Asteroids

2.14

Asteroids

TH Burbine, Mount Holyoke College, South Hadley, MA, USA ã 2014 Elsevier Ltd. All rights reserved.

2.14.1 2.14.2 2.14.3 2.14.3.1 2.14.3.2 2.14.3.3 2.14.3.4 2.14.3.5 2.14.3.6 2.14.3.7 2.14.3.8 2.14.3.9 2.14.4 2.14.4.1 2.14.4.2 2.14.4.2.1 2.14.4.2.2 2.14.4.3 2.14.4.4 2.14.4.5 2.14.4.6 2.14.4.7 2.14.4.8 2.14.4.9 2.14.4.9.1 2.14.4.9.2 2.14.4.9.3 2.14.4.9.4 2.14.4.9.5 2.14.4.10 2.14.4.11 2.14.4.12 2.14.4.12.1 2.14.4.12.2 2.14.4.12.3 2.14.5 2.14.5.1 2.14.5.2 2.14.5.3 2.14.5.4 2.14.5.5 2.14.5.6 2.14.5.7 2.14.5.8 2.14.5.9 2.14.6 2.14.6.1 2.14.6.2 2.14.6.3 2.14.6.3.1 2.14.6.3.2 2.14.6.3.3 2.14.6.3.4

Introduction Background Remote Observations Reflectance Spectroscopy Mineral Spectroscopy Electronic Transitions Vibrational Bands Visual Albedo Metallic Iron Organics Space Weathering Determining Mineralogies Taxonomy A-Types C-Complex B-types C-types D-Types K-Types L-Types O-Types Q-Types R-Types S-Complex S-types Sa-types Sq-types Sr-types Sv-types T-Types V-Types X-Complex (E-Types, M-Types, and P-Types) E-types M-types P-types Spacecraft Missions Galileo Deep Space 1 NEAR Shoemaker Stardust Hayabusa Rosetta Dawn OSIRIS-REx 2008 TC3 and Almahata Sitta Interesting Groups of Asteroids Earth and Martian Trojans Near-Earth Asteroids Asteroid Families Inner belt (
2.06–2.50 AU) Middle belt (
2.50–2.82 AU) Outer belt (
2.82–3.28 AU) Trojan region (
5.2 AU)

Treatise on Geochemistry 2nd Edition

http://dx.doi.org/10.1016/B978-0-08-095975-7.00129-7

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Asteroids

2.14.7 Taxonomic Distribution of Taxonomic Types 2.14.7.1 Earlier Work 2.14.7.2 Recent Work 2.14.7.3 Formation of Material in the Solar Nebula 2.14.7.4 Origin of the Taxonomic Distribution 2.14.8 Conclusions and Future Work Acknowledgments References

2.14.1

Introduction

Asteroids are the remaining planetesimals that helped form the terrestrial and Jovian planets. Asteroids, which are the parent bodies of almost all meteorites, are almost entirely observed remotely using Earth- and space-based telescopes with only a few bodies studied up close by spacecraft missions. Only one asteroid, (21543) Itokawa, has been the subject of a sample return, and only one other body, 2008 TC3, has been spectrally observed and had fragments recovered due to the explosion of the object in the Earth’s atmosphere over the Sudan. The chemical and isotopic compositions of meteorites can be determined with high precision in laboratories on Earth (e.g., MacPherson and Thiemens, 2011); however, remote observations of asteroids give geochemical information that meteorites cannot. Spacecraft studies of asteroids allow us to understand how the chemistries and mineralogies of asteroids vary on scales of meters to kilometers on the surface. Determining the surface mineralogies of asteroids allows us to gain insight on how the composition of the solar nebula varied versus the distance from the Sun. Comparisons between asteroid and meteorite spectra give insight on how the space environment changes the surface properties of asteroids. Asteroids also are important to characterize since an impacting nearEarth asteroid could potentially wipe out life as we know it. The asteroid impact that wiped out the dinosaurs (e.g., Schulte et al., 2010) is such an example. Any mitigation scenario would need to know the composition of the asteroid. Asteroids could also be potential sources of raw materials that are rare on Earth. Kargel (1994) estimates that the precious metals found in asteroids with LL chondrite mineralogies, which have the lowest metal abundances of the ordinary chondrites, would be worth hundreds of billions of dollars or more. Asteroids may have brought water (e.g., Morbidelli et al., 2000) and the ingredients of life (e.g., amino acids and sugars; e.g., Abramov and Mojzsis, 2011; Cooper et al., 2001; Glavin et al., 2011) to Earth. This chapter will discuss what is currently thought to be known concerning the chemistries of asteroids and what their chemistries tell us about the solar nebula from ground- and space-based telescopic observations and spacecraft missions.

2.14.2

Background

Asteroids are small solar system bodies that include main-belt objects, near-Earth asteroids (NEAs), or near-Earth objects (NEOs), Earth Trojans, Martian Trojans, and Jupiter Trojans. Approximately 600 000 asteroids are currently known. Small solar system bodies also include comets and bodies located in

399 399 400 400 401 402 403 404

the outer solar system such as Centaurs (semimajor axes between Jupiter and Neptune), Neptune Trojans, and transNeptunian objects. The dwarf planet (1) Ceres is located in the main belt and was known as an asteroid from its discovery (1801) until 2006. The term ‘asteroid’ usually refers to rocky or metallic iron bodies that are located in the inner solar system out to the orbit of Jupiter. Comets, Centaurs, Neptune Trojans, and trans-Neptunian objects appear more volatile-rich than asteroids and are thought to have formed in the outer solar system. All of these bodies grew initially through collisions of small grains (e.g., Weidenschilling, 2000) that stuck together and formed centimeter-sized bodies, which then collide to form larger bodies. When a small solar system body is first discovered, the object is given a provisional designation. Excluding comets, the provisional designation for a small solar system body is the year of the discovery plus a letter that indicates the half month of the discovery and another letter (e.g., 1977 RG) plus often a subscript number (e.g., 1987 DA6) that indicates the order within the half month of its discovery. When the orbit of a small solar system body (excluding comets) is confirmed, the object is given a number. The name of the small solar system body (excluding comets) will be either its provisional designation or a name approved by the Committee on Small Body Nomenclature of the International Astronomical Union. Comets have their own provisional designations and naming conventions. Main-belt objects have orbits that lie between the orbits of Mars and Jupiter. Six orbital elements at a particular epoch are needed to characterize an asteroid orbit (Bowell et al., 2002); however, the basic properties of an orbit can be given by three orbital elements (semimajor axis, inclination, and eccentricity). Proper elements are orbital elements that are nearly constant over time. NEAs have orbits that lie near the Earth’s orbit. NEAs are typically broken up into the Aten (semimajor axes smaller than Earth’s), Apollo (Earth-crossing NEAs with semimajor axes larger than Earth’s), and Amor (Earth-approaching NEAs with orbits exterior to Earth’s but interior to Mars’) groups depending on their particular orbits. Objects that cross Mars’ orbit are often called Mars-crossers. Mars and Jupiter Trojans lie at the L4 (60 ahead of the planet’s orbit) and L5 (60 behind) Lagrangian points around their respective bodies. The L4 and L5 Lagrangian points are two of the five positions where a small body can theoretically be stable in a configuration with two larger bodies (in this case the Sun and a planet). The only known Earth Trojan (Connors et al., 2011) occurs at the L4 Lagrangian point. Venus Trojans appear theoretically possible (e.g., Tabachnik and Evans, 2000). Trojan asteroids may be remnants (e.g., Dotto et al., 2008; Reilly, 2011) of the planetesimals that formed near where the planet

Asteroids

that they are in a stable orbit with is located; however, it is also possible that they have been captured into their current locations and formed far from their current locations. Another type of body is quasi-satellites, which are objects in a 1:1 orbital resonance with a planet or asteroid with the orbits being relatively short-lived (e.g., Christou and Wiegert, 2012; Kortenkamp and Joseph, 2011). Most objects in the main asteroid belt (Figure 1) have a semimajor axis (average distance from the Sun to the object) between
2.1 and
3.3 AU. When a histogram of asteroid semimajor axes is plotted, a number of gaps (called Kirkwood gaps) in the distribution are evident. The primary gaps occur at 4:1 (
2.06 AU), 3:1 (
2.50 AU), 5:2 (
2.82 AU), 7:3 (
2.95 AU), and 2:1 (
3.27 AU) mean-motion resonances with Jupiter, which supplies a regular and periodic gravitational influence on asteroids with those semimajor axes. The dynamical lifetimes of objects entering these resonances and becoming meteorites or NEAs are typically a few million years (Gladman et al., 1997). These resonances are used to divide the main belt into the inner belt (between the 4:1 and 3:1 resonances from
2.06 to 2.50 AU), the middle belt (between the 3:1 and 5:2 resonances from
2.50 to 2.82 AU), and the outer belt (between the 5:2 and 2:1 resonances from
2.82 to 3.27 AU). The inner part of the belt is also bounded by the n6 secular resonance with Saturn (Froeschle´ and Scholl, 1986; Yoshikawa, 1987), where the perihelion of the asteroid precesses at the same rate as the perihelion precesses for Saturn. These resonances supply asteroids into Earth-crossing orbits by changing their eccentricities (Wisdom, 1982, 1983, 1985). A few main-belt resonances cause asteroids to have stable orbits. One example is the Hungaria group (e.g., McEarchen et al., 2010; Warner et al., 2009), where
10 000 members are in a 9:2 orbital resonance with Jupiter and a 3:2 orbital resonance with Mars and are high-inclination objects located between
1.8 and
2.0 AU. Another example is the Cybele

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group (
3.3–3.7 AU; Vokrouhlicky´ et al., 2010), where
1500 members cluster around the 7:4 resonance with Jupiter. A third example is the Hilda group (e.g., Brozˇ and Vokrouhlicky´, 2008), whose
3000 members are in a 2:3 orbital resonance with Jupiter and are located at
4.0 AU. The Hungaria group is named after (434) Hungaria, the Hilda group is named after (153) Hilda, and the Cybele group is named after (65) Cybele. Currently there are
600 000 known main-belt asteroids,
9000 known NEAs, one known Earth Trojan, three known Martian Trojans, and
6000 known Jupiter Trojans. One quasi-satellite (2002 VE68) around Venus has been identified (Mikkola et al., 2004), while Earth has five known quasisatellites. The largest object, (1) Ceres, in the main belt has a diameter of
950 km. For diameters of 1 km or larger,
1.9  106 asteroids in the main belt (Tedesco et al., 2005),
1000–1200 asteroids (Mainzer et al., 2011b; Stuart, 2001) in the near-Earth population, and
1.3  106 Jupiter Trojans (Marzari et al., 2002) are estimated to exist. Except for the largest objects (e.g., (1) Ceres, (2) Pallas, (4) Vesta, (10) Hygiea) that are thought to have stayed relatively intact since their formation, asteroids are almost all the collisional fragments of disrupted planetesimals. There is a huge disparity between the millions of asteroids estimated to exist compared to the number of postulated parent bodies (
100–150; Burbine et al., 2002a; Wasson, 1995) for the
40 000 known meteorites. One possible explanation is that the delivery mechanisms of meteorites to Earth are extremely biased toward a relatively few parent bodies. However, this is in contrast to our understanding of the Yarkovsky effect, which is believed to be able to allow fragments from nearly any object in the main belt to drift to meteorite-supplying resonances and, therefore, add to the meteorite flux to Earth (Bottke et al., 2000). The Yarkovsky effect (e.g., Bottke et al., 2002b, 2006b; Hartmann et al., 1999) is the force acting on a rotating body in space caused by the anisotropic emission of thermal photons, which changes

0.8 0.7

Sine of proper inclination

0.6

5:2 3:1

Hungaria Region

7:3

4:1

0.5

2:1 Cybele Region

Hilda Region

0.4 0.3 ν6

0.2 0.1 0 1.5

2

2.5 3 Semimajor axis (AU)

3.5

4

Figure 1 Plot of semimajor axis (AU) versus sine of proper inclination for
300 000 asteroids. The positions of mean-motion resonances (4:1, 3:1, 5:2, 7:3, 2:1), the n6 secular resonance, and asteroid regions (Hungaria, Cybele, Hilda) are indicated with arrows.

368

Asteroids

the semimajor axes of small bodies. Maybe the relatively short time frame that most meteorites can survive on Earth (
50 000 years in warm deserts and
250 000 years for Antarctica; Jull, 2006) allows only a relatively small number of parent bodies to be actually sampled. Also, maybe fragments of most asteroids are not strong enough to make it through Earth’s atmosphere (Sears, 1998). Another possibility is that multiple asteroids could have formed with very similar mineralogies and isotopic compositions and, therefore, fragments from these bodies cannot be easily distinguished from each other. Asteroid families are clusterings of asteroids in proper orbital element (semimajor axis, inclination, and eccentricity) space (e.g., Cellino and Dell’Oro, 2010; Zappala` et al., 1990, 1995) and are named after the lowest numbered family member. Asteroid families are thought to be primarily due to the breakup of a much larger object. Exceptions include the Eunomia and Vesta families where apparently large impacts excavated a significant amount of material from the bodies but did not disrupt them. Estimated formation ages for asteroid families range from a few hundred thousand years (Nesvorny´ and Vokrouhlicky´, 2006) to
2 Ga (Marzari et al., 1995). Using Sloan Digital Sky Survey (SDSS) observation of
90 000 asteroids, Parker et al. (2008) estimated that
50% of objects in this data set belong to families, while Nesvorny´ (2010) places
36% of
300 000 asteroids in families. These percentages are much lower than the previous estimate (Ivezic´ et al., 2002) of 90% of observed asteroids in families, which was made using SDSS observations of only
10 000 asteroids. A few asteroids have also been cross-listed as comets, which is usually due to the later identification of a coma around the asteroid. These objects are called main-belt comets (e.g., Bertini, 2011; Jewitt et al., 2010). One exception is near-Earth asteroid 4015 Wilson–Harrington. Comet Wilson–Harrington was first identified as a comet in 1949 due to having a coma that then was lost. In 1992, the orbit of asteroid (4015) 1979 VA was found to be the same as that of comet Wilson–Harrington and the asteroid was named Wilson–Harrington. Except for the initial discovery images in 1949, no subsequent cometary activity (e.g., Ishiguro et al., 2011a) has been detected from Wilson–Harrington. Impacts have been proposed as the cause of these comas around main-belt objects. For example, a number of researchers (Ishiguro et al., 2011b; Jewitt et al., 2011; Moreno et al., 2011) have proposed that the coma around main-belt object (596) Scheila is due to an impact.

2.14.3 2.14.3.1

Remote Observations Reflectance Spectroscopy

The primary way of determining the chemistry of an asteroid is by determining the object’s reflectance spectrum in the visible and near-infrared wavelength regions with a telescope and a detector. The asteroid’s reflectance spectrum plots the relative fraction of light reflected as a function of wavelength. The photons measured by a detector are those emitted by the Sun that either have been reflected from the asteroid’s surface, are the result of the absorption and then emission by electrons associated with different transition elements in different minerals as the electrons change energy levels, or emitted thermal radiation, which is usually not significant for main-belt

asteroids in the visible and near-infrared (Harris and Lagerros, 2002). The Sun acts like a blackbody with the intensity of the emitted light peaking in the visible and decreasing in intensity for longer wavelengths (near-infrared to infrared). Therefore, the light from the asteroid also peaks in intensity in the visible and then decreases in intensity for longer wavelengths. The Earth’s atmosphere is relatively transparent in the visible and near-infrared so these are the wavelength regions that are primarily used for Earth-based observations of asteroids. Since the actual source of photons (the Sun) cannot be directly measured during the observations, the flux measured for the asteroid must be divided by the flux measured by a standard star (or the average of a series of standard stars) usually similar to the Sun (G2 spectral type) observed at a similar air mass and time as the asteroid. The ratio produces a reflectance spectrum since it removes the flux distribution of the Sun. Further discussion of asteroid reduction techniques can be found in Gaffey et al. (2002). Reflectance observations of asteroids are done using either spectrophotometry or spectroscopy. Spectrophotometric observations measure magnitudes (fluxes) using broadband filters, where the difference in magnitude between two filters produces the color index of the asteroid. The observations with each filter are made at different times. Until the advent of the CCD (charge-coupled device), spectrophotometry was the primary way of obtaining reflectance measurements of asteroids. In spectroscopy, reflectances from an asteroid are measured all at the same time. A spectrograph disperses radiation from an object into its component wavelengths (called a spectrum) onto a detector. The detectors are CCDs in the visible and infrared arrays for near-infrared wavelengths. A number of spectrophotometric surveys have made significant contributions to asteroid studies. The 24-color survey (Chapman and Gaffey, 1979) observed almost 300 asteroids using 25 filters from 0.32 to 1.08 mm. The eight-color asteroid survey (ECAS) (Zellner et al., 1985) observed almost 600 asteroids with eight filters from 0.34 to 1.04 mm. The 52-color survey (Bell et al., 1988, 1995) measured the reflectances of over 100 asteroids using 52 filters from 0.8 to 2.5 mm. Instituted in 2000, SDSS used five filters from 0.354 to 0.913 mm to obtain colors for over 100 000 asteroids (Carvano et al., 2010). SDSS was originally designed to map the sky and measure redshifts of galaxies. Since the late 1980s (e.g., Jewitt and Luu, 1990a, 1990b; Sawyer, 1991; Vilas and Gaffey, 1989), spectroscopy has been the primary way of obtaining asteroid reflectance spectra. Notable visible wavelength spectroscopic surveys include SMASS (Small Main-Belt Asteroid Survey) (
300 asteroids) (Xu et al., 1995), SMASS II (
1300 asteroids) (Bus and Binzel, 2002a), and S3OS2 (
800 asteroids) (Lazzaro et al., 2004). The primary telescope currently used for obtaining near-infrared spectra (e.g., Clark et al., 2010; DeMeo et al., 2009; Emery and Brown, 2003; Ostrowski et al., 2011; Popescu et al., 2011; Sunshine et al., 2004; Thomas and Binzel, 2010) is the infrared telescope facility (IRTF) on Mauna Kea, which houses the SpeX (spectral regions from
0.8 to 2.5,
1.9 to 4.2,
2.4 to 5.5 mm) spectrograph and imager (Rayner et al., 2003). SpeX has allowed near-infrared spectra of asteroids to be easily obtainable through in-person and remote (Binzel et al., 2004a) observations.

Asteroids

2.14.3.2

Mineral Spectroscopy

Many minerals (e.g., olivine, low-Ca pyroxene (e.g., orthopyroxene), high-Ca pyroxene, and serpentine) found in meteorites (Rubin, 1997) have characteristic absorption bands (e.g., Gaffey, 1976; Salisbury et al., 1975) in the visible (
0.4– 0.8 mm) and near-infrared (
0.8–5 mm), making this the prime wavelength region for mineralogical analysis. These absorption bands are also apparent in asteroid spectra and are due primarily to electronic transitions and vibrational bands where photons are preferentially absorbed by the minerals. Electronic transitions occur when an electron absorbs a photon and then moves to a higher energy orbital. These bands are due to crystal field transitions, charge transfer absorptions, and conduction bands. Vibrational absorption bands are due to the absorption of photons by molecules where the frequency of the absorbed radiation matches the vibrational frequency of the bond. The vibrational motion must cause a change in the dipole moment of the molecule for photons to be absorbed. Molecules can also have absorption bands associated with rotation and translation (movement of the molecule as a whole).

2.14.3.3

Electronic Transitions

The primary way of determining asteroid mineralogies is by analyzing crystal field transitions (e.g., Burns, 1993). Crystal field transitions are due to the absorption of photons by electrons in the partially filled inner (3d) orbitals of transition metal ions. The electrostatic interaction between the transition metal ion (cations) and the ions (called ligands, such as oxygen) bound to it causes the degenerate d-levels to split into different energy levels. The amount of splitting will be due to the arrangement of the ligands around the metal ion. Transition metals include Fe2þ, Fe3þ, Cr3þ, Mn2þ, and Mn3þ. Particular transition metals will occupy sites in minerals if their sizes and charges are appropriate. The most important transition metal for asteroid spectroscopy is Fe2þ due to its abundance in minerals such as pyroxene and olivine. Since different minerals have different chemical compositions and/or different

369

crystal structures, the energy level splittings will be different for different minerals, and, therefore, different minerals will have different characteristic crystal field absorptions. For example, olivine and pyroxene have very different absorption bands in the visible and near-infrared since they have different chemical compositions and crystal structures (e.g., Burbine et al., 2008). Olivine ((Mg,Fe2þ)2SiO4) has three bands at
0.9,
1.1, and
1.25 mm that form an asymmetric 1 mm feature (Figure 2). Forsterite (Fo) is the magnesium endmember of the olivine solid solution series, while fayalite (Fa) is the iron endmember. Brachinite meteorites have a prominent olivine feature (Figure 3). The bands move to longer wavelengths as Fe2þ substitutes for Mg. Pyroxene-group minerals have the general formula M2M1(Si2O6) where the smaller M1 site can be occupied by Al, Mg, Fe2þ, Fe3þ, Cr, and Ti while the M2 site may contain Ca, Na, Mg, Fe2þ, Mn2þ, Ni, and Li. Enstatite (En) is the magnesium endmember of the pyroxene solid solution series, ferrosilite (Fs) is the iron endmember, and wollastonite (Wo) (which is not a pyroxene) is the calcium endmember. Pyroxenes tend to have two symmetric bands at
0.9–1.0 and
1.9–2.0 mm (Figure 2). HEDs (howardites, eucrites, and diogenites) are meteorites that have prominent pyroxene bands (Figure 3; e.g., Beck et al., 2011a; Gaffey, 1976). The bands move to longer wavelengths as Fe2þ substitutes for Mg (e.g., Klima et al., 2007) or Ca substitutes for Fe2þ (e.g., Klima et al., 2011). Virtually FeO-free pyroxenes (enstatite) (Mg2Si2O6) do not have these absorption bands due to the absence of Fe2þ and, therefore, have featureless, flat spectra in the visible and near-infrared (Figure 2). Aubrites are composed primarily of enstatite and have featureless spectra (Figure 3). High-Ca pyroxenes (such as hedenbergite (CaFeSi2O6)) have Ca occupying the M2 site so Fe2þ in the M1 site causes absorption bands at
0.93–0.98 and
1.18 mm and no band at
1.9– 2.0 mm. Angrites have spectral features consistent with hedenbergite except that olivine in the angrite erases any distinction between the two bands (Figure 3). On asteroids, most minerals are expected to be in intimate mixtures on the surface. For minerals in the visible and nearinfrared, the strength of the absorption bands does not directly

4 High-Ca pyroxene (hedenbergite)

Normalized reflectance

3.5 3

Enstatite

2.5 Low-Ca pyroxene (hypersthene)

2 1.5 1

Olivine

0.5 0

0

0.5

1 1.5 Wavelength (mm)

2

2.5

Figure 2 Plot of reflectance spectra of olivine (Clark et al., 2007), low-Ca pyroxene (hypersthene) (Clark et al., 2007), enstatite from aubrite Pen˜a Blanca Spring (Burbine et al., 2002b), and high-Ca pyroxene (hedenbergite) (Clark et al., 2007). Spectra are normalized to unity at 0.55 mm and offset in reflectance.

370

Asteroids

4.5

Normalized reflectance

4

High-Ca pyroxene angrite D’Orbigny

3.5

FeO-free pyroxene aubrite Mayo Belwa

3 2.5 Low-Ca pyroxene eucrite Bouvante

2 1.5 1 0.5

Olivine-bearing brachinite EET99402

0 0

0.5

1 1.5 Wavelength (mm)

2

2.5

Figure 3 Plot of reflectance spectra of olivine-bearing brachinite EET 99402 (particle size <125 mm; Burbine et al., 2008), low-Ca pyroxene eucrite Bouvante (particle size <25 mm; Burbine et al., 2001b), FeO-free pyroxene aubrite Mayo Belwa (particle size <125 mm; Burbine et al., 2002b), and highCa pyroxene angrite D’Orbigny (particle size <125 mm; Burbine et al., 2006). Spectra are normalized to unity at 0.55 mm and offset in reflectance.

correspond to the abundances in the mixture. Photons are absorbed by a material (e.g., Clark, 1999) according to Beer’s law where I ¼ Ioekx where I is the observed intensity, Io is the original intensity, k is the absorption coefficient, and x is the distance traveled through the medium. The absorption coefficient is related to a material’s complex index of refraction, which measures the speed of light in that material. The absorption coefficient also varies versus wavelength. Pyroxenes tend to have larger absorption coefficients than olivine so pyroxenes tend to dominate the spectral properties of pyroxene–olivine mixtures since they are more absorbing. For mixtures of olivine and low-Ca pyroxene, Cloutis et al. (1986) found that the ratio of the areas of Band II to Band I (the band area ratio) is proportional to the relative abundances of olivine and low-Ca pyroxene. The widths and positions of crystal field absorption bands will also be a function of temperature (Burns, 1993). As the temperature decreases, the amplitude of the thermal vibrations of the cation about the center of its coordination site also decreases, which will reduce the width of the band. Since the coordination site is contracting and the distance between the cation and oxygen is decreasing, the energy level splitting will therefore increase, which results usually in a band that moves to shorter wavelengths. The band positions of olivines and low-Ca pyroxenes tend to decrease in width and move to shorter wavelengths as the temperature of the sample decreases (Hinrichs et al., 1999; Moroz et al., 2000; Reddy et al., 2012c; Roush and Singer, 1987; Sanchez et al., 2012; Schade and Wa¨sch, 1999; Singer and Roush, 1985). However, the Band II positions of Ca-rich pyroxenes do move to longer wavelengths (e.g., Schade and Wa¨sch, 1999; Singer and Roush, 1985) with decreasing temperature, while the Band I positions move to shorter wavelengths. A charge transfer absorption is where an electron absorbs a photon and then transfers from one ion to another ion. The two types typically found in asteroid spectra (e.g., Burns, 1993; Klima et al., 2007) are oxygen–metal charge transfers and intervalence charge transfers. The oxygen–metal charge transfers are high-intensity bands generally centered in the ultraviolet that extend into the visible. The sharp edge of this type of band is apparent shortward of
0.7 mm in many asteroid spectra and is attributed to an O–Fe2þ charge transfer transition. A feature centered at
0.7 mm in the spectra of many asteroids (e.g.,

Cloutis et al., 2011b; Vilas and Gaffey, 1989) has been attributed to an intervalence charge transfer (Fe2þ ! Fe3þ) transition in oxidized iron in phyllosilicates. Among meteorites, this feature at
0.7 mm tends to be found in the spectra of CM2 chondrites (Hiroi et al., 1993b, 1996) and is not readily apparent in the spectra of CR2 chondrites (Cloutis et al., 2012a) due to the lack of Fe3þ in their phyllosilicates. Conduction bands are present in some minerals such as sulfides. Electrons can reside in a higher energy level called the ‘conduction band’ where the electrons can move freely and a lower energy region called the ‘valence band’ where electrons are part of individual atoms. The energy that is able to excite an electron between the two bands is called the ‘band gap.’ Oldhamite (CaS), which is found in aubrites and enstatite chondrites (Watters and Prinz, 1979), has a strong conduction band that is centered at
0.5 mm (Burbine et al., 2002b) plus a feature at
0.9 that is due to a Fe2þ crystal field transition band (Figure 4).

2.14.3.4

Vibrational Bands

Minerals also have vibrational absorption bands (e.g., Salisbury, 1993), which primarily occur in the near- to mid-infrared. The positions of these vibrational bands are a function of mineralogy and temperature (e.g., Henning, 2010). For example, waterbearing and hydroxyl-bearing minerals can have very strong vibrational absorption bands (Figure 5) that are present in asteroid spectra (e.g., Rivkin et al., 2002). H2O is a V-shaped molecule with hydrogen atoms at the ends of each of the arms and an oxygen atom at the apex of the V. H2O can have three types of vibrations (Gaffey et al., 1989): a bend where the V opens and closes at a fundamental frequency corresponding to a wavelength of
6.1 mm, a symmetric stretch where the arms both lengthen and shorten together at
3.05 mm, and an asymmetric stretch where one arm lengthens or shortens at
2.9 mm. The bend and symmetric stretch fundamentals combine with the first overtone of the bend fundamental to produce a strong absorption feature in the 3 mm wavelength region. OH in minerals can only undergo an asymmetric stretch and causes a sharp absorption band at
2.7 mm (e.g., Rivkin et al., 2002). The structure of the 3 mm feature depends on what water-bearing and hydroxyl-bearing

Normalized reflectance

Asteroids

2.8 2.6 2.4 2.2 2 1.8 1.6 1.4 1.2 1 0.8 0.6 0.4

371

Oldhamite

~0.5 μm

0

0.5

1 1.5 Wavelength (μm)

2

2.5

Figure 4 Plot of reflectance spectrum of oldhamite (Burbine et al., 2002b). Feature at
0.5 mm is identified. Spectrum is normalized to unity at 0.55 mm.

3

Normalized reflectance

CR2 chondrite Renazzo 2.5 CM2 chondrite Murchison

2 1.5

CI1 chondrite Orgueil 1 Serpentine

0.5 0 2

2.5

3

3.5

4

4.5

Wavelength (μm) Figure 5 Plot of reflectance spectra in the 3 mm region for serpentine (Clark et al., 2007), CI1 chondrite Orgueil (particle size <100 mm) (Hiroi et al., 1993b), CM2 chondrite Murchison (particle size <63 mm) (Hiroi et al., 1993b), and CR2 chondrite Renazzo (particle size <100 mm) (Hiroi et al., 1993b). Spectra are normalized to unity at 2.5 mm and then offset in reflectance.

minerals are present as evident by the different 3 mm band shapes for different meteorite types (e.g., Sato et al., 1997). The aqueously altered CI, CM, and CR chondrites all have prominent 3 mm bands (Figure 5) with different shapes and intensities, while the much less aqueously altered CO, CV, and CK chondrites have very weak to no apparent 3 mm bands. Silicates (Henning, 2010) and organic material (Cloutis, 1989; Moroz et al., 1998) have a large number of vibrational bands in the near- and mid-infrared. Silicates have a large number of vibrational absorption bands such as the Si–O stretching, O–Si–O bending, metal–O stretching, and the asymmetric and symmetric stretching of the SiO4 tetrahedral group. Organic material has a large number of vibrational absorption bands due to stretches of bonds (e.g., C–H, C–O) and bends and stretches of groups (e.g., CH2, CH3).

2.14.3.5

Visual Albedo

Also important for understanding the mineralogy of the surface is the visual geometric albedo (e.g., Carvano, 2008; Hanner et al., 1981), which is the fraction of the incoming visible radiation that is reflected by a body. The visual albedo of an asteroid’s surface will depend on its mineralogy and

average particle size. For example, the FeO-free pyroxene enstatite has a very high albedo so aubrites, which are predominately enstatite, also have high albedos. FeO-containing pyroxenes (like those found in ordinary chondrites) tend to have lower albedos. Also carbonaceous chondrites (Johnson and Fanale, 1973), since they contain higher abundances of opaques, are much darker than ordinary chondrites. Eucrites, which contain much lower abundances of opaques, are brighter. Lowering the particle size of a mixture tends to increase the visual albedo of the surface (e.g., Johnson and Fanale, 1973). Visual albedos are usually determined using a thermal model (e.g., Delbo´ and Harris, 2002; Harris, 1998; Harris and Lagerros, 2002; Lebofsky and Spencer, 1989) that balances the energy absorbed by the asteroid in the visible with the thermal energy emitted by the object in the infrared. The energy absorbed in the visible is related to the asteroid’s absolute (H) magnitude (Bowell et al., 1989), which is the visual magnitude that would be recorded if the asteroid were placed at 1 AU away from the Earth, 1 AU from the Sun, and at a zero phase angle. For the infrared flux, observations from spacebased satellites such as the Infrared Astronomical Satellite (IRAS) (e.g., Ryan and Woodward, 2010; Tedesco et al., 2002b; Veeder et al., 1989b), Midcourse Space Experiment

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Asteroids

(e.g., Ryan and Woodward, 2010; Tedesco et al., 2002a), the Spitzer Space Telescope (e.g., Ryan and Woodward, 2011; Thomas et al., 2011b; Trilling et al., 2010), and the Widefield Infrared Survey Explorer (WISE) NEOWISE (NEO observations by WISE) project (e.g., Mainzer et al., 2011a,b,c) are typically used.

2.14.3.6

Metallic Iron

Common minerals found in meteorites that do not have absorption bands are kamacite and taenite, which are both composed of Fe and Ni. Kamacite has a lower concentration of Ni than taenite plus a different crystal structure. Cloutis et al. (2010a) found that compositional variations had no systematic effect on the spectral slopes of the spectra of metallic iron powders. However, radar analyses of asteroids can give an indication of metal content. The transmitted radar signal (Ostro, 1993) is usually circularly polarized and the reflected radar signal can have either the same circular polarization (called SC) or the opposite polarization (called OC). Normal reflection from a plane mirror yields echoes that are almost entirely in opposite polarization, while multiple scattering from a surface can cause echoes that are in the same polarization. The circular polarization ratio (SC/OC) measures the wavelength-scale roughness of the surface (Benner et al., 2008). The radar albedo of an asteroid is the ratio of a target’s OC radar cross section to its projected area (Ostro, 1993). Radar albedos are a complicated function of the near-surface bulk density (which is a function of the solid rock density and the porosity of the surface) and the mineralogy of the object. Increasing the metal content of the surface or decreasing the surface porosity will both increase the radar albedo of an object. Radar observations have been done for over 100 asteroids (Magri et al., 1999, 2007).

2.14.3.7

Organics

Organics in meteorites (e.g., Kerridge, 1999; Pizzarello et al., 2006) are usually found as a kerogen-like, insoluble material plus a less abundant soluble material. These organics are commonly found in carbonaceous chondrites. Objects with

extremely low visual albedos (
0.05) and an extremely red spectral slope in the visible and near-infrared (e.g., Dalle Ore et al., 2011) are assumed to be organic-rich. The Tagish Lake meteorite (C2-ungrouped chondrite) is an example of an extremely organic-rich carbonaceous chondrite (Herd et al., 2011; Pizzarello et al., 2001) and has an extremely red spectral slope (Figure 6; Cloutis et al., 2012c; Hiroi and Hasegawa, 2003; Hiroi et al., 2001b).

2.14.3.8

Space Weathering

The problem with deciphering the mineralogy of an asteroid is that its spectral properties are also a function of its surface’s interaction with the space environment. It has long been known that the spectral properties of the most common taxonomic type observed (now called S-complex objects) were redder than the most common meteorites to fall to Earth (ordinary chondrites; Figure 7; e.g., Chapman and Salisbury, 1973). Space weathering (e.g., Brunetto, 2009; Chapman and Salisbury, 1973; Hapke, 1973, 2001) is the term for the processes (e.g., Moretti et al., 2007) that can potentially alter the spectral properties of the surface of an ‘airless’ body. These processes include the interaction with galactic and solar cosmic rays, irradiation by solar wind particles, and micrometeorite and meteorite impacts. For silicates, these processes darken the surface, redden the spectral slope, and reduce the strength of absorption bands through the production of nanophase metallic iron (e.g., Loeffler et al., 2009; Sasaki et al., 2001, 2003). However, Gaffey (2010) argues that space weathering does not affect spectral parameters such as band center positions and band area ratios that have been used for determining mineralogy of asteroid surfaces. The effects of space weathering on organic material depend on the type of organics: space weathering appears to decrease the spectral slope of organic mixtures. Ion irradiation of simple organics (methanol, CH3OH; methane, CH4; and benzene, C6H6) (Brunetto et al., 2006) produced strong reddening and darkening. However, ion irradiation (Moroz et al., 2003, 2004) of complex hydrocarbons showed a decrease in the spectral slope with increasing irradiation.

2

C2-ungrouped chondrite Tagish Lake

Normalized reflectance

1.8 1.6

CI1 chondrite Orgueil

1.4 1.2 1

CM2 chondrite Murchison

0.8

∼0.7 μm

0.6 0.4 0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 6 Plot of reflectance spectra of CM2 chondrite Murchison (particle size <63 mm; Hiroi et al., 1993b), CI1 chondrite Orgueil (particle size <100 mm; Hiroi et al., 1993b), and C2-ungrouped chondrite Tagish Lake (particle size <125 mm; Hiroi and Hasegawa, 2003). Feature at
0.7 mm is identified in the Murchison spectrum. Spectra are normalized to unity at 0.55 mm.

Asteroids

373

1.6 1.5

Normalized reflectance

S-type (25143) Itokawa 1.4 1.3 1.2 1.1 1

LL4 chondrite greenwell springs

0.9 0.8

0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 7 Plot of reflectance spectrum of S-type (25143) Itokawa (Hayabusa sample return target) (squares) (Binzel et al., 2001b) versus spectrum of LL4 chondrite Greenwell Springs (line) (particle size <150 mm). Asteroid data are a combination of SMASS and SpeX spectra. Error bars are 1s.

Another process that can potentially alter the spectra of ‘primitive’ NEAs is heating of the surface (Marchi et al., 2009) as the object gets close to the Sun. Mueller et al. (2011) found that a number of NEOs had surfaces that could have been heated to temperatures high (
300 K) enough to destroy organics and dehydrate silicates (Delbo´ and Michel, 2011). Hiroi et al. (1993b, 1996) found that the heating of a CM chondrite to 673 K causes the 0.7 mm feature to disappear and the spectral slope to become less red. Wittmann et al. (2011) found evidence for transient post-metamorphic heating in a H/L chondrite and argue that its heating occurred when its source body traveled close to the Sun.

2.14.3.9

Determining Mineralogies

A number of different methods have been used to determine quantitative mineralogies of asteroids. Simple techniques (e.g., Hiroi et al., 1993a) use a checkerboard model, which linearly adds the spectra of a number of materials. These simple models assume that different materials are located on discrete areas on the surface. However, most asteroid surfaces are thought to be intimate mixtures of minerals. Most spectral mixing models used for asteroid incorporate intimate mixing based on Hapke (1981) theory. These Hapke (1981) theory models (e.g., Clark, 1995; Lawrence and Lucey, 2007) use the optical constants and particle sizes for specific mineral phases for an assumed geometry to calculate reflectance spectra that can be compared to asteroid spectra. The newest Hapke (1981) theory models (e.g., Lawrence and Lucey, 2007) consider the effects of nanophase iron on the surface to simulate space weathering. Another intimate mixture model (Shkuratov et al., 1999) has been also used to model (e.g., Vernazza et al., 2008) the reflectance spectra of asteroids. Cloutis et al. (1986) developed a formula for determining the relative abundances of olivine and low-Ca pyroxene from the Band Area Ratio of an olivine–low-Ca pyroxene mixture.

Gaffey et al. (2002) developed a series of formulas to derive the pyroxene mineralogies from the Band I and II centers of a pyroxene reflectance spectrum. Due to the Gaffey et al. (2002) formulas erroneously calculating high ferrosilite values from ordinary chondrite spectra using their Band II centers, Gaffey (2007) developed a method to correct for this error when analyzing ordinary chondrite spectra. Burbine et al. (2009) developed formulas for determining the Fs and Wo contents of pyroxenes on the surfaces of bodies with HED-like mineralogies. To better determine the mineralogy of ordinary chondrite assemblages using reflectance spectra, Dunn et al. (2010c) developed a number of formulas for determining the mineralogies of ordinary chondrites from their reflectance spectra using ordinary chondrite powders originally prepared by Jarosewich (1990) and mineralogically and spectrally characterized by Dunn et al. (2010a,b). The mineral percentages were determined using x-ray diffraction, while olivine and low-Ca pyroxene compositions were determined using microprobe analyses. Using methods derived from the work of Miyamoto and Zolensky (1994) and Sato et al. (1997) on carbonaceous chondrites, Rivkin et al. (2003) developed a formula for determining the equivalent water content (all the hydrogen is in the water) from 3 mm observations. The modified Gaussian model (MGM) (e.g., Sunshine and Pieters, 1993; Sunshine et al., 1990) was developed to resolve reflectance spectra into their constituent absorption bands by fitting absorption features with a series of modified Gaussians. Each individual absorption band is defined by three parameters (its center, its width, and its strength), which can be varied or kept constant. The positions of these bands can then be correlated with the positions of bands of actual minerals (e.g., olivine, pyroxene), and the strengths of the bands can be used to determine the abundance of the minerals. MGM has been used to estimate the mineralogies of a number of asteroids (e.g., de Leo´n et al., 2006; Sunshine et al., 2004, 2007).

374

Asteroids

2.14.4

Taxonomy

Asteroid taxonomies classify asteroids according to how their spectral reflectance changes versus wavelength. Sometimes these taxonomies use albedo to differentiate between classes with similar spectral properties. Each taxonomy was developed using a particular asteroid photometric or spectral survey but could be applied to later data sets. The taxonomy that is the basis for almost all asteroid classification schemes used today was developed by Tholen (1984). Using data from the ECAS (Zellner et al., 1985) plus visual albedo (when available), Tholen (1984) defined fourteen asteroid classes. These classes included the A-types, C-group (B, C, F, and G), D-types, Q-types, R-types, S-group, T-types, V-types, and X-group (E, M, and P). The members of the X-group can only be distinguished according to visual albedo, with the E-types having high albedos, M-types having moderate albedos, and P-types having low albedos. The P-types are also part of the C-group since they fall at one end of the C-class distribution. Near-infrared observations (
0.9–2.5 mm) of objects that were classified using visible data often show that asteroids with the same classification have a wide variety of mineralogies. Asteroids that were classified as S-types have near-infrared spectra (Gaffey et al., 1993) with olivine/(olivineþpyroxene) mineralogies that ranged from almost entirely olivine to almost entirely pyroxene. Gaffey et al. (1993) used a classification scheme (Figure 8) for S-types based on Band Area Ratios and Band I centers to create the S(I), S(II), S(III), S(IV), S(V), S(VI), and S(VII) subtypes, which have varying ratios of olivine to pyroxene. The Band Area Ratio tends to increase and the Band I Center tends to decrease as the Roman numeral increases in value, which tends to indicate an increasing pyroxene concentration. The S(I)subtype had olivine-dominated mineralogies, the S(IV)-subtype had olivine–pyroxene mineralogies similar to ordinary chondrites, and the S(VII)-subtype had pyroxene-dominated mineralogies. Observations of both M-type and E-type asteroids (Rivkin et al., 1995, 2000) show that these classes include objects with no detectable 3 mm absorption bands and objects with 3 mm absorption bands, implying objects with hydrated silicates on the surface exist among both M- and E-types. 1.15

Band I center (μm)

1.1 S(I)

1.05

S(II) S(III)

1

S(VII) S(V)

0.95

S(IV) S(VI)

HEDs

0.9

0

0.5

1

1.5

2

2.5

3

Figure 8 Distribution of Gaffey et al. (1993) subtypes on a plot of Band Area Ratio versus Band I center. Green line is olivine–orthopyroxene mixing line. Ordinary chondrite region is the same as the S(IV) region. HED region (rectangle) is also plotted.

Many of the letters used imply a composition even though the taxonomies only group objects according to similar spectral properties, which may not be indicative of a unique composition. ‘C’ was an abbreviation for carbonaceous, ‘E’ for enstatite, ‘M’ for metal, ‘O’ for ordinary chondrite, ‘S’ for stony, and ‘V’ for Vesta. For example, the M-types with 3 mm bands do not appear consistent with metallic iron surfaces. Rivkin et al. (1995) created the W-class for M-types with 3 mm bands. Shepard et al. (2010) proposed calling M-types with high radar albedos as Mm-types. Gaffey and Kelley (2004) proposed three subtypes for the E-types: E(I), E(II), and E(III). Members of the E(I)-subtype have a slightly reddish or curved spectrum and lack distinctive absorption bands. Members of the E(II)-subtype have a relatively strong feature centered near 0.49 mm and occasionally a weaker feature near 0.96 mm. Members of the E(III)-subtype have a flat or slightly reddish spectrum and have a weak but well-defined
0.88–0.90 mm band. Clark et al. (2004a) also separated E-types into three subtypes but with a slightly different membership in each subtype. Bus and Binzel (2002b) defined 26 classes on the basis of SMASS II (Bus and Binzel, 2002a) visible (
0.44–0.93 mm) CCD spectra (Figure 9). Bus and Binzel (2002b) defined more classes than Tholen (1984) due to their higher-resolution spectra and ability to observe smaller bodies. The classes included the A-types, C-complex (B, C, Cb, Cg, Cgh, and Ch), D-types, K-types, L-types, Ld-types, O-types, Q-types, R-types, S-complex (S, Sa, Sk, Sl, Sq, and Sr), T-types, V-types, and X-complex (X, Xc, Xe, and Xk). The lowercase letters indicate either a class that is intermediate between two types (Cb, Sa, Sk, Sl, Sq, Sr, Xc, and Xk) or the presence of a specific feature (Cg, Cgh, Ch, and Xe). The ‘g’ designation refers to a strong UV absorption, the ‘h’ refers to a feature at 0.7 mm, and ‘e’ refers to an absorption band shortward of 0.55 mm. DeMeo et al. (2009) defined 24 classes (Figure 10) on the basis of visible (primarily SMASS II data) and near-infrared spectra (SpeX) using the Bus and Binzel (2002b) taxonomic system as a guide. The Sl-, Sk-, and Ld-types were eliminated and one class (Sv) (intermediate between S- and V-types) was created. The ‘w’ notation was also added to identify objects with similar spectral features but had high spectral slopes. These high slope S objects include Sw, Sqw, Srw, and Svw. Originally SDSS colors were used to define three broad types (S-, C-, and V-types) (Ivezic´ et al., 2001; Juric´ et al., 2002), but each of these three types contained a number of different taxonomic classes. However, using SDSS colors for over 60 000 asteroids, Carvano et al. (2010) defined nine classes (Vp, Op, Qp, Sp, Ap, Lp, Dp, Xp, and Cp) where the lowercase p stands for photometric. Except for the Lp class (which includes the K, L, and Ld classes), the classes are intended to be directly comparable to the Bus and Binzel (2002b) classes. SDSS colors can also be used to identify interesting objects that can be designated for later higherresolution spectral observations. For example, SDSS colors have been very useful for identifying objects with V-type visible spectra (e.g., Binzel et al., 2006a; Hammergren et al., 2011; Roig and Gil-Hutton, 2006). The following sections will discuss the spectral properties and interpreted mineralogies of members of each asteroid class. To group objects with similar visible and near-infrared spectra (and, hopefully, mineralogies), the DeMeo et al.

Asteroids

Xe

Cb

B

X D Ch

T Xc

C

Xk

1.2 1.0 0.43

Cgh 0.93

Ld

Cg

L

K

Sk

SI S

Sq

So Sr

Q

O

A

R V Figure 9 Plots of the Bus and Binzel (2002b) classes. The wavelength region is
0.44–0.93 mm. Plot made available by S. J. Bus.

S-complex 1.5

S1

Sa

Sq

Sr

Sv

C

Cb

Cg

Cgh

Xc

Xe

Xk

D

K

L

T

A

O

Q

R

0.45

2.45

C-complex B

Ch

X-complex X Endmembers

V

Figure 10 Plots of the DeMeo et al. (2009) S-complex (S, Sa, Sq, Sr, Sv), C-complex (B, C, Cb, Cg, Cgh, Ch), X-complex (X, Xc, Xe, Xk), and endmember classes (A, D, K, L, O, Q, R, T, V). The wavelength region is
0.4–2.5 mm. Plot made available by F. E. DeMeo.

375

376

Asteroids

(2009) taxonomic classes will primarily be used. If an object was not classified by DeMeo et al. (2009), other taxonomic designations (e.g., Bus and Binzel, 2002b; Tholen, 1984) will be used instead. One exception will be the discussion of the E-, M-, and P-types since albedo information appears crucial for determining the mineralogies of these objects. Another is the discussion of (24) Themis with the B-types even though it is now classified as a C-type by DeMeo et al. (2009). A third exception is that the Gaffey et al. (1993) S(I)-subtype, which has interpreted olivine-dominated silicate mineralogies, will be discussed with the A-types.

2.14.4.1

A-Types

In the visible (Bus and Binzel, 2002b; Tholen, 1984), A-types have very strong UV features and moderately deep 1 mm bands that are similar to the spectra of olivine (Figure 9). Nearinfrared spectra (Cruikshank and Hartmann, 1984; DeMeo et al., 2009) of A-types tend to confirm their spectral similarity to olivine (Figure 11). Meteoritic analogs (Cruikshank and Hartmann, 1984; Sunshine et al., 2007) for A-types are olivine-dominated meteorites such as the pallasites, the brachinites, and the R chondrites. The fayalite content varies among these possible analogs, with most pallasites containing olivine with Fa121 (Mittlefehldt et al., 1998), brachinites containing Fa
30–35 (Mittlefehldt et al., 1998), and R chondrites containing Fa
37–40 (Brearley and Jones, 1998). A few pallasites have fayalite compositions as high as Fa18. Postulated linkages between A-types and pallasites tend to assume surface compositions that do not have a significant abundance of metallic iron. A-types tend to have much redder near-infrared spectra when compared to laboratory measurements of olivine, which is consistent with space weathering reddening their surfaces (Burbine and Binzel, 2002; Hiroi and Sasaki, 2001; Lucey et al., 1998). The average visible albedo of D-types from NEOWISE observations is
0.19 (Mainzer et al., 2011d). Sunshine et al. (2007) found that a number of A-types ((354) Eleonora, (446) Aeternitas, (863) Benkoela, (984) Gretia, (2501) Lohja, (3819) Robinson) had very magnesian (forsteritic) olivine compositions from MGM of A-type visible

and near-infrared spectra. Eleonora was classified by Gaffey et al. (1993) as an S(I)-subtype, which has interpreted silicate mineralogies that consist almost entirely of olivine. However, two A-types, (246) Asporina and (289) Nenetta, were found to be more ferroan (fayalitic) by Sunshine et al. (2007). The forsteritic A-types have mineralogies consistent with differentiation from an ordinary chondrite-like body, while the fayalitic A-types are consistent with the more oxidized R-types or melting from such a body. Five of these objects, (446) Aeternitas, (863) Benkoela, (984) Gretia, (2501) Lohja, and (3819) Robinson, have a Band II, which is consistent with 5–10% pyroxene on their surfaces. Lucey et al. (1998) had previously argued that the reflectance spectra of (246) Asporina, (289) Nenetta, (446) Aeternitas, and (863) Benkoela were consistent with low FeObearing olivines measured at asteroidal temperatures. Burbine and Binzel (2002) noted that a number of objects ((1126) Otero, (1600) Vyssotsky, (2732) Witt, (4142) DersaUzala, (4713) Steel) classified as A-types by Bus and Binzel (2002b) on the basis of visible spectra had much weaker bands in the near-infrared out to
1.65 mm than those commonly found for ‘typical’ A-types. These A-types with weaker Band Is have estimated diameters less than
13 km, while all the A-types with much stronger bands have diameters larger than 27 km. Using spectra out to 2.5 mm, Witt was classified as an Ltype and Steel was classified as an Sw-type by DeMeo et al. (2009) because they did not have strong Band I absorptions but did have apparent Band IIs. Most A-type bodies are commonly assumed to be remnants of mantle material from disrupted differentiated bodies. However, only six bodies (
2%) out of 371 objects were classified as A-types by DeMeo et al. (2009). Also, one object, (9) Metis, that was classified as ungrouped by Gaffey et al. (1993) was later classified by Kelley and Gaffey (2000) as an S(I)-subtype. Kelley and Gaffey (2000) noted that the calculated olivinedominated silicate mineralogy of Metis was similar to the assumed mineralogy of S(I)-subtype (113) Amalthea and argued that both objects may be fragments of a differentiated body with Metis containing more metallic iron. If differentiated bodies formed metallic iron cores, olivinedominated mantles, and basaltic crusts, the disruption of these

3

Normalized reflectance

A-type (863) Benkoela 2.5 A-type (354) Eleonora

2 1.5 1

Brachinite EET 99402 0.5 0 0

0.5

1 1.5 Wavelength (μm)

2

2.5

Figure 11 Plot of reflectance spectra of A-types (354) Eleonora (squares) and (863) Benkoela (squares) versus spectrum of brachinite EET 99402 (line) (particle size <125 mm; Burbine et al., 2008). Spectra are normalized to unity at 0.55 mm. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Error bars are 1s.

Asteroids

objects should produce significant numbers of olivinedominated fragments. The low number of identified A-types in the main belt that are currently present in the belt appears inconsistent (Chapman, 1986) with the large number of differentiated bodies that should have disrupted to produce the number of iron meteorites that are present in our meteorite collections. One scenario is that these olivine-dominated objects have been ‘battered to bits’ (Burbine et al., 1996) and are currently at sizes not observable in the main belt.

2.14.4.2

C-Complex

In the visible (Bus and Binzel, 2002b; Tholen, 1984), C-complex (or C-group) asteroids tend to have weak UV features shortward of 0.55 mm and relatively flat to bluish spectra longward of 0.55 mm (Figure 9). The C-complex objects of Bus and Binzel (2002b) included B, C, Cb, Cg, Cgh, and Ch objects, while the C-group of Tholen (1984) included B, C, F, and G objects. Objects with redder spectra would be classified as P-types. The C-complex of DeMeo et al. (2009) included the B, C, Cb, Cg, and Cgh objects. Approximately 13% of classified asteroids by DeMeo et al. (2009) are part of the C-complex. The average visible albedo of C-complex objects from NEOWISE observations is
0.06 (Mainzer et al., 2011d). Approximately 60% of observed C-complex objects (Rivkin et al., 2002) have 3 mm bands, indicating hydrated silicates on their surfaces. Using SDSS colors to estimate that 30  5% of C-complex objects have 0.7 mm bands (and, therefore, have 3 mm bands) and assuming half of the remaining C-complex bodies have 3 mm bands, Rivkin (2012) estimates that approximately two-thirds of C-complex asteroids have hydrated minerals on their surfaces. Traditionally, C-complex objects have been linked with carbonaceous chondrites (e.g., Feierberg et al., 1985; Johnson and Fanale, 1973) or thermally metamorphosed carbonaceous chondrites (Hiroi et al., 1993b; Ostrowski et al., 2010, 2011). The two largest objects in the main belt are C-type dwarf planet (1) Ceres (diameter of
950 km) and B-type (2) Pallas (diameter of
530 km).

377

2.14.4.2.1 B-types B-type asteroids have bluish spectral slopes longward of
0.5 mm. Yang and Jewitt (2010) noted that a number of asteroids that have been orbits consistent with comets or associated with meteor showers have been classified as B-types, implying that these objects may have originally incorporated significant amounts of water ice. This is consistent with the idea that there is a continuum between dark asteroids and comets (Gounelle et al., 2008). On the basis of visible and near-infrared spectra, Clark et al. (2010) broke up B-type spectra into three groups. Clark et al. (2010) looked at objects that were classified as a B-type by either Tholen (1984), Bus and Binzel (2002b), or DeMeo et al. (2009). One is the Pallas group, which have blue spectral slopes like Pallas (Figure 12). Another is the Themis group, which have concave curve shapes like 24 Themis (Figure 12). The other group have spectra unlike Pallas or Themis. Pallas has long been known (Feierberg et al., 1981; Jones et al., 1990; Larson et al., 1979, 1983; Rivkin et al., 2011b) to have a 3 mm band (Figure 13) that was similar in structure to the bands found for CM chondrites. Beck et al. (2010) noted the Pallas will even be a better match to CM chondrites in the 3 mm region if terrestrial adsorbed water is removed from CM chondrite spectra. Sato et al. (1997) found the best spectral match to Pallas in the 3 mm region was CR chondrite Renazzo. Pallas’ estimated bulk density of 2400–2600 kg m3 (Baer et al., 2011; Schmidt et al., 2009; Zielenbach, 2011) and the presence of a 3 mm band are both consistent with a body that formed from water-rich material. Modeling the accretion ages (which affects the initial 26Al concentration in the body) and sizes needed to result in internal differentiation and an unmelted crust, Elkins-Tanton et al. (2011) found that the shapes and masses of Pallas and Ceres are consistent with differentiated interiors with small iron cores but with hydrated silicate or ice–silicate mantles and an undifferentiated crust. Yang and Jewitt (2010) found that the concave shapes of a number of B-types in the near-infrared are consistent with a broad 1 mm absorption feature that is consistent with the mineral magnetite (Fe3O4). Magnetite is commonly believed to be

1.5 C-type (1) Ceres

Normalized reflectance

1.4 1.3 B-type (24) Themis 1.2 1.1 B-type (2) Pallas

1 0.9 0.8

0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 12 Plot of reflectance spectra of C-complex objects (2) Pallas, (24) Themis, and (1) Ceres. Data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Spectra are normalized to unity at 0.55 mm and then offset in reflectance. Error bars are 1s.

378

Asteroids

1.5 C-type (1) Ceres

Normalized reflectance

1.4

~3.06 μm

1.3 1.2

B-type (24) Themis

1.1

~3.4 μm

~3.1 μm

1 0.9

B-type (2) Pallas

0.8 0.7 0.6 0.5

2

2.5

3 3.5 Wavelength (μm)

4

4.5

Figure 13 Plot of 3 mm spectra of C-complex objects (2) Pallas (Lebofsky et al., 1995), (24) Themis (Rivkin et al., 2011a), and (1) Ceres (Milliken and Rivkin, 2009). Features at
3.1 and
3.4 mm are identified in the Themis spectrum and a feature at
3.06 mm is identified in the Ceres spectrum. Spectra are normalized to unity at 2.5 mm and then offset in reflectance. Extremely anomalous points have been deleted. Error bars are 1s.

the product of aqueous alteration. One object, (335) Roberta, a Themis group B-type (Clark et al., 2010), had a spectrum consistent with a mixture of a CI chondrite and magnetite. Both Roberta and the CI chondrite–magnetite mixture have a broad 1 mm band and a 3 mm feature. Themis was found to have spectral features in the 3 mm region (Figure 13) consistent with water ice (band at
3.1 mm) and organics (band at
3.4 mm) (Campins et al., 2010a; Rivkin and Emery, 2010). However, Beck et al. (2011b) have found that goethite also has absorption features similar to those found for Themis. Due to the absence of detectable CN emission lines, which are used as tracers of water in comets, and assuming cometary H2O/CN mixing ratios, an upper limit for surface water ice was determined for Themis by Jewitt and Guilbert-Lepoutre (2012) to be less than
10%.

2.14.4.2.2 C-types Ceres is the most studied C-type body due to its large size; however, the mineralogy of Ceres has been considerably debated (e.g., McCord et al., 2011; Rivkin et al., 2011d). In the visible and near-infrared (
0.4–2.5 mm), Ceres has a moderate strength UV features and a broad absorption band centered around 1.2 mm that has been attributed to magnetite (Figure 12; Larson et al., 1979; Rivkin et al., 2011d). With a spatial resolution of
75 km, Carry et al. (2012) found that Ceres was very homogenous spectrally between 1.17 and 1.32 mm and between 1.45 and 2.35 mm. However, the most prominent absorption bands for Ceres are in the 3 mm region (Figure 13), but these bands are unlike those found in carbonaceous chondrites. Lebofsky et al. (1981) argue that the 3 mm band was due to structural OH in hydrated silicates and a narrow absorption feature at
3.07 mm was due to water ice. King et al. (1992) believed that this feature was due to a NH4bearing phyllosilicate, which has an absorption feature due to NH4þ that is approximately at
3.07 mm and has a similar shape. Vernazza et al. (2005) also observed this feature at
3.06 mm and attributed it to water ice in a mixture of ion-irradiated organics. Milliken and Rivkin (2009) argued that the 3 mm absorption features for Ceres are consistent with the presence of the hydroxide brucite (Mg(OH)2), magnesium carbonates (MgCO3), and

serpentines ((Mg, Fe)3Si2O5(OH)4). This assemblage is not found in known carbonaceous chondrites. For an object with Ceres’ estimated bulk density of
2100 kg m3 (Baer et al., 2011; Thomas et al., 2005; Zielenbach, 2011), thermal modeling of Ceres by CastilloRogez (2011) finds that a significant amount of water ice must exist in the interior to produce Ceres’ density for the temperatures that Ceres must have experienced. A’Hearn and Feldman (1992) detected OH around Ceres using the International Ultraviolet Explorer measurements; however, Rousselot et al. (2011), using the more sensitive Ultraviolet and Visual Echelle Spectrograph on very large telescope, did not detect any OH emission around Ceres. As discussed earlier, ElkinsTanton et al. (2011) believe that Ceres may have a differentiated interior with a small iron core but with a hydrated silicate or ice–silicate mantle and an undifferentiated crust. C-types with a feature at
0.7 mm are given the ‘h’ designation. This band tends to be found in the spectra of CM2 chondrites (Hiroi et al., 1993b, 1996). Objects with a
0.7 mm band almost always have a 3 mm band (Howell et al., 2001; Rivkin et al., 2002; Vilas, 1994), which is consistent with this linkage to CM chondrites. Ch-types (13) Egeria and (19) Fortuna (Figure 14) were originally noted by Burbine (1998) to be spectrally similar to CM chondrites in the visible and near-infrared and to have 3 mm band strengths (Jones et al., 1990) consistent with CM2 chondrites. However, the calculated bulk density of Egeria,
3400 kg m3 (Baer et al., 2011), is much higher than the average value for CM chondrites,
2100 kg m3 (Britt and Consolmagno, 2003), while Fortuna’s bulk density,
1800 kg m3, is consistent with CM chondrites with some porosity. Egeria and Fortuna are both located near the 3:1 resonance, which may make it easier for the relatively weak CM fragments to make it into a meteorite-supplying resonance. CM fragments may not be strong enough to drift considerable distances in the belt and still survive. Cgh-types have stronger UV features than Ch-types and include (106) Dione, (706) Hirundo, and (776) Berbericia. Berbericia also has a 3 mm band (Jones et al., 1990). Cruikshank and Brown (1987) identified a 3.4 mm absorption band on Ch-type (130) Elektra, which was originally

Asteroids

classified as a G-type by Tholen (1984). This band was argued to be due to a C–H stretching mode in organics (e.g., Moroz et al., 1998). However, this feature was not found (Cruikshank et al., 2002) in a later spectrum of Elektra or in the spectra of a number of observed G-, P-, or D-types. Cruikshank et al. (2002) argue that this band may not be present in asteroid spectra due to space weathering, which could dehydrogenate C–H-bearing hydrocarbons over time through interaction with the solar wind.

2.14.4.3

D-Types

In the visible (Bus and Binzel, 2002b; Tholen, 1984), D-types have very red slopes (Figure 9). A number of D-types were found by Vilas et al. (1993) and Mothe´-Diniz (2010) to have weak absorptions between 0.60–0.65 and 0.80–0.90 mm that they attribute to charge transfer transitions in iron oxides and/ or sulfates (e.g., Cloutis et al., 2006b). In the near-infrared (DeMeo et al., 2009), D-types also have a very steep spectral

379

slope (e.g., Emery and Brown, 2003) with a slight curvature around 1.5 mm (Figure 15). Using SDSS colors and NEOWISE observations, Mainzer et al. (2012) found that red spectral slopes of Dp-types extend out to 3–4 mm. The spectral properties of D-types are usually thought to be due to organics (e.g., Gradie and Veverka, 1980), which would redden and darken the surface. Cruikshank et al. (1991a) identified a feature at
2.2 mm in the spectra of a number of D-types as being due to a CN stretching mode in organics. The average visible albedo of D-types from NEOWISE observations is
0.05 (Mainzer et al., 2011d). The presence of the 0.60–0.65 and the 0.80– 0.90 mm bands is consistent (Mothe´-Diniz, 2010; Vilas et al., 1993) with charge transfer absorptions of Fe3þ ! Fe2þ in minerals such as goethite (FeO(OH)), hematite (Fe2O3), and jarosite (KFe3(OH)6(SO4)2). D-types are commonly found in the Hilda group (e.g., Dahlgren and Lagerkvist, 1995; Grav et al, 2012a) and among Jupiter Trojans (e.g., Fornasier et al., 2007). About 2/3 of observed Trojans (Roig et al., 2008) using both SDSS colors

Normalized reflectance

1.3 CM2 chondrite LEW 90500

1.2 1.1

Ch-type (19) Fortuna 1 0.9 ~0.7 μm 0.8 0

0.5

1 1.5 Wavelength (μm)

2

2.5

Figure 14 Plot of reflectance spectrum of Ch-type (19) Fortuna (squares) versus spectrum of CM2 chondrite LEW 90500 (red line) (particle size <100 mm). Feature at
0.7 mm is identified in the asteroid and meteorite spectra. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

2

Normalized reflectance

1.8 D-type (3248) Farinella

1.6 1.4

C2-ungrouped chondrite Tagish Lake

1.2 1 0.8 0.6 0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 15 Plot of reflectance spectrum of D-type (3248) Farinella versus spectrum of C2-ungrouped chondrite Tagish Lake (particle size <125 mm; Hiroi and Hasegawa, 2003). Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra (Bus, 2011; Clark et al., 2009). Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

380

Asteroids

and reflectance spectra have reddish slopes compatible with D-type asteroids. A number of D-types, (884) Priamus, (1172) Aneas, (2207) Antenor, (2241) Alcathous, and (2357) Phereclos and one DU-type, (2223) Sarpedon, were observed by Howell (1995) to have 1 mm bands; however, observations by Yang and Jewitt (2011) could not confirm the presence of a 1 mm feature on any of these objects. Emery et al. (2006) found that D-type Trojans, (624) Hektor, (911) Agamemnon, and (1172) Aneas, had emission spectra from Spitzer Space Telescope observations that were consistent with fine-grained silicates. An emissivity high near 10 mm (assumed to be the Si–O stretch fundamental) and a broader high from
18 to 28 mm (assumed to be the Si–O bend fundamental) were apparent in the spectra. D-type asteroids also exist in the main asteroid belt. These objects include (336) Lacadiera (a ¼ 2.25 AU), (368) Haidea (a ¼ 3.07 AU), and (773) Irmintraud (a ¼ 2.86 AU). Hiroi et al. (2001b) noted the spectral similarities between main-belt D-types and the C2-ungrouped chondrite Tagish Lake. Kanno et al. (2003) found a 3 mm feature for Irmintraud, indicating either OH or H2O in hydrous minerals or water ice on the surface. Observations of Jones et al. (1990) did not detect the presence of such a band; however, Lebofsky (1991) preliminarily identified a 3 mm feature in the spectrum of Irmintraud.

2.14.4.4

K-Types

The K-class was first proposed by Bell (1988) for objects in the (221) Eos family (including Eos) that have spectral properties and visual albedos intermediate between S-types and C-types. The average visible albedo of K-types from NEOWISE observations is
0.13 (Mainzer et al., 2011d). In the visible (Bus and Binzel, 2002b), K-types have moderately strong UV features and a weak 1 mm feature (Figure 9). In the near-infrared (DeMeo et al., 2009), K-types have relatively wide but shallow 1 mm bands and spectral slopes (Figure 16) that vary from slightly bluish to slightly red. K-types have also been identified outside the Eos family. K-types have typically been linked to CO and CV chondrites (Bell, 1988; Burbine et al., 2001a; Clark et al., 2009) due to

their similar spectral properties in the visible and near-infrared and visual albedos. Clark et al. (2009) found that the best meteorite matches to K-types were CO and CK chondrites. Burbine et al. (2001a) noted that the best spectral match to Eos and (653) Berenike (also in the Eos family) out to
1.65 mm was a CO3 chondrite and that the best match to 599 Luisa (outside the Eos family) was a CV3 chondrite. However, Mothe´-Diniz and Carvano (2005) noted the spectral similarity of Eos to the differentiated meteorite Divnoe. Mothe´Diniz et al. (2008) using intimate mixing based on Hapke (1981) theory and MGM found that a number of K-types in the Eos family had calculated mineralogies consistent with forsteritic olivine (Fa
20), which they argue is consistent with either partial differentiation of an ordinary chondrite parent body or a CK chondritic mineralogy. Asteroid (15) Eunomia, which was classified as a K-type by DeMeo et al. (2009) but as an S-type by Tholen (1984) and Bus and Binzel (2002b), and other Eunomia family members also have had their mineralogies interpreted (Nathues, 2010; Nathues et al., 2005) as being consistent with the original parent body being partially differentiated.

2.14.4.5

L-Types

In the visible (Bus and Binzel, 2002b), L-types have very steep UV features shortward of 0.75 mm and then becoming flat longward of 0.75 mm (Figure 9). In the near-infrared (DeMeo et al., 2009), L-types often have a gentle concave down curvature with a maximum around 1.5 mm and often a slight 2 mm absorption band (Figure 17). Ld-types (Bus and Binzel, 2002b), which have visible spectra intermediate between L-types and D-types, were eliminated in the DeMeo et al. (2009) taxonomy. The average visible albedo of L-types from NEOWISE observations is
0.15 (Mainzer et al., 2011d). A number of L-types ((234) Barbara, (387) Aquitania, (599) Luisa, (980) Anacostia, (1858) Lobachevsky, (4426) Roerich, (5840) Raybrown) have absorption features consistent (Sunshine et al., 2008) with iron oxide-bearing aluminous spinel found in calcium–aluminum inclusions (CAIs) (Figure 17).

1.2

Normalized reflectance

K-type (221) Eos 1.1

1

0.9

0.8 0

0.5

1 1.5 Wavelength (μm)

2

2.5

Figure 16 Plot of reflectance spectrum of K-type (221) Eos. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra (Bus, 2011; Clark et al., 2009). The SpeX data shortward of
0.9 mm do not overlap well with SMASS II data. Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

Asteroids

1.4

L-type (387) Aquitania

1.4 1.2

Type B CAI

1 Fluffy Type A CAI 0.8

Normalized reflectance

1.6 Normalized reflectance

381

O-type (3628) Božn mcová 1.2

1

0.8

0.6 0

0.5

1 1.5 Wavelength (μm)

2

2.5

Figure 17 Plot of reflectance spectrum of L-type (387) Aquitania (squares) versus spectra of a Fluffy Type A CAI (line) (particle size <38 mm; Sunshine et al., 2008) and a Type B CAI (particle size <38 mm) (line) (Sunshine et al., 2008). Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra (Bus, 2011). Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

Aquitania and Anacostia were previously identified by Burbine et al. (1992) as possibly having spinel-rich surfaces. Cellino et al. (2006), Gil-Hutton et al. (2008), and Masiero and Cellino (2009) found that Barbara, Aquitania, and Anacostia have unusual polarimetric properties compared to other asteroids and was consistent with a mixture of high- and low-albedo particles in their regoliths. Polarimetric observations measures the degree of linear polarization versus phase angle for photons scattered off a surface (e.g., Muinonen et al., 2002). Sunshine et al. (2008) argue that these objects contain extremely high (30  10 vol.%) CAI abundances compared to meteorites and may be more ancient than anything found in our meteorite collections. They believe that these objects may have formed before the injection of the 26Al into the solar system since these objects would have melted with such high CAI abundances and canonical initial 26Al/27Al ratios. The radioactive isotope 26Al is the prime candidate for the heat source that causes asteroids to differentiate. However, Hezel and Russell (2008) argue that the actual CAI abundances are lower (10 vol.%) and the objects would not have melted if they had canonical 26Al/27Al ratios (see Chapter 1.11). Then, these bodies would not necessarily be more ancient than anything found in our meteorite collections.

2.14.4.6

O-Types

The O-type was first proposed by Binzel et al. (1993) based on the unusual visual spectra of (3628) Bozˇneˇmcova´. In the visible (Bus and Binzel, 2002b), O-types have moderately steep UV features shortward of 0.54 mm, becoming less steep over the interval from 0.54 to 0.7 mm, and a relatively deep 1 mm band (Figure 9). DeMeo et al. (2009) only classified one object (Bozˇneˇmcova´) as an O-type and found that it has a very rounded and deep, ‘bowl-shaped’ absorption feature at 1 mm (Figure 18) as well as a 2 mm band, consistent with the Burbine and Binzel (2002) observations out to
1.65 mm. Using visible and near-infrared spectra out to
1.65 mm, Cloutis et al. (2006a) interpreted the reflectance spectrum of Bozˇneˇmcova´ to be consistent with an assemblage of high-Ca pyroxene and plagioclase feldspar that would be

0.6

0

0.5

1

1.5

2

2.5

Figure 18 Plot of reflectance spectrum of O-type (3628) Bozˇneˇmcova´. Spectra are normalized to unity at 0.55 mm. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Error bars are 1s.

mineralogically similar to angrites. Angrites contain a high-Ca pyroxene (diopside–hedenbergite) (Ca(Mg,Al,Ti)(Si,Al)2O6) that typically does not a 2 mm feature (Burbine et al., 2006). However, near-infrared spectra out to 2.5 mm of Bozˇneˇmcova´ (Burbine et al., 2011; DeMeo et al., 2009) shows a distinctive 2 mm feature that is not consistent with an angrite composition. Burbine et al. (2011) found that even though the shapes of the absorption bands for Bozˇneˇmcova´ appeared characteristic of pyroxenes, the Band I and II centers were offset from the typical values found for pyroxenes. No known meteorite has a spectrum similar to that of Bozˇneˇmcova´.

2.14.4.7

Q-Types

In the visible (Bus and Binzel, 2002b; Tholen, 1984), Q-types have moderately strong UV features and moderately deep 1 mm bands that are similar to the spectra of ordinary chondrites (Figure 9). In the near-infrared (DeMeo et al., 2009), Q-types have distinct 1 and 2 mm bands that also resemble those of ordinary chondrites. All Q-types identified by DeMeo et al. (2009) were found among NEAs. Approximately 6% of NEAs (Binzel et al., 2004b) are classified as Q-types. Near-Earth asteroid (1862) Apollo was the first identified Q-type (Tholen, 1984) and has long been known to be spectrally similar to LL chondrites (Figure 19) in the visible (McFadden et al., 1985) and near-infrared (Binzel et al., 2006b). Other Q-types identified by DeMeo et al. (2009) include NEAs (3753) Cruithne, (4688) 1980 WF, (5143) Heracles, (5660) 1974 MA, (7341) 1991 VK, (66146) 1998 TU3, and (162058) 1997 AE12. Q-types are relatively rare in the main belt, but a few ((1270) Datura by Chapman et al. (2009), (23338) 2809 P-L by Vernazza et al. (2006a), and (203370) 2001 WY35 by Mothe´Diniz and Nesvorny´ (2008a)) have been tentatively identified on the basis of visible and/or near-infrared spectroscopy.

2.14.4.8

R-Types

In the visible (Bus and Binzel, 2002b; Tholen, 1984), R-types have very strong UV features and a deep 1 mm band (Figure 9). DeMeo et al. (2009) only classified one object, (349) Dembowska, as an R-type because of its deep 1 and 2 mm

382

Asteroids

1.3

Normalized reflectance

Q-type (1862) Apollo 1.2

1.1

1 LL4 chondrite Greenwell Springs 0.9

0.8 0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 19 Plot of reflectance spectrum of Q-type (1862) Apollo (squares) versus spectrum of LL4 chondrite Greenwell Springs (line) (particle size <150 mm). Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

Normalized reflectance

1.8 1.6

R-type (349) Dembowska

1.4 1.2 1 0.8 0.8

0.5

1 1.5 Wavelength (μm)

2

2.5

Figure 20 Plot of reflectance spectrum of R-type (349) Dembowska. Spectra are normalized to unity at 0.55 mm. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Error bars are 1s. Spectrum made available by S. J. Bus.

features (Figure 20). The 1 mm feature is much narrower than those found in Q-types but slightly broader than those found in V-types. Gaffey et al. (1989) interpreted the reflectance spectra of Dembowska as indicating a partially melted pyroxene–olivine assemblage with little to no metallic iron. Using Band Area Ratios and Band I centers, Abell and Gaffey (2000) and Abell (2003) also interpreted rotational spectra of Dembowska as indicating a partially melted low-Ca pyroxene–olivine assemblage. Hiroi and Sasaki (2001) find that Dembowska’s reflectance spectrum is consistent with a space-weathered orthopyroxene–olivine assemblage. Both Abell and Gaffey (2000) and Hiroi and Sasaki (2001) find Dembowska’s orthopyroxene/olivine ratio to be
45:55.

2.14.4.9

S-Complex

S-asteroids have been the most studied type of asteroids since they are so abundant in the inner main belt, making them easier to identity and observe from Earth. It has long been

argued whether S-types are primarily bodies with ordinary chondrite-like mineralogies (e.g., Feierberg et al., 1982) or are bodies that have experienced melting and differentiation. In the visible (Bus and Binzel, 2002b; Tholen, 1984), S-asteroids have moderately strong UV features and 1 mm bands of varying strength (Figure 9). Bus and Binzel (2002b) broke the S-types into the S-complex, which contained the S, Sa, Sk, Sl, Sq, and Sr classes, to show the continuum in the spectral properties between the S-types and other taxonomic types (A, K, L, Q, and R). In the near-infrared, S-asteroids are known to have a wide variety of spectral properties (Figure 10; e.g., DeMeo et al., 2009; Gaffey et al., 1993). Gaffey et al. (1993) broke the S-types into the S(I), S(II), S(III), S(IV), S (V), S(VI), and S(VII) classes, while DeMeo et al. (2009) broke the S-types into the S, Sa, Sq, Sr, and Sv (intermediate between S and V) types (Figure 10). Over 50% of asteroids classified by DeMeo et al. (2009) are in the S-complex. Most of the Gaffey et al. (1993) S-subtypes are classified as S-types by DeMeo et al. (2009). The average visible albedo S-complex objects from NEOWISE observations is
0.22 (Mainzer et al., 2011d). The interpreted mineralogies of S-asteroids (e.g., Gaffey et al., 1993) range from olivine-dominated to olivine–pyroxene mixtures to pyroxene-dominated.

2.14.4.9.1 S-types The S-type designation of DeMeo et al. (2009) is the typical classification for objects with absorption features due to pyroxene and olivine. Approximately 40% of asteroids classified by DeMeo et al. (2009) are designated as S-types. In the visible (Bus and Binzel, 2002b), S-types have moderately strong UV features and moderately deep 1 mm bands. In the near-infrared, S-types have moderately strong 1 and 2 mm features. DeMeo et al. (2009) S-types contain members that range from the S(II) to S(VII) classes of Gaffey et al. (1993). The S(II)subtype were interpreted as olivine-dominated mineralogies with accessory high-Ca pyroxene since they fall off the olivine– low-Ca pyroxene mixing line. The S(III)-subtype plots outside the ordinary chondrite region and were interpreted as having a high-Ca pyroxene component with substantial olivine.

Asteroids

S(IV)-subtypes have Band Area Ratios and Band I centers consistent with ordinary chondrites, but other meteoritic assemblages with the same mineralogies cannot be ruled out. Besides ordinary chondrites, possible analogs (Gaffey et al., 1993) for the S(IV)-subtype include ureilites, acapulcoites/lodranites, and winonaites, which all contain mixtures of olivine and pyroxene. Ureilites, acapulcoites/lodranites, and winonaites are classified as primitive achondrites and are all believed to have experienced partial melting but did not crystallize from a melt (e.g., Weisberg et al., 2006). The S(V)-subtype fall off the olivine–low-Ca pyroxene mixing line and were interpreted as having a significant high-Ca pyroxene component with some olivine. The S(VI)-subtype are consistent with a low-Ca pyroxene–metallic iron assemblage. The S(VII)-subtype was inferred to have a low-Ca pyroxene–high-Ca pyroxene–metallic iron mineralogy. Gaffey et al. (1993) also envisioned that a number of assemblages not present in out meteorite collections, such as high-Ca pyroxene basalt intrusions into an H chondrite matrix, could also be present on S-asteroid surfaces. Using a simple heating model using 26Al and incorporating specific heat and heat of fusion, Gaffey (2006) argues that most asteroids would spend most of their heating time in the partially melted phase, and these partially melted objects should be abundant among main-belt asteroids, like the S-types. From the analysis of their reflectance spectra, proposed partially differentiated or differentiated S-types include (8) Flora (Gaffey, 1984) and (29) Amphitrite (Hiroi and Takeda, 1990). S-type (Bus and Binzel, 2002a) and S(IV)-asteroid (Gaffey et al., 1993) (6) Hebe has been postulated as the parent body of the H chondrites (Farinella et al., 1993b; Gaffey and Gilbert, 1998). S(IV)-types have Band Area Ratios and Band I centers consistent with ordinary chondrites. The Band Area Ratios and Band I centers of rotational spectra of Hebe are consistent (Gaffey and Gilbert, 1998) with H chondrites. Hebe is also located near the 3:1 mean motion and n6 secular resonances and is argued to be a major potential source of meteorites because it should be supplying relatively large number of fragments into meteorite-supplying fragments (Farinella et al., 1993a). Bottke et al. (2006b, 2010) find that the distribution of cosmic-ray exposure ages for H chondrites is consistent with derivation from Hebe. Rivkin et al. (2001) detected a weak 3 mm band on Hebe, consistent with aqueous alteration products that have been detected (Rubin et al., 2002) in some H chondrites. Hebe’s calculated bulk density (
3800 kg m3) (Baer et al., 2011) is consistent with maximum porosity of some measured H chondrites (
3800 kg m3; Britt and Consolmagno, 2003). H chondrites have been linked with the IIE irons, unusual irons that contain abundant silicate inclusions. Similarities between oxygen isotopic compositions (Clayton and Mayeda, 1978; McDermott et al., 2011), silicate mineralogies (Casanova et al., 1995), and the nonmagmatic (did not form in a core) origin of the metal in IIE irons (Wasson and Wang, 1986) argue for a possible relationship. Gaffey and Gilbert (1998) proposed that impacts melted the surface and caused it to differentiate into layers of silicates overlaying layers of metallic iron. These metallic iron sheets were exposed by later impacts. However, other researchers believe that asteroid sized impacts on the surface of Hebe would not produce enough melt, and it would solidify too quickly to produce

383

large sheets of metal (Kerr, 1996). Evidence that some type of metal segregation can occur can be found in the H chondrite Portales Valley meteorite (Rubin et al., 2001), an impact melt breccia with large regions of coarse metal with a Widmansta¨tten structure. Widmansta¨tten structure indicates slow cooling. Fieber-Beyer and Gaffey (2011) also identified another S(IV)-asteroid, (974) Lioba, as a plausible parent body for the LL or L chondrites from the mineralogy interpreted from its Band I and II centers and Band Area Ratio. Like Hebe, Lioba is also located near the 3:1 resonance.

2.14.4.9.2 Sa-types Sa-types have intermediate spectra between S- and A-types in the visible (Bus and Binzel, 2002b) and near-infrared (Figure 10; DeMeo et al., 2009). Sa-types have distinctive olivine bands but are less red than A-types. Sa-types are relatively rare with only two (<1%) of bodies classified by DeMeo et al. (2009) have this designation. Sa-types include (984) Gretia and the Mars Trojan (5261) Eureka. Rivkin et al. (2007) observed Eureka and found that it had a very broad 1 mm but no 2 mm band and resembled the spectra of angrites (Burbine et al., 2006).

2.14.4.9.3 Sq-types Sq-types have intermediate spectra between S- and Q-types (Figures 9 and 10). Sq-types are relatively common (Binzel et al., 2004b) among NEAs but relatively rare among main-belt objects. Binzel et al. (2004b) group the Sq- and Q-types into the Q-complex since Q-types are also relatively common among NEAs but also extremely rare among main-belt objects. Mothe´-Diniz and Nesvorny´ (2008a) observed members of a few young (compact) asteroid clusters (estimated ages of hundreds of thousands of years) in the visible and found that they contained members with Sq, Sk, and O/Q spectra. Mothe´-Diniz et al. (2010) found that a vast majority of Sq- and Sk-bodies had visible spectra that matched ordinary chondrites. Main-belt objects classified as Sq by DeMeo et al. (2009) include (3) Juno (Figure 21) and (11) Parthenope. Both of these objects were classified as S(IV)-types by Gaffey et al. (1993). Gaffey et al. (1993) and Gaffey (1995) identified Juno as being a leading candidate as an ordinary chondrite parent body. Baliunas et al. (2003) imaged Juno at a variety of wavelengths and identified a low brightness area at 0.934 mm that may represent a recent impact on the surface that may have ejected fragments off the surface. Parthenope was identified by Farinella et al. (1993a) as a body that may be supplying relatively large number of fragments into meteorite-supplying resonances. Near-Earth asteroid (99942) Apophis (
300 m in diameter) (Delbo` et al., 2007), which could potentially impact the Earth (e.g., Giorgini et al., 2008) in the near future, is an Sqtype (Binzel et al., 2009). The visible and near-infrared spectrum of Apophis is best matched by a space-weathered LL chondrite. Apophis’ close encounter with the Earth in 2029 may be able to test the Binzel et al. (2010) model that ‘seismic shaking’ can change the spectral properties on NEAs so they appear less space-weathered.

384

Asteroids

1.8

Normalized reflectance

Sq-type (3) Juno 1.6

1.4

S-type (6) Hebe

1.2 H5 chondrite Allegan 1

0.8

0

0.5

1

1.5 Wavelength (μm)

2

2.5

Figure 21 Plot of reflectance spectra of S-type (6) Hebe (squares) and (3) Juno versus spectrum of H5 chondrite Allegan (line) (particle size <150 mm). Both asteroids appear spectrally similar to a reddened ordinary chondrite. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Asteroid spectra are normalized to unity at 0.55 mm and then offset in reflectance. Error bars are 1s. 2

Sr-types have intermediate spectra between S- and R-types in the visible (Bus and Binzel, 2002b) and near-infrared (DeMeo et al., 2009) with a narrow 1 mm band that is similar to but more shallow than the bands found in R-types (Figure 10). Using MGM on visible and near-infrared spectra, Sunshine et al. (2004) found that a number of S- or Sr-types, including (17) Thetis and members of the (808) Merxia and (847) Agnia families, had silicate mineralogies consistent with low- and high-Ca pyroxene with minor amounts of olivine (<20%). Approximately 40% of the total pyroxene on their surfaces is high-Ca pyroxene, which is only consistent with bodies that have undergone melting since partial melting of chondritic material tends to produce melts enriched in low-Ca pyroxene.

2.14.4.9.5 Sv-types Sv-types have intermediate spectra between S- and V-types in the visible (Bus and Binzel, 2002b) and near-infrared (DeMeo et al., 2009). Sv-types have a very narrow 1 mm band that is similar to but more shallow than the band found in V-types (Figure 10). Sv-types are relatively rare with only 2 (<1%) of objects classified by DeMeo et al. (2009) have this designation. Members include (2965) Surikov and (4451) Grieve (which was given the Svw classification). Binzel and Xu (1990) initially interpreted Surikov as possible mantle material from a differentiated body due to visible spectra that indicated a substantial olivine component; however, the near-infrared spectrum of Surikov from DeMeo et al. (2009) indicates a substantial pyroxene component due to the strength of its 2 mm feature.

Normalized reflectance

2.14.4.9.4 Sr-types

C2-ungrouped chondrite Tagish Lake

1.8 1.6 1.4

T-Type (308) Polyxo

1.2 1 0.8 0.6

0

T-Types

In the visible (Bus and Binzel, 2002b; Tholen, 1984), T-types have moderately strong UV features shortward of 0.75 mm and then flatten out longward of 0.85 mm. In the near-infrared (DeMeo et al., 2009), T-types have slopes redder than X-types, but not as red as D-types (Figure 22). Three out of four observed T-types were found to have 3 mm bands (Rivkin et al., 2002).

1 1.5 Wavelength (μm)

2

2.5

Figure 22 Plot of reflectance spectrum of T-type (308) Polyxo versus spectrum of C2-ungrouped chondrite Tagish Lake (particle size <125 mm; Hiroi and Hasegawa, 2003). Spectra are normalized to unity at 0.55 mm. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. The feature at
0.9 mm in the SMASS II spectrum is not evident in the SpeX spectrum. Error bars are 1s.

Britt et al. (1992) noted the red near-infrared spectral slopes and low albedos of T-types were similar to troilite. Currently, the spectral properties of T-types asteroids are usually thought to be due to organic-rich silicates that would redden and darken the surface. Hiroi and Hasegawa (2003) noted the spectral similarity between T-type (308) Polyxo and the C2ungrouped chondrite Tagish Lake.

2.14.4.11 2.14.4.10

0.5

V-Types

In the visible (Binzel and Xu, 1993; Bus and Binzel, 2002b; McCord et al., 1970), V-types have relatively strong UV features and deep 1 mm bands that are similar to the spectra of HEDs (howardites, eucrites, diogenites; Figure 9). Mineralogically, eucrites contain primarily anorthitic plagioclase and low-Ca pyroxene with augite exsolution lamellae, while diogenites are predominately magnesian orthopyroxene. Howardites are approximately 50:50 mixtures of eucritic and diogenitic

Asteroids material (see Chapter 1.6). Near-infrared spectra (Birlan et al., 2011; Burbine et al., 2001b; De Sanctis et al., 2011; DeMeo et al., 2009; Kelley et al., 2003) of V-types tend to confirm their spectral similarity to HEDs with evidence of two strong symmetric pyroxene features (Figure 10). The average visible albedo of V-types from NEOWISE observations is
0.36 (Mainzer et al., 2011d). Asteroid (4) Vesta was the first asteroid (McCord et al., 1970) to be identified with this unusual type of visible spectrum that was similar to HEDs (Figure 23). Vesta’s spectral similarity in the near-infrared to HEDs was then confirmed by Larson and Fink (1975). Vesta has rotational spectral variations (e.g., Binzel et al., 1997; Carry et al., 2010; Gaffey, 1997) consistent with HED mineralogies. Hiroi et al. (1994) found that the best spectral match to Vesta in the visible and near-infrared was a fine-grained howardite. Spectra of Vesta in the UV into the visible (0.22–0.95 mm) (Li et al., 2011) also match a finegrained howardite. From rotational spectral variations, Gaffey (1997) predicted that a large olivine unit exists on Vesta’s equator. Hasegawa et al. (2003) observed a 3 mm band on Vesta; however, Vernazza et al. (2005) and Rivkin et al. (2006) found no evidence for such a 3 mm feature. Stubbs and Wang (2012) predict that water ice could be stable in Vesta’s regolith on timescales of billions of years. Vernazza et al. (2006b) find that Vesta’s surface does not appear space-weathered and argues that Vesta has a magnetic field that is protecting Vesta’s surface from charged particles. Fu et al. (2012) found evidence from paleomagnetic studies of a eucrite that the HED parent body had a surface magnetic field, which suggests that this object had an advecting liquid metallic iron core. Vesta is commonly assumed to be the HED parent body (e.g., Consolmagno and Drake, 1977) due its presence as the only large (
525 km diameter) body with an HED-like visible and near-infrared spectrum. Originally Vesta was thought to dynamically unable (Wetherill, 1987) to supply large number of fragments to meteorite-supply resonances; however, the discovery of small V-type bodies (commonly called Vestoids) (Binzel and Xu, 1993) in the Vesta family and also between Vesta and the 3:1 meteorite-supplying resonance on the basis of visible observations and the ability of Yarkovsky effect to move small bodies over distances of tenths of an AU in the asteroid

385

belt (e.g., Bottke et al., 2006b) argues that Vesta could supply fragments to these resonances. Numerous V-types have been identified in the Vesta family (e.g., Mothe´-Diniz et al., 2005), in the inner main belt but outside the Vesta family (e.g., Carruba et al., 2005; Florczak et al., 2002), and the near-Earth population (e.g., Burbine et al., 2009; Cruikshank et al., 1991b). Near-infrared of inner-belt V-types (Burbine et al., 2001b; Kelley et al., 2003; Mayne et al., 2011; Moskovitz et al., 2010) confirm their spectral similarity to eucrites and howardites (Figure 24). A number of these inner-belt Vestoids also have a feature at
0.506 mm (Vilas et al., 2000), which has been also seen in the spectra of HEDs (e.g., Hiroi et al., 2001a). This feature has been attributed to a spin-forbidden Fe2þ absorption bands that are commonly found in the spectra of low-Ca pyroxenes (Cloutis et al., 2010c). Using SDSS colors, Marchi et al. (2010) found that the visible spectral slopes of V-types tend to be larger than HEDs. They believe that this reddening could be due to space weathering or systematic compositional differences. Binzel and Xu (1993) also designated some objects with visible HED-like spectra but very deep 1 mm bands as J-types, which they interpreted as having diogenitic mineralogies based on just visible spectra. However, near-infrared observations (out to 1.65 mm) of a number of J-types (Burbine et al., 2001b) appeared not to be consistent with diogenite spectra. However, a number of V-types with near-infrared spectra consistent with diogenites have recently been identified in the inner belt (De Sanctis et al., 2011; Mayne et al., 2011). Also, the near-infrared spectrum of one near-Earth asteroid ((237442) 1999 TA10) does have band positions consistent with diogenites (Reddy et al., 2011b). Lim et al. (2011) found that V-type (956) Elisa had a Spitzer Infrared Telescope emissivity (5–35 mm) spectrum that was consistent with an olivine–diogenite-like mineralogy. Besides identifying a number of diogenite V-types, Mayne et al. (2011) also identified one V-type, (2511) Patterson, as having a ferroan (Fe-rich) diogenite composition. Due to V-types having absorption features in the visible and near-infrared that are almost entirely due to pyroxene minerals, which tend to have band center positions that are a

2

Normalized reflectance

V-type (4) Vesta

1.2 1.1 1 0.9

Howardite EET 87503

0.8 0.7

Normalized reflectance

1.8

1.3

Eucrite Bouvante

1.6

V-type (4055) Magellan

1.4 1.2 1 0.8

Howardite EET 87503

0.6 0.4 0

0.6 0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 23 Plot of reflectance spectrum of V-type (4) Vesta (squares) versus spectrum of howardite EET 87503 (line) (particle size <25 mm). Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

0.5

1

1.5

2

2.5

Wavelength (μm)

Figure 24 Plot of reflectance spectrum of V-type (4055) Magellan (squares) versus spectra of eucrite Bouvante (line) (particle size <25 mm) and howardite EET 87503 (line) (particle size <25 mm). Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Spectra are normalized to unity at 0.55 mm. Error bars are 1s.

386

Asteroids

function of mineralogy, numerous Vestoids have had their pyroxene mineralogies estimated using formulas derived by Gaffey et al. (2002) and Burbine et al. (2009). Because the number of pyroxene bands can be easily estimated, MGM can be used to estimate the relative abundances of low- and highCa pyroxene. For their observed main-belt V-types, De Sanctis et al. (2011) found that the estimated ferrosilite contents ranged from Fs22–24 to Fs41–43 and the estimated wollastonite contents ranged from Wo0.5–2 to Wo9–12. For V-type NEAs, Burbine et al. (2009) found that the estimated ferrosilite contents ranged from Fs36 to Fs51 and the estimated wollastonite contents ranged from Wo7 to Wo13. Mayne et al. (2011) found that some V-types could be modeled with just a low-Ca pyroxene (diogenites) and some with two pyroxenes (low- and high-Ca pyroxene for eucrites and howardites). For objects that contain two pyroxenes, the estimated high-Ca pyroxene/(low- and high-Ca pyroxene) ratio from the work of Mayne et al. (2011) for main-belt and near-Earth V-types range from 18% to 59% using Band I strengths and 17–48% using Band II strengths. However, Vesta appears not to be the only body that formed in the asteroid belt with a basaltic crust. The first identified object was outer-belt body (1459) Magnya (Hardersen et al., 2004; Lazzaro et al., 2000), which was located at 3.14 AU and observed to have an HED-like spectrum in the visible and near-infrared. Band positions derived for Magnya were interpreted (Hardersen et al., 2004) as indicating a pyroxene mineralogy that is slightly less Fe-rich than Vesta. Magnya is located relatively far from Vesta (2.36 AU) and far past the 3:1 resonance, which makes it dynamically difficult (Michtchenko et al., 2002) to derive Magnya from Vesta. Magnya’s diameter (17  1 km) (Delbo´ et al., 2006) is larger than inner-belt Vestoids, but Magnya’s geometric visible albedo (0.37  0.06) is similar to Vesta. A number of V-types bodies have been identified in the middle and outer main belt (e.g., Roig and Gil-Hutton, 2006) on the basis of SDSS colors. Asteroid (21238) Panarea (a ¼ 2.54 AU), which is located past the 3:1 resonance, was found to have a typical visible V-type spectrum (Binzel et al., 2006a; Roig et al., 2008) and band positions in the nearinfrared consistent with diogenites (De Sanctis et al., 2011; Moskovitz et al., 2010). Asteroid (40521) 1999 RL95 is another identified V-type body (Roig et al., 2008) that is also located past the 3:1 resonance at 2.54 AU. Panarea and 1999 RL95 could be fragments of Vesta that luckily drifted past the 3:1 resonance (Roig et al., 2008) or fragments of another differentiated body like (15) Eunomia (Carruba et al., 2007). Outer-belt V-types have been also identified on the basis of SDSS colors. Two with visible and near-infrared spectra (Duffard and Roig, 2009) are (7472) Kumakiri (a ¼ 3.02 AU) and (10537) 1991 RY16 (a ¼ 2.85 AU). Both objects have visible spectra similar to other V-types except for an absorption band at
0.65 mm, which was interpreted as possibly being evidence of chromium in the pyroxenes. However, their near-infrared spectra (Burbine et al., 2011; Moskovitz et al., 2008b) have been identified as unusual compared to other V-types. Kumakiri has 1 and 2 mm bands that appear consistent with pyroxenes (Burbine et al., 2011), but the calculated band positions are offset from the values found for HEDs. Possible explanations for this anomalous behavior is that Kumakiri contains an unusual type of pyroxene not typically found in meteorites, or this object is a mixture of olivine with high-Ca pyroxene and

minor low-Ca pyroxene. No known meteorite has a spectrum similar to that of Kumakiri. Asteroid (10537) 1991 RY16 has a near-infrared spectrum (Moskovitz et al., 2008b) consistent with a substantial olivine component, which is unlike the spectral properties of other observed V-types. Consistent with the existence of V-types outside the inner man belt, a number of basaltic meteorites (Bland et al., 2009; Gounelle et al., 2009; Scott et al., 2009; Yamaguchi et al., 2002) have been found to have oxygen isotopic compositions distinct from ‘typical’ HEDs. Scott et al. (2009) argue that these ‘anomalous’ eucrites are evidence for five distinct Vesta-like parent bodies besides Vesta. Also consistent is the large number of iron meteorite groups, grouplets, and ungrouped irons that are found in our meteorite collections. Wasson (1995) argues that these irons are evidence of
70 disrupted differentiated parent bodies. Even though the Vesta–HED linkage is widely accepted, Consolmagno et al. (2011) note that this implied relationship is based on a series of assumptions and could just be a coincidence. Not all scientific analyses support this linkage. For example using magnesium isotopic analyses of diogenites, Schiller et al. (2011) found evidence for rapid (
2–3 Ma timescale) cooling and magma ocean crystallization, which could only occur on bodies smaller than
100 km in diameter according to recent thermal models. These bodies are much smaller than Vesta.

2.14.4.12

X-Complex (E-Types, M-Types, and P-Types)

In the visible (Bus and Binzel, 2002b; Tholen, 1984), X-complex (or X-group) asteroids tend to have relatively featureless spectra with a slight to moderately red spectral slope (Figure 9). Tholen (1984) used visual albedo for X-types to differentiate between E-types (albedos of >0.3), M-types (
0.1–0.3), and P-types (<0.1). E-types are usually linked with aubrites, M-types with metallic iron, and P with organicrich material due to these albedo differences. Bus and Binzel (2002b) did not use albedo in their taxonomy and instead broke the X-group into X-, Xc-, Xe-, and Xk-types. The Xe-types have an absorption band shortward of 0.55 mm, while the Xc-types have visible spectra intermediate between X- and C-types, and the Xk-types are intermediate between X- and K-types. In the near-infrared (Clark et al., 2004b; DeMeo et al., 2009), X-complex (or X-group) asteroids have a flat but slightly reddish slope (Figure 25).

2.14.4.12.1 E-types The E-types have been typically linked with the aubrites (e.g., Zellner, 1975; Zellner et al., 1977) since both types of objects have high albedos and relatively flat spectra longward of 0.55 mm. The aubrites are achondritic enstatite-rich meteorites (Watters and Prinz, 1979). Enstatite is a FeO-poor pyroxene that is whitish in appearance. Among meteorites (Gaffey, 1976), only aubrites have such high albedos and no prominent 1 and 2 mm absorption bands. The average visible albedo of E-types from NEOWISE observations is
0.43 (Mainzer et al., 2011d). Approximately 60% of identified E-types (Clark et al., 2004a) are found in the Hungaria region. Four of six observed E-types were found to have 3 mm absorptions (Rivkin, 1997; Rivkin et al., 1995, 2002), but no objects were observed in the

Asteroids

Normalized reflectance

2.4 2.2 P-type (153) Hilda 2 1.8 Iron meteorite Odessa

1.6

M-type (16) Psyche E-type (64) Angelina

1.4 ~0.5 μm

1.2 1

Aubrite Mayo Belwa

0.8 0

0.5

1

1.5

2

2.5

Wavelength (μm) Figure 25 Plot of reflectance spectra of E-type (64) Angelina (squares) (Bus, 2011; Clark et al., 2004b), M-type (16) Psyche (squares), and P-type (153) Hilda (squares) versus spectra of aubrite Mayo Belwa (line) (particle size <125 mm) and iron meteorite Odessa (unsorted particle size) (line) (Cloutis et al., 1990). Feature at
0.5 mm in the Angelina spectrum is identified. Asteroid data are a combination of SMASS II (Bus and Binzel, 2002a) and SpeX spectra. Spectra are normalized to unity at 0.55 mm and then offset in reflectance. The Odessa spectrum is offset so to directly compare with the Psyche spectrum. Error bars are 1s.

Hungaria region. If the observed E-types are igneous bodies, the presence of a 3 mm band due to hydrated silicates is surprising because these objects would not be expected to have undergone aqueous alteration. E-types do not have entirely featureless visible and nearinfrared spectra. For example, a number of weak bands have been found in the reflectance spectra of E-types. A high percentage (Clark et al., 2004a) of E-types were classified as Xe by Bus and Binzel (2002b) due the presence of a 0.55 mm feature (Bus and Binzel, 2002a) that was not present in the lower resolution ECAS data. Xe-types include (64) Angelina, (77) Frigga, (132) Aethra, (317) Roxane, (434) Hungaria, (1025) Riema, (2035) Stearns, and (3103) Eger. Angelina was noted by Kelley and Gaffey (2002) to have a number of weak bands in the visible and near-infrared, which they argue were due to accessory phases mixed with the dominant mineral enstatite. Near-Earth asteroid 3103 Eger, classified as an E-type by Veeder et al. (1989a) and an Xe-type by Binzel et al. (2004b), was postulated by Gaffey et al. (1992) to be the source body of the aubrites and also to be derived from the Hungaria region. Another NEA (4660 Nereus) is also an Xe-type (Binzel et al., 2004a) and could potentially be a possible source body for aubrites. Nereus has an extremely high visual albedo of
0.55 (Delbo´ et al., 2003). The Venus quasi-satellite (2002 VE68) is an X-type (Hicks et al., 2010) that may have a feature at 0.55 mm. Like E-types, this object also has a high circular polarization ratio (Benner et al., 2008) from radar observations. The
0.5 mm absorption band found in the spectra of Xe-types (Figure 25) was argued to be due to oldhamite (CaS) by Burbine et al. (2002a), a mineral (Watters and Prinz, 1979) commonly found in aubrites and enstatite chondrites but usually in very low abundances. Aubrites have much higher visual albedos (
0.42–0.48) (excluding the chondritic inclusion-rich Cumberland Falls) (Cloutis and Gaffey, 1993) than enstatite chondrites (
0.05–0.18) since enstatite chondrites have significant abundances of metallic iron and

387

opaques, which reduce the visual albedo. Due to these low concentrations in aubrites, Shestopalov et al. (2010) propose that the feature is instead due to Ti3þ crystal field bands or Fe2þ to Ti4þ charge transfer transitions in pyroxenes. However, oldhamite weathers very easily on Earth when exposed to atmospheric moisture (Okada et al., 1981) so the abundances found in meteorites may not be representative of the abundances on the surfaces of enstatite-rich asteroids. Asteroid (44) Nysa, classified as an E-type by Zellner et al. (1977) and Tholen (1984) and as an Xc-type by Bus and Binzel (2002a), has long been known to have weak bands at
0.9 and
1.8 mm, which have been attributed to a low-FeO pyroxene (Clark et al., 2004a; Gaffey et al., 1989). Cloutis and Gaffey (1993) argued that these weak bands could be due to carbonaceous chondritic inclusions that are found in some aubrites. Nysa’s 3 mm band (Rivkin et al., 1995) could be consistent with carbonaceous chondritic inclusions. Radar albedos of Nysa (Shepard et al., 2008) are higher than the average values for main-belt asteroids (Magri et al., 2007), indicating a high surface bulk density. This high surface density could be due to a low bulk porosity or a high metal content. Rosetta flyby target (2867) Sˇteins was well studied before the encounter. Sˇteins was classified as an E(II) object (Dotto et al., 2009; Weissman et al., 2008), indicating a 0.55 mm band. Ground-based observations of Sˇteins are consistent with an aubrite mineralogy (e.g., Nedelcu et al., 2007; Weissman et al., 2008) due to its high albedo and relatively featureless spectrum (except for the 0.55 mm absorption feature).

2.14.4.12.2 M-types M-type asteroids have been historically linked with iron meteorites due to both types of objects having relatively flat spectra in the visible and near-infrared and moderate visual albedos; however, enstatite chondrites are also known (e.g., Burbine et al., 2002a) to be spectrally similar to M-types and have similar albedos. The average visible albedo of M-types from NEOWISE observations is
0.13 (Mainzer et al., 2011d). A number of observed M-type asteroids ((16) Psyche, (69) Hesperia, (129) Antigone, (216) Kleopatra, (347) Pariana, (758) Mancunia, (779) Nina, (785) Zwetana) have been found to have high radar albedos (Shepard et al., 2010, 2011) that appear to almost certainly indicate metallic iron-rich surfaces. Shepard et al. (2010) call these Mm-types. However, most of these asteroids (e.g., Fornasier et al., 2010, 2011a; Hardersen et al., 2005, 2011; Ockert-Bell et al., 2010) have absorption features at
0.9 mm, which are consistent with the presence of Fe-bearing silicates. Birlan et al. (2007) observed a number of different M-types in the near-infrared and found them to be featureless with spectral slopes consistent with metallic iron surfaces. Rivkin et al. (1995) found that a number of M-asteroids had 3 mm absorptions and labeled them W-asteroids. Approximately 40% of observed M-types had 3 mm bands (Rivkin et al., 2002), which is consistent with hydrated silicates on the surface and not consistent with metallic iron-dominated assemblages. Psyche, classified as an M-type by Tholen (1984) and as an Xk-type by DeMeo et al. (2009), is one of the most studied asteroids due to its large diameter (
250 km) (Tedesco et al., 2002b). Psyche has long been known to have an extremely high radar albedo (0.42  0.10) (Ostro et al., 1985; Shepard

388

Asteroids

et al., 2010), which appears consistent with a metallic iron surface. A feature at
0.9 mm (Hardersen et al., 2005) in the spectra of Psyche is consistent with Fe-bearing silicates. Binzel et al. (1995) did not find any significant variation in spectral properties in the visible versus rotation for Psyche, implying that there are no significant surface variations in silicate abundances. Rivkin et al. (1995, 2000) did not identify a 3 mm band in Psyche’s spectrum, which is consistent with a metallic ironrich surface. Psyche’s estimated bulk density,
6700 kg m3 (Baer et al., 2011), is consistent with a metallic iron mineralogy with some porosity. Kleopatra, classified as an M-type by Tholen (1984) and as an Xe-type by Bus and Binzel (2002b) and DeMeo et al. (2009), has the highest radar albedo (0.60  0.15) (Ostro et al., 2000; Shepard et al., 2010) of observed main-belt asteroids. The 0.55 mm feature is very weak in the spectrum of Kleopatra (Bus and Binzel, 2002a). Kleopatra has an estimated bulk density of 3600  400 kg m3 (Descamps et al., 2011), which is consistent with an iron meteorite assemblage with a bulk density of
30–50% or an enstatite chondrite assemblage (
3600–3700 kg m3) (Britt and Consolmagno, 2003) with little to no porosity. Rivkin et al. (1995, 2000) did not identify a 3 mm band in Kleopatra’s spectrum, which is consistent with a metallic iron-rich surface. Rivkin et al. (1995, 2000) identified 3 mm bands in the spectra of M-types (21) Lutetia, (22) Kalliope, (55) Pandora, (77) Frigga, (92) Undina, (110) Lydia, (129) Antigone, (135) Hertha, and (201) Penelope. Shepard et al. (2010) identified Antigone as having a high radar albedo (0.36 0.09), while Lutetia (0.24  0.07), Kalliope (0.18  0.05), Lydia (0.20  0.12), and Hertha (0.18  0.05) have much lower radar albedos. Shepard et al. (2010) believe that the best meteoritic analog for Antigone is possibly the metal-rich CB chondrites, and the best analogs for Lutetia, Kalliope, Lydia, and Hertha are either the metal-rich CH or enstatite chondrites. If these features are due to hydrated silicates, CB- and Ch-like chondrites would appear to be better analogs since these meteorite types are known to contain much higher percentages of hydrated silicate components (e.g., Bischoff, 1998; Ivanova et al., 2008) than enstatite chondrites. Rosetta flyby target (21) Lutetia was well studied before the encounter. Lutetia was classified as an Xk by Bus and Binzel (2002b) and as an Xc by DeMeo et al. (2009), which indicates a relatively featureless spectrum in the visible and near-infrared. Ground-based spectral observations of Lutetia have typically been interpreted as indicating either a carbonaceous chondrite (e.g., Barucci et al., 2005) or enstatite chondrite (e.g., Vernazza et al., 2009b) mineralogy. Ground-based observations of Lutetia (Rivkin et al., 2011a) show a relatively weak 3 mm feature (strength of 3–5%) for the southern hemisphere and a weaker band for the northern hemisphere. It is unclear if the presence of this 3 mm band is consistent with an enstatite chondrite composition or if this feature indicates a very different mineralogy for Lutetia. Ground-based density calculations of Lutetia give values of 3500  1100 or 4300  800 kg m3 (Drummond et al., 2010), with the lower value only consistent with an enstatite chondrite composition (
3200–4100 kg m3) (Macke et al., 2010) or high-density carbonaceous chondrites (CH, CK) (
3400–3500 kg m3) (Britt and Consolmagno, 2003) with little to no porosity.

2.14.4.12.3 P-types P-types are commonly found in the Hilda group and among Jupiter Trojans. About 1/4 of observed Trojans (Roig et al., 2008) have spectral slopes compatible with P-type asteroids. Most P-types do not have a 3 mm band (Rivkin et al., 2002). The average visible albedo of Tholen (1984) P-types from NEOWISE observations is
0.04 (Mainzer et al., 2011d). The spectral properties of P-types asteroids are usually thought to be due to organics, which would redden and darken the surface. Hiroi et al. (2004) were able to duplicate the spectral properties of P-types by spectrally mixing CI/CM and Tagish Lake (C2-ungrouped) material that went through thermal metamorphism and/or space weathering. Schaefer et al. (2010) found that the opposition surges calculated from the phase curves (brightness versus phase angle) of Jupiter Trojans were consistent with the opposition surges of outer solar system bodies that have lost their ices (e.g., Centaurs, extinct comets) and not with main-belt C- and P-type asteroids, implying that Jupiter Trojans and main-belt C- and P-types have different surface compositions and/or textures. Main-belt P-type (65) Cybele (a ¼ 3.43 AU), which was classified as an Xc by Bus and Binzel (2002b) and an Xk by DeMeo et al. (2009), was found to have an absorption featured at
3.1 mm (Licandro et al., 2011a) that is consistent with water ice. Cybele has a visual albedo of 0.050  0.005 (Mu¨ller and Blommaert, 2004). This feature is similar to that found in the spectrum of Themis (Campins et al., 2010a; Rivkin and Emery, 2010). Previous observations of Cybele (Jones et al., 1990) did not find a 3 mm band. Due to the absence of detectable CN emission lines, which is used as tracer of water in comets, and assuming cometary H2O/CN mixing ratios, an upper limit for surface water ice was determined for Cybele by Jewitt and Guilbert-Lepoutre (2012) to be less than
10%. Other absorption features for Cybele between 3.2 and 3.6 mm appear consistent with organics (Licandro et al., 2011a) with no features detected that were consistent with hydrated silicates. The emissivity spectrum of Cybele (Licandro et al., 2011a) in the 5–14 mm range is similar to that found for Trojan asteroids (Emery et al., 2006). Main-belt object (50) Virginia is the only other observed P-type with a 3 mm band (Rivkin et al., 1995, 2002). Cybele has an estimated bulk density of
1000 kg m3 (Baer et al., 2011), which is consistent with a very porous material.

2.14.5 2.14.5.1

Spacecraft Missions Galileo

Galileo was the first spacecraft to fly by an asteroid. Galileo flew by S-asteroid (Tholen, 1984) (951) Gaspra (Belton et al., 1992) in October of 1991 and then S-asteroid (243) Ida (Belton et al., 1994) plus its satellite Dactyl, which was discovered during the encounter (Chapman et al., 1994), in August of 1993. Gaspra is a member of the Flora family, while Ida is a member of the Koronis family. Galileo imaged and obtained reflectance spectra of all three objects. Galileo used eight filters (0.40–0.99 mm) in the visible and a spectrometer from 0.7 to 5.2 mm to obtain reflectance spectra. Gaspra is a potato-shaped object with numerous craters (Figure 26). Ida is also irregularly shaped (Figure 27) but is

Asteroids

Figure 26 Picture of asteroid S-type (951) Gaspra, which is a mosaic of two images taken by the Galileo spacecraft. Image credit: NASA and JPL.

differences on the surface with the most spectrally distinct areas having stronger Band Is, higher albedos, and bluer spectral slopes than the average spectra for Gaspra. These spectrally distinct areas were often associated with small, fresh-appearing craters along ridges. This correlation is consistent with space weathering on the surface of Gaspra where older material tends to move downslope from the ridges (Belton et al., 1992; Chapman, 1996; Veverka et al., 1994). Evidence for space weathering is more prominent for Ida (Chapman, 1996). Two terrains (Veverka et al., 1996) were identified on Ida, which were called Terrain A (which covers most of Ida) and Terrain B. Terrain A has a shallower 1 mm absorption and a steeper red slope than Terrain B. Terrain B is predominately associated with well-defined small craters and the large crater Azzurra and its ejecta. Terrain B is associated with fresher, less space-weathered material, while Terrain A corresponds with the older, more space-weathered regolith. Veverka et al. (1996) proposed that Dactyl’s visible spectrum differed from Ida’s terrains by having a deeper 1 mm band and a redder slope, which is not the trend that would be expected using a typical space weathering model for material with an Ida-like mineralogy. However, Chapman (1996) argued that the spectral differences between Ida and Dactyl were relatively minor and no inference can be made concerning any inconsistency with the ‘typical’ space weathering model.

2.14.5.2

Figure 27 Image of asteroid (243) Ida and its moon Dactyl. Image credit: NASA and JPL.

more heavily cratered than Gaspra. Gaspra and Ida were also found to have relatively strong magnetic fields (Kivelson et al., 1993a,b). Kivelson et al. (1993a) argued that the Gaspra results were consistent with iron meteorites and magnetized chondrites, while Hood (1994) believed that the Gaspra results were consist with all meteorite types. Granahan (2011) found that Band Area Ratios and Band I centers derived from Gaspra’s visible and near-infrared reflectance spectra taken by Galileo indicated a body that was more olivine-rich than ordinary chondrites, indicating that Gaspra underwent some igneous differentiation. Granahan (2011) also saw evidence for a 2.7 mm feature that he argues is possibly due to OH on the surface. Analysis of visible reflectance spectra of Gaspra over its surface (Belton et al., 1992; Veverka et al., 1994) found subtle

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Deep Space 1

Deep Space 1 flew by Mars-crosser (9969) Braille in July of 1999. Deep Space 1 was the first spacecraft to use ion propulsion. Deep Space 1 obtained two images (Oberst et al., 2001) and three near-infrared (
1.2–2.6 mm) spectra of Braille (Buratti et al., 2004). A UV spectrometer failed before the encounter. The relatively fuzzy images combined with light curve data were consistent (Oberst et al., 2001) with a body with dimensions 2.1  1  1 km. The near-infrared spectra combined with ground-based visible data (Binzel et al., 2001a; Lazzarin et al., 2001) appear consistent with Q-type spectra. The geometric visual albedo (0.34  0.03) is consistent with a surface that has not been space-weathered to any degree. A magnetic field was also measured (Richter et al., 2001) around Braille, consistent with the magnetization found for meteorites (Fuller, 2007).

2.14.5.3

NEAR Shoemaker

The Near-Earth Asteroid Rendezvous-Shoemaker (NEAR Shoemaker) (Cheng, 2002) was the first spacecraft to orbit and then land on an asteroid, (433) Eros. NEAR Shoemaker first flew by main-belt C-asteroid (253) Mathilde in June of 1997 (Veverka et al., 1997). Mathilde has an irregular shape (66  48  46 km; Figure 28), is heavily cratered, and is very dark (visual geometric albedo between 0.035 and 0.050). The Mathilde has a very low density (1300  200 kg m3; Veverka et al., 1997; Yeomans et al., 1997), implying a very porous interior. Ground-based spectra of Kelley et al. (2007) found that Mathilde is a good spectral match to carbonaceous chondrites dominated by phyllosilicate minerals and argue that Mathilde has undergone an alteration process similar to CI1 or CM2 chondrites.

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Asteroids

NEAR Shoemaker orbited S-asteroid Eros from February of 2000 to February of 2001. Although the spacecraft was not designed to land, an intentional ‘crash’ landing at the end of the mission still resulted in the spacecraft landing intact due to Eros’ low gravity with its antenna still capable of sending data back to Earth and all sensors operational. Eros is a heavily cratered peanut-shaped object (Figure 29). Before the encounter, Eros was originally thought to have significant rotational spectral variations (Murchie and Pieters, 1996) and possibly have a surface mineralogy similar to the acapulcoites/lodranites (McCoy et al., 2000); however, Eros was found to be relatively homogenous spectrally with only modest spectral variability (e.g., Izenberg et al., 2003; Sullivan et al., 2002). The calculated average Band Area Ratio and Band I center for Eros (e.g., Izenberg et al., 2003) are consistent with ordinary chondrites (Figure 30). The brightest (highest albedo) areas on Eros are steep crater walls, which have lesser spectral slopes and deeper 1 mm bands. This trend is consistent with the exposure on the steep crater walls with material that is less space-weathered. The surface mineralogy of Eros was estimated using reflectance spectra, while its surface chemistry was estimated using the first x-ray and gamma-ray measurements of an asteroid. x-ray fluorescence is due to the absorption of x-rays or gamma rays by inner orbital electrons of atoms, which causes these

Figure 28 This picture of asteroid (253) Mathilde is constructed from four images acquired by NEAR Shoemaker. Image credit: NASA.

electrons to be expelled from these orbitals. As electrons from higher orbitals fill these electron ‘holes,’ x-rays are emitted. The strongest lines are those from the transition from the L orbital to the K orbital, which are called Ka lines. Since each element has orbitals with characteristic energies, the energies of the emitted x-rays are characteristic of different elements. The number of emitted characteristic x-rays is related to the concentration of the element on the surface. For asteroids, the x-ray source is the solar corona and the sampling depth is less than 100 mm. For Eros, the elements that were observed by x-ray fluorescence (Foley et al., 2006; Lim and Nittler, 2009; Nittler et al., 2001; Trombka et al., 2000) were magnesium, aluminum, silicon, sulfur, calcium, chromium, manganese, iron, and nickel. The newest (Foley et al., 2006; Lim and Nittler, 2009) calibrated bulk Mg/Si, Al/Si, Ca/Si, Fe/Si, Cr/Fe, Mn/Fe, and Ni/Fe elemental abundance ratios derived from Eros were consistent with ordinary chondrites, but S/Si was depleted relative to ordinary chondrites. Figure 31 plots the Mg/Si versus Al/Si and Fe/Si ratio for Eros versus values for ordinary chondrites and a number of other meteorite types. The best analog for Eros using these elemental ratios are H chondrites. The sulfur depletion is believed to be due to space weathering where troilite is vaporized (Loeffler et al., 2008; Trombka et al., 2000) through space weathering and sulfur is lost into space. Gamma rays from an asteroid (Evans et al., 2001) are emitted from the radioactive decay of very long-lived isotopes (e.g., 40K, 232Th, and 238U) or through the excitation of nuclei by galactic cosmic rays or solar particles. The galactic cosmic rays or solar particles strike nuclei, which emit neutrons that collide with other nuclei (Prettyman, 2007). These nuclei become excited and emit gamma rays when they return to their normal energy state. In orbit it was difficult to accumulate enough gamma-ray counts from Eros to detect any elements but when NEAR Shoemaker landed on the surface, numerous elements (magnesium, oxygen, silicon, potassium, and iron) were detected. The calculated Mg/Si and Si/O elemental abundance ratios and K abundances were consistent with chondrites and the Mg/Si ratio was consistent with the x-ray results. However, the Fe/Si and Fe/O ratios were low relative to chondrites 1.15

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Figure 29 This picture of (433) Eros is a mosaic of four images obtained by NEAR Shoemaker immediately after the spacecraft’s insertion into orbit. Image credit: NASA.

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Figure 30 Location of average Eros (red square) (Izenberg et al., 2003) on a plot of band area ratio versus band I center. Green line is olivine– orthopyroxene mixing line. Eros plots within the ordinary chondrite (S(IV)) region. HED region (rectangle) is also plotted. Error bars are 1s.

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Figure 31 Plot of calculated average Mg/Si versus average Al/Si and average Fe/Si (only for Eros) ratios derived from x-ray analyses of (433) Eros (Lim and Nittler, 2009) and (25143) Itokawa (Okada et al., 2006) versus data for H, L, and LL chondrites plus acapulcoites/lodranites, HEDs (howardites, eucrites, and diogenites), and carbonaceous chondrites (CI, CM, CO, and CV) (Nittler et al., 2004). Error bars are 1s.

and the Fe/Si ratio was low compared to the x-ray results. Possible reasons (Evans et al., 2001) for the nonchondritic values for the Fe/Si and Fe/O ratios are errors in the calculations, Eros is partially differentiated, or processes at the landing site have depleted iron. The magnetization of Eros (Acun˜a et al., 2002) was found to be much lower than values found for Gaspra and Braille and also lower than those found for ordinary chondrites. Wasilewski et al. (2002) argue that large ordinary chondrite bodies would have lower magnetic field intensities than meteorites since large bodies would have a juxtaposition of regions of randomly oriented material, which would lower the measured intensity. Using all available data, McCoy et al. (2001) found that the interpreted mineralogy and chemistry of Eros was best matched by an ordinary chondrite that was altered at the surface by space weathering or a primitive achondrite derived from material mineralogically similar to ordinary chondrites.

2.14.5.4

Stardust

In November of 2002, the Stardust spacecraft flew by S-type (Duxbury et al., 2004) (5535) Annefrank to practice for the later flyby of comet 81P/Wild (commonly called Wild 2). The asteroid was shaped like a triangular prism with dimensions of 6.6  5.0  3.4 km (Duxbury et al., 2004). No reflectance spectra were obtained for Annefrank, making any mineralogical analysis impossible. However, the geometric visual albedo (0.28  0.09 for all the phase angle data and 0.23  0.04 for phase angle data less than 90 ; Hillier et al., 2011) was measured, and it was consistent with S-types.

2.14.5.5

Hayabusa

Hayabusa was the first spacecraft to return samples of an asteroid back to Earth. Hayabusa rendezvoused with S-asteroid (25143) Itokawa (535  294  209 m) in September of 2005, landed twice briefly on the object to obtain samples, and then returned back to Earth where it landed in Australia in June of 2010. The sampler did not operate as planned; however, as the sampling

Figure 32 Image of (25143) Itokawa taken by Hayabusa. Very rocky surface areas and very smooth areas are evident. Image credit: ISAS and JAXA.

horn landed on the surface, small particles were ejected into sampler. At least 1500 grains of extraterrestrial origin were present in the sampler container (Nagao et al., 2011). Images of Itokawa (Saito et al., 2006) show an object shaped like a sea otter with both rugged (
80% of surface) and smooth terrains (Figure 32). Numerous boulders exist on the rough terrains. Less than 100 craters appear on the surface. Itokawa was an optimum target for sample return since its Earth-based visible and near-infrared reflectance spectrum resembled a space-weathered LL chondrite (Figure 7; Binzel et al., 2001b). Binzel et al. (2001b) analyzed the spectrum using a simple space weathering model and MGM. However, the Itokawa spectrum was also interpreted as indicating a partially melted surface (Abell et al., 2007) on the basis of its interpreted mineralogy (Fs435Wo145) that was derived from using the Gaffey et al. (2002) formulas with its calculated Band I and II centers. This mineralogy is much more iron-rich than the average pyroxene mineralogy of ordinary chondrites. Hayabusa also obtained reflectance spectra (0.76–2.1 mm) and did x-ray spectroscopy of the surface. The spectral shape of the 1 mm feature from Hayabusa observations were consistent with an LL chondrite (Abe et al., 2006). The average Mg/Si and

392

Asteroids

Al/Si ratios elemental abundance ratios (Okada et al., 2006) derived from x-ray analyses of Itokawa were similar to Eros and consistent with ordinary chondrites (Figure 31). The best ordinary chondrite matches to Itokawa are LL and L chondrites. The returned Itokawa particles were consistent with an LL4 to LL6 chondrite (Nakamura et al., 2011). The most abundant mineral in the fragments was olivine with a mineralogy of Fa28.61.1. Low-Ca pyroxene had a composition of Fs23.12.2Wo1.81.7, high-Ca pyroxene had a composition of Fs8.91.6Wo43.54.5, and plagioclase had a composition of Ab83.91.3Or5.51.2. FeNi inclusions within the silicates were kamacite (3.8–4.2 wt% for Ni, 9.4–9.9 wt% for Co) and taenite (42–52 wt% for Ni, 2.0–2.5 wt% for Co). Olivine, lowCa pyroxene, and kamacite compositions of most Itokawa particles were consistent with LL chondrites. Most particles were highly equilibrated with a narrow range of compositions that were consistent with LL5 and LL6 chondrites, while the unequilibrated fragments were consistent with LL4 chondrites. Tsuchiyama et al. (2011) found that the modal abundances calculated from their analyzed 40 particles (64% olivine, 19% low-Ca pyroxene, 3% high-Ca pyroxene, 11% plagioclase, 2% troilite,
0.02% kamacite,
0.2% taenite,
0.1% chromite,
0.01% Ca phosphate) were roughly similar to LL chondrites except for a slight depletion in troilite and metallic iron. Oxygen isotopic analyses of the returned fragments were consistent with either L or LL chondrites (Yurimoto et al., 2011). Noguchi et al. (2011) found that five of ten analyzed particles had evidence for space weathering with sulfur-bearing Ferich nanoparticles found in a thin surface layer (5–15 nm) on the olivine, low-Ca pyroxene, and plagioclase feldspar and sulfur-free Fe-rich nanoparticles were found deeper (<60 nm) in the ferromagnesian silicates. Noguchi et al. (2011) believe that two space weathering processes are at work with the surface layer of sulfur-bearing Fe-rich nanoparticles being the result of vapor deposition and the deeper layer of sulfur-free Fe nanoparticles being the result of solar wind irradiation. Nagao et al. (2011) measured large amounts of solar helium, neon, and argon at various depths in the Itokawa grains, indicating exposure to the solar wind. Using concentrations of cosmic-ray-produced 21Ne in the grains, Nagao et al. (2011) estimate an upper limit on the cosmic-ray exposure age of 8 Ma for Itokawa’s surface. This age suggest that Itokawa is continuously losing surface material into space at a rate of tens of centimeters per million years. Hayabusa unequivocally showed that at least one S-asteroid has an ordinary chondrite composition, that space weathering occurs on asteroid surfaces, and that it is possible to determine accurate asteroid mineralogies from ground-based observations. Binzel et al. (2001b) accurately predicted the mineralogy (LL chondrite) of Itokawa, while Abell et al. (2007) did not using the Gaffey et al. (2002) formulas. Gaffey (2007) has revised his method so as to better calculate the ferrosilite contents from ordinary chondrite spectra.

2.14.5.6

shape with an average diameter of
5.3 km (Keller et al., 2010). Sˇteins has a number of linear faults, a large
2.1 km crater near its south pole, and very few small craters. Lutetia (121  101  75 km) (Sierks et al., 2011) is an elongated body that is heavily cratered with a large bowl-shaped depression (
55 diameter Massilia crater in the Narbonensis region; Figure 33). Five separate regions have been identified on Lutetia’s surface using crater densities, crosscutting and overlapping relationships, and geologic features (faults, fractures, grooves). Sˇteins’ integrated geometric albedo (0.41  0.016 at 0.632 mm) from Rosetta is consistent with the high visual albedos of other E-type asteroids in the visible. The measured remnant magnetization of Sˇteins (Auster et al., 2010) was not found to be significant and appears consistent with aubrites. Lutetia’s visual geometric albedo is 0.19  0.01 from Rosetta. Ultraviolet observations (Stern et al., 2011) of Lutetia by Rosetta find a steep decline in reflectance from 1.85 to 1.60 mm, which they interpret as a strong solid state absorption feature in the surface. Rosetta’s visible and near-infrared spectra of Lutetia are relatively featureless from 0.5 to 2.5 mm. Using a combination of UV, visible, and near-infrared data, Stern et al. (2011) and Vernazza et al. (2012) find the best spectral fit to Lutetia to be enstatite chondrites. Barucci et al. (2012) argue that spectral variations on Lutetia seen by Rosetta were consistent with carbonaceous and enstatite chondrite areas on the surface, which would have formed from collisions of carbonaceous chondrite and enstatite chondrite asteroids. One meteorite (Kaidun) has an interpreted mineralogy similar to this postulated surface composition of Lutetia since Kaidun is a mixture of enstatite and carbonaceous chondrite material (e.g., Zolensky and Ivanov, 2003). Lutetia’s bulk density as measured by Rosetta was 3400  00 kg m3 (Pa¨tzold et al., 2011), which is higher than most measured carbonaceous chondrites and similar to enstatite chondrites. The density is also high than most measured asteroids (e.g., Britt et al., 2002).

Rosetta

The Rosetta spacecraft had a few hour flyby of E-type (2867) Sˇteins in September of 2008 (Keller et al., 2010) and of M-type (21) Lutetia in July of 2010, while on its way to rendezvous with comet 67P/Churyumov–Gerasimenko. Sˇteins has a diamond

Figure 33 Image of (21) Lutetia at closest approach taken by Rosetta. Image credit: ESA.

Asteroids

2.14.5.7

Dawn

The Dawn spacecraft orbited Vesta from July 2011 until September 2012 and is scheduled to reach Ceres in February 2015 (Russell et al., 2004, 2007). Dawn has a framing camera (with one clear filter and seven narrow band filters), visible and near-infrared spectrometer with a total coverage of 0.25 and 5.0 mm, and a gamma-ray/neutron detector. Gamma rays emitted by Vesta include those from naturally decaying radioactive elements (e.g., thorium, uranium, and potassium) and excited nuclei (e.g., iron, magnesium, silicon, aluminum, titanium, oxygen, and calcium; Prettyman, 2007). The energies of neutrons emitted from the surface due to incoming galactic cosmic rays and solar particles are sensitive to light elements such as hydrogen, carbon, and nitrogen and neutron absorbers such as gadolinium and samarium. Images of Vesta (Jaumann et al., 2012; Marchi et al., 2012; Schenk et al., 2012) show a heavily cratered northern hemisphere and a large impact basin (Rheasilvia) at the south pole (Figure 34). A crater at the south pole of Vesta was previously identified by Hubble Space Telescope images of Vesta (Thomas et al., 1997). Vesta also has a large number of equatorial troughs (Buczkowski et al., 2012), which have characteristics consistent with graben (a depressed part of the crust bounded by faults). The visual geometric albedo of Vesta varies from 0.10 to 0.67 (Reddy et al., 2012a), which is the largest variation observed on an asteroid. As expected, visible and near-infrared spectra of Vesta taken by Dawn (De Sanctis et al., 2012a; Reddy et al., 2012a) are consistent with HED meteorites and ground-based observations. Gamma ray-derived Fe/O and Fe/Si elemental ratios of Vesta’s surface (Prettyman et al., 2012), which are measurements that cannot be done from Earth, are also consistent with HED meteorites. Dawn also has found evidence for hydrated minerals on the surface of Vesta from the detection of widespread hydrogen on the surface from neutron measurements (Prettyman et al., 2012), detection of a widespread 2.8 mm feature (De Sanctis et al., 2012b), and the spectral similarity of Vesta’s dark material to carbonaceous chondrites (McCord et al., 2012; Reddy et al., 2012b). These Dawn results are consistent

with the Hasegawa et al. (2003) observation of a 3 mm band for Vesta and the identification of carbonaceous chondrite clasts in HEDs (e.g., Buchanan et al., 1993). A spectral study of Vesta using Dawn data by Pieters et al. (2012) found no evidence for the accumulation of lunar-like nanophase iron on the surface due to the absence of near-infrared spectral slope differences between freshly exposed and background regolith material. This lack of space weathering on Vesta’s surface could be due a magnetic field (Vernazza et al., 2006b; Fu et al., 2012) and/or compositional differences between Vesta and other spaceweathered bodies (e.g., Moon, Itokawa).

2.14.5.8

OSIRIS-REx

The Origins Spectral Interpretation Resource Identification Security Regolith Explorer (OSIRIS-REx) is a sample return mission to the NEA (101955) 1999 RQ36 that will be launched in 2016. OSIRIS-REx will first rendezvous with the asteroid and then gather a sample and return it back to Earth. This will be the first sample return from a C-complex asteroid. The target body is classified as a B-type (e.g., Clark et al., 2011) due to its blue spectral slope in the visible and nearinfrared (Figure 35). This object also has a low visual albedo of

Figure 34 Image of the south pole of 4 Vesta by NASA’s Dawn spacecraft. The Rheasilvia Basin is the large crater in the image. Image credit: NASA/JPL-Caltech/UCLA/MPS/DLR/IDA. 1.2

Normalized reflectance

Pa¨tzold et al. (2011) and Weiss et al. (2012) argue that Rosetta’s density measurements of Lutetia are consistent with a partially differentiated body. Lutetia’s bulk density is higher than most measured chondrites, but its surface properties resemble chondritic material. However, no conclusive signature of a magnetic field (Richter et al., 2012) around Lutetia was detected by Rosetta, whose signal would indicate a metallic iron core due to differentiation. Sierks et al. (2011) argue that Lutetia is a primordial planetesimal, and not a fragment of a larger body, due to its old surface age, complex surface geology, and high density. The mass spectrometer on Rosetta could not unambiguously detect an exosphere around Lutetia due to outgassing by the spacecraft, which was not constant during the encounter (Altwegg et al., 2012). Rosetta also did photometry of Vesta (Fornasier et al., 2011b) during its approach to Lutetia. Visible spectrophotometric (0.269–0.989 mm) measurements of Vesta by Rosetta were consistent with previous Earth-based observations and fine-grained howardites.

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Figure 35 Plot of reflectance spectrum of OSIRIS-REx B-type target (101955) 1999 RQ36. Spectrum is normalized to unity at 1.05 mm. Error bars are 1s.

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3–6% (Emery et al., 2010). A number of CI chondrites also have blue-sloped spectra (Cloutis et al., 2011a). Both Cloutis et al. (2011a) and Yang and Jewitt (2010) argue that this blue spectral slope is consistent with magnetite, which is commonly found in CI chondrites. Clark et al. (2011) found that the best spectral match to 1999 RQ36 was CI and/or CM chondrites and speculate that it is composed of CM1-like material. They also find that 1999 RQ36 has spectral properties intermediate between Pallaslike and Themis-like objects. Campins et al. (2010b) argue from dynamical and spectral arguments that it is derived from the Nysa/Polana family, which contains B-type objects. The top surface layer of 1999 RQ36 is most likely not pristine. Delbo´ and Michel (2011) argue that the surface of 1999 RQ36 is most likely dehydrated, and its organics have been degraded from solar heating. However, material slightly below the surface (
3–5 cm) should have experienced lower temperatures, which should have protected some of the organics from thermal breakup.

2.14.5.9

2008 TC3 and Almahata Sitta

Although not mission related, we include this unique event here. The first asteroid to be spectrally observed and have samples recovered on Earth was 2008 TC3 (Jenniskens et al., 2009, 2010). This small (
2–5 m) body was discovered 20 h before impact with the Earth’s atmosphere in October of 2008 and a visible spectrum was acquired. The visible spectrum was classified as an F-type. The F-class (Tholen, 1984) has a very weak UV feature and a blue spectral slope in the visible. B-types in the Bus and Binzel (2002a) taxonomy include objects classified as F-types in the Tholen (1984) taxonomy. F-types had been typically (e.g., Bell et al., 1989; Hiroi et al., 1993b) linked with thermally altered carbonaceous chondrite material. The object exploded and fragments rained down over the Sudan. Most of the recovered meteorites (called Almahata Sitta) were found to be polymict ureilites (Bischoff et al., 2010; Zolensky et al., 2010). Cloutis and Hudon (2004) previously had stated that ureilites were most similar spectrally to C-class asteroids. Ureilites generally spectra similar to carbonaceous chondrites with weak absorption bands and flat to blue spectral slopes (Cloutis and Hudon, 2004; Cloutis et al., 2010b). Oxygen isotopic compositions of the fragments (Rumble et al., 2010) were also consistent with ureilites. However, a number of fresh-looking fragments in the strewnfield were found to be chondritic (Bischoff et al., 2010; Shaddad et al., 2010). Most of these chondritic fragments are thought to have been previously incorporated in 2008 TC3 since these chondritic pieces are relatively unweathered and a few samples were found to have detectable short-lived cosmogenic isotopes (Bischoff et al., 2010). Two chondritic meteorites have cosmic-ray exposure ages and light noble gas concentrations similar to the ureilites (Welten et al., 2011) and one chondritic meteorite had amino acid compositions similar to one of the ureilites (Burton et al., 2011).

2.14.6 2.14.6.1

Interesting Groups of Asteroids Earth and Martian Trojans

There has only been one confirmed Earth Trojan (Connors et al., 2011) and four confirmed Martian Trojans. No spectral

information is currently available on the Earth Trojan. Rivkin et al. (2007) did a near-infrared spectral study of two Martian L5 Trojans, (5261) Eureka and (101429) 1998 VF31. Both objects had S-type spectra but very different spectral properties. Eureka had a very broad 1 mm but no 2 mm band, which resembles the spectra of angrites (Burbine et al., 2006), while 1998 VF31 was an S(VII) object with 1 and 2 mm bands typical of pyroxene. The different mineralogical interpretation for the two Martian Trojans argues that they have separate origins from each other.

2.14.6.2

Near-Earth Asteroids

NEAs range from the
32 km diameter (1036) Ganymed (Tedesco et al., 2002b) to meter sized. NEAs primarily sample a size range that is hard to observe in the asteroid belt. About two-thirds of over 400 observed NEAs that have been observed in the visible wavelength region (
0.4–1.0 mm) have been classified as S- or Q-types (Binzel et al., 2004b; Vernazza et al., 2008). Consolidating the results of a number of photometric and spectroscopic near-Earth asteroid surveys, Ye (2011) also finds that objects with silicate absorption features (e.g., S-types, Q-types) dominate the near-Earth population. Binzel et al. (2004) found that almost all of the Bus and Binzel (2002b) classes were represented among the NEAs. Binzel et al. (1996) found a continuum of spectral properties in the visible between Q-types and S-types. Binzel et al. (2010) found that NEAs classified as Q-types have experienced close approaches to Earth (passed within the Earth–Moon distance) within the last 0.5 Ma. Binzel et al. (2010) argue that seismic shaking during close encounters with Earth exposes unweathered surface material and changes the spectral properties of asteroids with ordinary chondrite compositions so that the objects appear less space-weathered. Nesvorny´ et al. (2005) originally postulated that planetary interactions could reset the surface and turn S-asteroids into Q-types and Nesvorny´ et al. (2010) were able to numerically simulate these effects on NEAs. Dell’Oro et al. (2011) argue that collisions may also refresh NEA surfaces. For Q-type and S-type NEAs, Delbo´ et al. (2003) and Thomas et al. (2011b) both found that that the albedos of these objects tended to increase with decreasing diameters. This result is consistent with space weathering on their surfaces since larger asteroids tend to have longer collisional lifetimes than smaller objects (e.g., O’Brien and Greenberg, 2005), and, therefore, older surfaces would tend to be more space-weathered. Vernazza et al. (2008) analyzed the spectra of
40S- and Qtype kilometer-sized NEAs. Using their radiative transfer model, they concluded that most NEAs have spectral properties similar to LL chondrites. This result was a surprise since H and L chondrites tend to fall to Earth at a much higher rate than LL chondrites. From their own spectral study, de Leo´n et al. (2010) also found that the calculated mineralogies of NEAs were most similar to LL chondrites and not H or L chondrites. Due to the spectral similarity between these ‘LL chondrite-like’ NEAs and Flora family members, Vernazza et al. (2008) link these ‘LL chondrite-like’ NEAs to the Flora family in the inner asteroid belt (near the n6 secular resonance). Vernazza et al. (2008) argue that kilometer-sized NEAs are preferentially derived from the inner-belt region, while meteorite source

Asteroids

bodies (assumed to be meter-sized) can come from much larger areas of the asteroid belt due to the Yarkovsky effect. Thomas and Binzel (2010) used MGM as a mathematical tool to fit
50 NEA (S, Sq, Q, V) plus
70 meteorite (H, L, LL, HED) spectra to determine the probability that each NEA was related to a particular meteorite class using derived 1 and 2 mm Geometric Band Centers and Band Area Ratios. (A geometric band center is the wavelength where the total area of the absorption feature is bisected and can be different than the actual band center.) The probabilities that each NEA was related to particular meteorite groups (H, L, LL, and HED) were then convolved with the probability for each NEA, using the source region model of Bottke et al. (2002a), that it originated a particular region in the solar system. These regions included the n6 secular resonance, the Mars-crossing region, and the 3:1 mean-motion resonance. Thomas and Binzel (2010) found that H chondrites appear to be preferentially derived from the 3:1 resonance. This preference appears consistent with the relatively short cosmic-ray exposure ages of H chondrites, which matches the calculated short delivery times expected from delivery from the 3:1 resonance.

2.14.6.3

Asteroid Families

The significance of studying asteroids families is that the interiors of asteroids can be observed and that these family members may have younger surface ages compared to typical main-belt asteroids so the effects and timescales of space weathering can possibly be discerned. Originally, it was postulated (Granahan and Bell, 1991) that a family that is composed predominantly of one taxonomic type would indicate the breakup of a homogeneous parent body. The breakup of a differentiated body could possibly produce a large metal iron core (M-type), lots of olivine-dominated mantle fragments (A-types), and a smaller number of basaltic ones (V-types). However, most asteroid families contain predominately members that are spectroscopically similar (e.g., Bus, 1999; Cellino et al., 2002; Mothe´-Diniz et al., 2005) and variances in spectral properties are consistent with slope and band strength differences that could be due to properties such as particle size and space weathering. This spectroscopic similarity is consistent with the work of Michel et al. (2003) who argued that families originate from the breakup of prefragmented parent bodies and that all large family members formed by the gravitational reaccumulation of smaller bodies. However, many families contain thousands of members but usually only tens of objects have actually been spectrally observed. Objects with spectral properties not consistent with other family members are usually called interlopers (e.g., Migliorini et al., 1995). Background objects are bodies that are nearby but not in the family and whose spectral properties can often be compared to the family members to see if the family members are spectroscopically different from the background members. NEOWISE observations of families (Mainzer et al., 2011a) find that most families have relatively uniform albedos. By looking at the assumed amount of spectral reddening of family members and the estimated ages for the families, the timescale for space weathering has been estimated (Jedicke et al., 2004; Nesvorny´ et al., 2005; Vernazza et al., 2009a; Willman et al., 2008, 2010). Vernazza et al. (2009a) argue

395

that solar wind irradiation rapidly weathers the surfaces of asteroids on timescales of 104–106 years (Strazzulla et al., 2005) when comparing the spectral properties of family members to unweathered meteorites. Willman et al. (2008, 2010) argue that the space weathering on asteroids nears completion between
500 Ma and
1 Ga. The number of identified families has changed over time. Hirayama (1918, 1928, 1933) initially identified three families (Eos, Koronis, and Themis) and then later identified the Flora and Maria families. Zappala` et al. (1990) identified 21 families, while Zappala` et al. (1994) identified 20 families. Twenty-six families were considered relatively robust by Bendjoya and Zappala` (2002) since they were identified by different clustering methods; however, the latest study by Mothe´-Diniz et al. (2005) identified 21 families in the main belt using the proper elements of more than 120 000 objects. Using SDSS colors for
90 000 asteroids, Parker et al. (2008) defined 37 statistically robust families with at least 100 members using a simple Gaussian distribution model in both orbital and color space. Twelve families have over 1000 members in the Parker et al. (2008) study. Nesvorny´ (2010) identifies 55 families (Figure 36) using the proper elements of almost 300 000 asteroids. Novakovic´ et al. (2011) also identified 38 highinclination families in the main asteroid belt. Two families (Schubart and Hilda) have been identified in the Hilda region (Brozˇ and Vokrouhlicky´, 2008; Brozˇ et al., 2011). A number of families have been proposed for the Trojan region (Roig et al., 2008); however, Brozˇ and Rozehnal (2011) argue that the Eurybates family is the only significant family that can currently be recognized in this region. A number of spectrally characterized families will be discussed in the inner, middle, and outer belt plus those in the Trojan region. These families include the ones defined by Nesvorny´ (2010) plus the high-inclination Pallas family and the Eurybates family in the Trojan region. Even though some families contain thousands of objects, the number of objects that have been spectrally observed are usually less than 50 (e.g., Mothe´-Diniz et al., 2005) with the exception being the Eos family that has
100 bodies that have been spectrally observed. However, SDSS photometric observations (e.g., Ivezic´ et al., 2002; Parker et al., 2008) have been done on thousands of family objects.

2.14.6.3.1 Inner belt (
2.06–2.50 AU) The (8) Flora family (
2.19–2.38 AU) is a large family in the inner main belt with over identified 10 000 members (Nesvorny´, 2010). The Flora family has also been called the Flora clan (e.g., Cellino and Zappala`, 1993; Zappala` et al., 1994) since the clustering techniques in the region found a very large grouping of objects with uncertain boundaries making it difficult to calculate a reliable list of members. Clans (Zappala` et al., 1994) may be produced by multiple cratering events or a cascade of fragmentation or could be a real family surrounded by a dense background. The Flora family is estimated to have an age of less than a billion years (Nesvorny´ et al., 2002). The Flora family contains primarily S-complex objects (e.g., Florczak et al., 1998). Flora was argued to be differentiated by Gaffey (1984); however, Vernazza et al. (2008) found that the spectral properties of observed Flora family members and also NEAs tended to be consistent

396

Asteroids

0.35

Tirela

3:1

Sine of proper inclination

0.3

5:2

7:3

Maria

2:1

0.25 Eunomia

0.2 Baptistina

Iannini Gefion

Adeona Vesta

Eos

Veritas

0.15 Datura

0.1

Hygiea

Merxia Flora

Koronis

0.05 Nysa/Polana Massalia

Karin

0 2

2.2

Themis

Agnia

2.4

2.6 2.8 3 Semimajor axis (AU)

3.2

3.4

3.6

Figure 36 Location of Nesvorny´ (2010) asteroid families in a plot of semimajor axis (AU) versus sine of proper inclination. Not plotted are the highinclination Pallas family and the Eurybates family in the Trojan region, which were defined by other researchers. Different colors are used to differentiate families that overlap in semimajor axis (AU) versus sine of proper inclination space. The 3:1 (
2.50 AU), 5:2 (
2.82 AU), 7:3 (
2.95 AU), and 2:1 (
3.27 AU) mean-motion resonances with Jupiter are labeled with gray lines and can be seen to set the boundaries for many of the families. Labeled families are discussed in the text.

with LL chondrites, which are the most olivine-rich ordinary chondrites. Vernazza et al. (2008) argue that the original Flora family parent body may be the source of the LL chondrites. Mothe´-Diniz et al. (2005) identified the (298) Baptistina family (
2.20–2.33 AU) from asteroids that were originally part of the Flora family and found that the observed family members had a variety of taxonomic classifications (Xc, X, C, L, S, V, A). However, Parker et al. (2008) found that Baptistina and Flora family members have very different SDSS color distributions and that the SDSS colors can be used to easily separate members of these two families. The Baptistina family has over identified 3000 members (Nesvorny´, 2010). Bottke et al. (2007) identified the breakup that led to the creation of the Baptistina family as the possible source body of the Cretaceous–Paleogene (or K–Pg) (historically known as the Cretaceous–Tertiary or K-T) extinction event (e.g., Schulte et al., 2010) that occurred
65 Ma. A fossil meteorite found at the K–Pg boundary (Kyte, 1998) appears most similar to a carbonaceous chondrite. Chromium isotopic measurements (Shukolyukov and Lugmair, 1998) of sediments at the K–Pg boundary also argue that the impactor was a carbonaceous chondrite. Bottke et al. (2007) estimated that the breakup occurred
160 Ma using the orbital and estimated size distribution of the Baptistina family; however, this age is a function of an assumed average visual C-type albedo of 0.04, which allows the sizes of the family members to be estimated, plus an assumed density and thermal inertia for a C-type body. The age of the family is estimated by the amount of Yarkovsky drift that should have occurred, which is a function of the diameter of an object since larger objects drift more slowly (e.g., Bottke et al., 2001). Carvano and Lazzaro (2010) initially put into question this relationship between the Baptistina and the K–Pg impactor through their calculation of a visual albedo of
0.35,

which is too high for an object with a carbonaceous chondrite composition and which would reduce the age of the family since it would decrease the sizes of the family members if they had similar albedos. However, the average visual albedo of the Baptistina family from NEOWISE data (Masiero et al., 2011) is 0.21, which is consistent with S-type bodies. Using this average visual albedo and a range of plausible densities and thermal conductivities for an S-type body, Masiero et al. (2012) found that the best fit formation age for the Baptistina family ranged from 140 to 320 Ma. Baptistina itself was classified as an Xctype (Lazzaro et al., 2004) on the basis of its visible spectrum; however, near-infrared data (Reddy et al., 2009, 2011a) of Baptistina found that it is an S-type. Near-infrared observations (Reddy et al., 2009, 2011a) of Baptistina family members found that most observed objects did not have spectral properties similar to carbonaceous chondrites. The (1270) Datura family (
2.23–2.24 AU) is a small asteroid family (six identified members) (Nesvorny´, 2010) that has an extremely young age (Nesvorny´ et al., 2006b), which is now estimated to be 530  20 thousand years (Vokrouhlicky´ et al., 2009). A visible spectra survey of Datura family members (Mothe´-Diniz and Nesvorny´, 2008a) identified two Sk-types (including Datura), one Sq-type, and one O/Q-type. A nearinfrared spectrum of Datura was identified as consistent with a Q-type (Chapman et al., 2009); however, Takato (2008) found that visible spectra of the body resembled an S-type. The (4) Vesta family (
2.26–2.48 AU) is composed of one large object (Vesta) plus over 13 000 identified smaller bodies (Nesvorny´, 2010). Observed members (Mothe´-Diniz et al., 2005) of the Vesta family are predominantly (
75%) V-types with the remainder primarily S-complex objects from visible observations. Ivezic´ et al. (2002) and Parker et al. (2008) found using SDSS colors that the Vesta family have relatively homogeneous colors that are distinct from the background

Asteroids

objects. The Vesta family is estimated to be at least 1 Ga (Nesvorny´ et al., 2008). All the V-types in the Vesta family have estimated diameters less than
5 km, while the Vesta family objects that do not have V-type taxonomies have estimated diameters greater than
10 km (Mothe´-Diniz et al., 2005). Moskovitz et al. (2010) and De Sanctis et al. (2011) found that almost all observed Vesta family members in the near-infrared have band parameters consistent with a eucrite and/or howardite mineralogy. Only one observed Vesta family member, (10349) 1992 LN (De Sanctis et al., 2011), had band parameters consistent with just a diogenite mineralogy. V-types objects have long been known to extend outside the Vesta family (e.g., Binzel et al., 1993; Carruba et al., 2005; Florczak et al., 2002; Moskovitz et al., 2008a; Nesvorny´ et al., 2008) implying that objects derived from Vesta extend outside the Vesta family. Binzel et al. (1999) found that the distribution of V-types in the inner main belt around Vesta is consistent with their derivation from Vesta. V-types in the inner main belt are widely spread in semimajor axis but more tightly constrained in orbital eccentricity and inclination, which is consistent with the distribution predicted from Lagrange’s equations for differential orbits. The (44) Nysa/(142) Polana family (
2.31–2.48 AU) is a very taxonomically diverse family. The Nysa/Polana family has over 15 000 identified members (Nesvorny´, 2010). Nysa is an E-type, while 142 Polana is a B-type. From visible data, observed members (Mothe´-Diniz et al., 2005) of the Nysa family are
40% C-complex (including B-types),
40% S-complex, and the remainder are X-complex. Based on spectroscopic observations, Cellino et al. (2001) argue that Nysa/ Polana family actually contains two families: the (878) Mildred family composed of S-types and the Polana family composed on B-types (originally called F-types). Nysa and (135) Hertha, an M-type, are relatively large (
70 km for Nysa and
80 km for Hertha; Tedesco et al., 2002b) and are often thought of interlopers due to their large sizes and classifications not consistent with most other family members. Gayon-Markt et al. (2011) postulate that the Almahata Sitta meteorite and near-Earth asteroid 2008 TC3 could be derived from the Nysa/Polana family since this family contains taxonomic classes that have been linked with the types of meteorites contained in Almahata Sitta. Almahata Sitta is mixture of ureilite material (linked with B-types) plus chondritic fragments such as ordinary chondrites (linked with S-types) and enstatite chondrites (linked with X-types). The (20) Massalia family (
2.33–2.48 AU) was identified by Mothe´-Diniz et al.(2005) as one of the least characterized large asteroid families since only a few members have been given taxonomic designations. Family members Massalia and (182) Elsa were classified as S-complex bodies (Bus and Binzel, 2002b) while family members (2946) Muchachos and (2316) Jo-Ann were classified as C-complex bodies (Bus and Binzel, 2002b; Xu et al., 1995). The Massalia family has over 5000 identified members (Nesvorny´, 2010) with Massalia being relatively large (diameter of
146 km; Tedesco et al., 2002b), Elsa having an intermediate diameter (
44 km; Tedesco et al., 2002b), and other family members being relatively small (< 10 km). The Massalia family is thought to have been created by an impact on Massalia less than 200 Ma ago (Vokrouhlicky´ et al., 2006a).

397

2.14.6.3.2 Middle belt (
2.50–2.82 AU) From visible data, observed members (Mothe´-Diniz et al., 2005) of the (170) Maria family (
2.52–2.62 AU) include L, Sl, S, and K-types. The Maria family has over 3000 identified members (Nesvorny´, 2010). From near-infrared observations, FieberBeyer et al. (2011) found that observed objects in this family had weak 1 and 2 mm and were red-sloped. They argued that these spectral properties were consistent with mesosiderites. From visible data, observed members of the (15) Eunomia family (
2.52–2.73 AU) are predominantly (
94%) (Nathues, 2010) composed of S-complex objects. As mentioned earlier, Eunomia was classified as a K-type by DeMeo et al. (2009) using near-infrared data. The Eunomia family has over 7000 identified members (Nesvorny´, 2010). Visible spectra (Nathues, 2010; Nathues et al., 2005) of Eunomia family members have been interpreted as indicating mineralogies consistent with the original parent body being partially differentiated due to the olivinerich nature of the observed members. From visible data, observed members (Mothe´-Diniz et al., 2005) of the (145) Adeona family (
2.56–2.71 AU) are predominantly composed of Ch-types in a background that is primarily composed of S-complex objects. The Adeona family has over 2000 identified members (Nesvorny´, 2010). Adeona is postulated (Mothe´-Diniz et al., 2005) to be due to the breakup of a CM-like parent body. The Adeona family is estimated to have an age of no more than 600 Ma (Carruba et al., 2003). From visible data, observed members (Mothe´-Diniz et al., 2005) of the (4652) Iannini family (
2.64–2.65 AU) are predominately S-complex objects with the best match being Sk-objects (Willman et al., 2008). The Iannini family has over 100 identified members (Nesvorny´, 2010). The Iannini family was estimated to have a relatively young age (5 Ma or less) (Nesvorny´ et al., 2003) and as a source of a dust band first identified by IRAS (Low et al., 1984) due to the similarity in proper inclinations. These dusts bands are located between Mars and Jupiter in the asteroid belt and thought to be related to asteroid collisions (e.g., Sykes and Greenberg, 1986) that have occurred relatively recently due to the short lifetimes of the dust particles (Nesvorny´ et al., 2003). From visible data, observed members (Mothe´-Diniz et al., 2005) of the (808) Merxia family (
2.68–2.82 AU) are predominantly S-complex objects. The Merxia family has over 1000 identified members (Nesvorny´, 2010). Using MGM on visible and near-infrared spectra, Sunshine et al. (2004) found that a number of Merxia family members have silicate mineralogies consistent with low- and high-Ca pyroxene with minor amounts of olivine (<20%) and the relatively high abundances of high-Ca pyroxene (
40% of the total pyroxene) was only consistent with bodies that have undergone melting. The Merxia family is estimated to have an age of
330 Ma (Vokrouhlicky´ et al., 2006a). From visible data, observed members of the (2) Pallas family (
2.71–2.79 AU) contains primarily B-type bodies (Carruba, 2010; Clark et al., 2010; Novakovic´ et al., 2011). Near-infrared observations of Pallas (Clark et al., 2010) and two Pallas members confirm their spectral similarity. The Pallas family has
60 identified members (Novakovic´ et al., 2011). Most likely the Pallas family is due to a cratering event that has ejected material off of Pallas’ surface (e.g., Novakovic´ et al., 2011) since the other family members have estimated

398

Asteroids

diameters much smaller than Pallas. Hubble Space Telescope images of Pallas (Schmitz et al., 2011) indicate a large impact basin on its surface. From visible data, observed members (Mothe´-Diniz et al., 2005) of the (848) Agnia family (
2.73–2.82 AU) are all S-complex objects. The Agnia family has over 1500 identified members (Nesvorny´, 2010). As with the Merxia family, MGM analyses of objects in the Agnia family have high abundances of high-Ca pyroxene, which are consistent with the original Agnia parent body undergoing igneous differentiation. The Agnia family is estimated to have an age of
100 Ma (Vokrouhlicky´ et al., 2006b). From visible data, observed members (Mothe´-Diniz et al., 2005) of the (1272) Gefion family (
2.74–2.82 AU) are predominantly composed of S-complex objects. The Gefion family has over 3400 identified members (Nesvorny´, 2010). Nesvorny´ et al. (2009) believe that the original Gefion parent body may be the parent body of L chondrites from its estimated age (
485 Ma), location, and interpreted mineralogy. Korochantseva et al. (2007) found an age of
470 Ma for the breakup of the L chondrite parent body from 39Ar–40Ar dating of shocked L chondrites and found that this age correlated with the age of fossil L chondrites located in Ordovician sediments and at the Lockne impact crater (e.g., Schmitz et al., 2001, 2011). The relatively short cosmic-ray exposure ages (<1 Ma) of some L chondrites (e.g., Eugster et al., 2006) is consistent with the Gefion family’s location near the 5:2 meteoritesupplying resonance.

2.14.6.3.3 Outer belt (
2.82–3.28 AU) From visible data, observed members (Mothe´-Diniz et al., 2005) of the (158) Koronis family (
2.82–2.96 AU) predominately contains S-complex objects in a background of C-complex asteroids. Ivezic´ et al. (2002) and Parker et al. (2008) found using SDSS colors that the Koronis family have relatively homogeneous colors that are distinct from the background objects. The Koronis family has
5000 identified members (Nesvorny´, 2010). The age of the Koronis family is estimated to be
2 Ga (Marzari et al., 1995). Rivkin et al. (2011c) found a number of small (<5 km) bodies had broadband visible colors consistent with Q-types. Also using broadband visible colors, Thomas et al. (2011a) found that the spectral slopes of Koronis members tends to increase with increasing size, consistent with a transition from Q- to Stypes due to space weathering. The (832) Karin family (
2.86–2.87 AU) is a small cluster of objects that is located in proper element space inside the much larger Koronis family (e.g., Cellino et al., 2010; Nesvorny´ et al., 2006a) and is commonly thought to be due to the breakup of a Koronis family member. From visible data, observed members (Vernazza et al., 2006a) of the 832 Karin family are almost all S-complex objects (primarily S- and Sk-types) plus one Q-type. The Karin family has
500 identified members (Nesvorny´, 2010). Near-infrared spectra of Karin family members find that that these bodies have spectra that resemble slightly reddened ordinary chondrites (Vernazza et al., 2006a) and are less red than typical Koronis family members (Chapman et al., 2009). This postulated small degree of space weathering is consistent with the estimated young formation age of the Karin family (5.75  0.05 Ma) (Nesvorny´ and Bottke, 2004). Sasaki et al. (2004) found that

visible and near-infrared rotational spectral variations for Karin indicated space-weathered (S-type) and fresh (Q-type) faces on the surface, however, Vernazza et al. (2007b) saw no indication of rotational variations for Karin. The mean visible geometric albedo (
0.22) (Harris et al., 2009) of Karin family members is also consistent with ordinary chondrite-like material. From visible data, observed members (Mothe´-Diniz et al., 2005) of the (221) Eos family (
2.99–3.03 AU) are spectrally relatively inhomogeneous (containing
30% T,
20% D, and
15% K) in a background of C- and X-complex asteroids. However, Ivezic´ et al. (2002) and Parker et al. (2008) found using SDSS colors that the Eos family have relatively homogeneous colors that are distinct from the background objects. The Eos family has over identified 6000 members (Nesvorny´, 2010). The Eos family has an estimated age of
1.3 Ga (Vokrouhlicky´ et al., 2006c). K-types have typically been thought to have mineralogies similar to the CO and CV chondrites (Bell, 1988; Burbine et al., 2001a; Clark et al., 2009) due to their similar spectral properties in the visible and nearinfrared and visual albedos. However, Mothe´-Diniz et al. (2008) argue that the calculated mineralogies of a number of Eos family members are consistent with forsteritic olivine surface assemblages, which they believe is consistent with either partial differentiation of an ordinary chondrite parent body or a CK chondritic mineralogy. Carporzen et al. (2011) believe that the apparent spectral diversity in the Eos family may be evidence for a differentiated body. Carporzen et al. (2011) and Sahijpal and Gupta (2011) argue that the paleomagnetism of the CV3 carbonaceous chondrite Allende (e.g., Weiss et al., 2010) is evidence that the CV3-parent body was partially differentiated with a convecting metallic iron core that resulted in a global magnetic field. The primitive meteorite Allende would be derived from an outer, unmelted crust of such a body, which could be compositionally similar to the original parent body of the Eos family. From visible data, the (1400) Tirela family (
3.03– 3.21 AU) contains primarily Ld-types (Mothe´-Diniz and Nesvorny´, 2008b). Ld-types have very steep UV slopes but become flatter past
0.75 mm (Bus and Binzel, 2002b). The Tirela family has over identified 2000 members (Nesvorny´, 2010). The near-infrared spectrum of Tirela (Mothe´-Diniz and Nesvorny´, 2008b) has weak 1 and 2 mm features, and the band positions of these features are best fit by the mineral pigeonite ((Ca,Mg,Fe)(Mg,Fe)Si2O6). From visible data, the (24) Themis family (
3.05–3.24 AU) contains primarily C-complex objects (Mothe´-Diniz et al., 2005). Ivezic´ et al. (2002) and Parker et al. (2008) found using SDSS colors that the Themis family have relatively homogeneous colors that are distinct from the background objects. The Themis family has over identified 6000 members (Nesvorny´, 2010). The age of the Themis family is estimated to be
2 Ga (Marzari et al., 1995). Observed members (Mothe´-Diniz et al., 2005) in the Themis family are primarily (
80%) C-complex objects with the remainder being X-types. B-types dominate among the C-complex objects in the Themis family. B-types dominate among the C-complex objects in the Themis family. Nearinfrared observations of Themis family members (Ziffer et al., 2011) found that Themis family objects tended to have concaveup shapes and red spectral slopes. DeMeo et al. (2011) observed a Themis family member (379 Huenna) and its satellite (S/2003

Asteroids

(379) 1) and found that both objects had similar near-infrared spectral properties that were consistent with other Themis family objects. A number of identified main-belt comets appear dynamically related (Hsieh and Jewitt, 2006; Licandro et al., 2011b) to the Themis family, which would argue that Themis family members are volatile-rich; however, no outgassing has been observed on Themis itself (Lovell et al., 2010). From visible data, the (10) Hygiea family (
3.03–3.24 AU) contains primarily C-complex objects (with a number of B-types) (Mothe´-Diniz et al., 2001, 2005) with a few S-types. The Hygeia family has
4000 identified members (Nesvorny´, 2010). Hygeia has a diameter of
400 km (Tedesco et al., 2002b). No feature at 0.7 mm was apparent in any of the spectra (Mothe´-Diniz et al., 2001). A 3 mm band was identified in the spectra of Hygeia (Jones et al., 1990) and Rivkin et al. (2011d) found that this feature resembles Ceres’ 3 mm band suggesting that Hygeia may have similar mineralogy to Ceres. Mothe´-Diniz et al. (2001) found no evidence for rotational visible spectral variations on Hygeia in contrast with other observational studies (Rivkin et al., 2002). From visible data, the (490) Veritas family (
3.17– 3.18 AU) contains primarily C-complex objects (Mothe´-Diniz et al., 2005). The Veritas family has
900 identified members (Nesvorny´, 2010). The estimated age of the Veritas family is 8.3 Ma (Nesvorny´ et al., 2003; Tsiganis et al., 2007). All observed members (Mothe´-Diniz et al., 2005) in the Veritas family are C-complex objects that are primarily Ch-types, which are usually associated with CM chondrites due to the presence of a 0.7 mm feature. Near-infrared observations of Veritas family members (Ziffer et al., 2011) found that Veritas family objects tended to have concave down shapes and flat spectral slopes. These spectral differences between Themis and Veritas family members could be due to composition and/or space weathering. The Veritas family is thought to be a source of one of the IRAS dust bands (Nesvorny´ et al., 2003) due to a similarity in proper inclination and the Veritas family’s relatively young age. Farley et al. (2006) attribute the disruption that created the Veritas family as increasing the interplanetary dust particle (IDP) flux on Earth
8.25 Ma, which also increases the 3He flux since IDPs are enriched in 3He.

399

2.14.6.3.4 Trojan region (
5.2 AU) The (3548) Eurybates family (
5.2 AU) is the only family that has been convincingly identified among the Jupiter Trojans (Brozˇ and Rozehnal, 2011). There are
200 identified members in the Eurybates family. C-types are abundant in the Eurybates family (Brozˇ and Rozehnal, 2011; Fornasier et al., 2007), which is located in the L4 region. Objects in the Eurybates family have weak UV features and less red spectra (Fornasier et al., 2007) in the visible compared to D- and P-types, which dominate the Trojan region (Roig et al., 2008). Eurybates family members also have distinct SDSS colors compared to the Trojan background (Brozˇ and Rozehnal, 2011). Yang and Jewitt (2011) find that the best match to Eurybates in the near-infrared to be a CM chondrite, which might imply that the parent body of Eurybates experienced aqueous alteration. However, a spectral study of Eurybates family members found no evidence of aqueous alteration (de Luise et al., 2010) due to their relatively featureless spectra in the visible and near-infrared.

2.14.7

Taxonomic Distribution of Taxonomic Types

2.14.7.1

Earlier Work

It has been known since the 1970s that the abundances of different taxonomic classes varied according to distance from the Sun. With
100 classified objects, Chapman et al. (1975) found that objects that were classified as C-asteroids become much more abundant in the outer belt. Gradie and Tedesco (1982) (Figure 37) and Gradie et al. (1989) looked at the biased-corrected distribution of taxonomic classes with a much larger sample and a larger number of taxonomic types. They found that E-types were preferentially found in the Hungaria group in the inner belt, the S-type distribution peaks at
2.2 AU, the C-type distribution peaks at
3 AU, P-types are preferentially found in the Hilda group at
4 AU, and D-types are preferentially found in the Trojan region at 5.2 AU. Using a much larger number of classified asteroids with diameters equal to or greater than 20 km from SMASS II, Bus and Binzel (2002b) noted that their taxonomic distribution matches pretty well the

1 0.9 0.8

C

E + ¢R¢

D

0.7 Fraction

0.6 0.5

S

0.4 0.3 P

0.2 M, F

0.1 0 1.5

2

2.5

3

3.5

4

4.5

5

5.5

Semimajor axis (AU) Figure 37 Plot of semimajor axis (AU) versus fraction of bias-corrected classified asteroids for a number of taxonomic types from Gradie and Tedesco (1982). The listed class ‘R’ is spectrally different from the currently used R-type spectral designation and instead includes objects with very red visible spectra with 1 mm deeper than typical S-types (e.g., A-types; Tholen and Barucci, 1989).

400

Asteroids

Gradie and Tedesco (1982) distribution. Bus and Binzel (2002b) did find a secondary S-complex peak at
2.85 AU and a secondary C-complex peak at
2.6 AU that were roughly evident in distributions plotted in Gradie et al. (1989). Bell (1986) and Bell et al. (1989) inferred degrees of metamorphic heating for each of the Tholen (1984) classes plus the K class (Bell, 1988) and placed the classes into ‘primitive’ (D, P, C, K, Q), ‘metamorphic’ (T, B, G, F), and ‘igneous’ (V, R, S, A, M, E) superclasses. The heat source was assume to be due to 26 Al and/or solar wind induction heating and had to vary in intensity with distance from the Sun. Excesses of 26Mg (daughter product of 26Al decay) in meteorites (e.g., Baker et al., 2005; Srinivasan et al., 1999) are evidence that 26Al was present. Grimm and McSween (1993) showed that 26Al could produces such a heating trend assuming that it takes more time for objects to accrete farther from the Sun (e.g., Chambers and Wetherill, 1998). Bodies that accreted later would have less 26 Al, due to its short half-life (
720 000 years). Solar wind induction heating is based on the premise that a T-Tauri (premain sequence) star generates a magnetic field due to its solar wind and an electric current is generated in the planetesimals that are moving through the disk, which heats the objects (Ghosh et al., 2006; Shimazu and Terasawa, 1995; Sonnett et al., 1968). Solar wind induction heating is not currently thought to be a dominant process for heating asteroids since direct observations of T-Tauri stars do not find (e.g., Ghosh et al., 2006; Wood and Pellas, 1991) that their solar wind flux in the disk is sufficient to heat planetesimals. With increasing distance from the Sun, the percentage of objects with
0.7 mm (Vilas and Gaffey, 1989) and 3 mm bands (Jones et al., 1990) tends to decrease, indicating that the percentage of objects with hydrated silicates on their surfaces tended to decrease with heliocentric distance. Bell et al. (1989) found that the igneous superclass objects dominated the main belt interior to
2.7 AU, the primitive superclass objects dominant the main belt past
3.4 AU, and the metamorphic superclass objects are relatively rare and peak in the belt from
3.0 AU. The degree of parent body heating appears to decrease with increasing distance from the Sun as seen by the increase in ‘primitive’ bodies; however, the general consensus that some S objects (e.g., Mothe´-Diniz et al., 2010) have ordinary chondrite mineralogies puts into question the actual shapes of these heating trends. Vernazza et al. (2007a) have estimated that over 70% of S-complex bodies have mineralogies consistent with ordinary chondrites.

2.14.7.2

Mothe´-Diniz et al. (2003) looked at the bias-corrected taxonomic distribution for
2000 main-belt objects classified in the visible (Bus, 1999; Bus and Binzel, 2002b; Lazzaro et al., 2004) from
2.1 to
3.3 AU. The bias correction corrects for observational biases such as tending to observe the brightest asteroids, which tend to have the highest albedos, largest diameters, and /or closest to Earth. Mothe´-Diniz et al. (2003) used the typical boundaries for the inner, middle, and outer belt to look at the distribution of X-, C-, and X-complex objects. As expected, Mothe´-Diniz et al. (2003) found that that S-complex objects are the most abundant asteroid class in the inner and middle belt while C-complex objects are the most abundant in the outer belt, especially past 3.05 AU. Between 2.1 and 2.5 AU, Alvarez-Candal et al. (2006) found that the concentration of S-complex objects decreases with increasing heliocentric distance and the concentration of C-complex objects increases. However, Michtchenko et al. (2010) found that S-complex bodies are evenly distributed in this inner main-belt region. X-complex objects are the least abundant complex in the inner and middle of the belt (Mothe´-Diniz et al., 2003) but become more abundant in the outer belt. Using SDSS colors, the taxonomic distributions of Carvano et al. (2010) were pretty consistent with the distributions seen by Mothe´-Diniz et al. (2003). On average, the spectral slopes of low-albedo asteroids in the visible and near-infrared tend to get redder with increasing distance from the Sun (e.g., Emery and Brown, 2003; Gradie and Tedesco, 1982; Jewitt and Luu, 1990b; Vilas and Smith, 1985). This trend is evident by looking at the concentrations of different taxonomic types for the Cybele, Hilda, and Trojan regions. From SDSS colors, the Cybele region (
3.3–3.7 AU) tends to contain more X-types (P-types if these objects are assumed to have low Cybele-like albedos) than D-types (GilHutton and Licandro, 2010). The average visual geometric albedo of bodies in the Cybele region is 0.06 (Kasuga et al., 2012). Using SDSS colors, Gil-Hutton and Brunini (2008) found approximately equal numbers of D- and X-types in the Hilda region. These Hilda X-types are most likely P-types since the Hilda region (
4.04 AU) tends to contain relatively lowalbedo objects (Grav et al., 2012a; Ryan and Woodward, 2011). Jupiter Trojans (
5.2 AU) tend to be dominated by D-types (Dotto et al., 2008; Fornasier et al., 2007; Grav et al., 2012b; Roig et al., 2008). The differences in spectral slopes (D-, P-, C-types) between these outer region objects have been argued to be due to space weathering (e.g., Gil-Hutton and Brunini, 2008) and/or compositional differences (e.g., Emery et al., 2011).

Recent Work

More recent work tends to confirm the Gradie and Tedesco (1982) taxonomic distribution of the belt. The Hungaria region (
1.8–2.0 AU) is located in the innermost part of the main belt. Using SDSS colors, Assandri and Gil-Hutton (2008) and Warner et al. (2009) found that
60% and
75% in this region, respectively, would be classified as X-types with the remainder being S- and C-complex objects. Known E-types (Clark et al., 2004a), which are usually linked with aubrites, are preferentially located in the Hungaria region. A large percentage of Hungaria objects have high albedos from NEOWISE observations (Masiero et al., 2011).

2.14.7.3

Formation of Material in the Solar Nebula

Mineralogical and isotopic studies of meteorites, observations of molecular clouds and young stellar objects, and theoretical models give insight on the history of the early Solar System (e.g., Alexander et al., 2001). Grossman (1972) calculated the order (Table 1) that minerals would condense out of the solar nebula with decreasing temperature. Later studies (e.g., Ebel, 2006) confirmed this basic sequence. The first minerals that are predicted to condense are oxides and members of the melilite groups, which are common constituents of CAIs. CAIs are the oldest components of meteorites and are commonly found in

Asteroids

Table 1 Condensation sequence for different minerals from Grossman (1972) for a gas of solar composition at 103 atm Temperature (K)

Mineral

1758 (1513) 1647 (1393) 1625 (1450) 1513 (1362) 1471 1450 1444 1362 1349 <1000 <1000

Corundum (oxide), Al2O3 Perovskite (oxide), CaTiO3 Melilite group, Ca2Al2SiO7–Ca2MgSi2O7 Spinel (oxide), MgAl2O4 FeNi metal Diopside (pyroxene), CaMgSi2O6 Forsterite (olivine), Mg2SiO4 Anorthite (feldspar), CaAl2Si2O8 Enstatite (pyroxene), Mg2Si2O6 Alkali-bearing feldspar, (Na,K)AlSi3O8–CaAl2Si2O8 Ferrous olivines, (Mg,Fe)2SiO4; ferrous pyroxenes, (Mg,Fe)2Si2O6 Troilite (sulfide), FeS Magnetite (oxide), Fe3O4

700 405

The temperatures in parentheses are the temperatures where the condensate disappears.

chondritic meteorites. The oldest known CAIs have formation ages of 4.567–4.568 Ga (Amelin et al., 2010; Bouvier and Wadhwa, 2010) with uncertainties of hundreds of thousands of years. After CAIs, metallic iron and a number of FeO-free minerals (diopside, forsterite, anorthite, enstatite) would condense. Enstatite chondrites, which are predominately enstatite and metallic iron, are examples of such material. As the temperature decreases, more alkali feldspars and FeO-bearing olivines and pyroxenes would condense. Then, troilite will condense at
700 K. At temperatures below
490 K (Lewis, 1997), phyllosilicates will form from the reaction of water vapor with Fe-bearing olivines and pyroxenes. Magnetite will condense at
405 K. Water ice will condense at temperatures of
160–170 K in the solar nebula (e.g., Lunine, 2006). Organics can form in a variety of ways. One way to form organics is through Fischer– Tropsch reactions (e.g., Kress and Tielens, 2001) where CO and H2 are converted to hydrocarbons through heating with a transition metal catalyst. Simple organics can also form (e.g., Ehrenfreund and Charnley, 2000; Nuevo et al., 2011) in extremely cold environments through irradiation by ultraviolet irradiation and cosmic-ray bombardment of H2O, CO, CO2, CH3OH, and NH3 ices on the surfaces of silicate and carbonaceous dust grains. Even though hydrated silicates should have formed in the solar nebula, the hydrated silicates in meteorites appear to have formed through aqueous alteration (e.g., Brearley, 2006; Tomeoka, 1990). For aqueous alteration to take place, the body must have liquid water and be exposed to a heat source (e.g., Grimm and McSween, 1989; McSween et al., 2002). Temperatures of at least
300 K (e.g., Jones and Brearley, 2006) are thought to be needed to aqueously alter asteroids. The temperature of the disk decreases with increasing distance from the Sun (e.g., Boss, 1998; Chapter 2.3) with the slope of the temperature decrease being very model-dependent. CAIs would condense closest to the Sun but have been shown to be able to dynamically diffuse outward (e.g., Boss, 2012; Ciesla, 2010; Cuzzi et al., 2003) so they can be incorporated into various chondritic meteorite classes (e.g., MacPherson and

401

Boss, 2011; Scott and Krot, 2005) and comet-forming regions (e.g., Simon et al., 2008). X-winds (high speed bipolar collimated jets) around young stellar objects have also been proposed (e.g., Hu, 2010; Shu et al., 1996, 2001) to be the mechanisms for producing CAIs (and chondrules) and transporting them outward from the Sun; however, some researchers (e.g., Desch et al., 2010) do not believe the X-wind model is viable for producing and transporting CAIs. Partially melted and fully differentiated parent bodies would form from the melting of chondritic planetesimals with the heat source usually assumed to be due to the decay of 26Al. Dating of differentiated and chondritic material find that the accretion of differentiated planetesimals occurred at an earlier time than the accretion of undifferentiated planetesimals (e.g., Baker et al., 2005; Qin et al., 2008). Aubrites are commonly assumed to be the result of melting of enstatite chondrite-like material (e.g., Keil, 1989, 2010). Experiments have shown (Jurewicz et al., 1991, 1993) partial melting of carbonaceous chondrites produce angrite- or eucrite-like depending on the oxygen fugacity. Ford et al. (2008) found that partial melting of ordinary chondrites does not appear to form primitive achondrite material (e.g., acapulcoites/lodranites) and argue that the precursor material for these meteorites was not as FeO-rich and oxidized as ordinary chondrites. Ebel and Alexander (2011) predict that enstatite chondrites would have formed under similar conditions that formed Mercury at 0.4 AU, which also argues that enstatite chondrites formed in the inner solar system. From MESSENGER’s x-ray measurements (Nittler et al., 2011; Weider et al., 2012) of its surface, Mercury’s elemental abundance ratios are best matched by a partially melted enstatite chondrite mineralogy (Burbine et al., 2002b). Wasson and Wetherill (1979) previously proposed that enstatite-rich meteorites formed near or interior to 1 AU. Depending on the model, water ice would start condensing between
2 and
5 AU (e.g., Lunine, 2006). FeO-bearing silicates would condense at intermediate heliocentric distances between the regions where enstatite and water ice condense.

2.14.7.4

Origin of the Taxonomic Distribution

The asteroid belt region is believed to have originally contained hundreds to thousands of times more mass than it currently does today (e.g., Lecar and Franklin, 1973; Wetherill, 1989) due to the extreme mass deficit (Weidenschilling, 1977) in the Mars and asteroid belt region compared to the relatively smooth decline from Venus to Neptune. Postulated scenarios (e.g., Petit et al., 2002) for this depletion are the sweeping of secular resonances and the gravitational scattering of planetary embryos that were later ejected from the solar system. O’Brien et al. (2007) argue that sweeping secular resonances and planetary embryos are not sufficient to dynamically excite asteroids in the timescales required. O’Brien et al. (2007) argue that the ‘Nice Model’ (Gomes et al., 2005; Morbidelli et al., 2005; Tsiganis et al., 2005), where the outer planets have significant migrations in the early Solar System, is needed to deplete the asteroid belt to the extremes needed. Using the ‘Nice Model,’ Walsh et al. (2011) postulated the ‘Grand Tack’ model, which has Jupiter initially forming at
3–4 AU and Saturn at
4–5 AU. These two planets then migrate inward and then outward. Jupiter’s has an inward

402

Asteroids

migration to
1.5 AU and then outward, which truncates the terrestrial disk at
1 AU. Saturn has a similar type of migration that starts
100 Ma later but moves much quicker inward. Jupiter’s migration initially empties the asteroid belt but then repopulates the belt with bodies from 1 to 3 AU, which are labeled as ‘S-types’ in the Walsh et al. (2011) simulation. Bodies that formed between the giant planets and from 8.0 to 13.0 AU are labeled as ‘C-types’ in the simulation. The movements of the giant planets scatter this ‘C-type’ material into the asteroid belt. ‘S-types’ are preferentially scattered into the inner asteroid belt, while ‘C-types’ are preferentially scattered into the outer asteroid belt, which is roughly similar to what is seen today. Other researchers have proposed that the main-belt asteroids were scattered into their present-day locations and did not form where they are currently located. Hartmann (1990) proposed that a large flux of C-asteroids were scattered in the early solar system to produce the relatively large number of CM clasts in HEDs and a number of satellite (like Phobos) that are similar to C-asteroids. Bottke et al. (2006a) proposed that the parent bodies of the iron meteorites formed in the terrestrial planet region and then were scattered into the main belt. Levison et al. (2009) argued that most primitive objects (Dand P-types) in the main belt, Hilda region, and Jupiter Trojans are scattered trans-Neptunian objects that formed past
15 AU and are scattered there through the migrations of the giant planets. Morbidelli et al. (2005) have previously argued that Jupiter Trojans formed in distant regions and were scattered there by giant planet migrations. This type of scattering of planetesimals is consistent with the distribution of taxonomic classes, for their assumed mineralogies, since the average final location is related to the initial formation location. Enstatite-rich material (e.g., aubrites, enstatite chondrites) would be preferentially scattered into the Hungaria region (
1.8–2.0 AU) and is now observed as E-types, which have high albedos and relatively featureless spectra. Without albedo information, asteroids with aubrite surfaces would be classified as X-types. Assandri and GilHutton (2008) argue that the C-complex objects in the Hungaria region either have been dynamically transported there or are the result of the space weathering of ‘fresh’ X-types. Lower albedo bodies in the Hungaria region may have enstatite chondrite mineralogies. FeO-rich material (e.g., ordinary chondrites, primitive achondrites) would be preferentially scattered into the inner and middle main belt and are now evident as S-types, which have distinctive bands due to Fe2þ. A small fraction of these FeO-rich bodies would have also been scattered into the outer belt and evidence for these bodies are outer-belt S- (e.g., Koronis family, Karin family) and V-types (e.g., (1459) Magnya, (7472) Kumakiri, (10537) 1991 RY16). With increasing distance from the Sun, the percentage of objects with
0.7 mm (Vilas and Gaffey, 1989) and 3 mm bands (Jones et al., 1990) tends to decrease, indicating that the percentage of objects that have undergone aqueous alteration also tends to decrease with heliocentric distance. Silicate bodies that formed with water ice would only aqueously alter if they accreted before 26Al started to significantly decay, which would occur if the accretion rate decreases with heliocentric distance (e.g., Chambers and Wetherill, 1998). During the scattering process, objects that aqueously altered (e.g., Ch-, Cgh-types) would have tended to form closer to the Sun and,

therefore, would have been preferentially scattered into locations closer to the Sun than those that did not aqueously alter. Organic-rich bodies (D- and P-types) would have formed in the outer solar system and would have been preferentially scattered into the Hilda region (
4 AU) and Jupiter region (
5.2 AU). A few of these organic-rich bodies would also been scattered into the main asteroid belt (Levison et al., 2009) and would be found as main-belt D- and P-types.

2.14.8

Conclusions and Future Work

From the start of the twenty-first century, the chemical and mineralogical characterization of asteroids has greatly expanded with the commissioning of SpeX on the IRTF, which has made near-infrared observations of asteroids relatively routine, the SDSS, which has obtained photometric colors in the visible of over 100 000 asteroids, and the numerous spacecraft missions, which have orbited and landed on asteroids and even returned samples of one to Earth. Since it started obtaining data in 2000, SpeX has allowed the mineralogies of faint asteroids to be derived by being able to obtain spectra out to 2.5 mm, which allows the olivine and pyroxene bands to be completely covered when combined with visible data, and across the 3 mm region, which allows for features due to water and OH to be identified. Also starting to collect data in 2000, the SDSS has allowed for the visible spectral properties of an extremely large set of bodies to be determined and to see whether trends apparent with much smaller sets of data are ‘real.’ Considering that SDSS was designed as a redshift survey, its results can be considered much more remarkable. Since the NEAR Shoemaker encounter in 2000, orbiting spacecraft missions have allowed for the bulk chemistries of asteroids to be directly calculated through x-ray and gamma-ray measurements and the spectral properties to be mapped in detail over the surface. However, the key event in determining asteroid compositions was the return of samples directly obtained from the NEA (25143) Itokawa by the Hayabusa spacecraft on 13th June of 2010. The returned grains allowed for the mineralogy and isotopic composition of a known asteroid and the effects of space weathering to be directly determined. The hypotheses that some S-types are the parent bodies of ordinary chondrites and that space weathering exists and changes the spectral properties of asteroid surfaces was conclusively confirmed. Orbiting spacecraft missions such as NEAR Shoemaker and Hayabusa that obtain reflectance spectra and have instruments for making chemical analyses (e.g., x-ray spectrometer) of the surface can argue strongly for particular compositions. McCoy et al. (2001) argued that all the data supported an ordinary chondrite mineralogy for Eros, but the conclusions were not conclusive. Abe et al. (2006) found that the reflectance spectra of Itokawa were consistent with an LL chondrite mineralogy and Okada et al. (2006) argued that x-ray data were consistent with Itokawa having an LL or L chondrite mineralogy. Without the Hayabusa sample return, it would have been impossible to know conclusively Itokawa’s surface mineralogy. Rosetta’s flyby of Lutetia, which obtained reflectance spectra but no geochemical information due to the instruments being designed for a comet encounter, was not able to conclusively argue for any surface mineralogy (Coradini et al., 2011; Sierks et al., 2011) due to the absence of any characteristic

Asteroids

Table 2

403

Postulated best taxonomic class analogs for each meteorite type

Type

Fall percentage

Best taxonomic class analogs

L H LL HED Iron CM L/LL Aubrite EH EL CV Mesosiderite CO Ureilite CI Martian Pallasite C2-ungrouped CR H/L Acapulcoite/lodranite CK Angrite C3-ungrouped CB K R Winonaite Brachinite CH Lunar

36.7 33.8 8.1 5.9 4.7 1.5 1.1 0.9 0.9 0.8 0.7 0.7 0.6 0.6 0.5 0.5 0.4 0.3 0.3 0.3 0.2 0.2 0.1 0.1 0.1 0.1 0.1 0.1 All finds All finds All finds

Q-type (e.g., Binzel et al., 2004b), S-complex (e.g., Gaffey et al., 1993) Q-type (e.g., Binzel et al., 2004b), S-complex (e.g., Gaffey et al., 1993) Q-type (e.g., Binzel et al., 2004b), S-complex (e.g., Nakamura et al., 2011) V-type (e.g., Consolmagno and Drake, 1977; McCord et al., 1970) M-type (e.g., Cloutis et al., 1990; Shepard et al., 2010) C-complex (e.g., Burbine, 1998; Cloutis et al., 2011b; Vilas and Gaffey, 1989) Q-type (e.g., Binzel et al., 2004b), S-complex (e.g., Gaffey et al., 1993) E-type (e.g., Clark et al., 2004a; Zellner, 1975; Zellner et al., 1977) M-type (e.g., Chapman and Salisbury, 1973; Shepard et al., 2010) M-type (e.g., Gaffey and McCord, 1978; Shepard et al., 2010) K-type (e.g., Bell, 1988; Burbine et al., 2001a) M-type (e.g., Shepard et al., 2010), S-complex (Gaffey et al., 1993) K-type (e.g., Bell, 1988; Clark et al., 2009) C-complex (e.g., Jenniskens et al., 2009), S-type (Gaffey et al., 1993) C-complex (e.g., Cloutis et al., 2011a; Johnson and Fanale, 1973) Mars (e.g., Bogard and Johnson, 1983) A-type (e.g., Cruikshank and Hartmann, 1984; Sunshine et al., 2007) D-type (e.g., Hiroi et al., 2001b), T-type (Hiroi and Hasegawa, 2003) C-complex (e.g., Hiroi et al., 1996; Sato et al., 1997) Q-type (e.g., Binzel et al., 2004b), S-complex (e.g., Gaffey et al., 1993) S-complex (e.g., Gaffey et al., 1993) K-type (e.g., Clark et al., 2009; Cloutis et al., 2012b) S-complex (e.g., Rivkin et al., 2007) K-type (e.g., Clark et al., 2009) M-type (e.g., Shepard et al., 2010) C-complex (e.g., Gaffey, 1980) A-type (e.g., Sunshine et al., 2007) S-complex (e.g., Gaffey et al., 1993) A-type (e.g., Cruikshank and Hartmann, 1984; Sunshine et al., 2007) M-type (e.g., Shepard et al., 2010) Moon (e.g., Marvin, 1983)

The fall percentages are calculated using data for classified meteorites from the Meteoritical Bulletin Database (2012).

absorption bands. Dawn, which orbited Vesta and will orbit Ceres, has instruments for obtaining chemical information (gamma ray/neutron detector) and will be observing bodies with strong absorption bands in their visible and infrared spectrometer’s wavelength range (
0.25–5 mm). The next step is the OSIRIS-REx sample return mission to Btype asteroid (101955) 1999 RQ36 (Clark et al., 2011). This body appears definitely to be primitive (limited degree of heating) due its extremely low visual albedo (Emery et al., 2010) and its blue spectral slope in the visible and nearinfrared, which is consistent with a magnetite-rich mineralogy similar to CI chondrites. NEA 1999 RQ36 appears to be the perfect target for sampling pristine organic material that may have seeded life on Earth. The OSIRIS-REx will launch in 2016 and return back to Earth with a sample in 2023. Ground-based spectroscopy and ground- (e.g., SDSS) and space-based (e.g., Hubble Telescope, NEOWISE) photometry will still be the backbone for trying to determine asteroid mineralogies since, in the best of times, only a few bodies will be directly sampled in the near future. Every meteorite type has plausible parent bodies in the main belt (Table 2). Sample return missions to a few different taxonomic types would confirm or refute many of these postulated linkages. Do E-types have aubrite mineralogies? Do V-types have mineralogies and oxygen isotopic compositions similar to most

HEDs? Do the high radar albedo M-types have compositions similar to iron meteorites or do they contain significant abundances of silicates on their surfaces? Do some M-types have mineralogies similar to Ch, CB, and enstatite chondrites? Do Ch-types have mineralogies similar to CM chondrites? Are D-, P-, and T-types organic-rich? For example, the primitive achondritic ureilites were commonly thought to be present among the S-types (e.g., Gaffey et al., 1993), but the recovery of samples from the 2008 TC3 (Jenniskens et al., 2009) showed that ureilites are present among C-complex bodies, as proposed by Cloutis and Hudon (2004). Currently it is only feasible to retrieve samples from NEAs, but this is not a serious problem since almost every taxonomic type is represented in the NEA population (Binzel et al., 2004b). Considering the advances in knowledge of asteroid chemistries since the start of the millennium, the rest of the century is expected to be even more productive. In the near future, astronauts may even physically grab rocks from NEA surfaces (e.g., Elvis et al., 2011).

Acknowledgments This publication utilizes meteorite and mineral reflectance spectra acquired by Takahiro Hiroi and Timothy McCoy

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using the NASA RELAB facility at Brown University and Roger Clark and coworkers at the USGS Denver Spectroscopy Lab. The bulk chemical meteorite database to compare to x-ray observations of asteroids was made available by Larry Nittler. Unless otherwise noted, near-infrared asteroid data utilized in this publication were obtained and made available by The MITUH-IRTF Joint Campaign for NEO Reconnaissance. The IRTF is operated by the University of Hawaii under Cooperative Agreement no. NCC 5–538 with the National Aeronautics and Space Administration, Office of Space Science, Planetary Astronomy Program. The MIT component of this campaign is supported by NASA grant 09-NEOO009-0001, and previously by the National Science Foundation under Grant No. 0506716. The author would like to thank Andrew Davis for many helpful comments.

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