Spectral reflectance properties of carbonaceous chondrites ... · Meteorites that ap-pear to be...

Icarus 212 (2011) 180–209

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Spectral reflectance properties of carbonaceous chondrites: 1. CI chondrites

E.A. Cloutis a,⇑, T. Hiroi b, M.J. Gaffey c, C.M.O’D. Alexander d, P. Mann a

a Department of Geography, University of Winnipeg, 515 Portage Avenue, Winnipeg, Manitoba, Canada R3B 2E9b Department of Geological Science, Brown University, Box 1846, Providence, RI 02912-1846, USAc Department of Space Studies, University of North Dakota, PO Box 9008, Grand Forks, ND 58202-9008, USAd Department of Terrestrial Magnetism, Carnegie Institution of Washington, 5241 Broad Branch Road, NW, Washington, DC 20015, USA

a r t i c l e i n f o

Article history:Received 24 August 2010Revised 9 December 2010Accepted 11 December 2010Available online 21 December 2010

Keywords:Asteroids, CompositionAsteroids, SurfacesMeteoritesSpectroscopy

0019-1035/$ - see front matter � 2010 Elsevier Inc. Adoi:10.1016/j.icarus.2010.12.009

⇑ Corresponding author. Fax: +1 204 774 4134.E-mail addresses: [email protected] (E.A. Clo

edu (T. Hiroi), [email protected] (M.J. Gaffey), alexaAlexander).

a b s t r a c t

Existing reflectance spectra of CI chondrites (18 spectra of 3 CIs) have been augmented with new (18spectra of 2 CIs) reflectance spectra to ascertain the spectral variability of this meteorite class and provideinsights into their spectral properties as a function of grain size, composition, particle packing, and view-ing geometry. Particle packing and viewing geometry effects have not previously been examined for CIchondrites. The current analysis is focused on the 0.3–2.5 lm interval, as this region is available forthe largest number of CI spectra. Reflectance spectra of powdered CI1 chondrites are uniformly dark(<10% maximum reflectance) but otherwise exhibit a high degree of spectral variability. Overall spectralslopes range from red (increasing reflectance with increasing wavelength) to blue (decreasing reflectancewith increasing wavelength). A number of the CI spectra exhibit weak (<5% deep) absorption bands thatcan be attributed to both phyllosilicates and magnetite. Very weak absorption bands attributable to otherCI phases, such as carbonates, sulfates, and organic matter may be present in one or a few spectra, buttheir identification is not robust. We found that darker spectra are generally correlated with bluer spec-tral slopes: a behavior most consistent with an increasing abundance of fine-grained magnetite and/orinsoluble organic material (IOM), as no other CI opaque phase appears able to produce concurrent dark-ening and bluing. Magnetite can also explain the presence of an absorption feature near 1 lm in some CIspectra. The most blue-sloped spectra are generally associated with the larger grain size samples. Forincidence and emission angles <60�, increasing phase angle results in darker and redder spectra, partic-ularly below �1 lm. At high incidence angles (60�), increasing emission angle results in brighter and red-der spectra. More densely packed samples and underdense (fluffed) samples show lower overallreflectance than normally packed and flat-surface powdered samples. Some B-class asteroids exhibitselected spectral properties consistent with CI chondrites, although perfect spectral matches have notbeen found. Because many CI chondrite spectra exhibit absorption features that can be related to specificmineral phases, the search for CI parent bodies can fruitfully be conducted using such parameters.

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1. Introduction

Carbonaceous chondrites (CCs) are an important group of mete-orites for understanding the origin and evolution of the solar sys-tem. They include a diversity of subgroups that exhibit differentdegrees of aqueous and/or thermal alteration, but are generallycharacterized by an overall dark appearance. They are importantfrom an astrobiological perspective because some of their car-bon-bearing phases are organic in nature and include many biolog-ical precursor molecules (e.g., Nagy, 1975). They may also havebeen an important source of water for the proto-Earth (e.g., Mor-bidelli et al., 2000). Their importance extends to understanding

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utis), [email protected]@dtm.ciw.edu (C.M.O’D.

the origin and evolution of the solar system, because they havebeen linked, largely through spectroscopic comparisons, to variousclasses of asteroids (e.g., Gaffey and McCord, 1977; Gradie andTedesco, 1982; Vilas and Gaffey, 1989; Gaffey et al., 1993; Burbineand Binzel, 1995; Burbine, 1998; Burbine et al., 2001).

CC-asteroid links have often been made on the basis of similar-ities in overall spectral shapes, which has often been necessitatedby the fact that many low albedo asteroid (and cometary nuclei)spectra are devoid of resolvable and/or diagnostic absorptionbands (e.g., Hiroi et al., 1993; Abell et al., 2005). The carbonaceouschondrite class consists of a number of groups, many of which havenot been spectrally characterized in a comprehensive way.

This paper is the first in a series of CC papers, and discusses thespectral reflectance properties of the CI carbonaceous chondrites.We have undertaken a spectral reflectance study of CCs for a num-ber of reasons: (1) to determine the range of spectral variabilitywithin CC classes; (2) to determine the spectral similarities or

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E.A. Cloutis et al. / Icarus 212 (2011) 180–209 181

differences that exist between CC classes; (3) to search for absorp-tion bands that are diagnostic of constituent phases; (4) and to bet-ter elucidate the relationships between CC mineralogy, structure,and spectra. In our analysis we include both existing CC reflectancespectra, new measurements made for this study, and new andexisting reflectance spectra of constituent minerals and phasesand mixtures. The goal of this study is to determine the range ofmineralogic and petrologic information that can be derived fromanalysis of CC reflectance spectra. This analysis can also informthe search for CC parent bodies among the asteroids.

2. Carbonaceous chondrite classifications

The carbonaceous chondrite class has been defined by a numberof criteria. They were defined by Mason (1962) as ‘‘stony meteor-ites characterized by the presence of an appreciable amount of car-bonaceous material other than free carbon (diamond andgraphite)’’ – but ‘‘appreciable’’ was not defined. In his classification,petrologic type 1 CCs are made up predominantly of largely amor-phous hydrated magnesium–iron silicate and hydrated magnesiumsulfate and magnetite. Type 2 CCs consist largely of a serpentinegroundmass impregnated with carbonaceous material enclosingvariable amounts of olivine and minor enstatite in the form ofchondrules. Type 3 CCs consist in large part (�70%) of olivine withminor amounts of pyroxene (5–10%), feldspar (5–10%), magnetite(�5%), and troilite (�6%); Fe–Ni metal may be present in accessoryamounts (Mason, 1962).

Newer components that have been developed for discriminat-ing CCs from other types of meteorites are more composition-based, and include elemental ratios, bulk composition, and oxygenisotopic ratios (Table 1; e.g., Van Schmus and Wood, 1967; VanSchmus and Hayes, 1974; Wasson, 1974, 1985; McSween, 1979;Dodd, 1981). While these classification schemes can provide somediscrimination between the major CC groups, whether these divi-sions are significant enough to translate into measurable spectraldifferences is not obvious, as the most volumetrically dominantphases are not necessarily the same ones that are spectrally dom-inant (e.g., Cloutis et al., 1990). Compositional parameters that areuseful for discriminating different CC groups and that may alsotranslate into measurable spectral differences include iron oxida-tion state (oxidation increasing in the general sequence:CH(?) < CR < CV < CO < CK < CM < CI (Brearley and Jones, 1998)),oxidation state of iron, water content, and abundance of reducedcarbon (Roy-Poulsen et al., 1981). For discriminating CIs from otherCCs, H2O, C, S, and matrix contents are all highest in CIs (Zandaet al., 2009).

Petrologic types are somewhat coupled to specific classes ofCCs. In the case of CI chondrites, they contain pervasive aqueousalteration products (and hence are all of petrologic type 1, and per-haps grading into type 2) (Van Schmus and Wood, 1967).

Some of the properties described in Table 1 that are used to de-fine and discriminate CC groups will likely not translate into spec-tral differences; these include chondrule size, texture, and matrix

Table 1Petrographic characteristics of C-chondrite groups. Source: Brearley and Jones (1998).

Group Chondrule abundance (vol.%) Matrix abundance (vol.%)

CI �1 >99CM 20 70CR 50–60 30–50CO 48 34CV 45 40CK 15 75CH �70 5

abundance. This is due to the fact that such textures would be de-stroyed during regolith evolution, as asteroid surfaces are generallyfine-grained (Dollfus, 1971; Dollfus et al., 1975). Other differencesbetween CC groups may, however, provide insights into whetherand how different C-chondrite groups may be spectrally distin-guishable. These include: the higher content of dark matrix mate-rial, lower bulk Fe, and �2 times more C and H2O in CV than COchondrites (Van Schmus, 1969; Van Schmus and Hayes, 1974;McSween, 1977a); lower mafic silicate Fe content and higher metalcontents in CR than other CCs (Bischoff et al., 1993); and higherabundance of refractory inclusions in COs and CVs versus otherCCs (Brearley and Jones, 1998). There are also correlations betweenorganic matter content (C, H, N abundances) and extent of pre-ter-restrial aqueous alteration, as expressed by clay mineral contents(Pearson et al., 2001). With the expectation that most CC parentbody asteroids have comminuted surfaces, this paper confines it-self largely to analysis of powdered CC reflectance spectra.

Classification schemes for CCs have evolved and multiplied overtime. Classification schemes now exist for CCs that show evidenceof thermal metamorphism following aqueous alteration (e.g., Akai,1994; Akai and Tari, 1997; Nakamura, 2005), for quantifying thedegree of shock metamorphism (Stöffler et al., 1991), and degreeof terrestrial weathering (e.g., Rubin and Huber, 2005). Recentsummaries of meteorite classifications can be found in Krot et al.(2003), Weisberg et al. (2006), and Brearley (2006).

3. Mineralogy of CI chondrites

Interpreting reflectance spectra of CCs requires knowledge ofthe spectral properties of the constituent phases, their composi-tion, and physical disposition. Previous studies (e.g., Cloutis et al.,1990) have shown how even small amounts of dark fine-grainedmaterials can strongly affect reflectance spectra when the opaquephases are fine-grained and effectively dispersed.

As mentioned, the known CI chondrites are confined to petro-logic type 1, and perhaps grading into type 2. Meteorites that ap-pear to be thermally metamorphosed CI-types (e.g., Akai andTari, 1997) are dealt with in a subsequent paper. In CI1s, pre-exist-ing anhydrous silicates have been almost completely aqueously al-tered to phyllosilicates (50–60 vol.%), predominantly serpentineand saponite, with intergrown ferrihydrite, abundant oxides (pri-marily magnetite and maghemite), accessory sulfides (primarilypyrrhotite and pentlandite and minor cubanite), carbonates, andsulfates (Richardson, 1978; Zolensky and McSween, 1988; Buseckand Hua, 1993; Zolensky et al., 1993; Brearley and Jones, 1998;Gounelle and Zolensky, 2001). CI1s have few recognizable lithiccomponents, being composed largely of matrix (fine-grained mate-rials), and have the highest matrix abundance of any CC group (Bu-seck and Hua, 1993).

The phyllosilicates-rich matrix of CI1s is dominated by differentproportions of intergrown Fe-bearing serpentine and saponite(also sometimes called montmorillonite) intimately intergrownwith ferrihydrite (McSween, 1987; Tomeoka, 1990; Brearley and

Refractory inclusionabundance (vol.%)

Metal meanabundance (vol.%)

Chondrulediameter (mm)

�1 0 –5 0.1 0.30.5 5–8 0.713 1–5 0.1510 0–5 1.04 <0.01 0.70.1 20 0.02

182 E.A. Cloutis et al. / Icarus 212 (2011) 180–209

Jones, 1998); some portion of the Fe in the phyllosilicates appearsto be in the Fe3+ state (Roy-Poulsen et al., 1981). CI serpentine ismore disordered than terrestrial equivalents, and with increasingaqueous alteration, CIs contain increasingly more ordered serpen-tine and a greater abundance of saponite (Howard et al., 2010).

Kerridge (1976) used wet chemical data, Mössbauer spectros-copy, and stoichiometric analysis of Orgueil phyllosilicates todetermine abundances of FeO (10.1 wt.%) and Fe2O3 (11.3 wt.%).Mössbauer analysis of Orgueil (CI1) indicates the presence ofnon-stoichiometric magnetite, ferrihydrite, Fe3+-bearing serpen-tine (cronstedtite), Fe2+-bearing serpentine, and Fe2+–Fe3+ smectite(Burns and Fisher, 1991). Bostrom and Fredericksson (1966) usedelectron microprobe, XRD, and wet chemical analyses and stoichi-ometric considerations to determine that the phyllosilicates inOrgueil contain both ferrous and ferric iron. Thermomagnetic anal-ysis indicates that magnetite content varies from 8.5 to 11.1 wt.%in CI1 chondrites (Hyman and Rowe, 1983).

Single olivine grains and grain fragments are very rare in CI1chondrites, with compositions similar to CO3 chondrites: a pro-nounced peak at Fa1 and a spread of more Fe-rich olivines to Fa55

(McSween, 1977b). Carbonates (mostly dolomite and breunnerite,and rare calcite) are also present (up to 5 vol.% in Orgueil) (Bostromand Fredericksson, 1966; Brearley and Jones, 1998).

Sulfur is present in a variety of forms: elemental, organic, sul-fide, and sulfates (primarily hexahydrite and epsomite and raregypsum, bloedite, and Ni-bloedite) (Bostrom and Fredericksson,1966; Burgess et al., 1991; Brearley and Jones, 1998). The bulk ofthe organic matter and carbon in CI1 chondrites is present in inti-mate association with the matrix phyllosilicates, and consists lar-gely of small heteroatom polycyclic aromatic hydrocarbons withvarious branches and side chains (Alpern and Benkhieri, 1973;Wdowiak et al., 1988; Pearson et al., 2002; Garvie and Buseck,2004, 2005, 2007; Garvie, 2007).

While this general overview suggests that the CI1 chondritesare a reasonably homogenous family, it should be noted that theyexhibit mineralogical and chemical heterogeneities at a variety ofscales. These variations reflect both brecciation and mineralogicaland chemical variations within the starting materials that persistthrough closed system alteration (Morlok et al., 2006). The appar-ent ‘‘threshold’’ for ensuring chemical heterogeneity appears to beon the order of 1–2 grams (Morlok et al., 2006). However, even inthe case where sample amounts on this order are available, the vol-ume of sample that will contribute to a reflectance spectrum willbe much less, as the penetration of incident light into such darksamples will be on the order of a few tens to hundreds of micronsat best.

4. Mineralogy and petrology of individual CIs

In order to assist the interpretation of the reflectance spectra ofthe three CIs included in this study (Alais, Ivuna, Orgueil), we sum-marize their mineralogy and petrology below. The state of knowl-edge is variable for each of these CIs but they appear to possessmineralogic and petrologic differences and similarities that mayhave implications for spectral analysis.

4.1. Alais (CI1)

Alais is composed largely of matrix, which consists mostly ofphyllosilicates, predominantly saponite with subordinate serpen-tine (Golden et al., 1994). Overall, the matrix averages Fe/(Fe+Mg)of 0.41, and 1.85 wt.% S (Zolensky et al., 1993), while the matrixphyllosilicates have an average Fe/(Fe+Mg) of 0.13 (Tomeoka,1990), or 0.30 (Tonui et al., 2003). Phyllosilicates in Alais range be-tween serpentine and smectite compositions (Tomeoka, 1990).

Mössbauer analysis shows that approximately 60% of the total Feis present in the phyllosilicates and it appears to contain a lowerproportion of magnetitic:serpentinitic Fe than Orgueil (Herr andSkerra, 1968). More recently, Bland et al. (2008) used Mössbaueranalysis to determine that the Fe in Orgueil is present in troilite(2%), silicates (8%), magnetite (34%), and a paramagnetic compo-nent (mostly Fe3+-bearing; 56%). The phyllosilicates in Alais are,on average, larger and more heterogeneous than those in Orgueil(Mackinnon and Kaser, 1988), and occur in two well-defined mor-phologies and compositional types: the larger size fraction is dom-inated by well-ordered phyllosilicates with a serpentine-type basalspacing, while the fine-grained (<100 nm) fraction has a poorly-de-fined crystallinity consistent with montmorillonite-type clays(Mackinnon and Kaser, 1988). Veins in Alais are dominated by car-bonates, Ca-sulfate, and Mg-sulfate (epsomite and hexahydrite),the latter being the volumetrically dominant species (Richardson,1978; Fredericksson and Kerridge, 1988). Thermomagnetic analy-sis indicates that magnetite content is 8.9 ± 0.9 wt.%, and is moreheterogeneously distributed in Alais than in Ivuna and Orgueil (Hy-man and Rowe, 1983). Bulk C content is 5.40 wt.% (Pearson et al.,2006). Alais likely suffered less aqueous alteration than Ivunaand Orgueil on the basis of less alteration of Fe–Ni sulfides (i.e.,troilite) to pyrrhotite and magnetite (Bullock et al., 2005), andother petrographic criteria (Petitat and Gounelle, 2010).

4.2. Ivuna (CI1)

Ivuna contains nearly 100 vol.% matrix and 2.25 wt.% C as insol-uble organic matter (IOM) (Alexander et al., 2007). The matrixaverages Fe/(Fe+Mg) of 0.41 (Zolensky et al., 1993) and consistsmostly of phyllosilicates, predominantly saponite with subordinateserpentine (Golden et al., 1994), or largely serpentine interlayeredwith minor saponitic smectite (Brearley, 1992a). The matrix iscompositionally homogeneous, with sulfides, oxides, and carbon-ates disseminated throughout (Brearley, 1992a). The matrix phyl-losilicates have an average Fe/(Fe+Mg) of 0.11 (Tomeoka, 1990),0.30 (Tonui et al., 2003), 0.2–0.4 (Brearley, 1992a), or 0.3–0.6(Zolensky et al., 1993). Compositionally, phyllosilicates in Ivunacluster near smectite (Tomeoka, 1990). Ivuna contains a higherproportion of a paramagnetic component (likely Fe3+-bearingphases) than Orgueil (Bland et al. 2008).

Organic matter is closely associated with the phyllosilicates(Pearson et al., 2002). Total S content is reported as 2.11 wt.%(Zolensky et al., 1993), or 2.94 wt.%, and is present in elemental, or-ganic, sulfide, and sulfate forms, with the sulfate form dominating(Burgess et al., 1991). Magnetite is sparsely distributed in the ma-trix and occurs in several morphologies (like Orgueil) but grainsizes are generally larger (up to 10 lm) than Orgueil (up to1 lm) (Brearley, 1992a). Thermomagnetic analysis indicates thatmagnetite content is 11.1 ± 0.3 wt.% (Hyman and Rowe, 1983).Unlike Orgueil, there is no evidence of intergrown ferrihy-drite + phyllosilicates (Brearley, 1992a). Veins in Ivuna are domi-nated by carbonates, Ca-sulfate, and Mg-sulfate (epsomite andhexahydrite), the latter being the volumetrically dominant species(Richardson, 1978; Fredericksson and Kerridge, 1988). The pres-ence of magnetite, gypsum, and possible carbonates were identi-fied by X-ray diffraction analysis (Bass, 1971).

4.3. Orgueil (CI1)

Orgueil contains nearly 100 vol.% matrix and 2.00 wt.% C as insol-uble organic matter (Alexander et al., 2007). Modal mineralogy isprovided in Table 2. Mössbauer spectra of Orgueil are consistentwith the occurrence of magnetite, ferrihydrite, Fe3+ serpentine(cronstedtite) and Fe2+ serpentine in the matrix (Fisher and Burns,1991). Mössbauer analysis shows that approximately 50% of the

Table 3Available spectral reflectance studies of CIs (�0.3–2.5 lm reflectance spectra).

Meteorite Sample type Source

Alais <150 lm (14 spectra) 1Alais <150 lm (3 spectra) 2Ivuna <125 lm (2 spectra) 3, 4Orgueil <45 lm 4Orgueil <40 lm 5Orgueil <100 lm 5Orgueil 40–100 lm 4Orgueil 100–200 lm 4Orgueil <150 lm (4 spectra) 1

Table 2Modal mineralogy of the Orgueil CI1 carbonaceous chondrite (Bland et al., 2004).

Phase Wt.% Vol.%

Olivine (Fo 100) 2.4 2.1Olivine (Fo 80) 3.3 2.6Olivine (Fo 60) 1.5 1.1Troilite 2.1 1.2Pyrrhotite 4.5 2.7Magnetite 9.7 5.1Serpentine 7.3 7.7Saponite–serpentine 64.2 73.8Ferrihydrite 5.0 3.7


total Fe (18.5 wt.%) is present in phyllosilicates (Herr and Skerra,1968). More recently, Bland et al. (2008) used Mössbauer analysisto determine that the Fe in Orgueil is present in silicates (5%), mag-netite (45%), and a paramagnetic component (mostly Fe3+-bearing;50%). The matrix consists of intergrowths of roughly equal propor-tions of Fe-bearing, Mg-rich serpentine and saponite (Tomeoka,1990; Bland et al., 2004), Bass (1971) reported the bulk of the crys-talline material as consisting of montmorillonite and serpentine,while Kerridge (1964) found the matrix to contain a phyllosilicatewith a serpentine layer spacing. Zaikowski (1979) identified cham-osite rather than any polymorphs of serpentine in Orgueil on the ba-sis of infrared spectral analysis. Phyllosilicates in Orgueil are largelyfine-grained (<100 nm) with a poorly-defined crystallinity consis-tent with montmorillonite-type clays (Mackinnon and Kaser,1988). The matrix also contains a poorly crystalline Fe-rich material,likely ferrihydrite (Tomeoka and Buseck, 1988; Tomeoka, 1990;Bland et al., 2004; Beck et al., 2010), which formed from pre-existingmagnetite and Fe–Ni sulfides (Tomeoka and Buseck, 1988).

Phyllosilicates comprise 60–65% of Orgueil and contain bothFe2+ and Fe3+ (10 and 11 wt.% FeO and Fe2O3, respectively) (Ker-ridge, 1976). The matrix averages Fe/(Fe+Mg) of 0.39 (Zolenskyet al., 1993), while the matrix phyllosilicates have an Fe/(Fe+Mg)of 0.15 (Tonui et al., 2003) (0.55 for the ‘‘chlorite’’ (Bostrom andFredericksson, 1966)). Olivine was found to be present by electrondiffraction analysis (Kerridge, 1964). Matrix olivine compositionsare heterogeneous and range from Fa5 to Fa75 (Peck, 1983).

Bulk C content is 4.88 wt.% (Pearson et al., 2006) or 2.80 wt.% C(Jarosewich, 1990). Organic matter is closely associated with thephyllosilicates (Pearson et al., 2002). Approximately 70–75% ofthe C is present as large polyaromatic units (Gardinier et al.,1999). The fraction of aromatic carbon is 0.49 (Cody et al., 2003).

Orgueil also contains abundant fine-grained (<1 lm size) mag-netite (Brearley, 1992a; Bland et al., 2004); thermomagnetic anal-ysis indicates that magnetite content is 11.0 ± 0.4 wt.% (Hyman andRowe, 1983). Total S content is 1.39 wt.% (Zolensky et al., 1993), or2.8–4.4 wt.%, which is present in elemental, organic, sulfide, andsulfate forms, with the sulfate form dominating (Burgess et al.,1991). Veins in Orgueil are dominated by carbonates, Ca-sulfate,and Mg-sulfate (epsomite and hexahydrite), the latter being thevolumetrically dominant species (Richardson, 1978; Frederickssonand Kerridge, 1988). The presence of magnetite and gypsum wasalso identified by X-ray diffraction analysis (Bass, 1971).

Orgueil 74–495 lm 3Orgueil 53–147 lm 3Orgueil <74 lm 6Orgueil <150 lm (5 spectra) 2

Sources of data: [1] This study – PSF; bidirectional: i/e: 0�/30�, 0�/60�, 13�/0�, 30�/�60�, 30�/0�, 30�/30�, 30�/60�, 60�/0�, 60�/30�, 60�/60�. Alais was also measured ati = 30� and e = 0� after fluffing with a needle, and after dense packing. Orgueil wasalso measured at i = 30� and e = 0� after fluffing with a needle. [2] Gaffey (1974) –available through RELAB; integrating sphere. [3] Johnson and Fanale (1973); inte-grating sphere. [4] RELAB archive; bidirectional: i = 30�, e = 0�. [5] Hiroi et al. (1993),<100 or <125 lm; bidirectional: i = 30�, e = 0�. [6] Salisbury et al. (1975); bidirec-tional: 15� phase angle.

5. Weathering effects

Terrestrial weathering is a factor which must be consideredwhen analyzing reflectance spectra of CCs. Terrestrial weatheringcan include the formation of brown (iron oxide) staining of silicates,cavities due to the loss of material by leaching, formation of sulfateveins, loss of sulfides, and oxidation of magnetite and Fe–Ni metal(e.g., Kallemeyn et al., 1991). Terrestrial weathering is often charac-terized by alteration of any pre-existing metal and sulfides to iron

oxides and hydroxides, particularly ferrihydrite (e.g., Brearley,1997) and limonite (Rubin and Kallemeyn, 1990; Brearley,1992b), which will appear as red-brown staining and/or diffuse‘‘halos’’ of rust, as well as the deposition of evaporates/bicarbonates(Grady et al., 1991). Goethite, limonite and ferrihydrite are typicalterrestrial weathering products of metal grains (Noguchi, 1994).Terrestrial oxidation can proceed rapidly (Lee et al., 2006) and issometimes difficult to separate from pre-terrestrial alteration(Gooding, 1986; Burns et al., 1995; Brearley, 1997; Bland et al.,2000). Even recent falls can undergo rapid terrestrial weathering,particularly if they contain susceptible minerals (e.g., Bass, 1971;Okada et al., 1981). There seems little doubt that weathering prod-ucts can affect reflectance spectra (Salisbury and Hunt, 1974), asiron oxides/hydroxides have intense absorption bands in the visibleregion (e.g., Morris et al., 1985) and commonly pervade the interiorof finds (e.g., Schwarz and Mason, 1988; Schwarz et al., 1989). In theparticular case of the CIs, the extensive sulfate veining may be theresult of terrestrial alteration (Gounelle and Zolensky, 2001; Blandet al., 2004). Terrestrial weathering seems to have little or no effecton bulk C content (Pearson et al., 2006).

6. Experimental procedure

This study focuses on analysis of reflectance spectra of pow-dered samples of various CIs. It includes existing spectra fromthe RELAB data base (http://relab.brown.edu), many of which havenot been compared or analyzed in detail (e.g., Hiroi et al., 1993,1994, 1997), CI reflectance spectra from other sources (Johnsonand Fanale, 1973; Gaffey, 1974; Salisbury et al., 1975), new spectraof Alais and Orgueil (Table 3), and reflectance spectra of most of theconstituent phases found in CIs and selected phyllosili-cate + opaque mixtures. Our analysis is focused on the 0.3–2.5 lm interval, as this region is available for the largest numberof CI spectra.

The reflectance spectra measured at the NASA-supported RELABfacility were measured at i = 30� and e = 0� in bidirectional reflec-tance mode relative to halon for the �0.3–2.5-lm region at 5 nmintervals, and corrected for minor irregularities in halon’s absolutereflectance in the 2.0–2.5-lm region. Details of the RELAB facilityare available at the RELAB web site (http://www.plane-tary.brown.edu/relab/). Standard reflectance spectra measured atthe University of Winnipeg Planetary Spectrophotometer Facility(PSF) were acquired with an ASD FieldSpec Pro HR spectrometer

http://relab.brown.edu

http://www.planetary.brown.edu/relab/

http://www.planetary.brown.edu/relab/

Table 4Compositional data for some of the minerals used in the analysis.

Sample Berthierine Chlorite Cronstedtite Greenalite NontroniteSample ID BER102 CLI101 CRO101 GRE001 NON101NMNH #a 93555 193910 158902

Wt.%SiO2 27.94 33.78 16.15 28.05 61.40Al2O3 7.92 21.63 0.24 0.62 9.49TiO2 0.52 0.22 0.00 0.03 0.38Fe2O3 30.59 4.23 63.51 25.05 22.08FeO 20.41 5.96 15.96 36.43 0.79Fe2O3

b (53.27) (10.85) (81.25) (65.54) (22.96)MnO 0.11 0.20 0.16 1.23 0.03MgO 3.42 32.18 0.22 3.06 1.07CaO 4.57 0.07 0.56 0.30 0.10Na2O 0.28 0.11 0.72 0.80 0.15K2O 0.05 0.00 0.00 0.00 3.67P2O5 1.81 0.02 0.12 0.05 0.51Totalc 99.89 99.38 99.42 99.68 99.76

LOId 6.45 11.89 8.64 10.81 8.56

ppmSr 167 8 5 50 40Zr 210 15 2 20 445V 1000 140 27 60 70Cr 390 375 6 50 15Ni 1120

XRDe ber chl cro gre nonmn cas mn spi mn qtz

Sample Saponite Saponite Saponite Serpentine SerpentineSample ID SAP101 SAP102 SAP103 SRP106 SRP117NMNH #a C3810 170279

Wt.%SiO2 62.89 56.31 46.92 45.30 40.74Al2O3 18.76 5.14 16.17 1.28 0.36TiO2 0.90 0.72 1.07 0.04 0.00Fe2O3 7.29 0.42 7.09 7.28 6.80FeO 0.39 0.54 3.67 1.60 2.51Fe2O3

b (7.72) (0.96) (11.17) (9.06) (9.59)MnO 0.04 0.03 0.29 0.12 0.15MgO 1.85 30.86 7.82 43.20 46.89CaO 2.30 2.18 8.68 0.58 0.09Na2O 2.12 2.87 2.84 0.02 0.15K2O 2.69 0.41 0.39 0.00 0.00P2O5 0.16 0.13 0.13 0.02 0.01Totalc 99.43 99.61 99.83 99.62 98.55f

LOId 12.27 14.36 6.16 14.86 17.23

ppmSr 386 70 299 58 36Zr 580 316 104 10 24V 117 50 204 10 67Cr 52 32 221 2820 5815Ni 2469 2570XRDe sap sap sap ser ser

fld mn amp mn alb mag

Sample Goethite Magnetite Hexahydrite Dolomite Siderite Magnesite TroiliteSample ID OOH003 MAG103 SPT143 CRB103 CRB108 CRB114 TRO201g

Wt.% Wt.%SiO2 1.02 0.00 0.00 0.02 0.00 31.21 Fe 61.50Al2O3 0.47 0.43 0.06 0.00 0.08 0.00 Ni 0.03TiO2 0.04 7.28 0.00 0.02 0.03 0.00 Na 0.01Fe2O3 96.17 60.38 0.00 n.d. 55.08 0.00 S 37.71FeO 1.02 28.80 0.00 n.d. 35.67 0.00 V 0.00Fe2O3

b (96.64) (83.13) (0.00) (6.84) (94.72) 0.00 Co 0.02MnO 0.23 2.00 0.01 0.14 3.83 0.01 P 0.00MgO 0.89 <0.01 25.09 32.42 0.76 53.79 Cr 0.32CaO 0.07 0.00 0.46 60.19 0.07 14.63 Mg 0.00Na2O 0.00 0.00 0.94 0.16 0.66 0.02 Mn 0.00K2O 0.02 n.d. 0.09 0.00 0.01 0.00 Cu 0.02P2O5 0.61 n.d. 0.04 0.02 0.03 0.04 Al 0.02SO3 n.d. n.d. 23.53 n.d. n.d. n.d. Zn 0.08


Table 4 (continued)

Sample Goethite Magnetite Hexahydrite Dolomite Siderite Magnesite TroiliteSample ID OOH003 MAG103 SPT143 CRB103 CRB108 CRB114 TRO201g

Totalc 99.95 99.37 50.22 99.79 100.19 99.70 99.69

LOId 10.36 n.d. 38.11 46.66 32.63 34.57

ppmSr 5 n.d. 90 144 25 327Zr 15 n.d. 25 5 5 20V 45 n.d. 30 20 45 10Cr 5 n.d. 15 <2 <2 12XRDe goe mag hex dol sid mgs tro

hun prh

Abbreviations: alb: albite; amp: amphibole; ber: berthierine; cas: cassiterite; chl: chlorite; cro: cronstedtite; dol: dolomite; fld: feldspar; goe: goethite; gre: greenalite; hex:hexahydrite; hun: huntite; mag: magnetite; mgs: magnesite; mn: minor; n.d.: not determined; non: nontronite; prh: pyrrhotite; qtz: quartz; sap: saponite; ser: serpentine(lizardite); sid: siderite; spi: spinel; tro: troilite.Sample descriptions and localities: BER102: berthierine (France). CLI101: clinochlore (near Murphys, Calaveras Co., CA, USA). CRB103: dolomite (Bamble, Telemark, Norway).CRB108: siderite (Ivigtut, Frederikshaab District, West Greenland). CRB114: magnesite (Currant Creek, NV, USA). CRO101: cronstedtite (411 level, Lallagua, Bolivia). GRE001:greenalite (La Union, Murcia, Spain). MAG103: magnetite (Langesundfjord, Norway). NON101: nontronite (Allentown, Lehigh Co. PA, USA). OOH003: goethite (Cary Mine,Ironwood, Gogebic Co., MI, USA). SAP101: saponite (Menzenberg, N. Hoffen, Rheinish Prussia, Germany). SAP102: Fe-poor saponite (Ballaray, CA, USA). SAP103: Ferroansaponite (Griffith Park, Los Angeles Co., CA, USA). SRP106: magnetite-bearing serpentine (var. lizardite) (Sand Dollar Beach, coast of S. Monterey Co., CA, USA). SRP117:magnetite-free serpentine (var. lizardite) (Fort Point, San Francisco Co., CA, USA). SPT143: hexahydrite (Basques Lakes, near Ashcroft, BC, Canada). TRO201: inclusion fromCanyon Diablo iron meteorite.

a NMNH = National Museum of Natural History sample ID.b All Fe reported as Fe2O3.c Total expressed on a volatile free basis and assuming all Fe as Fe2O3.d LOI = loss on ignition – weight loss after heating sample in air to 950 �C.e Phases identified by X-ray diffraction.f Includes 0.74 wt.% SO3.g Analysis performed at the University of Alberta.


from 0.35 to 2.5 lm at i = 30� and e = 0� relative to Spectralon� andcorrected for minor reflectance irregularities in the 2.0–2.5-lm re-gion as well as detector offsets at 1.00 and 1.83 lm. Details of PSFare available on the PSF web site (http://psf.uwinnipeg.ca). Stan-dard spectra of Orgueil and Alais were measured with a lamp-black-coated aluminum mask (with a 5 mm diameter centralhole) placed over the samples and Spectralon standard. Use of thismask was required as the sample did not always fill the full field ofview of the spectrometer. For Orgueil, insufficient sample wasavailable to fully fill the desired sample cup, as a result absolutereflectance values should be considered as approximate. Our sam-ple of Alais was also measured at a variety of viewing geometriesrelative to Spectralon measured at i = 13� and e = 0� (the smallestphase angle possible with our current set-up). Some of the spectraexhibit narrow features in the 0.62–0.67-lm region which are arti-facts from an order sorting filter in the ASD spectrometer.

Band minima wavelength positions were largely determinedusing visual criteria, supplemented by fitting horizontal chords atvarious positions across an absorption band and determining themid-point, and then averaging the mid-points or projecting themid-points to the band minimum. This latter approach is most use-ful for noisy data and weak absorption bands. Absorption banddepths were calculated using Eq. (32) of Clark and Roush (1984)after fitting a straight line continuum tangent to the spectrum oneither side of an absorption feature of interest.

Mineral samples that were used in our analysis were obtainedfrom various sources and ground by hand in an alumina mortarand pestle and dry sieved to obtain the requisite size fractions.Impurities were removed through a combination of visual inspec-tion with a binocular microscope and with a hand magnet. Thesamples were characterized using X-ray fluorescence (XRF) for ma-jor elements (Mertzman, 2000). Ferrous iron was determined usinga modification of the procedure outlined in Reichen and Fahey(1962). Ferric iron concentration was taken as the difference be-tween total and ferrous iron. Volatile contents were determinedby heating splits of the samples to 950 �C for one hour and measur-ing associated weight loss. The samples were structurally charac-terized using X-ray diffractometry (XRD) to ascertain the phases

present in the samples and to verify the phases determined byXRF. XRD analytical procedures are described in Cloutis et al.(2004). Sample descriptions, compositions and phases identifiedin the mineral samples through integrated XRF and XRD analysisare provided in Table 4.

6.1. Spectral metrics

In order to provide some measure of spectral slopes that couldprovide insights into how different CI spectra differ and how opa-que material may affect the spectra of constituent minerals, wedeveloped a number of spectral and sample metrics. For samplemetrics, we included in our analysis the minimum grain size ofthe CIs and mineral mixtures (set to 0 lm for samples for whichonly a maximum grain size is specified), and an average grain size(simply the average between the largest and smallest grain sizesfor sieved samples).

For spectral metrics we include the absolute reflectance of thepeak or most prominent slope break in the 0.6–0.8-lm region (‘‘peakreflectance’’) – we take this point as either the local or absolutereflectance maximum or the reflectance near the point of maximumslope change; we also use maximum reflectance of a spectrumregardless of wavelength (‘‘maximum reflectance’’) and absolutereflectance at 1.5 lm, which is outside the region of expectedabsorption bands. We use the ratio of reflectance at 2.2 lm (whichfalls outside the region of major absorption bands in both CI spectraand phyllosilicates and is not influenced by the H2O-induced down-turn toward longer wavelengths) to reflectance at the 0.6–0.8-lmregion peak or slope change, as a measure of overall slope (‘‘reflec-tance ratio’’). This ratio is <1 for blue-sloped spectra.

7. Spectral properties of CI constituent phases

7.1. Anhydrous silicates

Olivine is the only anhydrous mafic silicate identified in CIs.However it is rare (< few wt.%) and, when coupled with the low

http://psf.uwinnipeg.ca


albedo of CIs and more abundant Fe-bearing phyllosilicates, is notexpected to contribute to the spectral signature of CI chondrites.

7.2. Phyllosilicates

The nature of phyllosilicates in CIs is still imperfectly under-stood (Roy-Poulsen et al., 1981; McSween, 1987; Tomeoka, 1990;Brearley and Jones, 1998). This arises from the fact that they arecommonly finely intergrown with additional phases such as sul-fides, ferrihydrite, magnetite, and organic matter (which compli-cates phyllosilicate identification from compositionalinformation), and many are members of solid solution series(which complicates phyllosilicate identification from structuralinformation). It appears that the most common phyllosilicates inCIs are an Fe2+ and Fe3+-bearing serpentine and an Fe-bearingsmectite (saponite) (Roy-Poulsen et al., 1981; McSween, 1987;Tomeoka, 1990; Brearley and Jones, 1998). Matrix phyllosilicatesin CIs occur in two major forms: coarse-grained and opaque-poor,or fine-grained with abundant fine-grained and finely-dispersedopaques such as magnetite and ferrihydrite (Nagy, 1975; Tomeokaand Buseck, 1988; Brearley, 1992a).

Phyllosilicate spectra tend to be dominated by absorption bandsin the 0.3–1.2-lm region when Fe is present, OH absorption bandsin the 1.4-lm region, OH and H2O absorption bands in the 1.9-lmregion, and metal-OH stretching combination bands in the 2.2–2.5-lm region. Thus, multiple spectral regions are potentially availablefor phyllosilicate identification (e.g., Clark et al., 1990). Majorabsorption band positions for CI phyllosilicates are summarizedin Table 5.

The serpentine structural group of phyllosilicates encompassesa number of Fe-bearing members, including amesite (usuallyFe2+-bearing; Calvin and King, 1997), greenalite (Fe3+-bearing),cronstedtite (Fe2+- and Fe3+-bearing), and berthierine (Fe2+- andFe3+-bearing) (Roberts et al., 1990). The presence of Fe2+ in theoctahedral site leads to absorption bands near 0.90–0.98 and

Table 5Wavelength positions of major absorption bands of CI constituent minerals.

Phase Absorption band position(lm)

Cause

Serpentines �0.70–0.75 Fe2+–Fe3+ chargetransfer

�0.90–0.94 Octahedral Fe2+


�1.4 OH�1.9 OH/H2O�2.3 Mg–OH

Saponites �0.65–0.68 Fe2+–Fe3+ chargetransfer

�0.9 Octahedral Fe2+


�1.4 OH�1.9 OH/H2O�2.3 Mg–OH

Maghemite <�0.5 Fe–O charge transfer

Ferrihydrite <�0.5 Fe–O charge transfer�0.85–0.90 Fe3+ spin-forbidden�1.4 OH�1.9 OH/H2O

Magnetite �0.48 Fe3+ spin-forbidden�1–1.3 Octahedral Fe2+

Hexahydrite/epsomite

�1.45 H2O

�1.95 H2O�2.5 H2O

Carbonates �2.3 C–OP2.5 C–O

1.10–1.15 lm due to crystal field transitions (Sherman and Vergo,1988). The presence of Fe3+ results in lower overall reflectance, areduction in OH and MgOH absorption band intensities, and ared-sloped spectrum. When both Fe2+ and Fe3+ are present, anadditional charge transfer absorption band appears near 0.70–0.75 lm (Fig. 1a), and an additional weaker absorption band near0.65–0.66 lm may also be present in some serpentine spectra. Thepresence of intimately associated fine-grained magnetite can im-part an overall bluer spectral slope and lower overall reflectance(Fig. 1b). Very Fe-rich varieties (thuringite) may exhibit low overallreflectance (<5%) and a blue-sloped spectrum beyond �0.8 lm(Calvin and King, 1997). Among the serpentine-group phyllosili-cates, cronstedtite exhibits the most complex reflectance spec-trum, with absorption bands or prominent inflections appearingnear 0.71, 0.85, 0.92, 1.1, and 1.25 lm.

Reflectance spectra of smectite group phyllosilicates, which in-cludes saponite, often exhibit a more complex absorption featurein the 1.4-lm region, a broad H2O absorption feature near1.9 lm, and metal-OH absorption bands in the 2.3–2.4-lm region.The presence of Fe2+ leads to absorption bands near 0.90–0.94 and1.10–1.15 lm, overlapping the values for serpentines. The addi-tional presence of Fe3+ results in a charge transfer band near0.65–0.68 lm (vs near 0.70–0.75 lm in serpentines) (Shermanand Vergo, 1988) (Fig. 1c) (Table 5).

7.3. Oxides and hydroxides

Fe–Ni metal is essentially absent in CI chondrites, due to theextensive aqueous alteration they have undergone. Consequentlymagnetite (Fe3O4), maghemite (c-Fe2O3), and ferrihydrite(5Fe2O3�9H2O) are the dominant iron oxide/hydroxide species,and appear to be finely dispersed in the fine-grained matrix phyl-losilicates; some of the Fe-hydroxides may also be the result of ter-restrial weathering.

Spectra of submicron maghemite are characterized by an in-tense Fe–O charge transfer absorption feature below �0.5 lmand variable reflectance at longer wavelengths, depending onFe2+ abundance in the sample; some exhibit reflectance decreases(blue slope) beyond �1.1 lm (Sherman et al., 1982; Morris et al.,1985).

Ferrihydrite spectra also exhibit a region of strong absorptionbelow �0.5 lm, a generally well-resolved Fe3+ spin-forbiddenabsorption band near 0.85–0.9 lm, O–H stretching combinationand overtone bands in the 1.4-lm region, and O–H stretching plusH–O–H bending combinations in the 1.9-lm region (Shermanet al., 1982). Limonite, a probable weathering product, has spectralproperties that are broadly similar to ferrihydrite (Singer, 1980).Fig. 1d shows the reflectance spectrum of magnetite, maghemite,ferrihydrite, and goethite.

Magnetite spectra are characterized by low overall reflectance,a weak absorption band near 0.48 lm likely due to a spin-forbid-den Fe3+ absorption band, and a broad absorption band centeredbeyond �1.0 lm that is attributable to crystal field transitions inoctahedrally coordinated Fe2+ (Fig. 1e). In contrast to silicate min-erals, magnetite becomes darker and increasingly blue-sloped withdecreasing grain size, and reflectance spectra of submicron-size(cation-deficient) magnetites measured by Morris et al. (1985)show essentially no reflectance upturn longward of �1 lm (Huntet al., 1971; Morris et al., 1985); magnetite in CIs is generally cat-ion-deficient and of submicron size (Kerridge, 1970, 1976; Kerridgeet al., 1979; Hua and Buseck, 1998). Both the real and imaginaryparts of the dielectric function of magnetite increase sharply above�1 lm (Buchenau and Müller, 1972; Müller and Buchenau, 1975;Schlegel et al., 1979; Boppart et al., 1980). Major Fe oxide/hydrox-ide absorption bands are summarized in Table 5.

Fig. 1. Reflectance spectra (i = 30�, e = 0�) of <45 lm fractions (unless otherwise indicated) of various CI constituent minerals. (a) Serpentine-group phyllosilicates:berthierine (BER101), cronstedtite (CRO101), greenalite (GRE001), and serpentine (SRP117); (b) magnetite-bearing (SRP106) and magnetite-free (SRP117) serpentines; (c)saponite-group phyllosilicates: nontronite (NON101), Fe-poor saponite (SAP102), ferroan saponite (SAP103); (d) iron oxides/hydroxides: synthetic ferrihydrite andmaghemite (from RELAB archive), goethite (OOH003), and magnetite (MAG103); (e) 20-nm size synthetic magnetite; (f) synthetic polycyclic aromatic hydrocarbons; (g)amorphous and crystalline graphite and anthracite coal; (h): 21-nm size carbon lampblack; (i) <125 lm size insoluble matter from Murchison CM2 chondrite; (j) troilite fromthe Canyon Diablo iron meteorite; (k) sulfates: synthetic epsomite and hexahydrite (SPT143); (l) carbonates: dolomite (CRB103), magnesite (CRB114), siderite (CRB108).Spectra in 1a–d, g–i, and l measured at RELAB; spectra in 1e, f, j and k measured at PSF. See Table 4 for compositions.


Fig. 1 (continued)


7.4. Carbonaceous phases

CCs were initially defined as a distinct meteorite group on thebasis of their dark color, resulting at least in part from finely-dis-seminated carbonaceous material (Mason, 1962). This materialconsists predominantly (�70–90% of the total C) of a macromolec-ular, C- and aromatic-rich insoluble material (e.g., Vinogradov andVdovykin, 1964; Hayatsu and Anders, 1981; Wdowiak et al., 1988;Murae et al., 1990; Gardinier et al., 2000). Its structure and compo-sition can be affected by both aqueous alteration and thermalmetamorphism (Blanco et al., 1988; Yabuta et al., 2005; Busemannet al., 2007). Given its structural and compositional heterogeneity,it is difficult to define a ‘‘typical’’ spectrum to represent this mate-rial. Fig. 1f and g shows reflectance spectra of a number of C-richmaterials. The optical properties of C-rich materials are very sensi-tive to small changes in structure, composition, and grain size (e.g.,Draine, 1985; Bussoletti et al., 1987; Blanco et al., 1988; Rouleauand Martin, 1991). Carbonaceous materials containing some frac-tion of aliphatic C–H bonds can exhibit absorption bands in the1.7 and 2.3-lm region due to various C–H stretching overtonesand combinations. Complex aromatic-rich carbonaceous materialusually exhibits a dark, red-sloped, featureless spectrum (e.g.,Johnson and Fanale, 1973).

With increasing thermal processing, carbonaceous materialsgenerally undergo increasing aromatization and loss of aliphaticside chains, resulting in a darker and more featureless spectrum(e.g., Gavrilov and Ermolenko, 1975; Ito, 1992; Cloutis, 2003).Aqueous alteration generally leads to the formation of a more com-plex suite of organic molecules, including more soluble organicmolecules (Cruikshank, 1997). It appears that the insoluble carbo-naceous material in CCs may be relatively insensitive to heating (to300 �C) and exposure to vacuum (Orthous-Daunay et al., 2010).

The insoluble organic matter in both type 1 and type 2 CCs exhibitssimilar absorption features in infrared transmission spectra(Kebukawa et al., 2010).

Single phase C-rich materials most similar to the insoluble or-ganic matter in CI chondrites (in terms of being composed of afew condensed aromatic rings) generally show flat or slightlyblue-sloped spectra (Fig. 1f), a sharp absorption edge, and highoverall reflectance when pure. The high reflectance is due in partto photo-induced fluorescence, a phenomenon also seen in CIs(e.g., Vdovykin, 1962; Alpern and Benkhieri, 1973). Graphite,which is compositionally similar to anthracene and pyrene, exhib-its much lower overall reflectance, as does anthracite coal, a highlycondensed and aromatic-rich material (Fig. 1g). This illustrates theextreme spectral diversity of organic and C-rich materials, whereinsmall changes in structure and composition can have profound ef-fects on spectral reflectance. In a ‘‘typical’’ naturally occurring car-bonaceous material, a range of phases will be present; chargetransfers and conduction bands in these structurally complexmaterials will, collectively, result in strong absorption across awide wavelength range (e.g., Akamatu and Kuroda, 1963; Gavrilovand Ermolenko, 1975; Silverstein et al., 1991; Ito, 1992).

While the spectral reflectance properties of organic materialsare diverse, many are extremely dark and blue-sloped. These in-clude some carbon lampblacks (Fig. 1h), the insoluble organic mat-ter in Murchison (Fig. 1i) and a carbonaceous extract from theAllende CV3 chondrite (Johnson and Fanale, 1973).

7.5. Sulfur-bearing species

The sulfur in CI chondrites is present in a variety of forms,including elemental, organic, sulfide, and sulfates (Burgess et al.,1991). The sulfide form is most commonly troilite (FeS), which


exhibits a red-sloped spectrum, at least for <45 lm sized powder(Fig. 1j). When bound to carbon, sulfur does not contribute specificabsorption features in the 0.3–2.5-lm region (Silverstein et al.,1991). The most common sulfate forms in CIs, hexahydrite andepsomite, which could be the result of terrestrial weathering(Gounelle and Zolensky, 2001; Bland et al., 2004), exhibit highoverall reflectance, with H2O-associated absorption bands in the1.45-, 1.95-, and 2.5-lm regions (Fig. 1k; Table 5). The presenceof Fe2+ would lead to the appearance of additional absorptionbands below �1 lm (Cloutis et al., 2006).

7.6. Carbonates

Carbonates in CIs are present as intact or disrupted veins andindividual grains (Endress and Bischoff, 1994; Brearley and Jones,1998). The most common carbonates are dolomite and Fe2+-bear-ing magnesite (breunnerite) (Johnson and Prinz, 1993). Reflectancespectra of these minerals exhibit an absorption feature near 2.3 lmdue to C–O stretching overtones and the short wavelength wing ofan additional overtone located just longward of 2.5 lm (Fig. 1l; Ta-ble 5). When Fe2+ is present it results in two absorption bands near1 and 1.3 lm due to crystal field transitions in octahedrally coordi-nated Fe2+ (Gaffey, 1987). While carbonate composition and abun-dance may be correlated with degree of aqueous alteration (Petitatand Gounelle, 2010), it is unlikely that carbonate abundances aresufficient to result in strong, readily-resolvable absorption bandsin CI spectra.

8. CI reflectance spectra

Three CIs have been spectrally characterized by multiple inves-tigators, in some cases involving multiple grain sizes. These results,the spectral properties of CI constituent phases, phyllosili-cate + opaque mixtures, and analysis of CI mineralogy and petrol-ogy, can collectively provide insights into the spectral propertiesof CI chondrites, as the known CI chondrite group includes only se-ven members, all of petrologic type 1 (e.g., Weisberg et al., 2006),and the three included CIs exhibit differences in degree of aqueousalteration, composition, and petrology (see Section 4).

8.1. Alais

Only a single reflectance spectrum of Alais was previously avail-able (Fig. 2a). The spectrum is characterized by an overall red slopebetween 0.35 and 1.85 lm, with a distinct slope change near0.6 lm. There is a shoulder near 0.45 lm, and an H2O-associatedabsorption band centered near 1.92 lm, and a gradual reflectancedecline beyond �2.2 lm. There is a suggestion of a shallow broadabsorption feature in the 0.8–1.3-lm region. The shoulder near0.45 lm is coincident with an absorption band in the insoluble or-ganic matter isolated from Murchison (Fig. 1i), graphite (Fig. 1g),some serpentines (Fig. 1b), and magnetite (Fig. 1e).

The absorption feature near 1.92 lm, and the reflectance down-turn beyond 2.2 lm can plausibly be ascribed to a number of thehydrated phases present in Alais, such as saponite, ferrihydrite,or the hydrated sulfates. The 0.8–1.3-lm region absorption featureappears to consist of two absorption bands centered near 0.95 and1.10–1.15 lm, with a maximum depth of 3% (Fig. 2b). This featureis best attributed to contributions from the Fe2+ in the phyllosili-cates, specifically serpentine. A likely Fe3+–Fe2+ charge transfermay account for the slope break near 0.7 lm, although this is nota unique assignment. The absorption band in the 0.78–0.83-lm re-gion may be attributable to ferrihydrite and/or phyllosilicates.

The new spectra of Alais (Fig. 2c) are broadly similar to theexisting spectrum: visible region reflectance of �5%, a red-sloped

spectrum, and absorption features near 0.5, 0.9–1.2, and 1.90–1.95 lm. The two continuum-removed spectra show absorptionfeatures near 0.95, 1.02, and 1.17 lm, although there are notequally resolved in the two spectra (Fig. 2d). The overall profileof this region of absorption is similar to cronstedtite, but can alsobe attributed to Fe2+-bearing phyllosilicates + magnetite.

To further assess how phase angle changes affect CI spectra(e.g., French and Veverka, 1983; Capaccioni et al., 1990), we mea-sured the powdered sample of Alais under a variety of viewinggeometries. Where i is fixed (at either 0� or 30�) and e increases,we observe decreases in absolute reflectance (Fig. 2e and f) andredder overall slopes (Fig. 2g and h). When i was fixed at 60� ande increases, we observed increasing overall reflectance (Fig. 2i)and redder overall slopes (Fig. 2j). For a fixed emission angle (0�,30�, or 60�), increasing incidence angle, overall reflectance de-creases and the spectra become redder. The change in overall slopewas most apparent below �1 lm. Overall reflectance at 0.56 lmvaried from 1.8% (i = 60�, e = 0�) to 4.2% (i = 13�, e = 0�), while the2.4/0.56 lm reflectance ratio varied between 1.94 (i = 30�,e = �60�) and 2.79 (i = 60�, e = 60�).

Particle packing, which appears to have an effect on the spectralproperties of CIs (Salisbury et al., 1975) was also investigated.Fig. 2k shows reflectance spectra of the regularly packed sampleof Alais (lightly tamped with a flat surface), the same sample‘‘fluffed’’ with a needle (and consequently an irregular surface),and a more densely packed and tamped sample with a flat surface(all measured at i = 30� and e = 0�). The regularly packed sampleshows the highest overall reflectance, followed by the fluffed anddense sample spectra. In terms of overall spectral slope, the fluffedsample is slightly redder, while the dense sample is slightly bluerthan the regular packed sample (Fig. 2l). These results are consis-tent with those of Salisbury et al. (1975), Capaccioni et al. (1990),and Sakai and Nakamura (2004).

8.2. Ivuna

Two previously measured reflectance spectra of Ivuna are avail-able, both for <125 lm fractions (Fig. 3a). They differ in terms ofoverall slope and presence of absorption bands. The c1mb60 spec-trum shares some characteristics with Alais, specifically a red-sloped spectrum, a shoulder near 0.5 lm, and a broad absorptionband near 1.95 lm. There is no resolvable absorption feature inthe 0.8–1.3-lm region, however.

The second Ivuna spectrum (c1mp18) has lower overall reflec-tance (Fig. 3a). Once again, the shoulder near 0.5 lm is present,and the spectrum is blue-sloped beyond �0.65 lm. There is evi-dence for an Fe3+–Fe2+ charge transfer in the 0.75–0.8-lm region(flattening of the spectrum in this region), and 0.95–1.15-lm re-gion absorption bands are somewhat resolved (Fig. 3b); these fea-tures are most consistent with Fe2+-bearing phyllosilicates. Theconcave overall shape of the absorption feature is similar to cron-stedtite or Fe2+-bearing phyllosilicates + magnetite. The aliphaticportion of the organic component may be expressed as a subtleabsorption feature near 1.7 lm (see Fig. 1f), and some or all ofthe phyllosilicates, carbonates and sulfates may be contributingto the 1.95-lm region feature and possible weak absorption fea-tures in the 2.3–2.5-lm region.

8.3. Orgueil

Orgueil is the most spectrally characterized CI chondrite, withfive spectra in Gaffey (1974), two in Johnson and Fanale (1973),one in Salisbury et al. (1975), and six in the RELAB data base, span-ning a range of grain sizes. With increasing particle size (53–147vs. 74–495 lm) the two spectra from Johnson and Fanale (1973)show a reflectance decrease and steeper blue slope.

Fig. 2. Reflectance spectra of <150 lm size sample of Alais. (a) (i = 30�, e = 0�; spectrum mgp084 from RELAB archive); (b) same as (a) after continuum removal; (c) measuredat PSF (i = 30�, e = 0�) with repacking of sample between the two measurements; (d) same as (c) after continuum removal; (e) fixed i (0�) and variable e; (f) fixed i (30�) andvariable e; (g) same as (e) with spectra normalized at 1.87 lm; (h) same as (f) with spectra normalized at 1.87 lm; (i) fixed i (60�) and variable e; (j) same as (i) with spectranormalized at 1.87 lm; (k) spectra measured with regular packing, after fluffing the sample surface with a needle, and after compressing the sample into the sample cup(dense packing); (l) same as (k) with spectra normalized at 0.56 lm. Spectra in 2e–l are all referenced to the standard measured at (i = 13� and e = 0�). Spectra in 2c–lmeasured at PSF. See text for details of continuum removal procedure.


Fig. 2 (continued)

Fig. 3. Reflectance spectra of <125 lm size sample of Ivuna (i = 30�, e = 0�; spectra from RELAB archive). (a) Absolute reflectance; (b) spectrum c1mp18 after continuumremoval. See text for details of continuum removal procedure.


The reflectance decrease with increasing particle size is alsopresent in the RELAB Orgueil spectra (Fig. 4a). Spectral slopeschange from flat or red overall for the finest grain sizes to blue-sloped for the coarser-grained sorted powders (40–100 and 100–200 lm) (Fig. 4b and c). Beyond these general observations, thereis some diversity among individual spectra.

The <40 lm-size spectrum (cdms03) is red-sloped. It exhibits adistinct shoulder near 0.45 lm (attributable to phyllosilicates,magnetite, and/or insoluble organic matter), a slope change near0.6 lm (similar to Alais), the non-diagnostic H2O band near1.95 lm, and the likely phyllosilicate MgOH and/or carbonateabsorption band near 2.32 lm.

By contrast, the <45 lm spectrum (s1rs44) has lower overallreflectance and is not red-sloped. It does exhibit the shoulder near0.5 lm, the slope change near 0.6 lm, and the H2O absorptionband near 1.95 lm. There is also a weak H2O/OH feature near1.45 lm, close to the position expected for the hydrated Mg-sul-fates. The 1-lm region is characterized by three well-resolvedabsorption bands attributable to phyllosilicate Fe2+ crystal fieldtransitions (0.95 and 1.15 lm) and probably magnetite (1.03 lm)(Fig. 4d). The wavelength position of the main magnetite absorp-tion band varies around 1.0–1.3 lm for different magnetites.

The <100 lm Orgueil spectrum (c1mb57; Fig. 4b) has higheroverall reflectance than the <40 and <45 lm grain size spectra,

Fig. 4. Reflectance spectra (i = 30�, e = 0�) of Orgueil. (a–c) Different size fractions (from RELAB archive); (d) spectrum s1rs44 after continuum removal; (e): spectrum mgp080after continuum removal; (f) spectrum ccms03 after continuum removal; (g) <150 lm size sample mgp004 after emptying and repacking the sample between measurements;(h) same as (g) after continuum removal; (i) <150 lm size sample mgp004 as regularly packed and after fluffing the sample surface with a needle (j) same as (i) with spectranormalized at 0.56 lm. Spectra in 4a–f measured at RELAB; spectra in 4g–j measured at PSF. See text for details of continuum removal.


Fig. 4 (continued)


and is most similar to the <40 lm grain size (cdms03) spectrum inhaving an overall red-sloped spectrum, a recognizable 0.45-lm re-gion shoulder, a slope change near 0.6 lm, and weak H2O/OHabsorption features in the 1.4 and 1.95-lm regions. No absorptionfeature is seen in the 0.8–1.3 lm interval. This spectrum has noabsorption features that could be used for robust mineralidentification.

The spectra of the two <150 lm samples (mgp078, mgp080;Fig. 4c) are from Gaffey (1974). The mgp078 spectrum is spectrallyquite flat beyond 1.4 lm, has a possible 0.45-lm region shoulder,and an absorption feature in the 0.9-lm region. It should be notedthat this spectrum is affected by digitization limits (evident as stepfunctions) inherent to the spectrometer on which the data was ac-quired, but is included here for completeness. It does not exhibit aresolvable second absorption band in the 1.15-lm region, althoughit should be noted that some of the Fe-bearing phyllosilicate spec-tra (e.g., greenalite – Fig. 1a) show a more resolved 0.95 lm bandthan the accompanying 1.15 lm band. It also shows a possible2.35-lm region phyllosilicate MgOH and/or carbonate band.

The mgp080 spectrum exhibits the 0.45-lm region shoulder,H2O/OH bands in the 1.38 and 1.93-lm region and a broad0.9-lm region absorption feature. The 1.38 lm band position isconsistent with serpentine rather than saponite. The continuum-removed 0.9-lm region absorption feature (Fig. 4e) shows thestrongest band near 0.92 lm, with additional possible bands near0.98 and 1.02 lm. Evidence for a feature near 1.10 lm is indirect– the gradual rise in reflectance toward longer wavelengths be-yond 1 lm, similar to the mgp078 spectrum. The 0.92 and1.02 lm features are consistent with either Fe2+-bearing saponiteor serpentine. The 0.98–1.02 lm features may be attributable tomagnetite; alternatively, all three features could be attributableto cronstedtite.

Moving to spectra for more constrained grain sizes, the 40–100 lm Orgueil spectrum (cfms03; Fig. 4a) is blue-sloped beyond0.62 lm and shows only minor evidence for an Fe3+–Fe2+ chargetransfer band near 0.76 lm, and a weak H2O band in the 1.95-lm region. Overall reflectance is generally lower than for the finergrain sample spectra.

The largest grain size spectrum (100–200 lm, ccms03) has thelowest overall reflectance and is also blue-sloped. It exhibits a un-ique concave shape beyond �0.9 lm. There are weak absorptionfeatures in the 1.45 and 1.95-lm regions and possibly near2.37 lm. An absorption feature near 2.37 lm is seen in the berthi-erine and Fe-poor saponite spectra, and is likely due to an MgOHcombination. The 0.9-lm region shows evidence for contributionfrom Fe2+ in serpentine: regions of absorption between 0.92 and0.98 and in the 1.1-lm region (Fig. 4f). There is also a possibleabsorption feature at 0.41 lm in this spectrum. It may be attribut-

able to V/Ni-bearing porphyrins that are present in Orgueil (Hodg-son and Baker, 1964; Holden and Gaffey, 1987).

The new spectra of Orgueil (Fig. 4g) are broadly similar to someof the existing ones: low overall reflectance and a flat to blue over-all slope, a broad weak absorption feature in the 0.9–1.2-lm re-gion, and an H2O absorption band in the 1.90–1.97-lm region.The continuum-removed 0.9–1.2-lm region is similar to Alais(Fig. 2d), showing evidence of single absorption bands near 1.00and 1.15 lm and perhaps an additional band near 0.90 lm(Fig. 4h).

Fig. 4i shows reflectance spectra of the regularly packed sampleof Orgueil compared to the same sample fluffed with a needle(measured at i = 30� and e = 0�). As with Alais, the fluffed sampleshows lower overall reflectance than the regularly packed sampleand both have similar overall spectral slopes (Fig. 4j).

9. Discussion

CI spectra exhibit a variety of spectral shapes, and their mostcommon characteristics are low overall reflectance (<10%), a slopechange or peak near 0.6 lm, and weak (<4% deep) absorptionbands in the 0.8–1.2-lm region. Absorption bands that do appearare of variable utility for detecting the presence of specific compo-nents, and in many cases confident detections of absorption fea-tures are hampered by noise. Here we examine the utility of thevarious absorption bands that have been detected, and discussthe lack of expected absorption bands. We have also producedand measured spectral reflectance properties of a number of min-eral mixtures designed to provide insights into the nature of CIreflectance spectra.

9.1. Phyllosilicate + opaque mineral mixtures

A major control on the appearance of absorption features asso-ciated with the non-opaque minerals in CIs (e.g., phyllosilicates,sulfates, carbonates) are the fine-grained opaque phases (e.g., mag-netite, organics) that pervade the matrices of CIs. Laboratory stud-ies of intimate mixtures involving opaque materials have shownthat finely-dispersed fine-grained opaques are very effective atsuppressing even strong absorption bands in phyllosilicates andother minerals (e.g., Clark, 1983; Cloutis et al., 1990; Millikenand Mustard, 2007).

Insoluble organic matter is the dominant organic component inCCs. Its initial abundance in chondrite matrices seems to have beenfairly constant across all chondrite groups (Pearson et al., 2001,2006). As a result, the initial abundance is directly related to thematrix abundance. The variations in abundance that are now seen


within a meteorite group are mostly due to the degree of heating,either hydrothermal or anhydrous. In enstatite meteorites, forexample, heating under very reducing conditions has producedgraphitic material with some loss of C. In the CO, CV, and OCs, con-ditions were more oxidizing and destruction (presumably to COand CO2) competed with condensation/graphitization. The CV3chondrite Allende, for example, may have lost �2/3 of its initialIOM C and what was left is highly condensed. Hydrothermal heat-ing of IOM, even below 300�C can lead to some destruction/alter-ation of the IOM. But heating at 300 �C in vacuum or in an inertgas for short periods probably has little effect, although some Ofunctionality may be lost (Orthous-Daunay et al., 2010).

The IOM in the most primitive meteorites is made of small aro-matic moieties that are highly substituted and cross linked byshort, highly-branched aliphatic chains. In terms of H/C ratio, prim-itive IOM resembles a fairly mature coal, but structurally differsfrom it. The IOM is dispersed in meteorite matrix as discrete parti-cles that are a few microns across or less (at least down to 50–100 nm). Some organics may be intergrown with phyllosilicates(e.g., Garvie and Buseck, 2004, 2005, 2006, 2007).

To provide some insights into how fine-grained opaques affectmineral reflectance spectra, we produced intimate mixtures of var-ious phyllosilicates (serpentine, saponite, and chlorite) with differ-ent opaque materials (carbon black, graphite, and magnetite)(Table 6). These mixtures supplement those available from previ-ous studies (Johnson and Fanale, 1973; Singer, 1980, 1981; Clark,1983; Milliken and Mustard, 2007). We focused on mixtures withopaque abundances similar to those found in CIs:�5 wt.% carbona-ceous materials (represented by lampblack and graphite), and �10wt.% magnetite. What are not well simulated are grain sizes andthe physical distribution of these opaques relative to the abundantphyllosilicates.

Table 6Mineral mixtures produced to examine slope-altering effects of opaque materials.

Phase 1 Wt.% Phase 2 Wt.%

Serpentine (SRP117) 100 Carbon black (LCA101) 0Serpentine (SRP117) 99.9 Carbon black (LCA101) 0.1Serpentine (SRP117) 99.5 Carbon black (LCA101) 0.5Serpentine (SRP117) 98 Carbon black (LCA101) 2Serpentine (SRP117) 95 Carbon black (LCA101) 5Serpentine (SRP117) 0 Carbon black (LCA101) 100Serpentine (SRP106) 100 Carbon black (LCA101) 0Serpentine (SRP106) 99.9 Carbon black (LCA101) 0.1Serpentine (SRP106) 99.5 Carbon black (LCA101) 0.5Serpentine (SRP106) 98 Carbon black (LCA101) 2Serpentine (SRP106) 95 Carbon black (LCA101) 5Serpentine (SRP106) 0 Carbon black (LCA101) 100Saponite (SAP101) 100 Carbon black (LCA101) 0Saponite (SAP101) 99.9 Carbon black (LCA101) 0.1Saponite (SAP101) 99.5 Carbon black (LCA101) 0.5Saponite (SAP101) 98 Carbon black (LCA101) 2Saponite (SAP101) 95 Carbon black (LCA101) 5Saponite (SAP101) 0 Carbon black (LCA101) 100Chlorite (CLI101) 100 Carbon black (LCA101) 0Chlorite (CLI101) 99.9 Carbon black (LCA101) 0.1Chlorite (CLI101) 99.5 Carbon black (LCA101) 0.5Chlorite (CLI101) 98 Carbon black (LCA101) 2Chlorite (CLI101) 95 Carbon black (LCA101) 5Chlorite (CLI101) 0 Carbon black (LCA101) 100Serpentine (SRP117) 100 Graphite (GRP102) 0Serpentine (SRP117) 99.5 Graphite (GRP102) 0.5Serpentine (SRP117) 98 Graphite (GRP102) 2Serpentine (SRP117) 95 Graphite (GRP102) 5Serpentine (SRP117) 0 Graphite (GRP102) 100Serpentine (SRP117) 100 Magnetite (MAG102) 0Serpentine (SRP117) 95 Magnetite (MAG102) 5Serpentine (SRP117) 90 Magnetite (MAG102) 10Serpentine (SRP117) 0 Magnetite (MAG102) 100

All the samples used in these mixtures had a grain size of <45 lm except the carbonblack (<0.021 lm).

Spectra of the mixtures generated for this study are shown inFig. 5. Figs. 5a–d shows reflectance spectra of intimate mixturesof <45 lm magnetite-free (neutral-sloped) (SRP117) or magne-tite-bearing (SRP106) serpentine with submicron (<0.021 lm) car-bon black. With increasing lampblack abundance, the spectrabecome darker, as expected, and less blue-sloped, approaching aflat overall spectral slope. This trend is systematic for the SRP106sample and nearly so for the SRP117 sample. The spectral slopesof the mixtures become less blue (redder) even in the wavelengthrange where the lampblack is blue-sloped (�0.5–1.1 lm).

Even with 5 wt.% lampblack (similar to the average C content ofCIs), absolute reflectance of the lampblack-containing mixtures(>13%) still exceeds the maximum reflectance of the CI spectra(10%), suggesting that mixing <45 lm phyllosilicates with thismaterial is not effective for reproducing the low reflectance ofCIs. The Fe-associated serpentine absorption bands below 1.2 lmare not entirely suppressed with 5 wt.% lampblack.

Reflectance spectra of other phyllosilicate + lampblack mixturesshow more variable spectral changes. Montmorillonite + lamp-black mixture spectra of Clark (1983) become more blue-slopedand darker with increasing lampblack content; at 5 wt.% lampblackthe mixture has a 2.2/0.8 lm reflectance ratio of 0.82 (0.98 for puremontmorillonite) and absolute reflectance of 11%, which is bright-er than comparably blue-sloped CI spectra. Milliken and Mustard(2007) showed similar results for their montmorillonite + lamp-black series: the reflectance ratio became bluer sloped withincreasing lampblack (from 0.98 to 0.84 at 5 wt.% lampblack) anddarker (8% maximum reflectance with 5 wt.% lampblack); againbrighter than comparably blue-sloped CI chondrites. Johnson andFanale (1973) also measured a 95–5% mixture of montmorilloniteand lampblack which gave a reflectance ratio of 0.76 (versus 0.86for the pure montmorillonite) and 8.6% reflectance; again this isbrighter than comparably blue-sloped CI spectra (see Fig. 6).

Saponite (SAP101) and chlorite (CLI101) were also mixed withcarbon black to assess their behavior in such mixtures (Fig. 5e–h). The saponite was selected to represent an Fe3+-bearing smectiteand the chlorite to represent a red-sloped Fe3+/Fe2+-bearing phyl-losilicate. The saponite spectrum exhibits the expected downturntoward the UV due to metal-O charge transfers and weak to non-existent Fe absorption bands. The various absorption bands inthe saponite spectrum, located in the 1.4-, 1.9-, and 2.2-lm regionsare almost completely suppressed with 5 wt.% lampblack (Fig. 5e).The lampblack also reduces the steepness of the red spectral slopebelow �0.8 lm, and introduces a very slight reddening beyond0.8 lm (Fig. 5f). The initially neutral saponite spectrum does notbecome blue-sloped with increasing lampblack – the reflectanceratio remains �1 regardless of lampblack content, and maximumreflectance (13%) is higher than the CI chondrites.

The chlorite spectrum exhibits strong Fe absorption bands near0.7, 0.9 and 1.1 lm, which are not entirely suppressed by even 5wt.% lampblack (Fig. 5g). The pure chlorite is strongly red-slopedand increasing lampblack reduces the overall spectral slope, butagain does not make the spectrum blue-sloped (Fig. 5h).

It should be noted that lampblack spectra measured by differentinvestigators show diverse spectral behavior. The spectra of<1 lm-sized lampblack of Milliken and Mustard (2007) and lamp-black of Johnson and Fanale (1973) have absolute reflectance <2.2%and are both blue-sloped across the entire 0.35–2.5 lm interval.The <1 lm-sized lampblack spectrum of Clark (1983) has anabsorption feature near 0.65 lm, is blue-sloped beyond �1.2 lm,and has absolute reflectance <3.2%. Our lampblack spectrum hasabsolute reflectance <2%, is blue-sloped from 0.4 to �1.2 lm, andred-sloped from 1.2–2.5 lm. Thus, even a material as ‘‘simple’’ ascarbon lampblack can show diverse spectral shapes and slopes.

When the SRP117 serpentine is mixed with <45 lm amorphousgraphite, the spectra become darker (Fig. 5i) and increasingly blue-

Fig. 5. Reflectance spectra (i = 30�, e = 0�) of various intimate mineral mixtures. (a) Serpentine (SRP117, <45 lm) and lampblack (LCA101, <0.021 lm); (b) same as (a) withspectra normalized at 0.63 lm; (c) serpentine (SRP106, <45 lm) and lampblack (LCA101, <0.021 lm); (d) same as (c) with spectra normalized at 0.56 lm; (e) saponite(SAP101, <45 lm) and lampblack (LCA101, <0.021 lm); (f) same as (e) with spectra normalized at 0.56 lm; (g) clinochlore (CLI101) and lampblack (LCA101, <0.021 lm); (h)same as (g) with spectra normalized at 0.56 lm; (i) serpentine (SRP117, <45 lm) and graphite (GRP102, <45 lm); (j) same as (i) with spectra normalized at 0.63 lm; (k)serpentine (SRP117, <45 lm) and magnetite (MAG102, <45 lm); (l) same as (k) with spectra normalized at 0.63 lm. Weight percent of opaque associated with each spectrumis indicated on the figures. Spectra measured at PSF.


Fig. 5 (continued)

Fig. 6. Reflectance ratio (2.2/0.8 lm) versus reflectance at 0.6–0.8 lm peak or slopebreak for the various CI chondrite spectra (filled squares), the fields occupied bynanophase magnetite (Morris et al., 1985), and carbon black (Johnson and Fanale,1973; Clark, 1983; Milliken and Mustard, 2007; this study). The letters indicate thepoints occupied by 95–5% mixtures of montmorillonite and lampblack from: A –Milliken and Mustard (2007); B – Clark (1983); C – Johnson and Fanale (1973); D –this study (serpentine SRP106); E – this study (serpentine SRP117); and F – thisstudy (saponite SAP101). Points that fall outside the field of this plot include 95%serpentine (SRP117) and 5% graphite (0.89 reflectance ratio and 34% reflectance;this study), 95% montmorillonite and 5% charcoal (1.01 reflectance ratio and 36%reflectance; Clark, 1983), and 90% serpentine (SRP117) and 10% magnetite(MAG102) (0.87 reflectance ratio and 37% reflectance; this study). The verticaldashed line separates blue-sloped from red-sloped spectra. The arrows indicatedirection of movement from the pure phyllosilicate end members (which fall off theplot) as follows: longest arrow for points A–C; medium arrow for point D, andshortest arrow for points E and F.


sloped (Fig. 5j) in spite of the red overall slope of the pure graphite.Overall reflectance is not dramatically reduced (34% maximumreflectance for 5 wt.% graphite), likely due to the fact that the

graphite and serpentine have comparable grain sizes. As fine-grained opaque minerals reduce overall reflectance more thancoarse-grained opaques, finer-grained graphite should have a morepronounced effect on lowering overall reflectance than found inthese mixtures.

When SRP117 is mixed with <45 lm magnetite (Figs. 5k and l),the spectra become darker and distinctly blue-sloped, with thereflectance ratio decreasing from 1.00 for the pure serpentine to0.87 for the mixture containing 10 wt.% magnetite. This decreaseoccurs in spite of the overall red slope of the magnetite beyond�1 lm. These results probably represent a ‘‘worst case’’ scenarioin that while magnetite abundance (10%) is comparable to thatin the CIs, the magnetite used in these mixtures is coarser-grainedand likely redder-sloped than the magnetite in CIs. The overallreflectance of the 10 wt.% magnetite mixture is still appreciablyhigher than CI chondrites (37% maximum reflectance), again likelydue to the limited dispersal of magnetite in the mixture.

While some mixtures of phyllosilicates + lampblack, and allmixtures involving magnetite and graphite can create increasinglyblue-sloped spectra, none of the mixtures involving opaque abun-dances comparable to CI chondrites are as dark as the CI spectra(Fig. 6). There are a number of explanations for the mismatch be-tween the mineral mixtures and CI spectra. It is possible, and per-haps likely, that lampblack is not a suitable spectral analogue for CIcarbonaceous material. It is also probable, but perhaps less impor-tant, that these intimate mixtures do not accurately reproduce theoptical properties of intimately dispersed carbonaceous material.Finally, carbonaceous matter is not the only abundant opaquepresent in CI chondrites – magnetite is a volumetrically importantopaque that is also finely dispersed (e.g., Hyman and Rowe, 1983;Brearley, 1992a; Bland et al., 2004). Thus, to better reproduce CIspectra it is probably necessary to include multiple opaques, en-


sure that the opaque materials are spectrally comparable to CI opa-ques, and that they are fine-grained and finely dispersed. Interest-ingly, the CI data points define a trend that points towardnanophase magnetite rather than carbon black, suggesting thatnanophase magnetite is the dominant opaque for both reducingreflectance and imparting a blue spectral slope.

9.2. Formation and alteration of carbonaceous material

As the carbonaceous material in CIs likely has a measurable ef-fect on their spectral properties, it is useful to consider the com-plexities inherent in determining the genesis and subsequentalteration of this material. This information also provides a basisfor comparison with other CC groups. Both organic matter and or-ganic precursor molecules will be subjected to pre-accretionaryprocesses that can fragment, alter, and polymerize them (e.g., An-ders et al., 1974; Robert and Epstein, 1982; Greenberg, 1989;Moore et al., 1991; Kaiser and Roseeler, 1997; Vuong and Foing,2000; Powles et al., 2005). In parent bodies, thermally-inducedpolymerization and aromatization of organic molecules throughsurface catalysis could occur (e.g., Hayatsu et al., 1977). CI parentbody processes, which are aqueous and/or hydrothermal, resultin preferential removal of an aliphatic component (Alexanderet al., 2005, 2007), and both synthesis and cracking likely occurred(Sephton et al., 2000). These processes are reflected in the fact thatdifferences with other CCs that exist in terms of fraction of aro-matic carbon (Cody and Alexander, 2005).

The CI insoluble carbonaceous matrix consists of a poorly or-dered, extensively substituted and cross linked framework of smallaromatic molecules (Gardinier et al., 2000; Sephton et al., 2000;Derenne et al., 2002; Cody and Alexander, 2005; Rouzaud et al.,2005; Garvie and Buseck, 2006; Derenne and Robert, 2009). Someportion of it may be unaltered presolar material (e.g., De Gregorioet al., 2010). It is intimately associated with the phyllosilicates(Pearson et al., 2002; Garvie, 2007; Garvie and Buseck, 2007). UVand ion irradiation of presumed organic matter on interstellar dustgrains causes changes in optical properties that mimic those due toheat treatment of amorphous carbon films (Jenniskens, 1993), fur-ther complicating the derivation of unique relationships betweenstructure, composition, and formation/modification processes.

9.3. Specific wavelength regions of absorption features

An absorption feature that appears in the 0.45–0.50-lm regionof many of the CI spectra (all but one spectrum of Ivuna and two ofOrgueil) is also seen in the insoluble organic matter from Murchi-son (and tentatively assigned to bonding–antibonding transitionsin aromatic molecules). It also appears in spectra of magnetiteand a number of the phyllosilicates, where it is attributed tospin-forbidden crystal field transitions in tetrahedrally coordinatedFe3+ (Burns, 1993). Consequently, an absorption feature in this re-gion is not considered to be particularly diagnostic of a particularcomponent, but does seem to be a common feature of CI spectra.

One feature that is nearly ubiquitous in the available CI spectraoccurs at or near 0.6 lm, where the CI spectra exhibit a distinctchange in spectral slope. In some of the spectra this wavelengthcorresponds to the start of a region of generally constant reflec-tance, while in others it is the region of maximum reflectance.Absolute reflectance in this region ranges between 2% and 7% forthe available CI spectra. The cause of this feature is not well con-strained, but may represent the onset of longer wavelength Feabsorption bands.

A few of the CI spectra, particularly Orgueil, exhibit an absorp-tion feature in the 0.65–0.8-lm region which is best attributed toFe3+–Fe2+ charge transfers in the phyllosilicates. However, it doesnot exhibit a consistent shape and is generally absent from the fin-

est fractions, suggesting that it will likely not be detected in spec-tra of CI parent bodies with fine-grained regoliths. It is likely thatcomminution of samples for spectral measurements results inthe destruction of any opaque-free phyllosilicate enclaves thatmay have existed, effectively dispersing opaques throughout thesample and suppressing spectral contributions to this feature fromany opaque-free phyllosilicate lithologies. If this feature is well re-solved, it may be possible to determine whether it is due to sapo-nites (band closer to 0.65–0.68 lm) or serpentines (band closer to0.70–0.75 lm).

The 0.9–1.2-lm region of a number of the CI spectra exhibits acomplex absorption feature. When examined in detail, this featureoften appears to consist of two or three absorption bands in the0.9–0.95-, 0.98–1.03-, and 1.10–1.15-lm regions. The 0.90–0.95and 1.10–1.15 lm bands are attributable to Fe2+ crystal field tran-sitions in Fe2+-bearing phyllosilicates, and hence the appearance ofthese two features is strong evidence for the presence of suchphases. The band in the 0.98–1.03-lm region is most consistentwith magnetite, although it should be noted that this band in sub-micron magnetites appears as a broad region of absorption with noupturn in reflectance beyond 1 lm (Morris et al., 1985), but is seenas a distinct absorption feature in coarser-grained (<45 lm) mag-netite spectra. A composite absorption feature with shoulders near0.9 and 1.15 lm is also seen in cronstedtite spectra. Consequently,the 0.9–1.2-lm region can be diagnostic of the presence of Fe2+-bearing phyllosilicates and the possible presence of magnetite.However, it does not appear that the type of phyllosilicate (serpen-tine vs saponite) can be distinguished using this wavelength regionalone.

The 1.4-lm region is where absorption bands attributable toovertones of O–H stretching from both H2O and OH are expected.In serpentines, which contain only OH, this band appears near1.38 lm. In saponites, which contain both structural OH and inter-layer H2O, there are sharp OH bands in the 1.39–1.42 lm interval,with H2O likely contributing to absorption starting near 1.42 lmand extending to beyond 1.5 lm as a long wavelength wing (e.g.,ferrihydrite, hydrated sulfates). Thus, this wavelength region hassome potentially diagnostic value: an absorption feature short-ward of �1.42 lm would be evidence for phyllosilicates, while afeature longward of �1.42 lm would be characteristic of H2O(Clark et al., 1990). A possible OH feature near 1.38 lm is seen ina couple of the Orgueil spectra (cfms03, mgp080: Fig. 4a and c).

The 1.7-lm region is where absorption bands due to C–Hstretching combinations and overtones are expected, but onlyone CI spectrum (Ivuna, c1mp18) has a suggestion of an absorptionfeature in this region. The lack of such a feature is consistent withsuppression of this band by the low overall reflectance, and thehighly condensed nature of the organic matter in CIs. A well-re-solved absorption band in this region is also not seen in the spec-trum of the insoluble organic matter from the Murchison CM2chondrite (Fig. 1i).

The 1.9-lm region is where a combination band due to H–O–Hbending plus O–H stretching is expected, and would be attribut-able to the presence of H2O. It is centered near 1.90–1.93 lm inthe various phases present in CIs. It can also appear when adsorbedwater is present, which could happen to CIs exposed to the terres-trial environment, and hence such a feature, which appears inmany of the CI spectra, is not a reliable indicator of specific CImineralogy.

A few of the CI spectra seem to exhibit weak absorption bandsnear 2.32 lm (Orgueil – cdms03), 2.38 lm (Orgueil – ccms03), orboth regions (Ivuna – c1mp18). These bands are uniformly weak(<2% deep). Similar features are seen in the spectra of some ofthe serpentines and saponites (Fig. 1), and are attributed to me-tal-OH combinations (Clark et al., 1990; Calvin and King, 1997).An absorption feature in this region is also seen in carbonates


(Fig. 1). The current data suggest that this wavelength region mayhave some potential for detection of phyllosilicates and/or carbon-ates on dark asteroids. It is unlikely that the low carbonate abun-dances in CIs will either translate into readily detectableabsorption bands or allow for discrimination of carbonate compo-sitional differences that exist between CIs and other CCs (e.g., deLeuw et al., 2010; Petitat and Gounelle, 2010).

Overall, the CI spectra provide some interesting insights into thesearch for possible parent bodies among asteroids. In spite of theirdark overall reflectance, a number do show resolvable absorptionbands attributable to the presence of phyllosilicates and magnetite,and detailed analysis of relevant absorption features, particularlythe 0.6–0.8- and 0.9–1.2-lm regions, suggests that the nature ofthe phyllosilicates (serpentines vs smectites) can be discriminated.However, some of the CI spectra do not show any phyllosilicateabsorption bands or show only a subset of the expected suite ofabsorption bands. This is likely attributable to a combination ofinadequate signal-to-noise and low overall reflectance.

9.4. Heterogeneity between CIs and alteration

The three CIs which have been most extensively studied and in-cluded in this study (Alais, Ivuna, and Orgueil) exhibit importantdifferences in mineralogy and texture, which likely relate to differ-ences in the style and degree of pre-terrestrial aqueous alteration(e.g., Bullock et al., 2005). These differences are likely responsiblefor some of the spectral differences between these three CIs. Firstand foremost, the three CIs included in this study are brecciatedto varying degrees (Orgueil being the most brecciated (Endressand Bischoff, 1993)), containing a variety of distinct types of clasts(e.g., McSween and Richardson, 1977; Richardson, 1978; Tomeokaand Buseck, 1988; Endress and Bischoff, 1994; Morlok et al., 2006).The matrices of different CIs also exhibit significant mineralogicaldifferences that reflect the degree, and likely oxidation state, ofaqueous alteration. Largely on the basis of textural evidence, ithas been suggested that the degree of aqueous alteration in CIs in-creases in the sequence Ivuna < Alais < Orgueil (Endress and Bisc-hoff, 1993). Alternatively, Alais may be more ‘‘evolved’’ thanOrgueil (Mackinnon and Kaser, 1988).

Increasing aqueous alteration appears to result in decomposi-tion of ferrihydrite-free crystalline coarse-grained phyllosilicates(serpentine- and saponite-type) to clusters of fine-grained inter-growths of less well-ordered phyllosilicates (serpentine- and sapo-nite-type) and disordered and fine-grained S-bearing ferrihydrite,dispersal of magnetite framboids, and alteration of magnetite andpyrrhotite to ferrihydrite. Coarse-grained phyllosilicates are domi-nated by serpentine-groups minerals, while the finer-grainedaggregates contain both serpentine- and smectite-group minerals(Mackinnon and Kaser, 1988). Orgueil appears to have undergonethe greatest degree of aqueous alteration based on these criteria(Tomeoka and Buseck, 1988; Endress and Bischoff, 1993; Tomeokaet al., 1989), resulting in the formation of a more Mg-rich saponite(Brearley and Prinz, 1992). Alais has a wider range of phyllosili-cates grain sizes than Orgueil, likely reflecting the less pervasivenature of aqueous alteration to which it was subjected (Mackinnonand Kaser, 1988).

Comparing Ivuna to Orgueil, Ivuna contains little fine-grainedmagnetite, and the dominant fine-grained phyllosilicates aremostly serpentine with minor saponite and, unlike Orgueil, containno ferrihydrite (Brearley, 1992a). The less common coarse-grainedphyllosilicates are dominated by saponite with rare serpentine(Brearley, 1992a,b). Textural evidence suggests that the fine-grained phyllosilicates in Orgueil formed from alteration of coarsesilicates (Tomeoka and Buseck, 1988), while in Ivuna, coarse-grained phyllosilicates may have grown from alteration of fine-grained phyllosilicates (Brearley, 1992a). Unraveling the results

of aqueous alteration is hampered by the fact that it appears thatfluid compositions were both spatially and temporally variable(e.g., Richardson, 1978; Fredericksson and Kerridge, 1988;McSween et al., 1993).

Translating these mineralogical characteristics and differencesinto expected spectral behavior is difficult. The larger average mag-netite grain size in Ivuna, for instance (up to 10 lm vs. up to 1 lmin Orgueil) might be expected to lead to a steeper spectral slope forIvuna, but its spectrum is not appreciably redder than Orgueil. It islikely that other factors, such as the lack of ferrihydrite (a red-sloped material) in the matrix of Ivuna could be overriding anyslope changes induced by differences in magnetite grain size.

The two spectra of <125 lm fractions of Ivuna are instructive interms of emphasizing the heterogeneity of CI chondrites (Fig. 3a).Their overall slopes are different, and one of the spectra showsclear evidence of an absorption feature in the 1-lm region. Grainsize differences cannot plausibly account for these differences. Itis more likely that the two specimens sample different portionsof the meteorite. Similar, but less dramatic differences are seenin two fine-grained (<40 and <45 lm) spectra of Orgueil (Fig. 4b).Significant spectral slope differences were also found in our dupli-cate spectra of both Alais (Fig. 2c) and Orgueil (Fig. 4g). The 2.4/0.56 lm reflectance ratio for the two Alais spectra is 1.9 and 2.4,while for Orgueil, the ratios are 0.84 and 0.96. These results rein-force the notion that CI chondrites are compositionally and spec-trally heterogeneous at the mm–cm (1–2 g) scale (Morlok et al.,2006). Compositional and textural differences between differentsubsamples also likely affect other carbonaceous chondrite spectra.

9.5. Spectral metrics

The various spectral metrics previously described are useful forcomparing CIs, and mineral end member and mixture spectra,although they only imperfectly capture spectral diversity. As notedabove, CI spectra range from blue- to red-sloped, with variableabsorption bands, particularly in the 1-lm region. The spectralmetrics described above were compared to search for useful spec-tral–compositional-petrographic trends.

It was found that increasing reflectance at the 0.6–0.8-lm re-gion peak or slope break, as well as reflectance at 1.5 lm and max-imum reflectance, all correlate with increasing red slope for Ivunaand Orgueil. Alais shows no dependence between these parameters(Fig. 7a and b). When these reflectance maximum measures arecompared to both minimum and average grain size, the largergrain size samples (those with a minimum grain size >0 lm) aredarker than most of the finer-grained samples (Fig. 7c and d), whiledecreasing average grain size shows a weak positive correlationwith increasing reflectance (Fig. 7e and f). When spectral slope isused as a metric, all of the samples with a minimum grain size>0 lm are blue-sloped (Fig. 7g), while the blue-sloped spectra in-clude only those samples with an average grain size >70 lm(Fig. 7h). The positive correlation between decreasing reflectanceand bluer slope may be attributable to increasing abundance ofthe opaques which can both darken a spectrum and impart anincreasingly blue overall slope. This is discussed in more detailbelow.

Even with this limited sample size, the following trends seem toemerge. Increasing overall reflectance (brightness) is roughly cor-related with increasing red slope. This is consistent with the ex-pected behavior of the opaque phases. As discussed below, someof the opaques in CIs (magnetite, carbonaceous material) are ex-pected to be blue-sloped, and decreasing opaque abundance wouldresult in both increasing reflectance and less blue slopes, due to agreater contribution from the red-sloped phyllosilicates and ironhydroxides. Increasing reflectance and increasing red slope areroughly correlated with decreasing grain size. These same trends

Fig. 7. Spectral metrics of CIs, nanophase magnetites and carbon blacks. (a) Spectral slope (as measured by the reflectance ratio at 2.2 lm to the local maximum or slopebreak in the 0.6–0.8-lm region) versus absolute reflectance of local maximum or slope break in the 0.6–0.8-lm region; (b) same as (a) for spectral slope versus absolutemaximum reflectance; (c) minimum grain size versus absolute reflectance of local maximum or slope break in the 0.6–0.8-lm region; (d) minimum grain size versus absolutemaximum reflectance; (e) average grain size versus reflectance of 0.6–0.8-lm region local maximum or slope break; (f) average grain size versus maximum reflectance; (g)minimum grain size versus spectral slope; (h) average grain size versus spectral slope. Squares: CIs sieved to only a maximum grain size and no minimum grain size; circles:CIs sieved to a minimum grain size >0 lm. Data that define the regions occupied by nanophase magnetite are from Morris et al. (1985) and this study; for carbon black, thedata are from Johnson and Fanale (1973), Clark (1983), Milliken and Mustard (2007) and this study.



were found by Johnson and Fanale (1973) for a variety of CCs, andthey attributed this, at least for C2 (CM2) and C3 (CV3, CO3) chon-drites, to comminution causing the breakup of chondrules which,being brighter than the matrix, contributing proportionately moreto the finer-grain size spectra. While CIs lack chondrules, albedoincreases with decreasing grain size may also occurs in CIs dueto the breakup of bright sulfate and carbonate clasts and veinsand coarse-grained opaque-poor phyllosilicate enclaves.

9.6. Blue spectral slopes

As discussed above, a number of the CI spectra exhibit blueslopes beyond �0.7 lm. There are a number of possible explana-tions for this behavior. One factor which likely plays a role is thegrain size of the magnetite, which is a major opaque phase. Mostof the magnetite in CIs is of submicron size, although there are dif-ferences in magnetite grain size and spatial distribution betweenvarious CIs (Hyman and Rowe, 1983; Brearley, 1992a). Submicronmagnetite is blue-sloped beyond �0.7 lm (Morris et al., 1985) incontrast to coarser magnetite (Fig. 1d and e). Comparison of thetwo serpentine spectra in Fig. 1b shows that the presence of mag-netite, likely finely dispersed in one of the serpentines, imparts anoverall blue slope to the serpentine spectrum. The same behavior isseen in the serpentine + magnetite mixture spectra. The fact thatmagnetite can impart a blue overall spectral slope likely relatesto the fact that its absorption coefficient and refractive index in-crease dramatically above �1 lm (Schlegel et al., 1979). The differ-ence in magnetite grain size between Ivuna (generally <10 lm)and Orgueil (generally <1 lm), which have similar magnetite con-tents (10–11%; Hyman and Rowe, 1983), does not appear to trans-late into a significant difference in overall spectral slope,suggesting that average grain size variations below a few lm isnot a significant factor in spectral slope.

Calvin and King (1997) found that many Fe-bearing phyllosili-cates are red-sloped over the 0.6–2.5-lm region, with overallreflectance generally being highest for the most Fe-poor samples.One exception was found – thuringite – a very Fe-rich sample withan Fe/(Fe+Mg) ratio of �0.95. Its Fe abundance is much higher thanCI phyllosilicates, which generally fall in the range 0.15–0.6 (Brear-ley, 1992a; Zolensky et al., 1993; Tonui et al., 2003). Phyllosilicateswith Fe/(Fe+Mg) ratios in this range, and including samples withvariable Fe2+/Fe3+ ratios (Table 4; Calvin and King, 1997) exhibitred-sloped spectra across the 0.35–2.5 lm interval, due largely toFe–O charge transfers, and variable strength Fe absorption bands.Thus it seems unlikely that the phyllosilicates on their own can ac-count for the blue slopes of CI spectra. Our phase angle measure-

Fig. 8. Reflectance spectra (i = 30�, e = 0�) of different size fractions of the acid-insoluble fraction of the Murchison CM2 chondrite; measured at RELAB.

ments, at least for Alais, suggest that overall spectral slope canvary somewhat as a function of phase angle, but does not changefrom strongly red- to strongly blue-sloped.

Reflectance spectra of the acid-insoluble matter in CCs are notuniform. Previous investigations of acid- insoluble matter in CCshave found that it is either strongly blue-sloped (Allende: Johnsonand Fanale, 1973) or red-sloped (Murchison: Gaffey, unpublished).Fig. 8 shows reflectance spectra of four size fractions of the acid-insoluble fraction from the Murchison CM2 chondrite. This mate-rial is composed largely of carbonaceous material but also includes�8% acid-insoluble oxides, such as chromite. The finest fraction(<125 lm) has the highest overall reflectance, and all the spectraexhibit a local reflectance peak in the 0.36–0.40-lm region. Thisis close to the value for reflectance peaks in some aromatic mole-cules. It also exhibits a reflectance minimum in the 0.42–0.46 lminterval, similar to the reflectance minimum seen in graphite. Atlonger wavelengths there is a reflectance maximum in the 0.8–1.1-lm region. Overall, reflectance declines slightly beyond this re-gion and the decline is most pronounced in the 90–106 lm sizefraction. The overall blue slope and spectral features suggest thatthe carbonaceous phases, specifically a diversity of condensed aro-matic molecules, graphite-like phases, and a more amorphous car-bon-rich component (probably required to produce low overallreflectance), are the dominant contributor to this spectrum.

Blue spectral slopes can also be produced from nominally neu-tral or red-sloped materials. Singer (1980) found that some mix-tures of limonite (very fine powder) and magnetite (<45 lm),both of which were red sloped or flat across the 1.2–2.5 lm range,could produce overall blue-sloped spectra when intimately mixed.He attributed this behavior to the change in optical density withincreasing wavelength, where the fine-grained limonite coats thelarger magnetite grains. With increasing wavelength, the limonitebecomes increasingly transparent. Ferrihydrite, which is present insome CIs (Orgueil) but apparently not others (Ivuna) (Brearley,1992a), and which is spectrally similar to limonite, could also ac-count for the blue slope of some of the CI spectra via this mecha-nism. Fischer and Pieters (1993) found that the thickness of aferric oxide-bearing cinder coating on a basalt slab affects the de-gree of bluing of a spectral slope. However, this bluing mechanismonly operates at wavelengths >�1 lm, where the absorption coef-ficient decreases.

Development of a blue slope is seen in intimate mixtures ofmontmorillonite and carbon lampblack (Clark, 1983; Millikenand Mustard, 2007). Reflectance ratios taken at 0.8 and 2.15 lm(2.15/0.8 lm) are 0.98 for the pure montmorillonite in bothstudies, and decrease (become bluer) with increasing lampblack

Fig. 9. Normalized (at 0.56 lm) reflectance spectra of olivine (Fa10, <45 lm) poorlycrystalline graphite (GRP102, <45 lm), and a 95:5 by weight mixture of the two.


content (to 0.82 and 0.84 for mixtures containing 5% lampblack).The reflectance ratio for the pure lampblack is 1.0 (Milliken andMustard, 2007) or 0.93 (Clark, 1983), well above the values forthe 5% mixtures. However, for mixtures of carbon black and anLL6 chondrite, Miyamoto et al. (1982) found that overall spectralslope became redder with increasing lampblack content. These re-sults indicate that the reflectance spectra of pure end memberscannot be used to confidently predict the spectral properties ofintimate mixtures and that in the case of carbonaceous materials,small changes in composition or structure could have large effectson spectral reflectance. Some ways in which the addition of opaquematerials can affect absorption features in reflectance spectra arediscussed in Moroz and Arnold (1999).

An example of the need to exercise caution in predicting thespectral properties of mixtures from end member spectra is shownin Fig. 9, which shows reflectance spectra of fine-grained olivine(<45 lm), fine-grained ‘‘amorphous’’ graphite (<45 lm) and a mix-ture of 95 wt.% olivine and 5 wt.% graphite. The pure graphitereflectance spectrum is red-sloped longward of �0.6 lm, the oliv-ine spectrum is flat beyond �1.8 lm and its 0.56 lm peak is lowerthan the reflectance in the 1.8–2.5 lm interval. However the mix-ture spectrum exhibits a higher reflectance for the 0.56 lm peakrelative to the 1.8–2.5 lm interval and a blue-sloped spectrumacross the 1.8–2.5 lm interval.

The reason for the bluing of the mixture spectrum is not fullyunderstood. The real part of the dielectric constant of graphite in-creases toward longer wavelengths beyond �0.4 lm, while theimaginary part exhibits more complex behavior – different investi-gators have found that is can be relatively constant, increase con-stantly, or shows both increases and decreases across the 0.4–2.5 lm interval (Taft and Phillipp, 1965; Foster and Howarth,1968; Draine, 1985); conductivity shows a decrease between�0.6 and 1 lm, a slight increase to a peak near 1.5 lm, and a fur-ther decrease longward of this value (Taft and Phillipp, 1965). Thevalue of 1.5 lm also coincides with a maximum in absorption(McCartney and Ergun, 1967), and is attributed to an overtone ofelectronic valence band transitions (McClure, 1957; Boyle and Noz-ières, 1958). The extinction coefficient for thin sections of graphitealso decreases over the 0.25–0.65 lm interval (McCartney and Er-gun, 1967). Changes in both the real and imaginary parts of thedielectric function depend on grain orientation (Draine, 1985),and some of the differences in results between different investiga-tors are due to the large spectral changes that can accompanysmall variations in graphite structure (Taft and Phillipp, 1965).Coals, which are also dominated by condensed and cross-linkedaromatic structures, broadly similar to the insoluble organic mat-

Fig. 10. Band depth at 1 lm for CI spectra versus: (a) spectral slope, and (b) maximum resize; circles: CIs sieved to a minimum grain size >0 lm.

ter in CCs, have optical properties similar to graphite (McCartneyand Ergun, 1967; Foster and Howarth, 1968).

Amorphous carbon shows decreasing extinction efficiency withincreasing wavelength from 0.3 to 0.8 lm (Colangeli et al., 1986)and 0.3 to 2.5 lm (Maggipanto et al., 1985; Bussoletti et al.,1987), while clusters of amorphous carbon grains show increasingrefractive index and oscillating absorption coefficient with increas-ing wavelength across the 0.3–2.5 lm interval (Rouleau and Mar-tin, 1991).

It appears likely that the relative changes in, and importance ofabsorption versus scattering as a function of wavelength in graph-ite and related materials, as well as how these materials are dis-persed in a matrix, will all control the spectral reflectanceproperties of a carbonaceous-bearing assemblage.

Regardless of the details of the optical properties at play, thereappear to be several mechanisms that could account for blueslopes in CI spectra. These include the presence of very Fe-richphyllosilicates, magnetite, some forms of carbonaceous matter,and possibly the presence of dispersed ferrihydrite. The spectralreflectance properties of individual components from CIs are notnecessarily a good predictor of the spectral reflectance propertiesof a more complex assemblage. The fact that an increasingly blueslope correlates with a decrease in reflectance does not particularlyprovide insights into possible mechanisms, as an increasingly blueslope coupled to a decrease in reflectance is seen in LL6 chon-drite + magnetite mixtures (Miyamoto et al., 1982), montmorillon-ite + carbon black mixtures (Clark, 1983; Milliken and Mustard,2007), and serpentine + graphite and serpentine + magnetite mix-tures (this study). We tend to favor the cause of simultaneous blu-ing and reflectance decrease, in large part, to the magnetite in CIs.Fine-grained (<5 lm) magnetite has overall reflectance of a fewpercent (Hunt et al., 1971; Morris et al., 1985), and most of theblue-sloped spectra also exhibit the expected magnetite absorp-tion band near 1 lm (in addition to the phyllosilicates bands oneither side of this feature). There are suggestions of rough correla-tions of band depth at 1 lm with both increasingly blue slope(Fig. 10a) and decreasing maximum reflectance (Fig. 10b), as wouldbe expected if magnetite is contributing to both darkening and blu-ing. The fact that the correlation is not perfect and that spectrawhich lack a 1-lm region absorption band exhibit a range of max-imum reflectance values and blue slopes indicates that other CIproperties besides magnetite abundance contribute to changingoverall reflectance and blue slope. Some of the scatter is also dueto uncertainties in measuring accurate band depths for some ofthe analogue spectra from Johnson and Fanale (1973) and Gaffey(1974).

flectance. Squares: CIs sieved to only a maximum grain size and no minimum grain

Fig. 11. Reflectance spectra of 55–63 lm fraction of the acid-insoluble matter fromthe Murchison CM2 chondrite as a function of viewing geometry; spectra measuredat RELAB.


10. Implications for remote sensing of CI parent bodies

10.1. Phase angle effects

While the RELAB and PSF reflectance spectra that form the bulkof this analysis were acquired at i = 30� and e = 0�, asteroidal obser-vations are often acquired at different phase angles. Coupled tothis, disk-unresolved asteroid spectra will include contributionsfrom the subsolar point to the limb, thus spanning a wide rangeof incidence, emission, and phase angles. This difference in observ-ing conditions between the laboratory and telescopes may trans-late into differences in spectral slope and the intensity ofabsorption bands, which must be accounted for, or at least recog-nized, when utilizing laboratory spectra for analysis of observa-tional data. This effect has repeatedly been recognized as onethat should be accounted for in quantitative spectral analysis(e.g., Adams and Filice, 1967; Egan et al., 1973; Veverka et al.,1978; Hapke et al., 2007; Pommerol and Schmitt, 2008).

Phase angle effects on reflectance spectra have been examinedby a number of investigators, with CC (and CI) meteorites includedin some studies (e.g., French and Veverka, 1983; Capaccioni et al.,1990; Kamei and Nakamura, 2002). Particulate samples of the Bru-derheim L6 chondrite, measured across the 0.3–1.1 lm intervalshowed increasingly red slopes with increasing phase angle, anda reversal at phase angles >�70� (Egan et al., 1973); there also ap-peared to be some wavelength dependency for coarser grain sizes.Adams and Filice (1967) showed that red/blue ratios as a functionof phase angle differ for different materials, and are also affected byparticle packing. Tomita and Nakamura (2002) and Kamei andNakamura (2002) measured phase functions for a number of ordin-ary chondrites and the Allende CV3 chondrite and found differ-ences in reflectance with phase function both between differentmeteorites and for incidence angles of 0� vs. 75�.

A number of photometric studies have included CCs, and in par-ticular the Allende CV3 chondrite. The major results arising fromthese studies include the fact that dark asteroids likely show limbdarkening near opposition, suggesting that measured geometricalbedos are lower limits to the normal reflectance (French and Ve-verka, 1983). Reflectance spectra measured in bidirectional modeare redder than those measured with an integrating sphere, withthe degree of difference increasing with increasing absolute reflec-tance (Gradie and Veverka, 1986). The integrating sphere provides,in general, a better match to a planetary body at small phase angles;as phase angle increases, bidirectional reflectance curves provide abetter match. In the case of the Allende CV3 measurements (0.4–1.2 lm), the difference between bidirectional (4� phase angle) andintegrating sphere measurements is < 2% relative, and increases to�10% relative at 60� phase angle (Gradie et al., 1980; Gradie andVeverka, 1986). Results for Allende also suggest that the degree ofreddening is mostly a function of wavelength and emission angle,with reddening increasing toward longer wavelengths and largerphase angles (Gradie et al., 1980; Gradie and Veverka, 1986). For aconstant phase angle and variable emission angle, the degree of red-dening is less than that due to variations in phase angle, but couldstill lead to reddening toward the limb and terminator (Gradieet al., 1980). The opposition surge (increase in brightness at lowphase angles (<12�) relative to a linear extrapolation of larger phaseangle results) is greatest for the CCs, at least in the visible wave-length region (Capaccioni et al., 1990).

Capaccioni et al. (1990) found that reflectance of Orgueil de-creases with increasing phase angle (for i = 0�), similar to our results.We also find differences in degree of overall reflectance changes as afunction of incidence, emission, and phase angle variations. Our re-sults also suggest that the degree of spectral variability is wave-length dependent, with the greatest variability appearing atshorter wavelengths (<�1 lm).

Phase angle measurements have been conducted on the acid-insoluble material from the Murchison CM2 chondrite, which islargely composed of carbonaceous materials. Three grain sizes ofthis material were measured at phase angles of 10�, 30� and 50�(e = 0� in all cases). The phase angle effects are similar for the dif-ferent size fractions: a reflectance peak near 0.40 lm is most pro-nounced in the 10� phase angle measurements (Fig. 11). There is noconsistent systematic change in overall spectral slope when thevalues for the reflectance ratio are fixed to the peak near 0.8 lmand at 2.2 lm. When the true maximum is used, the 10� phase an-gle measurements show the most blue-sloped spectra followed bythe 50� and then the 30� phase angle spectra. The reason for theappearance of a pronounced peak near 0.4 lm in the smallestphase angle spectra and its relative absence in the larger phase an-gle spectra is unknown. However, dark materials such as carbonand CCs show greater sensitivity, at least in terms of overall reflec-tance in the visible region, to variations in phase angle (Capaccioniet al., 1990), and phase angle variations are also known to be wave-length dependent (Gradie et al., 1980; Gradie and Veverka, 1986);both of these observations have been verified in our phase anglestudy of Alais.

10.2. Packing and porosity and roughness

Salisbury et al. (1975) noted that a loosely packed sample ofMurchison (CM2) was darker than a more densely packed sample(5% vs. 4% reflectance at 1.5 lm). Sakai and Nakamura (2004)found that ‘‘fluffy’’ samples of the Allende CV3 chondrite were dar-ker and redder than compacted and ‘‘knocked’’ samples over the0.4–0.8 lm interval. These observations may have implicationsfor detection of CI parent bodies. Asteroids, being low-gravity ob-jects, may have an underdense surface layer, which could resultin a lower overall reflectance than equivalent laboratory spectra.Lower overall reflectance (at least in the visible region) has alsobeen noted for loose versus compacted fine-grained samples ofdiabase and peridotite (Capaccioni et al., 1990) and the differencebetween the two increases with increasing phase angle. There isless of a difference between loose and compacted samples for coar-ser grain sizes (Capaccioni et al., 1990).

Roughening of a surface, at the scale of individual grains (butstill larger than the wavelength of light), results in a reduction inlimb darkening for a variety of materials (Veverka et al., 1978;French and Veverka, 1983). The change in normal reflectance withphase angle also appears to be wavelength dependent (Gradie andVeverka, 1986). It appears that regolith scattering parameters may


be relatively insensitive to particle size when the particles aresmaller than the wavelength of light (Piatek et al., 2004).

Reflectance spectra of normally packed and densely packedAlais (Fig. 2k and l) show that increasing sample density does re-duce overall reflectance, consistent with the observations of Salis-bury et al. (1975), but overall slope is largely unchanged. Fluffing asample (by disrupting the sample surface with a needle) causes anoverall darkening relative to a powdered sample with a flat surfaceand either a reddening of the slope (Alais; Fig. 2k and l), or nochange in overall slope (Orgueil; Fig. 4i and j). It is also worth not-ing that some of these effects are wavelength dependent. Forexample, the Alais fluffed sample is darker than the densely packedsample only below 0.5 lm.

10.3. Space weathering

Space weathering, defined here as changes in the spectroscopicproperties of asteroid surfaces due to exposure to the space envi-ronment, is expected to have an effect on the spectral propertiesof CIs, although the nature of this effect is not fully known. Giventhat the opaque minerals dominate the overall spectral signatureof CIs, space weathering effects in these materials may dominateany spectral changes that CIs may undergo.

Moroz et al. (2004a) found that ion irradiation of strongly red-sloped organic materials resulted in a reduction in slope, althoughno blue-sloped spectra were produced. Microsecond pulse laserirradiation of Mighei (CM2), designed to simulate micrometeoritebombardment, caused an increase in overall reflectance and a red-der slope (Moroz et al., 2004b), while the microsecond pulsed laserirradiation of a CC simulant resulted in a decrease in overall reflec-tance but no change in overall slope (Hiroi et al., 2003). On theother hand, nanosecond pulse laser irradiations (Yamada et al.,1999) on the Tagish Lake meteorite produced a bluer (flat) contin-uum (Hiroi et al., 2004). Ion implantation into organic materialscan result in loss of spectral structure due to disruption of bondsand formation of a more amorphous material (Powles et al.,2005). Ion irradiation experiments conducted on an altered perido-tite, which contains serpentine, show that irradiation leads to bothdarkening and reddening of the spectrum over the 0.7–2.2 lminterval (Brunetto and Strazzulla, 2005).

Exposure of the phyllosilicates present in CIs to the vacuum ofspace is also expected to result in spectral changes. Serpentine andnontronite that were exposed to low atmospheric pressures andUV irradiation showed a decrease in the depth of the H2O-associatedabsorption feature in the 1.9-lm region. Features associated withOH, including the O–H stretching overtone and metal-OH combina-tion bands were relatively unaffected. Fe-related absorption bandsin the 0.4–1.2-lm region in most minerals are usually also unaf-fected by such exposure, but nontronite did exhibit changes in itsFe absorption bands (Cloutis et al., 2008).

It appears that the spectrum-altering effects of space weather-ing and exposure to the space environment, which encompass awide range of processes, some of which cannot be well simulatedin the laboratory, are not yet well defined. The degree and typeof spectral alteration that may arise will be a complex functionof the physical and chemical makeup of the starting materials,and the relative dominance of various space weathering processesthat are operating.

10.4. Impact effects

In addition to space weathering processes that operate at themicro scale, asteroid surfaces will also be affected by larger-scaleimpact processes. These can lead to excavation of fresh material,burial of space-weathered surficial materials, heating and/or melt-ing, release/mobilization of volatiles, formation of shock features

and possible high pressure phases, comminution, and redistribu-tion of materials. Multiple processes operate during impact eventsand their relative importance changes depending on the nature ofthe impact event, thus it is difficult to enumerate and quantify therelative importance of processes that will consistently operate inall impact events. Below we review some studies relevant to howimpact processes may affect CI spectra.

Heating experiments have been performed on a number of CCs(e.g., Miyamoto et al., 1982; Hiroi et al., 1996a,b). In the case ofIvuna, heating resulted in a number of spectral changes in the0.3–2.5-lm region, including shifting of the position of slopechange/reflectance maximum in the visible region toward shorterwavelengths, and changes in overall spectral slope, with slopesbecoming bluer or redder (Hiroi et al., 1996b). Simulation of CCmelting due to impacts, using Murchison, showed that meltingcauses an increase in overall reflectance, formation of a glass-asso-ciated Fe2+ absorption band near 1.1 lm, and a slight increase inoverall slope (Clark et al., 1993). Vacuum melting, vaporization,and condensation experiments on Murchison (CM2) and Allende(CV3) produced condensates that showed extremely diverse spec-tra, ranging from slightly red-sloped with or without resolvableabsorption bands (similar to the starting material) to extremelyred-sloped and spectrally featureless (King et al., 1983). If theseresults are representative of impact melt products on asteroid sur-faces, and if impact melt production is a significant process on CIparent bodies, there may be little likelihood of being able to relatethe spectra of asteroid with significant impact melts on their sur-faces to meteorite spectra.

Comparison of light and dark lithologies in ordinary chondrites,where the dark lithologies result from impact shock, shows thatimpact shock lowers reflectance but has little to no effect on over-all spectral slope (Britt and Pieters, 1992, 1993); other studies sug-gest that shock results in somewhat redder overall spectral slope(Keil et al., 1992). The significance of these results to CIs is not cer-tain, as production and dispersal of fine-grained opaques, which ispresumed to account for much of the darkening, may not be signif-icant in CI chondrites, which already possess fine-grained and dis-persed opaques.

Impact experiments using serpentine, subjected to shock pres-sures up to 41 GPa, suggest that impact shock has little effect onOH-related absorption bands but may result in a redder spectrum,at least above 1 lm (Rivkin et al., 2003). Shock-loading of a red-sloped C-rich organic, asphaltite, and a 50:50 asphaltite:kamacitemixture showed that increasing shock pressure resulted in the for-mation of a more C- and aromatic-rich and ordered material,approaching graphite. These transformations were accompaniedby a decrease in reflectance (at least above 2 lm) and loss of C–H absorption bands (Korochantsev et al., 1997). These results reit-erate that our understanding of how impact processes will affect CIchondrites is incomplete, but that a range of spectrum-altering ef-fects may occur and that these processes will be a function of tar-get composition and impact characteristics.

10.5. Grain size effects

The reflectance spectra of the powdered CI samples show differ-ences in overall slope (e.g., Fig. 3a) in spite of the fact that they allconsist of fine-grained powders. There is a weak correlation be-tween minimum and average grain size and spectral slope, withthe larger grain size samples generally exhibiting bluer spectralslopes (Fig. 7). Further evidence that grain size affects spectralslope comes from a study of the Almahata Sitta ureilite. Hiroiet al. (2010) found that reflectance spectra of chips of this ureiliteshowed uniformly blue slopes, while spectra of <125 lm powderswere red-sloped. This suggests that compositionally identical CIsmay show a wide range of spectral slopes that are due solely to


variations in grain size. Similarly, rotational spectral variations thatmay be seen in dark asteroid spectra could plausibly be related tograin size variations across a surface rather than compositionalvariations. Opposition effect variations for dark asteroids also sug-gests that the magnitude and shape of the opposition surge are afunction of spectral slope (color index) and albedo, which may betied to grain size variations on asteroid surfaces (Shevchenkoet al., 2010).

11. Identification of CI parent bodies

A number of spectroscopic studies have forged tentative linksbetween various asteroids and CCs. Here we review some of theseresults in light of the CI laboratory spectra. In spite of the fact thatCI chondrites contain abundant Fe2+-Fe3+-bearing phyllosilicates,their spectra are largely devoid of an expected Fe3+–Fe2+ chargetransfer absorption band in the 0.7-lm region. This is likely attrib-utable to suppression by the various opaque phases. Consequently,the presence of such an absorption feature in asteroid spectra (e.g.,Barucci et al., 1998) is not very compatible with the available CIlaboratory spectra, suggesting that CCs other than CIs may be bet-ter matches. The near ubiquity of a slope change near 0.6 lmseems to be the most consistent feature in the CI spectra, andhence represents a potentially usable criterion for identifying pos-sible CI parent bodies. As a cautionary note, its cause is not fullyknown and could be a manifestation of terrestrial weathering(Salisbury and Hunt, 1974).

While tentative links between other CCs and various taxonomicclasses of asteroids have been made (e.g., Burbine et al., 1992; Bur-bine, 1998; Carvano et al., 2003), attempts to match CIs to darkasteroids often fail, usually on the basis of presence/absence ofabsorption bands, wavelength position of the UV–visible reflec-tance downturn, and overall spectral slopes (e.g., Gaffey andMcCord, 1977; Hiroi et al., 1993; Kelley et al., 2007). While possiblelinks between some C-/B-class asteroids and CIs have been ad-vanced, it appears that modifications to CI mineralogy must be in-voked to match CIs to members of other asteroid taxonomic classes(Larson et al., 1983; Gaffey et al., 1993).

An absorption complex in the 0.95–1.2-lm region, attributableto Fe2+-bearing phyllosilicates ± magnetite, present in some of theCI spectra, suggests that it may be fruitful to search for similar fea-tures in asteroid spectra. However, a similar absorption complex(bands near 0.95 and 1.15 lm) can also be attributed to Ca-richspectral type A pyroxenes (Cloutis and Gaffey, 1991; Schadeet al., 2004). If these bands are accompanied by additional absorp-

Fig. 12. B-class asteroid versus CI chondrite spectra. (a) Same as Fig. 6 with the poinAbbreviations: Hi: 426 Hippo, Li: 213 Lilaea, Pa: 2 Pallas, Ph: 3200 Phaethon, Ra: 2100 Rclass Asteroid 1508 Kemi and the 100–200 lm size sample of Orgueil (spectrum ccms0

tion bands attributable to the presence of OH, or the 0.7-lm regionFe3+–Fe2+ charge transfer band, this would be strong evidence forassigning the bands to phyllosilicates rather than high-Capyroxene.

Examination of blue-sloped B-class asteroid spectra (Clark et al.,2010) reveals a number that share many common features withsome CI spectra (Fig. 12a). Asteroid 2 Pallas, 213 Lillaea, and3200 Phaethon all have blue-sloped spectra characteristic of someCI chondrites but have higher reflectance (pv = 0.09–0.16) than anyof our CI spectra. The spectral slope of 3200 Phaethon is also bluerthan any of our measured CI spectra. Asteroid 229 Adelinda has avisible albedo (pv) of 0.05, a local reflectance maximum near0.55 lm, a red-sloped spectrum with a 2.2/0.6-lm region reflec-tance ratio of 1.50, a possible absorption band in the 0.8-lm re-gion, and no evidence of a 1-lm region absorption band. All ofthese features are present in at least one of the CI spectra: the vis-ible albedo is well within the CI range, the position of the localmaximum is close to the value for CIs (the position of the peakin the asteroid spectrum cannot be entirely constrained becauseof coarse spectral resolution), the reflectance ratio is close to thatof Alais (mgp084), and a 0.8-lm region absorption band with aweak or non-existent 1-lm region band is present in one Orgueilspectrum (mgp078). However, no one CI spectrum has all of thesefeatures.

Asteroid 379 Huenna (Clark et al., 2010) is red-sloped (2.2/0.6-lm region reflectance ratio of 1.21) with a reflectance plateau inthe 0.6–0.7-lm region, a visible albedo of 0.06, and an absorptionband in the 1-lm region. Again, while no one CI spectrum matchesall of these features, they all match or appear in at least one CIspectrum.

Asteroid 426 Hippo (Clark et al., 2010) has a visible albedo of0.05, a reflectance peak near 0.6 lm, a blue slope (2.2/0.6-lm re-gion peak ratio of 0.85), a possible absorption band near 0.9 lm,and a flattening out of the spectrum beyond �1.5 lm. The blueslope is present in a number of CI spectra, and its value falls withinthe range for CIs (as low as 0.79), while a prominent absorptionband near 0.9 lm is present in one Orgueil spectrum (mgp080),as is a flattening of the spectral slope near 1.5 lm (ccms03). Onceagain, no one CI spectrum exhibits all of these features, althoughblue-sloped CI spectra usually exhibit the 0.9–1.2-lm regionabsorption feature.

B-class asteroid 1508 Kemi (Clark et al., 2010) does not have adetermined visible albedo, is blue-sloped with a peak near0.67 lm, has a probable 1-lm region absorption band, and showsa flattening out of the blue slope beyond �1.3 lm. It resembles the100–200 lm grain size spectrum of Orgueil (ccms03; Fig. 12b) in

ts for a number of blue-sloped B-class asteroids added from Clark et al. (2010).a Shalom (2 spectra available); (b) normalized (at 0.56 lm) reflectance spectra of B-3).


these respects; both also exhibit the shoulder near 0.5 lm. Thisasteroid provides the best overall spectral match to a CI chondriteof the B-asteroid spectra that were examined.

12. Summary and conclusions

Individual CI chondrites possess differences in mineralogy andpetrology that are significant enough to result in wide variationsin spectral properties, particularly in terms of spectral slopes. Smalldifferences in composition can cause significant changes in spec-tral slope, as well as the appearance of phyllosilicate-associatedabsorption bands. Significant spectral differences are seen evenin duplicate spectra of the same sample after repacking.

Most of the CI spectra exhibit weak but resolvable absorptionbands in the 1-lm region that are best attributed to the abundantphyllosilicates that they contain, as well as magnetite. Howeverboth absolute reflectance and absorption band depths are lowerthan for pure phyllosilicates due to the presence of fine-grained, fi-nely-dispersed opaques. The presence of other CI phases, such ascarbonaceous phases, sulfates, and carbonates, is rarely derivablefrom the CI reflectance spectra, likely due to the fact that thesephases are either spectrally quite featureless (i.e., carbonaceousmatter), or their abundances (e.g., carbonates, sulfates) are lowand absorption bands are weak compared to the other phases.

No one mechanism can be uniquely ascribed to explain the blueslope of CI spectra. Analysis of the spectral properties of the variousCI constituents as well as mineral mixtures suggests that magne-tite, the insoluble organic matter, and perhaps ferrihydrite, couldall impart a blue slope. Of these, magnetite has been directlyshown to be able to impart a blue slope to phyllosilicate spectra,while the bluing effect of other phases (insoluble organic matterfrom Murchison, lampblack, graphite, ferrihydrite, very Fe-richphyllosilicates (thuringite)) and their appropriateness as proxiesfor opaques in CIs, are less well constrained. The overall trend ofdecreasing spectral slope with decreasing absolute reflectancetends toward the region occupied by nanophase magnetite, butthat an additional darkening agent (likely the carbonaceous mat-ter) may be required to lower overall reflectance below that attain-able with nanophase magnetite. The fact that many of the CIspectra exhibit an absorption band in the 1-lm region that canbe attributed to magnetite further strengthens the assignment ofthe blue slope in large part to magnetite. The rough similaritiesin magnetite and Fe3+-bearing phase abundances between Alaisand Orgueil (Bland et al., 2008), and their different spectral slopes,suggests that abundances of opaque phases is not sufficient to ex-plain differences in CI spectral slope, and other factors, such asgrain size and dissemination of opaques may be important contrib-uting factors to overall spectral slope.

Phase angle studies of Alais indicate that for incidence andemission angles below �60�, increasing phase angle results in low-ering of overall reflectance and reddening of spectra slope, partic-ularly below �1 lm. At larger incidence and emission angles,increasing phase angle results in redder slopes, but the change inabsolute reflectance is not as systematic.

For CIs and ureilites (Hiroi et al., 2010) it appears that increasinggrain size results in lower overall reflectance and generally moreblue-sloped spectra. However, increasing nanophase magnetiteabundance also produces the same trends, and hence it appearsunlikely that these two effects can be discriminated by reflectancespectroscopy alone. Polarimetric measurements of asteroids maybe required to resolve this ambiguity (Dollfus, 1971; Dollfuset al., 1975).

The effects of shock, heating, and space weathering on CI chon-drites were not specifically examined as part of this study, butexisting data suggest that these processes will likely have an effect

on the spectral properties of CI parent bodies, if only through com-minution of regolith. Most studies done to date suggest that theseprocesses will redden carbonaceous asteroid reflectance spectra,although at least one study showed bluing associated with heating(Hiroi et al., 1993, 1996b).

The results of this study suggest that at least one, and likelymore, plausible explanations exist for the cause of blue slopes inCI spectra (i.e., magnetite). What is less well-known is why CI spec-tra exhibit the diversity of spectral shapes that they do. Futurestudies of CI chondrites and intimate mineral mixtures that betterreproduce the textures of CIs will shed new light on these intrigu-ing meteorites and hopefully forge robust links to their asteroidalparent bodies.

Acknowledgments

We wish to thank the invaluable and generous assistance pro-vided by many individuals which made this study possible. Weparticularly thank Dr. Jeffrey Post of the Smithsonian InstitutionNational Museum of Natural History and Dr. Linda Reinen of Pomo-na College for providing a number of the mineral samples used inthis study; Mr. Neil Ball and Dr. Frank Hawthorne of the Universityof Manitoba for acquisition of XRD data for the mineral samples,Dr. Stanley Mertzman for XRF analysis of the mineral samples,and Dr. Richard Binzel of the Massachusetts Institute of Technologyfor providing some of the B-class asteroid spectra. The establish-ment and operation of the Planetary Spectrophotometer Facility(PSF) at the University of Winnipeg was made possible throughthe assistance of the Canada Foundation for Innovation, the Mani-toba Research Innovations Fund, the Canadian Space Agency, theNatural Sciences and Engineering Research Council of Canada(NSERC), and the University of Winnipeg. The RELAB facility atBrown University is a multi-user facility operated with supportfrom NASA Planetary Geology and Geophysics grant NNG06GJ31G,whose support is gratefully acknowledged. Finally we thank thetwo anonymous reviewers for their many helpful and cogent com-ments that improved the quality of this paper.

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