 not detected sp-py inclusions spine&bearing

U80-1,

isolated

U80-4, olivines:

U36,

U38;

UI 1, U65

sp-py-01 inclusions:

U81,

U2;

-

4

4

m

4

0

4

4

m

0

4

I

I

I

I

I

I

I

I

I

1924

R. C. Greenwood et al.

FIG. 8. (a) Bright-field TEM image of coarse cronstedtite crystals rimming a sp-py inclusion. in this image, the darker the crystal, the more strongly it is diffracting electrons. (b) Selected area diffraction (SAD) pattern of rim cronstedtite comparable to that in (a). Spots along the basal row have a d-spacing of -0.7 I nm (001 )_ The continuous streaking of k # 3n spots parallel to c* is due to stacking faults in the lattice. (c) Bright-field TEM image and SAD pattern (inset) of cronstedtite and tochilinite in the Fe-serpentine rim of a sp-py inclusion. The strongly diffracting (dark) crystal is cronstedtite, with tochilinite just below. This is also shown by the SAD pattern which has reflections from both cronstedtite (~0.7 1 nm) and tochilinite ( m 1.08nm). (d) High-resolution, bright-field TEM image and SAD pattern (inset) of the interface between ~0.7 1 nm cronstedtite and = I .08 nm tochilinite in (c).

The sp-py inclusions COLDBIT and, to a lesser extent, COLDB7 display a range of FMgvalues. Spinel-rich areas in both have intricate shapes, so some secondary ions possibly were sputtered from the surrounding pyroxene and phyllosilicate-rich material.

Table 6. Phyllosili~ate

DISCUSSION We examined, in situ, 345 inclusions in Cold Bokkeveld. The textural environment of each has been preserved to yield

mineral analyses (ATEM data).

Fe-serpentine

MS-Fe serpentine

Mg-serpentine

PCB.07

P5.01

9PY.01

PY6.01

7PY.01

MgO

9.35

9.96

7 1.07

25.27

36.21

A’P,

3.25

3.08

9.57

5.71

27.33

27.10

27.20

35.38

42.08

1.39

1.86

3.74

0.41

0.33

0.34

SiO,

0.79

SO, CaO

0.55

TiO,

0.32

MnO Fe0

60.07

59.07

50.35

30.90

NiO

Total

1.57

14.20 1.55

7 00.00

All analyses in wt%;

100.00

100.00

data normalised to 100 wt%

100.00

100.00

CAIs in Cold Bokkeveld

Fe

Fe

1925

Fe

Si FIG. 9. Analytical TEM analyses (wt%) of Cold Bokkeveid CA1 phy~losilieates. (a) Fe-rich phyl~osi~~~ateCA1 rims (filled triangles), phyllosilicates in dust mantles surrounding CAfs (open circles). Mg-serpentines in sp-py-01 inclusions (filled squares). (b) Fe- and Mg-serpentines in granular sp-py inclusions (filled stars). (c) Fields for the phyllosilicates plotted in (a) and (b) compared to bulk CB matrix (inverted triangle). and endmember tochilinite, cronstedtite, and Mg-serpentine (filled stars). Field 1: Mg-serpentine in sp-py-01inclusions, field 2: dust mantle phyllosilicates, field 3: Fe-rich phyllosilicate CA1 rims, field 4: Fe- and Mg-serpentmes in granular sp-py inclusions.

information about its post-formational history, so that we can construct an outline chronology for the CAIs, namely: 1) Formation of CAIs, 2) Fragmentation of CAIs and formation of dust mantles, 3) incorporation and reprocessing within a parent body; 3a) Aqueous alteration and formation of Fe-rich phyllosilicate rims. Fe- and Mg-serpentines in granular sp-py inclusions, Mg-serpentines in sp-py-of inclusions, calcite and calcium sulphate; 3b) Regolith deformation. In the following discussion, each of these events will be examined in reverse chronological order, our aim being to “remove” sequentially the effects of secondary processing to gain an understanding of the conditions that prevailed during the formation of CAIs. Incor~~tion and Reprocessing Aqueous Alteration

within a Parent Body:

On the basis oftheir mineralogy and chemical composition, three different groups of phyllosilicates have been identified in Cold Bokkeveld CAIs. Phyllosilicates very similar to those of the Fe-serpentine group have been described by MACDOUGALL(1981) and MACPHEKSONet al. (1983, 1984) in the rims of Murchison CAls. The other two groups, Fe- and Mg-serpentine and Mg-serpentine, have never previously been described from CAIs in CM2 meteorites. The mineralogy and chemical composition of the Fe-serpentine and Feand Mg-serpentine groups (predominantly serpentine with minor to~hiiinite) is very similar to that of phyllosilicates which constitute much of the dust mantles and matrix in

Cold Bokkeveld and many other CM meteorites. Thus, the origin of the phyllosilicates within CAIs has to be considered with regard to these other contexts. There are two conflicting views for the origin of phyllosilicates in CM meteorites: ( 1) hydrous alteration in the solar nebula ( METZLERet af., 1992) and (2) parent body aqueous processes (KERRIDC~Eand BUNCH, 1979; MCSWEEN, 1979, 1987: BUNCHand CHANG, 1980;BARBER,198 1; TOMEOKAand BUSECK,1985; ZOLENSKY and MCSWEEN, 19S8; ZOLENSKYet al., 1993). METZLERet al. ( 1992) have argued that dust mantle phyllosilicates were formed by hydration of fine silicate grains in a nebular environment. The evidence they cite in favour of this model includes the juxtaposition within dust mantles of phyllosilicates and phases susceptible to alteration, such as forsterite, mesostasis glass, and tiny metal grains. Such petrographic relationships are interpreted by METZLER et al. ( 1992) to indicate that dust mantle phyllosilicates must have formed before they came into contact with the unstable phase. i.e., hydration of the precursor to the phyllosilicates took place in the solar nebuia prior to accretion. That CM meteorites comprise a disequilibrium assemblage of hydrous and anhydrous objects and particles is not in question. However, we do not consider that the coexistence of hydrous and anhydrous minerals, however intimate. says anything significant about the location in which aqueous alteration took place. It simply demonstrates that alteration did not reach completion. The controls on which materials were altered, and to what extent, were no doubt highly complex. and would have included factors such as micron-scale differences in porosity and permeability. and spatial and temporal fluctuations in the chemical composition of the fluid. We consider that most phyllosilicates in Cold Bokke-

1926

R. C. Greenwood

et al.

FIG. 10. (a) Bright-field TEM image of Fe-Mg serpentine phyllosilicates from the interior of a granular sp-py inclusion. In between the paraliel, coarse, strongly diffracting phyilosilicate laths is a more finely crystalline phylIosiljcate. (b) High-resolution, bright-field TEM image ofthe finely crystalline phyllosilicate from (a). The phyllosilicate is made up of numerous short, wavy crystals. (c) High-resolution, bright-field TEM image of the finely crystalline phyllosilicate (granular sp-py inclusion). Cross-sections of two chrysotile fibres can be seen (00 I = -0.73 nm) in addition to other phyllosilicate crystals. (d) High-resolution, bright-field TEM image showing a cross-section through one chrysotile fibre bright(granular sp-py inclusion). Note the concentric ~0.73 nm lattice fringes and an axial hole. (e) ~iigh-re~lution. field TEM image of Mg-serpentine phyllosilicates from the interior of a granular sp-py-01 inclusion. ( f) SAD pattern of finely crystalline Mg-serpentine (granular sp-py-01 inclusion). The ring closest to the central spot has a d-spacing of -0.73 nm (001).

veld, including those in CAIs, formed by aqueous alteration in a parent body environment. No unequivocal petrographic evidence has been found to support this assertion. However, the chemical and mineralogical composition of CA&hosted phyliosilicates is very similar to that of matrix phyllosilicates, for which there is a wealth of evidence that they formed by parent body aqueous processes. Formation of Fe-Rich PhyIlosiiic~te Rims

Spine1 and sp-py inclusions have rims whose phyllosilicates make sharp contact with the enclosed spine], perovskite, and

hibonite, even on the scale of TEM. The sharp contact indicates that the phyllosilicates did not form by alteration of these phases. The contact between the rim phyllosilicates and those of the dust mantle or meteorite matrix is also sharp (Fig. 3b), indicating that there was negligible reaction between them. The Mg-rich phyllosilicates that dominate the dust mantles and matrix [ Fe/Mg ratios of 1.84 and 2.40, respectively f METZLERet al., 1992; MCSWEENand RICHARDSON, 1977)] contrast strongly in composition with the Fe-rich phyllosilicates of the CA1 rims (mean Fe/Mg = e7.95) (Table 6). Given the rarity of metal in CA1 rims (LEE and BAR-

CAIs in Cold Bokkeveld

Table 7. Magnesium

1927

isotope data.

Fh%Jb tVo0) Spine1

Day

wo)

spherule 0.12543

CBSPI Spiflef

dz8Mg”

rt 20

0.13672

f 30

-2.1 k2.4

-0.1

4

inclusions

COLDBl3A

0.12613+17

0.13813+15

56rtr1.6

-0.7

3

COLDBI

0.12641

0.13889*16

7.8~~~1.6

0.3

3 3

3B

f 16

COLDB5A

0.725551t16

0.13698&21

1.01t:1.6

0.2

COLDB24

0.12581

0.13756rt16

3.0rt1.6

0.4

3

CBSPG

0.12605*73

0.13769&14

2.6f2.1

-2.5

4

f15

CBSP2

O.l2615z!zl4

0.13820*17

3.4k2.1

-0.4

4

CBSP7

0.12579&14

0.13717522

1.5*1.9

-2.2

5

0.12576&15

0.13736&19

1.2f2.0

-0.3

5

CBSPI 1 B

0.12585~~20

0.13732-+31

-2.0

3

COLDB7A

0.12618rt19

0.13847

1.1

2

COLDB7B

0.12547~:13

0.13682rt14

-0.2~1~1.5

0.5

2

COLDBl7A

0.12619~13

0.13801+20

5.5zt1.5

-2.4

2

COLDBl7B

0.12589

0.13763kl6

3.2~~1.4

-0.5

2

COLDBl7C

0.72632+?8

0.13871

6.6f1.8

0.4

2

COLDBl7D

0.12514+20

0.136lOk35

-2.8~1~1.9

0.5

2

COLDBI

0.12587zklO

0.13783*12

3.ozk1.4

1.3

2

0.13705*17

0.5

0.7

2

0.137OOkl6

0.81t:l.l

-0.5

1

CBSP5 Sp-p y inclusions

7E

0.12555

COLDB77F

zt 12

-c 12

l.Qrt2.3 f 20

k29

5.4k

1.9

Lk 1.4

Sp-p y-01 inclusion 0.1256OttlO

COLDB2 Spinel-bearing

olivine

CBOLl-Olivine

0.12580t13

CBOL2-Olivine

0.12591

CBOL-Spinet

0.12673k21

ztz21

0.13750+16

2.9kl.4

0.1

3

0.13818+21

3.8kl.9

3.2

3

0.13918rt29

10.341.9

-2.6

5

a- Error is given as twice standard error i.e. 2 Jz where n is number of cycles per analysis

(normally 50) A*%?g = ~(25Mg/24Mg)/0. 12663-l 11000 (O/~/amu); h- P&Q = A25Mg - A25Mg,s,,: 0. f 2663 is the accepted terrestrial value for (*5Mg/24Mg) (Catanzaro et al., 1966); Error calculated as J(ERROR A25Mg)2 + (ERROR A25Mg_STo)2; 25Mg.s,, is the average value of the spinal standard calculated on a daily basis i.e. Day 1 = -8.9kO.8 (n=13), Day 2 = -9.Okl.l (n=19), Day 3 = -9.521.0 (n=16), Day 4 = -7.2 + 1.8 (n =5), day 5 = -8.1 + 1.6 (n =7). c- d2’Mg = Az6Mg _ 2Az5Mg; Az6Mg = [(26Mg/24Mg)/0.1 3932-1]1000(0/~/amu); 0.13932 is the accepted terrestrial value for (2eMg/24Mg) (Catanzaro et a/., 1966); dz6Mg for the spine1 standards was Day 1 = -0.04, Day 2 = -0.18, Day 3 -0.55, Day 4 = -2.7, Day 5 = -2.3.

BER, I99

I ) it is unlikely that Fe-serpentine formed from such

composition of these phyllosi~icates reflects the chemistry of the fluids from which they formed and that these had a high Fe/ Mg ratio. TOMEOKA and BUSECK (1985) and MCSWEEN ( 1987) have proposed that during alteration of CM2 chondrites, early-formed phyllosilicates were Fe-rich from the breakdown ofmetals such as kamacite, but became enriched in Mg as olivine and pyroxene reacted. This suggests that the precursor to Fe-serpentine phyllosilicates was replaced at an early stage during alteration and was unreactive thereafter. The Fe-rich phyllosilicate rim invariably occurs exterior to spine! and immediately inside the pyroxene layer (i.e.. Fig. a precursor.

It is much more likely that the Fe-rich

4a). In the CAls of CV3s and CO3s ( ~ACPHERSON et al., 1988), and rarely in CM& this position is occupied by melihte, the primary phase most susceptible to alteration (BARBER et al., 1984). The volatile-poor carbonaceous chondrite, ALH 85085, has CAls that are texturally similar to those in CM2s, but are less altered and contain abundant melilite (GROSSMAN et al., 1988b). In a number of the ALH 85085 CAIs illustrated by KIMURA et al. ( 1993), melilite forms the inner rim layer separating the oxide-rich core from the outer pyroxene rim. In the Murchison spherule B6 (SIMON et al., 1993) melilite also forms the inner rim layer, but in this inclusion the rim sequence also includes a layer of anorthite sandwiched between melilite and pyroxene. ZOLENSKY

phyllosilicates. The latter may have formed foIlowing alteration of Fe-rich phyllosihcate by Mg-rich fluids (BARBER, 1981). On a Si-Fe-Mg rtemary diagram, s0me ana& crfehc Feand Mg-serpentine group pIot ciose to endmember croastedtite, but most lie close to the phyllosihcates of the associated dust mantles and matrix, plotting along a parallel trend of corn~~jt~on~~ variation (Fig. 9&c). Thns, dust mant3e ai& matrix phyllasilicates are minerafogicahy and chemictiy similar to those within granular sp-py inclusions. The Feand Mg-serpentines are therefore inferred to have crysttihsed at the same time ;ISthose wit&t the matrix f albeit ~n~tip~~~ and from arjueous solutions of a compambie Eh, pH, annd chemical composition. The open structure of granular sp-py inclusions allowed matrix-derived fluids to gain access to the interior of the CAIs. Formation of Mg-Serpentirres in Sp-Py-01 Inclusions

( 1984) estimate& that t~h~~~n~~~was pre~p~~ate~ from an a4ueous Quid of pH l&-E?. UncEet such extreme alkaline conditions, alumina is highly soluble (Fig. 6,h of MASON, 1966), and is transported as AlO; (BROOKINS 1988) or AtfOH); (V&XLAST and CHOU, I988). We propose, thercfore, that the Fe-&h ~~~~~o~~~~~ein the rims resuhed from the early, selective replacement of melilite in the presence of fluid with a high Fe/Mg ratio. The mare Mg-rich phyllosilicats that dominate the dust mantles and meteorite matrix wouid have formed I&X* from aqueous so&tions of a fewer Pe/Mg ratio which resuhed from the alteration afless reactive minerals, such as olivine and pyroxrne.

Cronstedtite and tochilinite are common to both the Fe= serpentine and Fe and ~~~se~ntj~~ phy~~os~~~c~te groups, but in gram&r sp-py ‘inclusions. cronsiedtite cry&& are rmaller and tochilinite is always coherently interstratified with serpentine, forming a mineral with ml.78 nm basal layer spacing. The Povien, tui’xdar chrysotih and poorly crystaihne ~by~~#~~~~t~ are unique to the Fe= and ~g~e~~t~~e group. PovTen crystaTs are common in C?vt matrices t %NER, 1% 1), and are analogous to terrestrial ““polygonal serpentineYY (~RESSEY and ZEISSMAN,8%%I. %3Wever, m0S-k %&I CryxUds in CXd ~kk~ve~d CA& are ~n~~~p~ete_ being semicirc&~ rather than circular in cross-section. Between the cronstedtite laths and Povlen crystals is a poorly crystalline phyllosilicate intergrown with chrysotile &es. Simiktr textures (Fig. 5c of BARBER,198 3fare ~n~~~~~d to have resnited from the separation of pan&Tel platy phy~iosi~~c~t~crystals, followed by occlusion ofthe resultant voids by porous, poorly+crystallina

The Mg-serpenlines constitute much of the interiors of fiheir host sp-py-oi indusiong so they must haye formed by replacement ofone or more precursors. Mineratogicaily, the Mg_serpentines differ from other phyllosihcates in CATs in that most crystals are only a few nm in size. The circular crystal structures are ~rn~n~Sce~~ of sections of chrysotiie ftbres, but Most discernibfe ctystafs are straight Or curved. Such textural characteristics are very similar to those of lizarditc described by WICKS f 19861. Despite Iizardite being an endmember ~~~~~~n~~~, it may contain rrp to I6 wt% Fe0 [ WrcKS and PLANT, f$79 1.Thus, brrth chemica’tly and texturally this M&serpentine is closely comparable to terrestrial lizardite. The distincZ chemical and ~~~e~j~~a~ natnre ofthe Ezardife indicates thsBeafter ~~~~~t~on by replacement ofa precutsor, it equilibrated little with the more Fe-rich dust mantles and meteorite matrix. The most likely prerursor is an olivine ( z Fo& which was more Fe-rich than the near pine fomterire which remains in the sp-py-ol indusions. This ~n~e~~~at~~n is supported by BUNCH and CHANG ( 198U) who described chondrutes and aggregates from Nogoya (CM2), which contain c&cite. PCP, and Mg-rick phyl~~j~~~~tes. In addition. rfrese authors observed sever& chon&uIes in which rehc forsteritic olivine is surrounded by fibrous Mg-serpentine in a textural relationship like that in sp-py+ol inclusions. The presence within Cold Bokkeveld sp-py-ol inchtsions of tochifiaite, ~nt~a~~~te, and cr~~s~edt~~e~and 53and Mi associated with the lizardite indicate that Fe-Hi metal and sulphides may have accompanied the olivine precursor. Ah three presumably were ahered together_ The simifarity between the a!reratfon products of sp-py-ol indtts%ms and those Found within the CoTd Bokkeveld matrix strongly suggests that both formed under similar conditions. We propose that this took place within the Cotd Bokkeveld parent body.

Calcite has been described fram a number of Murchison CA&, and is directly associated with gypsum in two of them, Sine An& fARMSTRONGet al., 19XZf and MtrCE-I (MHZPWER$ON et al., 1983). MACPHERSON et ai, (1983) suggest that both calcite and calcium sulphate formed in the solar

CAls in Cold Bokkeveld nebula, whereas ARMSTRONG et al. ( 1982) advocate an origin by parent body aqueous alteration. The calcium sulphate in Cold Bokkeveld CAIs is interpreted to have formed in the parent body. This is based on the observations of LEE ( 1993) that calcium sulphate-filled fractures crosscut the Cold Bokkeveld matrix, which is an unambiguous indication of a parent body origin. Calcite in Cold Bokkeveld CAIs, as with the matrix calcite, is also inferred to be the product of aqueous alteration. This is supported by GRADY et al. (1988), who found a correlation between the degree of alteration of CM2 meteorites, the concentration of carbonate minerals within them, and the isotopic composition of their carbon and oxygen. Within CAIs, calcite occurs as relatively coarse, twinned, inclusion-free single crystals with planar intercrystalline boundaries, and is interpreted as having precipitated as a cement. ARMSTRONG et al. ( 1982) likewise interpreted calcite in Murchison Blue Angel to be pore-filling cement. No remnants of a precursor are preserved, and so the mineral that formerly occupied the space now filled by calcite in Cold Bokkeveld CAIs must have been more reactive than diopside, hibonite, spine], or perovskite within the prevailing diagenetic environment. ARMSTRONG et al. (1982) suggested that in Blue Angel, melilite was the most likely precursor and MACPHERSON et al. ( 1983) described a calcite-bearing CA1 from Murchison ( MUM- I ) with a core of corroded melilite. Thus, a minority of Cold Bokkeveld CAIs (
Incorporation and Reprocessing Regolith Deformation

Within a Parent Body:

Many inclusions in Cold Bokkeveld show varying degrees of flattening and elongation parallel to the planar fabric in the meteorite (Fig. 5c 1. while others show evidence of rotation (Fig. 5d). As with other macroscopic objects (Fig. 1c,e) the dust mantles that surrounded many CAIs have been partially or completely removed. Questions that arise are, where and how did this deformation take place, and what was its timing relative to aqueous alteration? Some olivine grains in Cold Bokkeveld contain solar flare particle tracks recorded when the grains were within a few hundred microns of the surface of an unconsolidated regolith

1929

( HOHENBERG et al., 1990). Most of these grains (80%) are enriched in cosmogenic Ne. Assuming a particle flux similar to that of today, HOHENBERG et al. ( 1990) derive a minimum precompaction exposure time for Cold Bokkeveld of 45 Ma, compared with 14.5 Ma for Murchison and Murray. Approximately 20% of the track-rich grains analysed by HOHENBERG et al. ( 1990) contain no cosmogenic Ne. This indicates that near-surface irradiation was followed by burial to a depth of a few metres or more, to provide shielding from galactic cosmic rays, and suggests that the unconsolidated Cold Bokkeveld regolith was actively churned, presumably in response to impacts. The complex secondary mineral assemblages in the CAIs indicate that aqueous alteration on the parent body took place over a protracted interval. TOMEOKA and BUSECK (1985) argued that the Type-l PCP was produced initially by alteration of kamacite in chondrules and became mixed into the matrix during regolith “gardening.” Simultaneous aqueous alteration and regolith deformation were also advocated by RICHARDSON ( I98 1) to account for the breakdown of CM2 chondrule mesostases to “spinach.” It therefore seems likely that aqueous alteration and regolith deformation occurred together. GRIMM and MCSWEEN ( 1989) suggest that postaccretional impact heating would have melted ice within an asteroidal regolith to a depth of several kilometres, and the water produced would have been available over a sufficient period to produce the aqueous alteration observed in CM2s.

Fragmentation of CAIs and Formation of Dust Mantles Virtually all the CAIs examined in this study are fragments, yet despite their generally small size, most are fully enclosed by dust mantles (Table 3b). Thus, the main episode of CAI fragmentation clearly predates the formation of their dust mantles. Fragmentation probably occurred within a particleladen gas cloud. Large lateral movements of turbulently convecting gasses are a characteristic feature of the viscous accretion disc from which the solar system is believed to have formed (WOOD and MORFILL, 1988). CAIs within such a disc would at first have undergone fragmentation by highenergy collisions, but when particle velocities decreased, dust would have been deposited on CA1 cores to form dust mantles. The importance of this pre-dust mantle fragmentation episode in modifying the structure and composition of CAIs seems until now to have gone largely unnoticed. The only difference between most spine1 and sp-py inclusions is the thin, and often highly disrupted pyroxene selvedge of the latter (Fig. 4a). Thus, spine1 inclusions may have formed by the removal of the pyroxene layer from sp-py inclusions. It is possible that spine1 and sp-py inclusions were formed by varying degrees of fragmentation of sp-py-01 inclusions. Even where an intricate surface morphology is preserved (Fig. 4d), the CA1 had been abraded before the development of its dust mantle. The small size of the CAIs in CM2 chondrites compared with CV3s may be due to the more intense fragmentation experienced by the former. If fragmentation did take place in the nebula, it follows that the CAIs in CM2 meteorites were derived from a more turbulent region than those in cv3s.

1930

R. C. Greenwood et al.

Formation of CAIs

In common with Allende fine-grained inclusions (GROSSand GANAPATHY, 1976; MASON and TAYLOR, 1982) many CM2 spinel-rich inclusions display group II REE patterns ( EKAMBARUM et al., 1984; MACPHERSON et al., 1988; DAVIS and MACPHERSON, 1992). A two-stage process seems to be required to explain these patterns with the more refractory elements (Gd, Tb, Dy, Ho, Er, and Lu) being fractionated into an initial solid phase, which was then removed before the residual gas condensed to form the inclusion ( BOYNTON, 1975; DAVIS, 1991 ). The presence of group-11 patterns in many CM2 inclusions demonstrates clearly that condensation played a role in the formation of these objects. Evaporation, as opposed condensation, has been widely proposed for the formation of CAIs ( KLJRAT, 1970; WOOD, 198 1; KORNACKI and FEGLEY, 1984). In the WOOD ( 198 1) model, CAIs originated as interstellar dust which was heated so that the more volatile elements, such as Mg and Si, were distilled off to leave a residue enriched in Ca, Al, and other refractories. From the experiments of HASHIMOTO et al. ( 1979), which used powdered Murchison as a starting composition, evaporation of Mg and Si occurs above 1500°C (P < 10m5 Torr), and is accompanied by significant partial melting. Under these conditions, mass fractionation of Mg isotopes is produced (DAVIS et al., 1990). With the notable exception of COLDB 13 (Fig. II), spinel, sp-py, and sp-py01 inclusions have normal Mg isotopic compositions, so they were not the products of simultaneous melting and evaporation of more primitive materials. As already noted, the sppy inclusions COLDB17 and COLDB7 show a range of FM, values, some of which lie above the upper limit of the terrestrial standard. Portions ofthese inclusions may be enriched in the heavy isotopes of Mg, but the variation more likely results from inadvertently sputtering material from silicates surrounding the spinels. With FM8 values of 7.8 f 1.6%0 and 5.6 f 1.6%0, both analyses of the spine1 inclusion COLDB 13 are outside the range of the terrestrial standard, so this inclusion is enriched in the heavy isotopes of Mg. COLDB13 is a fragmental inclusion, 40 pm dia. and composed only of spinel. It has no voids and is enclosed only by a thin, discontinuous rim of Mg-rich phyllosilicate. It is probably a portion of a larger inclusion that became detached during regolith gardening. The Mg isotopic composition is, in this case, consistent with formation of the parental inclusion by simultaneous melting and evaporation of primitive dust (WOOD 198 1; KORNACKI and FEGLEY, 1984), but there is no conclusive evidence. Pyroxenes in sp-py and sp-py-01 inclusions are strongly zoned from Al- and Ti-poor diopside at the margins to Aland Ti-rich fassaite towards the interior. Both groups of inclusions show overlapping trends in Al vs. Si and Ti vs. Si plots (Fig. 7a,b), suggesting that they formed under similar conditions from precursors of related composition. Allende Type B inclusions contain pyroxenes (SIMON et al., 199 1) which display similar compositional variations to those in the Cold Bokkeveld CAIs. The experimental work of BECKETT and GROSSMAN ( 1982 ), STOLPER ( 1982 ), and STOLPER and PAQUE ( 1986 ) demonstrates that the mineral chemistry and crystallisation sequences seen in Type B inclusions are consistent with their formation by crystal-liquid fractionation processes. MAN

Spinel, sp-py, and sp-py-01 inclusions Spinel-bearing inclusions in CM2 chondrites, with the exception of refractory spherules (see below), are generally viewed either as the products of vapour/solid condensation ( MACPHERSON et al., 1983; DAVIS and MACPHERSON, 1992 ), or as refractory residues produced during the evaporation of primitive nebula dust ( KURAT, 1970; WOOD, 198 1; KORNACKI and FEC;LEY, 1984). In a general sense, the equilibrium condensation model (LORD, 1965; GROSSMAN, 1972) is successful in explaining many features of refractory inclusions, in particular their oxide-rich mineral assemblages and refractory bulk compositions. However, in detail there is a less than perfect match between the order in which phases should appear in a cooling gas of solar composition, as predicted by condensation models and the crystallisation sequence observed in CAIs. The most significant discrepancy is that melilite should condense before spine], irrespective of gas pressure ( KORNACKI and FEGLEY, 1984; WOOD and HASHIMOTO, 1993). Yet in Murchison melilite-rich inclusions, such as MUM- 1, large melilite crystals contain abundant euhedral spinels ( MACPHERSON et al., 1983), consistent with the reverse ofthe predicted sequence. The general lack of melilite in spinet-rich inclusions in CM2s has also been used as an argument against condensation ( KORNACKI and FEGLEY, 1984), but in our view this merely reflects its efficient dissolution during aqueous alteration. Another potential problem with a condensation origin for these inclusions is the absence of CaA1407. Thermodynamic calculations predict that early formed hibonite reacts with nebular gas to form CaA&O,, which in turn reacts to form gehlenite, then spine1 ( KORNACKJ and FEGLEY, 1984; FEGLEY, 199 1). According to these calculations, spine1 inclusions with accessory hibonite (Table 3b), ifformed by condensation from a cooling nebular gas, should also contain CaAl+,O, ( KORNACKI and FEGLEY, 1984; FEGLEY, 199 1). In the view of some workers, CAIs without CaA&O, are not pristine condensates (FEGLEY, 1991; B. Fegley, 1993, pers. commun.). Although it has been argued on thermodynamic grounds that CaA140, would not have been stable in cooling solar nebular gas ( MACPHERSON and GROSSMAN, 1984; GEIGER et al., 1988), this view has been strongly contested (B. Fegley, 199 1, 1993 pers. commun.). That CaAl,,O, was a stable primary phase in the nebula is clear from recent studies reporting its presence in the CAIs of a diverse range of carbonaceous chondrites ( CHRISTOPHE MICHEL-LEVY et al., 1982; PAQUE, 1987; DAVIS et al., 1987; GROSSMAN et al., 1988b; GREENWOOD et al., 1992; Huss and HUTCHEON, 1992; BISCHOFF et al., 1993; KIM~JRA et al., 1993), including the CM2 Murchison (SIMON et al., 1993). The spine1 hibonite-bearing CAIs examined in this study do not contain CaA1407, and hence, cannot be regarded as pristine condensates. However, it is possible that they originally contained CaA140,, which was subsequently removed during secondary alteration. That CaA140, may be more susceptible to alteration than other primary phases, including melilite, has recently been demonstrated in the case of VICTA, a Vigarano Type A inclusion (GREENWOOD et al., 1993b).

CAIs in Cold Bokkeveld The mineralogy and texture ofsp-py-ol inclusions indicate that they are the most complete CAIs in Cold Bokkeveld, and many sp-py and spine1 inclusions may be fragments from them. Outermost rims of forsterite on many spinel-rich inclusions in Murchison (MACPHERSON et al., 1983) appear to support this view. Because the compositional trends of their pyroxenes are consistent with crystal/liquid fractionation, we now examine whether the bulk composition and texture of sp-py-01 inclusions corroborate this conclusion. The least altered sgpy-01 inclusion US 1 (Fig. 4e) has the following bulk composition (wt%): SiOz 40.0, Ti02 0.6, A&OS 17.2, MgO 3 1S, CaO 10.8 (rastered beam microprobe analysis, recalculated to 100 wt%). From phase relations in the system CaO-MgO-A1203-Si02 ( OSBORNet al., 1954), U5 1 liquid crystallises forsterite as the liquidus phase at approximately 16OO’C. The residual liquid would become enriched in A&O3 and spine1 would coprecipitate with forsterite until approximately 1350°C. Spine1 would then react with liquid to form anorthite (assuming 20 wt% A1203 in the liquid). After spine1 had disappeared, anorthite and forsterite coprecipitate until approximately 1325”C, when pyroxene joins them and the residual liquid is consumed. The texture of U5 1 (Fig. 4e) is reasonably consistent with this crystallisation sequence. Olivine occurs around the margin ofthe inclusion, and spine1 and pyroxene are present in the interior. The only significant difference from the prediction is that AI-rich pyroxene is in direct contact with spine1 and anorthite is absent. The observed relationship is consistent with dynamic rather than equilibrium crystallisation (STOLPERand PAQUE, 1986); with the former, pyroxene appears before plagioclase, which in many cases fails to form. Finally, if CAIs formed by the processing of primitive CIlike material, single-stage melting cannot account for their refractory bulk compositions. As noted earlier, our Mg-isotopic data preclude simultaneous melting and evaporation, except for a few examples such as COLDB 13. For most CAIs, a two-stage model is required in which a refractory bulk composition is produced before the onset of melting. The mechanism that formed the refractory compositions is unknown. Vapour/solid condensation or slow evaporation of a solid precursor are possibilities. Refractory spherules A chondrule-like morphology and the inward radiating habit of hibonite in a few examples led MACDOUGALL( 198 1) to conclude that refractory spherules in Murchison formed from liquid droplets. Because many spherules are monominerahc spinel, MACDOUGALL ( 198 1) estimated that their melting temperature at I atmosphere might have been as high as 2135°C. Murchison refractory spherules have also been described by MACPHERSONet al. ( 1983). In agreement with MACDOUGALL( 198 1) they suggested that these crystallised from molten droplets with an estimated liquidus temperature of > 1550°C. A value of >2OOO”C has recently been claimed for the formation temperature of the Murchison CaAl407-bearing spherule B6 (SIMON et al., 1993). If these estimates are valid, refractory spherules in CM2 chondrites were produced in the highest temperature events in the early solar system for which we have evidence.

1931

Signifi~nt amounts of hibonite occur in only one of four spherules in Cold Bokkeveld. Its estimated mode is 89% spinel, 10% hibonite, and 1% perovskite, giving a bulk composition of 7 1.7% AlzOr , 26.0% MgO, 1.2% CaO, and 1.1% TiOz. In the system MgO-CaO-A&O3 (RANKIN and MERWIN, 19 16) this analysis, as well as the bulk estimates for Murchison spherules ( MACDOUGALL, 198 1; MACPHERSON et al., 1983; SIMONet al., 1993) all plot well within the primary phase field of spine], and should therefore crystallise spinet before hibonite. For Murchison inclusion MCI-H ( MACDOUGALL,1981), textural evidence indicates the reverse order ofc~stalli~tion. Sprays of hibonite crystals radiate from the perimeter of the spherule towards its interior and there is no evidence their habit was modified by contact with spine1 grains, as would be expected if spine1 crystallised first. On the contrary, Fig. 3a of MACDOUGALL( 198 1) shows clearly that spine1 is interstitial to hibonite. Despite the presence of inwardly radiating hibonites, SIMONet al. ( 1993) claim that in Murchison B6, spine1 crystallised first. Their argument is based on the observation that a 7 Mm-thick, nearly pure layer of spine1 forms its outer margin. However, this spine1 rim is not continuous, so that an inwardly radiating spray of hibonites is present at one point, its apex in direct contact with the inclusions outer margin. This texture suggests that hibonite crystallised prior to spinel, rather than in the reverse order as claimed by SIMON et al. ( 1993). In the Cold Bokkeveld, hibonite-bearing spherule and Murchison B 1 and I32 ( MACPHERSONet al., 1983 ) inwardly radiating sprays of hibonite are not developed. However, in these inclusions hibonite occurs as tabular or lath-shaped crystals and as noted by MACPHERSONet al., 1983 “where hibonite is intergrown with spine& the shapes of the spinels are controlled by those of the hibonites”. Such textures indicate that if CM2 spine1 spherules formed by crystallisation of molten droplets. either hibonite crystallised before spine1 or the two phases crystallised simultaneously. This textural evidence is crucial because it demonstrates that the present bulk composition of the spherules cannot be that of the parental liquid from which they formed. Formation temperatures derived by plotting bulk analysis on phase diagrams will therefore be erroneously high, perhaps by as much as 300-400°C. Most refractory spherules in Cold Bokkeveld and Murchison (MACDOIJGALL, 1981; MACPHERSON et al., 1983) contain numerous voids which constitute up to 30% of their volume. The observed highly refractory compositions of the spherules may result from the removal of less refractory material which originally filled the voids. Phases known to occur in CAIs, susceptible to aqueous alteration, and whose removal would increase the refractory character of the remaining assemblage include melilite, CaAi407, and Si-bearing glass. However, addition of any combination of these phases in the quantity indicated by the volume of void space (up to 30% ) would be insufficient to move the bulk compositions into the primary phase field of hibonite in the system CaO-MgOA1203-Si02 (DE VRIES and OSBORN, 1957; GUTT, 1963). The main reason for this is the restricted stability of hibonite with respect to MgO compared to the relatively MgO-rich compositions of the spherules. In view of this evidence, an

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origin for CM2 refractory spherules by crystallisation of molten droplets seems unlikely. The inwardly radiating hibonites may be a relic from an earlier melting event, represent regrowth of relic hibonite during partial melting, or have formed by solid-state recrystallisation. In CV3 chondrites, individually rimmed, spinel-rich (+ hibonite) objects, similar in size, morphology, and mineralogy to CM2 refractory spherules, are found in “fluffy” Type-A CAls (MACPHERSON and GROSSMAN, 1984). amoeboid olivine aggregates (HASHIMOTO and GROSSMAN, 1987), and various fine-grained inclusions ( KORNACKI and WOOD, 1985 ). In CM3 chondrites. refractory spherules are also known to occur within larger silicate-rich inclusions (GREENWOOD,1992). If spine]-rich spherules in CM2s were released by the breakup of larger inclusions like those in CV3s, they need not have formed by isochemical droplet crystallisation. Instead, the question that must be addressed is how did the spinel-rich spherical structures in CV3 inclusions form? Framboids are spherical clusters of spine] crystals commonly found in Type B2. more rarely in Type Bl, coarsegrained inclusions (EL GORESY et al., 1979; WARK and LovERIIVG, 1982a,b). Experiment ( WARK and LOVERING, 1982a; BECKETT and GROSSMAN, 1982) suggests that these structures form when Type-B precursors incompletely melt, the framboids being residual solid material. Rimmed, spinel-rich (i hibonite) objects found in fluffy Type As, and other large CV3 inclusions, may have formed in a similar manner to framboids and represent refractory material that remained in a solid state while the enclosing silicate material undenvent melting. One objection that has been raised to the derivation of CM2 spine1 spherules by fragmentation of larger CV3-type inclusions is the high Vz03 contents (up to 4.9 1 wt%) recorded by MACPHERSON and GROSSMAN ( 1984) in spinels from Allende fluffy Type As. Our measurements of the V203 content of spine1 in a cm-sized Allende fluffy Type A gave values of 0.2-0.5 wt%‘. Values of up to 0.7 wt% V20, were obtained by KORNACKI and WOOD ( 1985) in spinels from a range of Allende Type A inclusions. These are comparable with the values recorded in Cold Bokkeveld refractory spherules. and consequently do not pose a problem for our model. KORNAC‘KI and FEC;LEY ( 1984) proposed that spine1 spherules formed by evaporation and melting of primitive dust to produce a spinel-rich residue and Ca-rich liquid. Experiments by DAL IS et al. ( 1990) show that under such conditions, the spinel-rich residue would become enriched in the heavy Mg isotopes. Our spine1 spherule CBSPI has a normal Mg isotopic composition with an FMg value of -2.1 (Table 7: Fig. 11) which indicates that simultaneous evaporation and melting was not an important process in its formation.

In this study, a number of isolated olivine grains were located which contain small Mg-rich spine1 inclusions. These ohvines are distinct objects, unrelated to the other groups of refractory inclusions. One isolated olivine (CBOL) proved to have normal F ~~ values, whereas the enclosed spine1 is significantly enriched in the heavier Mg isotopes. Isolopic

heterogeneity between coexisting spine1 and olivine of similar magnitude to that in CBOL is a common feature of a group of objects in CV3 and CM2 chondrites. termed plagioclaseolivine inclusions ( POls) ( SHENG et al., I99 1 ). The Mg isotopic data indicate that at least one isolated olivine may have been derived from a fragmented POI. rather than from a chondrule (JONES, 1991) OJ by direct condensation from nebular gas (STEELE,1986). Although POIs have an igneous texture, the survival of isotopic heterogeneity suggests that they were never completely molten ( SHENG et al., 199 1). The isotopic composition of spine1 appears to have been fixed before it was incorporated into the PO1 precursor. POls, and olivines derived from them, preserve unambiguous evidence of a multistage formational history. They are, therefore, potentially powerful probes of the processes that occurred before the ubiquitous late-stage melting event. CONCLUSIONS CAIs in Cold Bokkeveld underwent extensive aqueous alteration and contain an assemblage of secondary minerals, including calcite and various phyllosilicates. We conclude that aqueous alteration took place within the parent body regolith. The Fe-phyllosilicate Jim layer present in many Cold Bokkeveld CAls formed early in the alteration sequence by the dissolution of melilite, and subsequent precipitation of cronstedtite from fluids with a high Fe/Mg ratio. CAIs experienced two episodes of deformation. an earlier nebula phase which predates the formation of their dust mantles, and a later phase which occurred in an asteroidal regolith. The compositional trends of pyroxenes in sp-py and sppy-01 inclusions indicate that the members of both groups crystallised from liquids. The crystallisation sequence deduced from the textures of sp-py-01 inclusions is compatible with that predicted by experiment. The textures displayed by refractory spherules indicate that hibonite crystallised either before or simultaneously with spine]. As a consequence, these objects could not have crystallised from molten droplets with a composition close to that of the bulk inclusion. The normal Mg isotopic composition displayed by most spine]-bearing inclusions in Cold Bokkeveld suggests that they did not form by simultaneous melting and evaporation of primitive dust. To form most of the inclusions two (or more) events seem to be required. The first event produced the refractory compositions, and in the second this precursor material was melted.

A~,linoM,/~d~mcnr.r-We thank Tony Wighton and Chris Jones for their help with sample preparation, Terry Williams for his assistance with mineral analysis. and Richard Hinton and John Craven for instruction and supervision during analytical work on the Edinburgh ion microprobe. This manuscript was improved as a result of comments from Richard Hinton and Bob Symes. and from critical reviews by A. El Goresy, D. Weber, and a third anonymous referee. RCG was supported from SERC grant GRF80074 to R.H. and M.R.L. from NERC grant GR3/7207 to D.J.B. R.C.G. received NERC support for ion microprobe work at Edinburgh University.

~~ilorialhonc~lilin~: H. Palme

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