Modal mineralogy of CM2 chondrites by X-ray diffraction (PSD-XRD). Part 1: Total phyllosilicate abundance and the degree of aqueous alteration

Available online at www.sciencedirect.com

Geochimica et Cosmochimica Acta 73 (2009) 4576–4589 www.elsevier.com/locate/gca

Modal mineralogy of CM2 chondrites by X-ray diffraction (PSD-XRD). Part 1: Total phyllosilicate abundance and the degree of aqueous alteration K.T. Howard a,*, G.K. Benedix a,b, P.A. Bland a,b, G. Cressey c a

Impacts and Astromaterials Research Centre (IARC), The Natural History Museum, Mineralogy Department, London SW7 5BD, UK b IARC, Dept. of Earth Sci. & Eng., Imperial College, London SW7 2AZ, UK c The Natural History Museum, Mineralogy Department, London SW7 5BD, UK Received 23 February 2009; accepted in revised form 29 April 2009; available online 12 May 2009

Abstract CM carbonaceous chondrites are samples of incompletely serpentinized primitive asteroids. Using position sensitive detector X-ray diffraction (PSD-XRD) and a pattern stripping technique, we quantify the modal mineralogy of CM2 chondrites: Mighei; Murray; Murchison; Nogoya and Cold Bokkeveld. There is a narrow range in the combined modal volume (vol%) of the most abundant phases Mg-serpentine (25–33%) and Fe-cronstedtite (43–50%). Cold Bokkeveld is anomalous in containing more Mg-serpentine (49–59%) than Fe-cronstedtite (19–27%). Even including Cold Bokkeveld, the range in modal total phyllosilicate is 73–79% (average = 75%). Total phyllosilicate abundance provides a non-ambiguous measure of the degree of aqueous alteration and indicates that these meteorites have all experienced essentially the same degree of aqueous alteration. This reflects pervasive hydration of matrix across CM2 samples. Apparent differences in the alteration of chondrules observed in petrographic studies represent various stages in the progression towards complete hydration of all components but are not manifest in significant differences in modal mineralogy. For all samples there is a limited range in olivine (6.9%) and pyroxene (5%) abundances. Modal abundances of the remaining identified phases also show a limited range: calcite (0–1.3%); gypsum (0–1.6%); magnetite (1.1–2.4%); pentlandite (0–2.1%) and pyrrhotite (1–3.8%). As expected, we observe a strong negative correlation in the modal abundance of anhydrous Fe–Mg silicates (olivine + pyroxene) and total phyllosilicate (Mg-serpentine + Fe-cronstedtite) consistent with the idea that phyllosilicate is forming by aqueous alteration of the anhydrous components. The negative correlation in the modal abundance between Mg–serpentine and Fe-cronstedtite indicates: (a) mineralogic transformation of Fe-cronstedtite to Mg-serpentine by fluid driven recrystallisation or (b) that these meteorites had different initial abundances of olivine and pyroxene. The observed positive correlation in the relative proportion of Mg-serpentine with increasing total phyllosilicate abundance reflects the evolution of increasingly Mg-rich phyllosilicate during aqueous alteration. Fe-cronstedtite is the dominant phyllosilicate, while CM chondrule olivines are forsteritic and will form Mg-serpentine during aqueous alteration. This implies that matrix olivine was more Fe-rich than chondrule olivine prior to aqueous alteration. Ó 2009 Elsevier Ltd. All rights reserved.

1. INTRODUCTION CM carbonaceous chondrites sample incompletely serpentinized fragments of some of the most primitive aster*

Corresponding author. E-mail address: [email protected] (K.T. Howard).

0016-7037/$ - see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2009.04.038

oids known. These meteorites preserve a record of nebular and parent body processes, evident in a range of complex petrographic and chemical features that must be unravelled to understand their origin. The mineralogy of CM’s has been extensively studied; they consist of high-temperature anhydrous phases and low-temperature serpentine family phyllosilicates. While the mineral phases present in CM’s

Modal Mineralogy of CM2 chondrites

have been characterised in detail, modal mineralogy remains poorly constrained. Deriving the modal mineralogy of CM2’s will allow for a greater understanding of their petrogenesis – as for all the carbonaceous chondrites. Critical to understanding the origin of CM chondrites is formation of the abundant phyllosilicate. The solar nebular was first suggested as the site of phyllosilicate formation by direct condensation from a cooling gas (Grossman and Larimer, 1974), this was later challenged by models that indicated silicate hydration would be kinetically inhibited in a nebular setting (Prinn and Fegley, 1987). Most workers have concluded from extensive petrographic studies that phyllosilicate formed by low-temperature aqueous alteration on the CM parent body asteroid[s] and this is currently the most accepted model (McSween 1979; Bunch and Chang, 1980; Tomeoka and Buseck, 1985; Zolensky et al., 1993; Browning et al., 1996; Hanowski and Brearley, 2001; Rubin et al., 2007). However, Metzler et al. (1992) concluded that fine-grained phyllosilicate-dominated rims around CM chondrules formed in the solar nebular. They suggest chondrules and other coarse-grained high-temperature components were partially hydrated in an uncompacted regolith pile, before being ejected during impact and passing through a nebular dust cloud(s) composed of mostly fine-grained hydrated phases that were accreted to form the rims (Metzler et al., 1992). The CM parent body[ies] then accreted a heterogeneous agglomeration of these rimmed components with little subsequent aqueous alteration (Metzler et al., 1992). Observing massive mineralogical heterogeneity, Lauretta et al. (2000) also concluded fine-grained, phyllosilicate-dominated rims in one CM (ALH 81002) were best explained by formation in the solar nebular. More recent kinetic models have also supported the possibility of phyllosilicate formation by low-temperature aqueous alteration in the solar nebular (Ciesla et al., 2003). Clearly, the origin and complex evolutionary history of CM’s remains a matter of debate. Many studies of CM’s have also focussed on defining the degree or extent of aqueous alteration (e.g. McSween, 1979; Bunch and Chang, 1980; Tomeoka and Buseck, 1985; Zolensky et al., 1993; Browning et al., 1996; Rubin et al., 2007). One reason this is of interest is because CM’s are less pervasively altered than the most primitive CI carbonaceous chondrites, but far more altered than the few CV chondrites containing phyllosilicate, so the possibility exists that there might be a mineralogic relationship between these groups related to the extent of hydration that is yet to be revealed. In addition, an alteration scale allows other relevant data to be placed in some context. Defining the modal mineralogy and relative degree of aqueous alteration will also allow characterisation of the source asteroid[s] and may provide insights to the long unanswered question of whether these meteorites are samples from a one or more parent bodies. There has been little agreement on the relative degree of aqueous alteration in CM’s and many approaches have been explored in an attempt to determine this (Table 1) – these are discussed in detail later. Important to this study is the early work of McSween (1979) who was first to explore a relationship between modal mineralogy and alter-

4577

ation. However, whereas McSween (1979) was limited to arbitrary identification of ‘modal matrix content’ by optical microscopy, we use PSD-XRD to accurately quantify the modal mineralogy of CM2 chondrites (Mighei, Nogoya Murray, Murchison, Cold Bokkeveld). The modal abundance of total phyllosilicate provides a non-ambiguous determination of the degree of aqueous alteration each meteorite has experienced. 1.1. Classification and formation of CM phyllosilicate CM phyllosilcates are serpentines and in this family of minerals there is a solid solution series between Fe and Mg end-members. Studies of CM chondrite serpentines reveal that end-member compositions are rarely, if ever, reached (Tomeoka and Buseck, 1985; Zolensky et al., 1997; Lauretta et al., 2000). TEM studies have also shown that CM phyllosilicate morphology is intimately related to composition, with Fe-rich serpentines typically occurring as relatively large platy crystals (up to 2500 nm) and more Mg-rich serpentines as smaller curved cylindrical or fibrous crystals (Lauretta et al., 2000). By PSD-XRD we make a distinction between highly crystalline Fe-rich serpentines or cronstedtite [(Fe,Mg)3(Fe,Si)2O5(OH)4] – although technically the terrestrial nomenclature reserves the term cronstedtite for the pure Fe end-member – and Mg-rich serpentines [(Mg,Fe)3Si2O5(OH)4] with much smaller crystalline domain sizes. Hereafter, we refer simply to Fe-cronstedtite and Mg-serpentine but we are not implying endmember compositions in either phase. The formation of serpentine phyllosilicate from aqueous alteration of anhydrous olivine is described in Eq. (1). Important to note is that the starting composition of the anhydrous component controls the composition of phyllosilicate formed by reaction with water. This will have important implications to later discussion. ðMg; FeÞ2 SiO4 þ SiO2 ðaqÞ þ 4H2 O ! 2ðMg; FeÞ3 ðFe; SiÞ2 O5 ðOHÞ4

ð1Þ

CM phyllosilicate may be associated with intergrowths of Tochilinite [(FeNi)SH2O] and clumps of intergrown phyllosilicate and tochilinite have previously been described as a ‘‘Poorly Characterised Phase (PCP)”. Tochilinite is extremely rare in terrestrial settings. It appears to form in the earliest stages of aqueous alteration in CM chondrites (Tomeoka et al., 1989). From TEM studies, Tomeoka et al. (1989) defined a four-stage model for the sequence of aqueous alteration and formation of phyllosilicates in CM’s that shows a progressive evolution to more Mg-rich phyllosilicate during aqueous alteration and is summarized as follows (with schematic reactions from Lauretta et al. (2000)): 1. Tochilinite formation in a reaction between kamacite and S – rich fluid: Fe; Ni þ H2 OðlÞ þ SðaqÞ ! FeðOHÞ2  ðFe; NiÞS

ð2Þ

2. Fe-rich matrix olivine alteration to form Fe-rich cronstedtite in an oxidizing reaction (as the cronstedtite contains Fe2+ and Fe3+):

4578

K.T. Howard et al. / Geochimica et Cosmochimica Acta 73 (2009) 4576–4589

Table 1 Previous alteration sequences for CM chondrites. Authors

Alteration sequence

Approach

McSween (1979)

Murray < Mighei < Murchison < Cold Bokkeveld < Nogoya

Modal content (optical microscopy) and Fe/Si ratio of matrix ( alteration) by SEM

Bunch and Chang (1980)

Murchison < Murray < Cold Bokkeveld < Nogoya

Petrography (optical microscopy, SEM), Cl abundance (altered)

Tomeoka and Buseck (1985)

Mighei < Murchison < Murray

Petrography (optical microscopy, SEM, TEM), matrix composition by SEM and TEM (>Mg = >alteration)

Burgess et al. (1991)

Mighei < Murray < Murchison < Nogoya

Bulk oxidised/reduced S ratios (>ox/red = >alteration)

Zolensky et al. (1993)

Nogoya < Murchison < Cold Bokkeveld < Mighei < Murray

Modal matrix content and Feo/(FeO + MgO) ratio of phyllosilicate (alteration)

Browning et al. (1996)

Murchison < Murray < Mighei < Nogoya < Cold Bokkeveld

Average matrix phyllosilicate composition (>Mg = >alteration); vol% isolated matrix silicate; vol% chondrule alteration

Rubin et al. (2007)

Murchison < Murray < Nogoya < Cold Bokkeveld

Petrography (optical microscopy, SEM); FeO/SiO2 ratio in PCP (alteration); S/SiO2 ratio in PCP (alteration); carbonate mineralogy (>dolomite and alteration)

ðFe; MgÞ2 SiO4 þ H2 OðlÞ ! ðFe; MgÞ3 ðFe; SiÞ2 O5 ðOHÞ4 ð3Þ 3. Alteration of chondrule olivine and pyroxene (Mgrich) releasing Si and Mg that reacts with tochilinite to produce Mg-rich cronstedtites, MgSiO4 þ 2FeðOHÞ2  ðFe; NiÞS þ H2 OðlÞ ! ðMg; FeÞ3 ðFe; SiÞ2 O5 ðOHÞ4 þ FeðOHÞ2  ðFe; NiÞS

ð4Þ

4. Tochilinite consumption continues, producing Mgrich cronstedtites and MgFe serpentine; S is released from tochilinite and redeposited as Fe–Ni sulfides ðFe; MgÞ3 ðFe; SiÞ2 O5 ðOHÞ4 þ FeðOHÞ2  ðFe; NiÞS ! ðMg; FeÞ3 Si2 O5 ðOHÞ4 þ ðFe; NiÞ9 S8

ð5Þ

Recent thermodynamic modeling also supports progressive evolution of more Mg-rich phyllosilicate during aqueous alteration (Zolotov and Mironenko, 2008). As expected, this modeling shows faster alteration of metal compared to silicates. Correspondingly, a low Mg/Fe ratio in early solutions leads to formation of Fe-rich serpentine. At later stages increasing supply of Si and Mg from dissolving silicates leads to deposition of increasingly Mg-rich serpentine. The Tomeoka et al. (1989) model describes evolution of phyllosilicate being deposited from a single fluid as anhydrous components are consumed cf. equilibrium conditions. However, studies such as that of Hanowski and Brearley (2001) show that Fe-rich phyllosilicate is forming at the same time as Mg-rich phyllosilicate within a single sample. This likely reflects the fact that there is limited opportunity for flow and homogenisation of fluid compositions in CM samples (Corrigan et al., 1997). As such, phyllosilicate of different compositions could be depositing simultaneously in different regions of the same sample, because without extensive fluid flow and mixing, the initial composition of the anhydrous component may control the composition of the fluid and resulting phyllosilicate.

1.2. Defining the extent of aqueous alteration in CM chondrites (progessive alteration sequences) In the history of research into CM chondrites a main concern of many studies has been in attempting to define the degree of aqueous alteration that has affected these meteorites. The focus has been on moving beyond broad classification of a CM sample as CM1 (almost completely hydrated) or CM2 that are extensively hydrated samples which still contain anhydrous components in chondrules and matrix (as we are studying here). Authors have been attempting to define a sequence of alteration that ranks CM according to the amount of hydration of anhydrous components that has occurred. The concept of an aqueous alteration sequence implies the progressive transformation of an initially anhydrous mineralogy to phyllosilicate. If a sequence is to be validly used to compare the extent of alteration that has taken place, an obvious requirement is that the starting point was identical between the samples being compared e.g. a uniform modal mineralogy and composition at the point of accretion. For any sequence to be valid requires not only a uniform abundance of constituent mineral phases but also components (e.g. chondrule and matrix) grain size and texture – all of these things will affect the extent to which the same amount of fluid will affect any rock. In more than 30 years of research many approaches have been taken in an attempt to resolve this sequence and there has never been any consensus amongst authors (Table 1). Initially, McSween (1979) made the assumption that matrix was predominantly hydrated (phyllosilicate) and therefore modal matrix abundance could be used to estimate the extent of aqueous alteration. He demonstrated that the matrix Fe/Si ratios from different CM chondrites appeared to decrease with increasing matrix abundance and suggested Fe/Si ratios could be used as proxy for the degree of alteration. This led many subsequent authors to use various other geochemical indicators to assess the

Modal Mineralogy of CM2 chondrites

degree of aqueous alteration (Table 1). All schemes using geochemical indicators to compare the degree of alteration are making the above assumptions that the initial mineralogy and composition of the samples being compared was identical. The most comprehensive of these approaches was by Browning et al. (1996) who define the CM Mineralogic Alteration Index (MAI). This index relied heavily on the Tomeoka et al. (1989) model of progressive alteration defined earlier and used an algorithm to processes microprobe data in order to monitor the Si and Fe3+ substitutions that accompany the transition from Fe-cronstedtite to Mg-serpentine. In accordance with the Tomeoka et al. (1989) model, samples being composed of more Mg-rich phyllosilicate were considered ‘more altered’ and assigned a higher MAI value. The MAI was later shown to exhibit a strong correlation with year of fall – the more altered meteorites being older falls (Benedix and Bland, 2004). This suggests the possibility that the MAI is actually, and inadvertently, a measure of terrestrial contamination and highlights the difficulties in understanding the meaning of geochemical proxies used to define the degree of aqueous alteration. Rubin et al. (2007) provide the most recent progressive alteration sequence for CM chondrites. Removing the need for the use of the complex algorithm employed to process SEM data by Browning et al. (1996), these authors utilise a combination of modal mineralogy determined by detailed petrography and geochemical proxies (Table 1) to gauge the extent of aqueous alteration. These authors not only re-define the petrographic sub-division of CM chondrites but also provide another progressive alteration sequence different from that of previous workers. Rubin et al. (2007) classify CM chondrites in their study from petrographic subtype 2.0 to 2.6. In this scheme completely hydrated samples that were traditionally termed CM1 are designated as petrographic type 2.0, and from here all other CM samples are then designated petrographic subtypes from 2.1 (most altered) to 2.6 (least altered). The CM2 samples we are studying here are classified as petrographic subtypes 2.5 (Murchison); 2.4/5 (Murray) and 2.2 (Nogoya and Cold Bokkeveld) by Rubin et al. (2007). Determined by microscopy and point counting, the modal mineralogy defined by Rubin et al. (2007) estimated the abundances of ‘‘silicate and oxide”; ‘‘metal in chondrules”; ‘‘metal in matrix”; ‘‘sulfide” and ‘PCP’, they also focus on carbonate but as a minor component do not estimate its abundance. Rather, Rubin et al. (2007) are concerned with the Ca and Mg (calcite/dolomite) content of CM carbonates and take more Mg-rich compositions to indicate a greater degree of alteration. Rubin et al. (2007) define essentially an identical abundance of silicate and oxide across the entire range of petrographic sub-types. As such they rely heavily on metal and sulfide abundances and carbonate composition in defining the degree of alteration (Rubin et al., 2007). Regardless of the potential difficulties in using SEM or optical methods to estimate the abundances of metal and sulfide, that are typically intergrown with phyllosilicate, the origin of these phases and their relationship to aqueous alteration is poorly understood. Similarly, the initial abun-

4579

dance of metal and sulfide prior to parent body processing is poorly constrained; based on studies of other primitive carbonaceous chondrites that have been relatively unaffected by aqueous processes (e.g. CV3 samples) these abundances appear to vary significantly (Howard et al., 2009). Thermal processes could also affect such abundances as well as aqueous alteration and deconvolving these effects is very difficult (Howard et al., 2009). Similarly, the origin of carbonates in CM and the exact controls on compositional variations are also not well understood. For these reasons we do not rely on sulfide, metal or carbonate abundance as indicators of aqueous alteration, nor any geochemical proxy, but rather focus on the total abundance of phyllosilicate. It is remarkable that no previous workers have attempted to define the extent of aqueous alteration by the abundance of phyllosilicate. Essentially this reflects the inability of most techniques to resolve this abundance. What is clear is that phyllosilicate forms by aqueous alteration of the anhydrous components. Therefore, the total abundance of phyllosilicate should provide a more assumption free and less ambiguous indicator of the degree of aqueous alteration each meteorite has experienced. This is the approach we explore throughout the following pages. 2. SAMPLE SELECTION AND ANALYTICAL PROCEDURES We selected a suite of some of the most extensively studied CM2 falls for study: Mighei, Nogoya, Murchison, Murray and Cold Bokkeveld. In each case care was taken to select sub-samples from fresh interiors of fragments catalogued at the Natural History Museum in London. Samples were prepared by grinding in a mortar and pestle to a uniform 35 lm size so as to avoid grain size artefacts during analyses. 2.1. Modal mineralogy by PSD-XRD X-ray diffraction analyses were conducted on an INEL diffractometer with a curved position sensitive detector (PSD) described in detail in Cressey and Schofield (1996). A Ge monochromator was used to select only Cu Ka radiation and post monochromator slits were used to restrict the beam to 0.24  5.00 mm during analyses to ensure irradiation of equal sample volumes. For phase quantification, powdered samples were packed into circular aluminium wells with a volume of 180 mm3. To avoid inducing preferred alignments of platy crystals parallel to the top surface of the sample, each well was packed using the sharp edge of a spatula in an attempt to produce a high-degree of randomness in grain orientations as in Bland et al. (2004). To monitor potential heterogeneity in samples and to ensure resolution of all phases, small sub-samples were also mounted on single crystal quartz substrates for analyses under identical conditions. Silicon and silver behenate were used as external standards; calibration and data collection was performed using Diffgrab . Patterns were collected with the beam at an incident angle of 4.6° to the flat-topped sample that was rotated continuously during TM

4580

K.T. Howard et al. / Geochimica et Cosmochimica Acta 73 (2009) 4576–4589

analyses. For single-phase standards, patterns were collected over 15 min. Meteorites were analysed over extended periods up to 16 h in order to produce patterns with minimum noise and allow accurate determination of all phases present. Phase quantification was performed in GUFI software using a refinement of the method developed previously by our laboratory (Cressey and Schofield 1996; Cressey and Batchelder, 1998) and applied to complex ultra-basic terrestrial samples by Schofield et al. (2002) and to meteorites by Bland et al. (2004). The PSD-XRD phase quantification method is described in the following expression:  0  0   t I 0 ðl=qÞi Xi ¼ Wi ð6Þ t I 0 ðl=qÞ0 TM

where Xi is the pattern intensity fraction from phase i in the mixture relative to the intensity of a standard pattern of pure i (intensity – 1.0), t and t0 are counting times for the standard and mixture, I0 and I0 0 represent the incident beam flux during acquisition of standard and mixture patterns, (l/q)i and (l/q)0 are mass absorption coefficients for the standard phase and mixture, and Wi is the actual weight fraction of phase i present in the mixture. Differences in the incident beam flux over time need to be monitored and corrected to allow standard patterns to be used in pattern fitting of mixtures analysed at different times from the standards. Cressey and Schofield (1996) described a method for correcting flux variations using the background intensity of the silicon standard calibration patterns to monitor X-ray flux. We now use a polished Fe-metal block to monitor flux as this avoids potential flux variations resulting from differences in packing density of the silicon standard over time. Phase quantification proceeds by initial identification of the phases present in the mixture pattern using the STOE reference database. Phase-pure samples of the mineral phases identified are then selected from our databank and their diffraction patterns are used in pattern fitting. The standard diffraction pattern (100% single-phase) is reduced in intensity by a proportion/pattern fit factor to resolve a best-fit match to its intensity in the mixture pattern and is subtracted to leave a residual pattern. This process is repeated for all phases identified; a residual of zero counts indicates that all phases in the mixture are accounted for. The determined pattern fit proportions are not the actual phase proportions in the mixture until the component phases in the mixture are corrected for their relative differences in X-ray absorption. Mass absorption coefficients for the phases present are calculated using values for the constituent elements given in International Tables for X-ray crystallography (Table 2). The actual weight fraction of each phase is a function of its measured pattern fit fraction modified by calculated mass absorption coefficients:

TM

X i =ðl=qÞi Wi ¼P i X i =ðl=qÞi

ð7Þ

This technique has been demonstrated accurate to 1–3% in an international round robin to test phase quantification commissioned by the International Union of Crystallography (Madsen, 1999; Madsen et al., 2001). The specific error

bars stated in the manuscript were determined as follows. When fitting each phase, once the best pattern fit value is arrived at, the value of this factor is varied until the resulting fit is noticeably ‘‘worse” – thereby giving the uncertainty of the fit that is then expressed as percentage of the best-fit value. This yields errors of 2–4% for the anhydrous phases and 3–5% for the phyllosilicate – in all cases the maximum error is stated and plotted. For each meteorite the determined modal mineralogy is also converted to an estimate of the bulk Fe content that is compared to measured literature values in order to provide an independent means of checking the accuracy of these data. Our calculated bulk Fe values will tend to underestimate measured values because we use an Mg-rich serpentine composition [Mg2.9Fe0.1Si2O5(OH)4] in our calculation and CM phyllosilicates may contain significantly more Fe (Tomeoka and Buseck, 1985). The accuracy of our approach as applied to meteorites is also evident in earlier works from our laboratory. Menzies et al. (2005) defined modal mineralogy in a suite of unequilibrated ordinary chondrites. Comparing a calculated bulk chemistry (based on the modal data) to INAA bulk chemistry by Jarosewich (1990), Menzies et al. (2005) observed a 1:1 correlation. A similar correspondence between calculated and measured chemistry was observed in carbonaceous chondrites by Bland et al. (2004). 3. RESULTS 3.1. Modal mineralogy Our modal data accounts for all crystalline phases present in abundances >1 wt%. Results are presented here in vol% so reported abundances may therefore be slightly less than 1%. There is no single-phase standard for tochilinite so we do not determine the modal abundance of this phase. That we still arrive at a zero residual implies we are accounting for X-ray counts from tochilinite when pattern stripping the phyllosilicate phases in which it may

Table 2 Cu Ka mass absorption coefficients for identified phases. Phase

Mass absorption coefficient (l/q)iCu

Olivine (Fo100) Olivine (Fo90) Olivine (Fo80) Olivine (Fo60) Olivine (Fo40) Enstatite (En98) Calcite Gypsum Magnetite Pentlandite Pyrrhotite Mg-serpentine Fe-cronstedtite

32.2 52.3 71.3 104.5 132.9 36.4 74.4 46 223.3 146.7 138.1 35.1 103.5

To calculate the absorption coefficients for phyllosilicate we assume and Mg-rich serpentine [Mg2.9Fe0.1Si2O5(OH)4] and the weight fraction of Fe in cronstedtite (.272) was calculated from Browning et al. (1996).

Modal Mineralogy of CM2 chondrites

4581

serpentine, we select a fine-grained and poorly crystalline phyllosilicate from our reference database that represents the best structural match possible to use as an analogue for the material in CM’s during pattern stripping (Fig. 2). It is also worth noting that we could determine the abundance of the fine-grained phyllosilicate, or indeed total phyllosilicate, by the residual difference after subtraction of all crystalline phases during pattern fitting and this would not change the resulting modal abundances we determine. Our technique provides a robust means of determining the modal abundance of total phyllosilicate even in complex samples such as these.

be intergrown. After analyses of CM matrix, McSween (1987) used mass balance calculations to estimate an average tochilinite abundance of 6–7 vol% in the samples being studied here. McSween (1987) notes that the presence of magnetite and sulfides in matrix would significantly affect Fe contents; meaning tochilinite abundances are over estimated. As described below these samples all contain significant magnetite (1.1–2.4%) and sulfide (1–4%), so based on mass balance calculations tochilinite is expected to be a minor component. This is consistent with point counting by Zolensky et al. (1993) who estimated a modal abundance of 2% tochilinite in Murray and <1% for Murchison and Nogoya. Fig. 1 shows annotated PSD-XRD patterns for Mighei and Cold Bokkeveld. These patterns show highly crystalline and ordered phyllosilicate producing sharp, high-intensity peaks at 12° and 25° (2h Cu Ka). This is obviously distinct from the broad peaks that are most pronounced at 19° and 61° (2h Cu Ka) that correspond with a much finer grainsize phyllosilicate. TEM shows that Fe-cronstedtite in CM chondrites is highly crystalline, well ordered and generally larger in grainsize than more Mg-rich serpentines (Lauretta et al., 2000; Barber, 1985 and others). Therefore, the highly crystalline phyllosilicate peaks evident in the XRD patterns are from Fe-cronstedtite, while the intergrown Mg-serpentines, present as platy crystals consisting of only a few structural layers along [0 0 1], are too thin to diffract X-rays coherently to produce 0 0 l reflections but these layers are extensive enough to result in broad hk reflections in the diffraction patterns. We do not attempt to further deconvolve the Mg-rich polytypes because the focus of our work is on total phyllosilicate abundance. To allow pattern stripping and resolution of the abundance of the intergrown and disordered phyllosilicate, without deconvolving the polytypes of Mg-

3.1.1. Mighei An XRD pattern recorded from Mighei is shown in Fig. 3. Peak stripping using patterns collected from single-phase standards leaves a residual at zero (Fig. 3). Correcting for absorption the pattern fits are converted to wt%, then to vol% (from well-known density data). The following modal mineralogy results were obtained for Mighei (n = 3): olivine (11–14%); pyroxene (6–8.4%); calcite (1.2– 1.3%); magnetite (2.1–2.4%); Fe-cronstedtite (46–50%); Mg-serpentine (25–27%); and pyrrhotite (1.2–3.1%). Repeat analysis and pattern fitting of separate Mighei samples produces results that show far less variation than might be expected from the heterogeneous appearance of separate thin sections, with the only significant differences reflecting heterogeneity in olivine compositions (Fig. 4 and Table 3). We calculate a bulk Fe of 20%, slightly lower than 21.8– 24.1% in Metzler et al. (1992). 3.1.2. Nogoya For Nogoya (n = 2) we determine a modal mineralogy of: olivine (15%); pyroxene (4%); calcite (1.1–1.2%); magnetite (2.2%); Fe-cronstedtite (43–44%); Mg-serpentine Artefact (detector glitch)

40000

Crystalline phyllosilicate

36000

Fine grained/ poorly crystalline phyllosilicate 32000

Absolute Intensity

Olivine

Olivine

Fine grained/ poorly crystalline phyllosilicate

Pyroxene

Pyroxene

Tochilinite

28000

Magnetite Enstatite + sulphide possibly but this can be x-ray amorphous

Calcite 24000

Magnetite Olivine

Crystalline phyllosilicate

Probably Tochilinite, could also be sulphide or Fe-Mg silicate

20000

Probably magnetite, could also be a phyllosilicate or FeMg silicate.

16000

12000 10.0

20.0

30.0

40.0

50.0

60.0

2Theta

Fig. 1. Annotated PSD-XRD pattern for Mighei (top, bold line) and Cold Bokkeveld showing main identified phases. Note: Cold Bokkeveld is offset on the Y-axis for clarity and contains a reduced proportion of highly crystalline phyllosilicate. See text for detailed description.

4582

K.T. Howard et al. / Geochimica et Cosmochimica Acta 73 (2009) 4576–4589

Mighei

Fine grained/ poorly crystalline phyllosilicate standard

0.0

20.0

40.0

60.0

80.0

100.0

2Theta

Fig. 2. PSD-XRD patterns for Mighei and the standard analogue we use to model fine-grained phyllosilicate (Mg-serpentine) in CM chondrites. The main broad reflections used in pattern stripping are indicated.

(32–33%); pentlandite (1.2–1.6%) and pyrrhotite (0.75– 1.1%). Repeat analysis of separate samples again produces consistent results (Fig. 4 and Table 3). For a meteorite that is typically considered to be highly brecciated this is perhaps unexpected. These results may provide evidence that our sample size is representative of the modal mineralogy of a parent body that has been randomly mixed during impact brecciation. The calculated bulk Fe of 18.5% is slightly lower than the 19 and 21 wt% measured by Rubin et al. (2007). 3.1.3. Murchison Murchison (n = 1) consists of: olivine (15%); pyroxene (8.3%); calcite (1.2%); magnetite (1.1%); Fe-cronstedtite (50%); Mg-serpentine (22%); pentlandite (0.65%) and pyrrhotite (1.2%) (Fig. 4 and Table 3). Our bulk Fe estimate of 19% is close to the 19.8–22% in Rubin et al. (2007). 3.1.4. Murray Murray (n = 1) is composed of: olivine (17%); pyroxene (5.6%); magnetite (1.4%); Fe-cronstedtite (49%); Mg-serpentine (25%); and pyrrhotite (1.8%). PSD-XRD data for olivine in this sample predominantly match with pure forsterite (Fig. 4 and Table 3), as our standard database is limited to 10 mol% intervals along the solid solution series. However, a low abundance of more Fe-rich olivines may be present (as indicated by SEM) but these are not apparent when using a low angle beam incidence necessary to record low angle phyllosilicate peaks. The modal abundance is accurate but SEM data provides information on the range of olivine compositions (if not their abundances) and suggests variation in Murray from at least Fo40–Fo100 (Leshin et al., 2000). Our data may indicate that this varia-

tion is weighted towards the forsterite end of the solid solution. The calculated bulk Fe is 19.4%, only slightly lower than the average measured value of 21.3% (Rubin et al., 2007). 3.1.5. Cold Bokkeveld Analyses of Cold Bokkeveld (n = 2) show the greatest variation in modal determinations for the samples studied. We determine a modal mineralogy of: olivine (10– 13%); pyroxene (3.3–6.6%); magnetite (1.7–2.2%); Fecronstedtite (19–27%); Mg-serpentine (49–59%); pentlandite (1.3–2.1%) and pyrrhotite (0.95–1.6%) (Fig. 4 and Table 3). The calculated bulk Fe of 15% is too low relative to measured value of 19.8% (Rubin et al., 2007) because this sample is dominated by Mg-serpentine and, as described above, we are using an Mg-rich composition in our calculation. 4. DISCUSSION The CM’s in this study have a very similar modal mineralogy. Excluding Cold Bokkeveld, there is a narrow range in the modal abundance of Mg-serpentine (25– 33%) and Fe-cronstedtite (43–50%). Cold Bokkeveld appears to be anomalous in containing more Mg-serpentine (49–59%) than Fe-cronstedtite (19–27%). Across the whole sample, the range in olivine and pyroxene abundance is less than 7% and 5%, respectively. Again, if Cold Bokkeveld is excluded, the ranges in olivine and pyroxene abundances are reduced to less than 3% and 4%, respectively. There is also only a limited range in the abundances of the remaining phases identified: calcite (0–1.25%); gypsum (0–1.6%); magnetite (1.1–2.4%);

Modal Mineralogy of CM2 chondrites

Intensity (counts)

30000

4583

actual (offset by 10 000 counts)

20000

calculated

10000

0

residual

0

80

40

120

2θ Fig. 3. Mighei PSD-XRD quantitative phase analysis. The actual collected pattern for Mighei (count time of 16 h) is shown with a calculated pattern fit constructed by summing single-phase standard patterns. For the actual task of pattern stripping to determine modal mineralogy, we subtract single identified phases individually. This figure is illustrative of the result where the intensity of each single-phase standard pattern (collection time of 15 min) is normalised to the collection time of Mighei, then by a pattern fit fraction to fit the proportions of the single-phase in Mighei. Subtracting the composite pattern results in a zero residual indicating that all phases have been identified. This is illustrative of our technique but in reality we strip each phase individually until a zero residual is reached. The calculated pattern appears noisy because it is the sum of many Fe-rich phases analysed only for long enough to allow their use in pattern stripping individually (e.g. 15 min). When many short analyses of Fe rich phases are combined the noise appears great, this is why we analyse the meteorites over such long periods.

pentlandite (0–2.1%); and pyrrhotite (1–3.8%). At present, we have no way of explaining why Cold Bokkeveld has a different modal mineralogy relative to the other samples. However, this difference is unlikely to relate to any terrestrial process since the sample was recovered shortly after fall. Perhaps Cold Bokkeveld derives from a different region of the same parent body as the other CM’s studied, or somewhat different conditions (w/r ratio, pH, etc.), or from a separate parent body. New modal determinations for other CM’s will further elucidate this.

4.1. Correlations in modal mineral abundances and origin of CM phyllosilicate Modal data allow examination of the relationship between mineral phases and this can have important implications for petrogenesis. There is expected to be a strong inverse relationship in the modal abundance of anhydrous phases and phyllosilicate if the latter forms by aqueous alteration of the anhydrous phases. On a plot of total anhydrous Fe–Mg silicate vs. total phyllosilicate our data do

4584

K.T. Howard et al. / Geochimica et Cosmochimica Acta 73 (2009) 4576–4589

100%

Mg-Serpentine

95% 90%

22%

25%

85%

27%

28%

Fe-cronstedtite

25% 33%

32%

Pyrrhotite

80%

49%

75%

Magnetite

vol%

65% 60% 55%

Pentlandite

59%

70%

Gypsum

50%

49%

50%

46%

48%

50% 43%

45%

Calcite

44%

40%

27%

35%

Forsterite 40

19%

30%

Pyroxene

Forsterite 60

25% 20%

Forsterite 80

15% 10%

Forsterite 90

5%

Forsterite 100

0% Murchison Murray

Mighei

Mighei [2] Mighei [3] Nogoya Nogoya [2]

Cold Cold Bokkeveld Bokkeveld [2]

Fig. 4. Stacked column showing modal mineralogy of CM chondrites by PSD-XRD. Each analysis characterises a sample volume of 180 mm3. All samples were fresh interior fragments ground to a 35 lm powder. For phases other than phyllosilicate, errors are 2–4%, for phyllosilicate 3–5%.

tionship between modal abundance of Fe-cronstedtite and Mg-serpentine. Fe lost to fluid during this recrystallisation will be available for incorporation into other Fe bearing phases, such as sulfide, as is evident in our data that shows a general increase in the modal abundance of sulfide with increasing abundance of Mg-serpentine. Upon alteration the fluid and phyllosilicate composition is largely controlled by the initial anhydrous component. Therefore, another way of explaining the observed variation in the modal abundance of Mg-serpentine and Fe-cronstedtite would be that the studied meteorites had different initial modal abundances of Fe and Mg rich anhydrous components (e.g. different forsterite–fayalite solid solution and Fe–Mg pyroxene solid solution) that

indeed show a strong inverse relationship indicating that phyllosilicate is forming from aqueous alteration of the anhydrous components (Fig. 5). While expected, this is the first time it has been quantitatively demonstrated in CM’s. Such a relationship has previously been shown by XRD for separate aliquots of Orgueil (Bland et al., 2004). In the plot of Fe-cronstedtite vs. Mg-serpentine data define an almost perfect negative slope 1 line (Fig. 6). Fe-cronstedtite is an unstable structure because octahedral Fe3+ is too small and tetrahedral Fe3+ too large for the ideal layered structure. In the presence of water, Fe-cronstedtite will re-crystallise to form Mg-serpentine and during this process – which may be repeated many times – Fe is lost to fluid. This recrystallisation may explain the observed inverse rela-

Total anhydrous silicate (olivine + pyroxene) vol%

35%

Murchison Murray Mighei Nogoya Cold Bokkeveld

30%

25%

20%

15%

10%

5%

0% 70%

72%

74%

76%

78%

80%

82%

84%

Total Phyllosilicate (Mg-serpentine + Fe-cronstedtite) vol%

Fig. 5. Plot showing variation in modal abundance of total anhydrous Fe–Mg silicates (olivine + pyroxene) vs. total phyllosilicate (Mgserpentine + Fe-cronstedtite). The strong inverse relationship indicates that phyllosilicate is forming from aqueous alteration of olivine and pyroxene. Error bars are 4% for olivine + pyroxene and 5% for Mg-serpentine + Fe-cronstedite.

100 100 100 100 100 100 100 100 100 22.2 25 27.2 27.6 25.4 33.2 31.7 59.3 48.8 50.3 49.1 46.3 47.7 49.6 42.7 44.1 19.2 27.3 1.2 1.8 3.1 1.2 1.6 1.1 0.7 1.5 0.9 0.7 – – – – 1.2 1.6 2.1 1.3 1.1 1.4 2.1 2.3 2.4 2.2 2.2 1.7 2.2 – – – – – – – 1.6 – 1.2 – 1.2 1.2 1.3 1.1 1.2 1 1 8.3 5.6 6 6.2 8.4 3.8 3.9 3.3 6.6 3.9 – 2.4 4.4 4 – – 1.2 – – – – – – 1.2 1.2 5.1 – 4 – – – – 2.5 6.5 – 8.8 – – – – – – – 2.2 – 7.2 17.3 11.7 9.5 7.3 11 6.8 1.9 3.8

Olivine (Fo100) Olivine (Fo90) Olivine (Fo80) Olivine (Fo60) Olivine (Fo40) Enstatite Calcite Gypsum Magnetite Pentlandite Pyrrhotite Fe-cronstedtite Mg-serpentine Total

4585

are now reflected in the relative modal abundances of Fecronstedtite and Mg-serpentine. This is supported by a comparison of the most pristine carbonaceous chondrites such as Acfer 094 and ALH 77307. These meteorites have experienced minimal aqueous processing, but they contain anhydrous Fe–Mg silicates with a range of compositions (Greshake, 1997; Greshake et al., 2004; Nagahara and Kushiro, 1982), most likely indicating variable compositions at the time of accretion. There is also the possibility that anhydrous Fe–Mg silicate components were variably mixed during impact brecciation that appears to have affected CM’s. Fe-cronstedtite is the dominant phyllosilicate in all of these samples except Cold Bokkeveld. Chondrule olivines in CM’s are forsteritic and will form Mg-serpentine during aqueous alteration. This implies that matrix olivine was more Fe-rich (fayalitic) than chondrule olivine prior to hydration, as is consistent with the model of Tomeoka et al. (1989). Models described earlier showed the evolution of increasingly Mg-rich serpentine during aqueous alteration (Tomeoka et al., 1989; Zolotov and Mironenko, 2008). These models suggest that Fe-rich anhydrous components (metal, fayalite) are hydrated first, followed by Mg-rich components (forsterite, pyroxene). These models also suggest that as alteration progresses, the relative proportion of total phyllosilicate that is Mg-serpentine will increase. This is evident in our data where as the modal abundance of total phyllosilicate increases, so does the relative proportion of the total phyllosilicate that is Mg-serpentine. However, Cold Bokkeveld contains an average of 54% Mg-serpentine, twice the average volume of the other samples (24%), but only an average of 4% less olvine + pyroxene. This indicates that the progressive evolution of more Mg-rich phyllosilicate is not controlling the much greater abundance of Mg-serpentine in this meteorite e.g. alteration in Cold Bokkeveld has not progressed further from the same initial mineralogy as the other CM samples studied, because if this were the case the abundance of olivine + pyroxene would necessarily also be much more greatly reduced. Either the initial mineralogy of Cold Bokkeveld differed or this meteorite has experienced different alteration conditions and processes that remain to be defined. 4.2. Degree of aqueous alteration

Murchison Murray Mighei Mighei [2] Mighei [3] Nogoya Nogoya [2] Cold Bokkeveld Cold Bokkeveld [2]

Table 3 PSD-XRD quantitative modal mineralogy of CM2 chondrites (vol%).

Modal Mineralogy of CM2 chondrites

Previous alteration sequences for CM samples were described earlier in Section 1.2 and are summarized in Table 1. While we concur with evidence for evolution to more Mg-rich serpentines during aqueous alteration of any individual meteorite, the possibility that CM’s had different initial abundances of anhydrous Fe–Mg silicates complicates the use of geochemical comparisons of phyllosilicate composition in defining the degree of alteration. This is because given different initial amounts of Fe and Mg rich components, fluid and phyllosilicate compositions may evolve to any given composition after different amounts of aqueous alteration (time, volume of fluid, etc.). Earlier we also described how the severely limited fluid flow in CM chondrites

4586

K.T. Howard et al. / Geochimica et Cosmochimica Acta 73 (2009) 4576–4589 65%

Murchison Murray Mighei Nogoya Cold Bokkeveld

60%

Mg-serpentine vol%

55% 50% 45% 40% 35% 30% 25% 20%

15%

25%

35%

45%

55%

Fe-cronstedtite vol%

Fig. 6. Plot of Mg-serpentine vs. Fe-cronstedtite illustrating variation in modal abundance. This may reflect fluid driven re-crystallisation of Fe-cronstedtite to form Mg-serpentine, or different initial abundances of Fe–Mg rich anhydrous components that are now reflected in the relative modal abundance of Fe-cronstedtite and Mg-serpentine. Error bars are 5%.

might allow heterogenous fluid compositions, controlled by the anhydrous component being attacked, to deposit Fe and Mg rich phyllosilicate simultaneously in different regions of the same sample. This also presents a potential complication to geochemical comparisons of phyllosilicate composition in different meteorites as indicators of the extent of aqueous alteration. We suggest that since phyllosilicates formed by aqueous alteration, the degree of aqueous alteration in CM chondrites should be defined by the total abundance of phyllosilicates. No other petrographic or chemical parameter can unambiguously and without assumption define the degree of hydration. Our PSD-XRD derived total phyllosilicate abundances in the studied CM chondrites are shown in Fig. 7. The studied CM’s all have nearly identical modal

abundances of total phyllosilicate. Across the entire sample the mean is 75 ± 3.4% (2r) and the range from 73% to 79%. Our results indicate these CM2 meteorites are essentially the same in terms of the amount of aqueous alteration that they have experienced. It is not our assertion that relict chondrules in the studied samples will all necessarily show the same degree of alteration as many petrographic studies suggest this is not the case (e.g. Rubin et al., 2007). For example, Rubin et al. (2007) suggest that mafic silicate phenocrysts in Murchison chondrules are completely intact, although the chondrule mesostases have been altered. This is in contrast to Nogoya in which they estimate 20–50% of the mafic silicate phenocrysts in chondrules have been altered. Subsequently, Rubin et al. (2007) use their scheme to assign petrographic

100% 90%

Total Phyllosilicate vol%

80%

Murchison, 73%

Murray, 74%

Mighei, 73%

Mighei, 75%

Mighei, 75%

Nogoya, 76%

Nogoya, 76%

Cold Bokk. Cold Bokk. 79% 76%

70% 60% 50% 40% 30% 20% 10% 0%

Fig. 7. Total modal abundance of phyllosilicate (Fe-cronstedite + Mg-serpentine) in CM chondrites. Since phyllosilicate forms by aqueous alteration of anhydrous Fe–Mg silicates, the total abundance of phyllosilicate defines the degree of aqueous alteration. The average volume of total phyllosilicate is 75 ± 3.4% (2r) and the range from 73% to 79%. Based on the total modal abundance of phyllosilicate, these meteorites have all experienced essentially the same degree of aqueous alteration. Error bars are 5%.

Modal Mineralogy of CM2 chondrites

subtype 2.5 to Murchison and 2.2 to Nogoya. Our quantitative modal data indicate the difference in the modal abundance of anhydrous mafic silicates between Murchison (23%) and Nogoya (19%) is 4%. Therefore, the difference in alteration observed by Rubin et al. (2007) in chondrules represents a maximum of 4% of the bulk modal mineralogy, keeping in mind that chondrules are a minor component and by optical means only the coarse grained mafic silicate can be characterised. This difference is reflected in a 3% difference in total phyllosilicate abundance between Murchison (73%) and Nogoya (76%). We suggest that to define Murchison as petrographic type 2.5 and Nogoya as 2.2, on the basis of such minor differences in modal mineralogy, is to belie the fact that these samples have experienced essentially the same degree of hydration. We also consider that it is difficult to envisage a situation in which olivine phenocrysts in a chondrule could remain completely intact, while mesostasis was completely altered. It is important to recognize that alteration in chondrules, at least prior to pervasive hydration, may only be visible by TEM. Most likely in chondrule olivine is a type of alteration frequently observed in terrestrial settings: topotactic alteration, where the primary mineral is replaced without substantially disturbing the original crystal structure, with widespread inheritance of structural polymers by the weathering product (Eggleton, 1984; Smith et al., 1987; Banfield et al., 1990; Hochella and Banfield, 1995). In this situation, an olivine grain may be chopped-up by lamellae of weathering products, each only a few 10’s nm wide and only visible by TEM. If that were the case, the degree of alteration in chondrules could be much more similar than obvious by optical microscopy or SEM. Given the similarity in the modal mineralogy defined in this study, we suggest there is no need for further definition of petrographic subtypes beyond CM2 – leaving CM1 to refer to essentially completely hydrated samples as previously. Nevertheless, apparent differences in the extent of chondrule alteration do provide a potentially important basis for classification of CM meteorites by the study of thin sections and allow definition of alteration sequences on this basis as Rubin et al. (2007) demonstrate. Based on our data, the Rubin et al. (2007) alteration sequence probably reflects varied stages of progression in the formation of a few volume percent phyllosilicate by alteration of minor coarse-grained components in chondrules - after pervasive and uniform hydration of matrix – as these authors suggest. This relative uniformity in the extent of aqueous alteration in CM2 chondrites we are describing, on the basis of modal mineralogy, is entirely consistent with the near homogenous chemical compositions of these samples (Kallemeyn and Wasson, 1981; Bland et al., 2005). The O-isotopic compositions of CM2 are also very similar (Clayton and Mayeda, 1999) although Rubin et al. (2007) show a weak inverse correlation between D17O and their petrographic subtypes. This indicates apparent petrographic differences not evident in bulk modal data may be resolvable in isotopic systematics as has also recently been suggested by de Leuw et al. (2009) using the Mn/Cr system to study CM carbonates. Geochemical proxies used in previous alteration sequences (Table 1) are possibly also measuring vari-

4587

ous stages in the progression towards complete hydration of all components. However, as our modal data demonstrates, the CM2 samples are at very similar stages in this progression that has been truncated-conflicting sequences may result simply from the use of different geochemical indicators to compare relatively uniform samples. 5. CONCLUSION Our PSD-XRD and pattern stripping technique determines an accurate modal mineralogy for all phases present in CM chondrites at abundances greater than 1 vol%, other than tochilinite for which there is no standard. This technique improves greatly on modal estimates determined by optical and SEM based point counting. Our modal determinations are also consistent with literature values for the measured bulk chemistry of CM chondrites and this further suggests our results are accurate. The sample volume we are characterising is significantly larger than in previous studies, resulting in repeatable modal determinations for the studied samples. We therefore have good reason to believe our results are representative of the meteorites and potentially their parent body. Our modal data reveal very limited mineralogic heterogeneity between these CM2 meteorites, despite significant textural variation in thin section. We observe a strong inverse relationship in the modal abundance of anhydrous Fe–Mg silicates and total phyllosilicate indicating that the latter is forming from aqueous alteration of the anhydrous components. The variation in the modal abundances of Mg-serpentine and Fe-cronstedtite can be explained by (a) mineralogic transformation of Fe-cronstedtite to Mg-seprentine by fluid driven recrystallisation or (b) the studied CM’s having had different initial amounts of anhydrous Fe–Mg silicates prior to aqueous alteration. Our data suggest that matrix olivine must have been more Fe-rich than chondrule olivine prior to alteration. As phyllosilicate forms by aqueous alteration, its total abundance provides a non-ambiguous measure of the degree of aqueous alteration that each meteorite has experienced. This can be determined by PSD-XRD and our data show there is a very limited range in the modal abundance of total phyllosilicate across the whole sample set from 73% to 79% (average = 75%). This indicates that all of the studied CM2 meteorites have experienced essentially the same degree of aqueous alteration. This implies pervasive hydration of matrix across CM2 samples and incomplete alteration of chondrules, apparent differences in the extent of chondrule alteration observed in petrographic studies may reflect various stages in the progression towards complete hydration of all components. This study demonstrates that in the studied CM2 samples these apparent differences in alteration extent are very minor and are not manifested in significant variations in modal mineralogy. ACKNOWLEDGMENTS We thank Alan Rubin and an anonymous reviewer for detailed and constructive reviews that benefited this manuscript greatly.

4588

K.T. Howard et al. / Geochimica et Cosmochimica Acta 73 (2009) 4576–4589

Associate Editor Anders Meibom is also thanked for handling this work. Caroline Smith is thanked for providing NHM samples used in this study. Dominik Hezel and Hazel Hunter (NHM) were involved in useful discussions. This work was supported by STFC through the UK Cosmochemistry Analysis Network (UK-CAN).

REFERENCES Banfield J. F., Veblen D. R. and Jones B. F. (1990) Transmission electron microscopy and subsolidus oxidation and weathering of olivine. Contrib. Mineral. Petrol. 106, 110–123. Barber D. J. (1985) Phyllosilicates and other layer structured materials in stony meteorites. Clay Miner. 20, 415–454. Benedix G.K. and Bland P.A. (2004) What does the CM chondrite mineralogic alteration index really mean. Meteorit. Planet. Sci 39 (Suppl.). #5184 (abstr.). Bland P. A., Cressey G. and Menzies O. N. (2004) Modal mineralogy of carbonaceous chondrites by X-ray diffraction and Mossbauer spectroscopy. Meteorit. Planet. Sci. 39, 3–16. Bland P. A., Alard O., Benedix G. K., Kearsley A. T., Menzies O. N., Watt L. and Rogers N. W. (2005) Volatile fractionation in the early Solar System and chondrule/matrix complementarity. Proc. Natl. Acad. Sci. USA 102, 13755–13760. Browning L., McSween H. Y. and Zolensky M. (1996) Correlated alteration effects in CM carbonaceous chondrites. Geochim. Cosmochim. Acta 60, 2621–2633. Bunch T. E. and Chang S. (1980) Carbonaceous chondrites. II. Carbonaceous chondrite phylloilicates and light element geochemistry as indicators of parent body processes and surface conditions. Geochim. Cosmochim. Acta 44, 1543–1577. Burgess R., Wright I. P. and Pillinger C. T. (1991) Determination of sulfur-bearing components in C1 and C2 carbonaceous chondrites by stepped combustion. Meteoritics 26, 54–64. Ciesla F. J., Lauretta D. S., Cohen B. A. and Hood L. L. (2003) A nebular origin for chondritic fine-grained phyllosilicates. Science 299, 549–552. Corrigan C. M., Zolensky M. E., Dahl J., Long M., Weir J., Sapp C. and Burkett P. J. (1997) The porosity and permeability of chondritic meteorites and interplanetary dust particles. Meteorit. Planet. Sci. 32, 509–515. Clayton R. N. and Mayeda T. K. (1999) Oxygen isotope studies of carbonaceous chondrites. Geochim. Cosmochim. Acta 63, 2089– 2104. Cressey G. and Batchelder M. (1998) Dealing with absorption and micro absorption in quantitative phase analysis. Int. Union Crystallogr. Newslett. 20, 16–17. de Leuw S., Rubin A. E., Schmitt A. K. and Wasson J. T. (2009) Mn–Cr systematics for the CM2.1 chondrites QUE93005 and ALH83100: implications for the timing of aqueous alteration. Lunar Planet. Sci. 40, Lunar and Planetary Institute, Houston #1794 (abstr.). Eggleton R. A. (1984). Formation of iddingsite rims of olivine; a transmission electron microscope study Clays Clay Min. 32, 1–11. Greshake A. (1997) The primitive matrix components of the unique carbonaceous chondrite Acfer 094: a TEM study. Geochim. Cosmochim. Acta 61, 437–452. Greshake A., Krot A. N. and Keil K. (2004) Mineralogy and chemistry of fine-grained matrices, rims, and dark inclusions in the CR carbonaceous chondrites Acfer/El Djouf 001 and the ungrouped carbonaceous chondrites Acfer 094 and Adelaide. Workshop on Chondrites and the Protoplanetary Disk. #9041 (abstr.). Cressey G. and Schofield P. F. (1996) Rapid whole-pattern profile stripping method for the quantification of multiphase samples. Powder Diffraction 11, 35–39.

Grossman L. and Larimer J. W. (1974) Early chemical history of the solar system. Rev. Geophys. Space Phys. 12, 71–104. Hanowski N. P. and Brearley A. J. (2001) Aqueous alteration of chondrules in the CM carbonaceous chondrite, Alan Hills 81002: implications for parent body alteration. Geochim. Cosmochim. Acta 65, 495–518. Hochella M. F. and Banfield J. F. (1995) Chemical weathering of silicates in nature: a microscopic perspective with theoretical considerations. In Chemical Weathering Rates of Silicate Minerals (eds. A. F. White and S. L. Brantley). Min. Soc. Am, Washington, DC, USA, pp. 353–406. Howard K. T., Benedix G. K., Bland P.A. and Cressey G. (2009). Modal mineralogy of CV3 chondrites by PSD-XRD: mineralogic insights into a complex evolutionary history. Lunar Planet. Sci. 40, Lunar and Planetary Institute, Houston. #1235 (abstr.). Jarosewich J. (1990) Chemical analyses of meteorites: a compilation of stony and iron meteorite analyses. Meteoritics 25, 323–337. Kallemeyn G. W. and Wasson J. T. (1981) The compositional classification of chondrites. I. The carbonaceous chondrite groups. Geochim. Cosmochim. Acta 45, 1217–1230. Lauretta D. S., Hua X. and Buseck P. R. (2000) Mineralogy of finegrained rims in the ALH81002 CM chondrite. Geochim. Cosmochim. Acta 64, 3263–3273. Leshin L. A., McKeegan K. D. and Benedix, G. K. (2000) Oxygen isotope geochemistry of olivine from carbonaceous chondrites. Lunar Planet. Sci. 31, Lunar and Planetary Institute, Houston. #1918 (abstr.). Madsen I. C. (1999) Quantitative phase analysis round robin. Int. Union Crystallogr. Newslett. 22, 3–5. Madsen I. C., Scarlett N. V. Y., Cranswick L. M. D. and Lwin T. (2001) Outcomes of the international union of crystallography commission on powder diffraction round robin on quantitative phase analysis: samples 1a to 1h. J. Appl. Crystallogr. 34, 409–426. McSween H. Y. (1979) Alteration in CM carbonaceous chondrites inferred from modal and chemical variations in matrix. Geochim. Cosmochim. Acta 43, 1761–1770. McSween H. Y. (1987) Aqueous alteration in carbonaceous chondrites: mass balance constraints on matrix mineralogy. Geochim. Cosmochim. Acta 51, 2469–2477. Menzies O. N., Bland P. A., Berry F. J. and Cressey G. (2005) A Mossbauer spectroscopy and X-ray diffraction study of ordinary chondrites: quantification of modal mineralogy and implications for redox conditions during metamorphism. Meteorit. Planet. Sci. 40, 1023–1042. Metzler K., Bischoff A. and Stoffler D. (1992) Accretionary dust mantles in CM chondrites: evidence for solar nebular processes. Geochim. Cosmochim. Acta 56, 2873–2897. Nagahara H. and Kushiro I. (1982) Petrology of chondrules, inclusions and isolated olivine grains in ALH-77307 (CO3) chondrite. Natl. Inst. Polar Res. Memoirs 25, 66–77. Prinn R. G. and Fegley B. J. (1987) The atmospheres of Venus, Earth and Mars: a critical comparison. Ann. Rev. Earth Planet. Sci 15, 171–212. Rubin A. E., Trigo-Rodriguez J. M., Huber H. and Wasson J. T. (2007) Progressive aqueous alteration of CM carbonaceous chondrites. Geochim. Cosmochim. Acta 71, 2361–2382. Schofield P. F., Knight K. S., Covey-Crump S. J., Cressey G. and Stretton I. C. (2002) Accurate quantification of the modal mineralogy of rocks when image analysis is difficult. Min. Mag. 66(1), 173–184. Smith K. L., Milnes A. R. and Eggleton R. A. (1987) Weathering of Basalt: formation of Iddingsite. Clays Clay Min. 35, 418–428. Tomeoka K. and Buseck P. R. (1985) Indicators of aqueous alteration in CM carbonaceous chondrites: microtextures of a layered mineral containing Fe, S, O and Ni. Geochim. Cosmochim. Acta 49, 2149–2163.

Modal Mineralogy of CM2 chondrites Tomeoka K., McSween H. Y. and Buseck P. R. (1989) Mineralogical alteration of CM chondrites: a review. Proc. NIPR Symp. Antarct. Meteorites 2, 221–234. Zolensky M., Barrett R. and Browning L. (1993) Mineralogy and composition of matrix and chondrule rims in carbonaceous chondrites. Geochim. Cosmochim. Acta 57, 3123–3158. Zolensky M., Mittlefehldt D. W., Lipschutz M. E., Wang M-S., Clayton R. N., Mayeda T. K., Grady M. N., Pillinger C. and

4589

Barber (1997) CM chondrites exhibit the complete petrologic range from type 2 to 1. Geochim. Cosmochim. Acta 61, 5099–5115. Zolotov M. Yu. and Mironenko M. V. (2008) Aqueous alteration of CM2 chondrites evaluated with kinetic models (abstract). Meteoritics Planet. Sci 43, A177. Associate editor: Anders Meibom