Remote Detection of Clay Minerals

Janice L Bishop; Joseph R Michalski; John Carter

doi:10.1016/b978-0-08-100355-8.00014-x

. Author manuscript; available in PMC: 2021 May 26.

Published in final edited form as: Dev Clay Sci. 2017 Oct 13;8:482–514. doi: 10.1016/b978-0-08-100355-8.00014-x

Remote Detection of Clay Minerals

Janice L Bishop ^1,^*, Joseph R Michalski ², John Carter ³

PMCID: PMC8152337 NIHMSID: NIHMS1048790 PMID: 34045934

Abstract

Spectral remote sensing in the visible/near-infrared (VNIR) and mid-IR (MIR) regions has enabled detection and characterisation of multiple clays and clay minerals on Earth and in the Solar System. Remote sensing on Earth poses the greatest challenge due to atmospheric absorptions that interfere with detection of surface minerals. Still, a greater variety of clay minerals have been observed on Earth than other bodies due to extensive aqueous alteration on our planet. Clay minerals have arguably been mapped in more detail on the planet Mars because they are not masked by vegetation on that planet and the atmosphere is less of a hindrance. Fe/Mg-smectite is the most abundant clay mineral on the surface of Mars and is also common in meteorites and comets where clay minerals are detected.

Keywords: IR, Raman, IR spectroscopy, Raman spectroscopy, Clay mineral, Clay minerals, Remote sensing, VNIR, TIR, Reflectance, Emission, Spectroscopy, Mars, Ceres, Meteorites, Comets

14.1. PRESENCE OF CLAY MINERALS IN OUR SOLAR SYSTEM

Clay minerals are not only found on Earth, but also on our neighbouring planet Mars and in meteorites, asteroids and comets. These minerals can be identified through several techniques on Earth, most commonly by X-ray diffraction (XRD) or visible/near-infrared (VNIR) reflectance spectroscopy. However, identification and characterisation of clay minerals on Mars is largely carried out using orbital VNIR spectra from ~0.4 to 5 μm. Thermal infrared (TIR) spectra collected in the mid-IR (MIR) region from ~200 to 2000 cm⁻¹ (~5–50 μm) also provide orbital information about mineralogy including clay minerals. The Mars Science Laboratory (MSL) rover uses an XRD instrument called CheMin to identify clay minerals on the surface at the landing site inside Gale Crater (Blake et al., 2013) and several martian surface missions have measured surface chemistry and this data was then used to infer mineralogy (Baird et al., 1977; Toulmin III et al., 1977; McSween et al., 2009). Detections of clay minerals on asteroids such as Ceres (De Sanctis et al., 2015) and Jupiter-family comets (Lisse et al., 2006) are typically achieved through NIR and MIR spectroscopy because these small bodies are too far away for other techniques. Many more detections of clay minerals in meteorites have been made than in asteroids, likely due to the ability to study meteorites in the lab and challenges resulting from space weathering on asteroid surfaces. Clay minerals are frequently identified in meteorites using microprobe analyses of thin sections.

The history of clay minerals on the planet Mars has experienced a rather disjointed path towards our current understanding. Elemental analyses of surface material by Viking first led scientists to believe clay minerals were likely present on Mars. Modelling of the major element data was found to be consistent with 60–80 wt.% smectite (Baird et al., 1977; Toulmin III et al., 1977). This was based on the major element chemistry of the surface material and also the expectation that alteration of volcanic ash and tephra on Mars would form clay minerals as observed on Earth. Telescopic VNIR and TIR spectra of Mars followed in the 1980s and 1990s without clear evidence of any clay minerals (McCord et al., 1982; Singer, 1985; Bibring et al., 1990; Bell et al., 1994). As the signal to noise ratio of these data improved and the spot size became smaller and no clay minerals were found, scepticism about clay minerals on Mars arose. Further remote sensing of Mars with the Thermal Emission Spectrometer (TES) instrument on Mars Global Surveyor beginning in the late 1990s (Christensen et al., 2001) mapped the martian surface at 3–5 km spot size and still did not find evidence of clay minerals. It was not until the Observatoire pour la Minéralogie, l’Eau, les Glaces et l’Activité (OMEGA) VNIR imaging spectrometer on Mars Express started mapping Mars at 1–3 km spot sizes in 2004 that clay minerals could be definitively identified on that planet (Poulet et al., 2005). As the Compact Reconnaissance Imaging Spectrometer for Mars (CRISM) instrument on the Mars Reconnaissance Orbiter began collecting targeted high resolution (18 m/pixel) hyper-spectral imagery in 2006, windows of small and large clay mineral-bearing outcrops were discovered nearly everywhere that the ancient rocks were exposed on the surface (Mustard et al., 2008; Murchie et al., 2009).

14.2. REMOTE DETECTION OF CLAY MINERALS

Remote sensing of clay minerals is primarily performed using VNIR and TIR instruments. Similar instruments are used on Earth, Mars and other planetary bodies. Remote detection of clay minerals in the MIR region with TIR spectra uses Si–O stretching (v(SiO)) and bending (δ(SiO)) vibrations, and M–OH bending δ(MOH) vibrations (where M is a metal cation such as Al, Fe³⁺, Fe²⁺ or Mg), while NIR clay mineral detections focus mostly on M–OH stretching plus bending combinations ((v + δ)MOH) and stretching overtones (2v(OH)). Similar modes also occur for H₂O in phyllosilicates. These include a stretching plus bending combination band ((v +δ)H₂O) and a stretching overtone (2v (H₂O)). The extended visible region covers iron absorptions and is particularly helpful in discriminating between Fe³⁺ and Fe²⁺ in clay minerals. The MIR region has the advantage that the spectral features are fundamental vibrations and the bands are strong; however, a disadvantage is that other silicates exhibit bands in the same wavelength range and identifying clay minerals in the presence of quartz, feldspar and pyroxene can be challenging. An advantage of the NIR region is that there are few other spectral features at these wavelengths so detection is easier. In addition, there are bands near 1.4, 1.9 and 2.1–2.5 μm that can be evaluated jointly for added confidence in the clay mineral detection. Band assignments were confirmed using comparison of the NIR overtones and combinations with the MIR fundamental stretching and bending vibrations from MIR transmittance spectra (Bishop et al., 2002c; Petit et al., 2004b). While MIR transmission spectroscopy uses an active radiation source, MIR remote sensing is passive. The thermal emission spectra used in remote sensing studies includes components from absorption as well as scattering, so they are more complex than transmittance spectra, which include only the absorption component. Reflectance spectra are often easier to measure in the laboratory, but they are not always directly comparable to thermal emission data. Further complicating matters is the reality that all VNIR and MIR spectra of planetary surfaces contain particle size scattering effects that cannot be completely reproduced in any laboratory setting. However, the spectral contrast, shapes and absorption positions measured on planets provide strong clues regarding the effects of particle size, lending confidence to clay mineral detections using purely IR techniques.

14.2.1. VNIR Bands Used for Detection of Clay Minerals

Detection of clay minerals in the VNIR region varies with mineral structure and cation type (Fig. 14.1). Several studies have investigated the VNIR reflectance spectra of pure clay minerals for remote sensing studies. These include studies of smectites (Cariati et al., 1981; Post, 1984; Clark et al., 1990; Post and Noble, 1993; Bishop et al., 1994, 1999, 2002a,c, 2008a, 2011; Post et al., 1997; Gates, 2005; Decarreau et al., 2008), kaolin-serpentine group minerals (King and Clark, 1989; Clark et al., 1990; Petit et al., 1999b; Bishop et al., 2002c, 2008a,b), chlorites (King and Clark, 1989; Bishop et al., 2008a), micas (Bishop et al., 2008a), talc (Petit et al., 2004a), and nanophase/poorly crystalline clay minerals such as allophane and imogolite (Bishop et al., 2013a). In remote sensing studies the materials measured are always mixtures of multiple components. Thus, studies of the spectral properties of natural mixtures of clay minerals (Cuadros et al., 2015; Michalski et al., 2015), mixtures prepared from clay minerals and other minerals (e.g. Madejová et al., 2002b; McKeown et al., 2011), and clay mineral-bearing rocks (e.g. Bishop et al., 2002d; Hamilton et al., 2008b; Ehlmann et al., 2012) are important as well.

14.2.1.1. VNIR Characterisation of Smectites

Smectites have received the most study of clay minerals in this spectral region and have clearly identifiable bands for their identification. All smectite spectra exhibit an H₂O stretching overtone (2v(H2O)) at 1.41 μm and an asymmetric H₂O combination ((v + δ)_asH₂O) band centred at 1.91 μm (Fig. 14.1, Bishop et al., 1994). The MOH features depend on the presence of Al, Fe³⁺ or Mg cations in the octahedral sites, as well as Al and Fe³⁺ substitutions in the tetrahedral sites, as observed in spectra of the Clay Mineral Society standards SAz-1, SWy-1, SWa-1, SHCa-1, SapCa-1 and others. Montmorillonites with high Al abundance in the octahedral sites (e.g. SAz-1) exhibit an OH combination band near 2.20 μm, but this can be extended toward 2.21 μm for montmorillonite with some Fe³⁺ or Mg in the octahedral sites (e.g. SWy-1). The AlOH stretching overtone (2v(Al₂OH)) lies at 1.41 μm, at the same wavelength as the 2v(H₂O), which was a source of confusion for decades in making band assignments. Beidellite spectra exhibit 2v(Al₂OH) vibrations at 1.40 and 2.18 μm (Bishop et al., 2011), which are shorter wavelengths than the bands observed in montmorillonite spectra due to increased tetrahedral Al substitution for Si. Nontronite spectra include a 2v(Fe³⁺₂OH) band at 1.43 μm and a (v + δ)Fe³⁺₂OH band at 2.29 μm. Spectra of nontronite containing some Al³⁺ in octahedral sites (e.g. SWa-1) have an additional shoulder or weak (v + δ)AlFe^{3 +}OH band near 2.23 μm. Hectorite (e.g. SHCa-1) and saponite (e.g. SapCa-1) spectra include an MgOH stretching overtone (2v(Mg₃OH)) band at 1.39 μm and an MgOH combination ((v + δ)Mg₃OH) band at 2.31 μm. For spectra of Fe-and Mg-rich smectites, the 2v(OH) bands are distinct from the 2v(H₂O) bands, which aides in their identification in remote sensing data. Smectites also include an additional OH band at longer wavelengths (~2.4–2.5 μm) that depends on the octahedral cations. This occurs at 2.38 μm for saponite, 2.39 μm for hectorite, 2.41 for nontronite, 2.44 μm for beidellite, and 2.45 μm for montmorillonite.

14.2.1.2. VNIR Characterisation of Kaolinites-Serpentines

VNIR spectra of kaolinite and serpentine are readily distinguished from spectra of smectites because the kaolin-serpentine group clays do not have H₂O bands, except for halloysite. Kaolinite spectra exhibit a sharp pair of bands due to the 2v(Al₂OH) stretching overtone at 1.396 and 1.416 μm and a (v + δ)Al₂OH doublet at 2.17 and 2.21 μm. The halloysite spectrum is similar to that of kaolinite, but also includes a water band near 1.9 μm due to H₂O in the structure (e.g. Clark et al., 1990). Serpentine spectra include a 2v(Mg₃OH) overtone at 1.39 μm and a (v + δ)Mg₃OH combination band near 2.33 μm. Many serpentine spectra also include bands near 2.52–2.57 μm.

14.2.1.3. VNIR Characterisation of Chlorites

VNIR spectra of chlorites include the expected OH bands near 1.39 and 2.33–2.35 μm and the band positions vary depending on the relative abundance of Fe³⁺ and Mg; however, they are distinguished from serpentines by the presence of an additional (v + δ)AlFe²⁺MgOH band near 2.25 μm. The Mg-rich chlorites such as clinochlore have spectral bands near 1.39, 2.25 and 2.33 μm, while Fe²⁺-rich chlorites such as chamosite have spectral bands near 1.42, 2.26 and 2.36 μm. Also, the 2v(Fe²⁺₃OH) band for chamosite is weaker in intensity, likely because of the strongly absorbing Fe²⁺ absorptions from ~0.7 to 1.2 μm. Sometimes H₂O bands are observed near 2.0 μm in chlorite spectra, rather than near 1.9 μm as in most hydrated clay minerals. Spectra of some chlorites also contain a band near 2.48–2.57 μm.

14.2.1.4. VNIR Characterisation of Micas, Illites and Other Clay Minerals

Micas tend to exhibit much weaker NIR bands from 1.4 to 2.5 μm than smectites, kaolin-serpentines, and chlorites, and the bands near 1.4 and 1.9 μm are generally too weak to be useful for remote detection. The OH combination vibrations typically occur as multiple bands in micas and vary with the octahedral cation composition. Al-rich micas like zinnwaldite have spectral bands near 2.20 and 2.25 μm, while celadonite has bands near 2.25, 2.30 and 2.35 μm due to multiple cations. Biotite with Al and Fe³⁺ cations has bands near 2.25 and 2.35 μm, while glauconite with Fe³⁺ and Fe²⁺ cations includes spectral bands near 2.30 and 2.36 μm. Additional clay minerals such as illite, palygorskite and sepiolite all exhibit spectral features due to both OH and H₂O in their structures. Illite spectra have bands near 1.41, 1.91, 2.22, 2.35 and 2.44 μm. Palygorskite spectra contain bands near 1.41–1.42, 1.91, a doublet at 2.18 and 2.22, and a weak band near 2.25 μm. The VNIR spectra of sepiolite include features near 1.38, 1.42, 1.91, 2.18 and 2.31 μm due to both Al and Mg in its structure. The VNIR spectra of Mg-talc include a band at 1.39 μm, a doublet at 2.29 and 2.31 μm, and additional bands near 2.39 and 2.46 μm.

14.2.1.5. VNIR Characterisation of Poorly Crystalline Clay Minerals

The importance of poorly crystalline clay minerals including allophane and imogolite as soil components and proto-clays was recognised long ago (Wada et al., 1972; Farmer et al., 1979, 1983; Parfitt and Furkert, 1980). More recently, the importance of these for remote sensing on Mars has been recognised (Rampe et al., 2012; Bishop and Rampe, 2016). The nanophase aluminosilicates allophane and imogolite have VNIR spectral features similar to those of Al-rich smectites (AlSm) and opal. The primary difference is broadened spectral features for the nanophase clay minerals due to a distribution of Al–OH and Si–OH sites in the structure. Another difference is that the (v + δ) H₂O band is centred at 1.92 μm instead of 1.91 μm. The corresponding MOH bands are also shifted. The 2v(Al₂OH) bands occur at 1.38 and 1.40 μm for allophane and at 1.37 and 1.39 μm for imogolite, while the (v + δ)Al₂OH band occurs near 2.19 μm for both of these materials.

14.2.2. MIR Bands Used for Detection of Clay Minerals in TIR Spectra

The properties of transmittance spectra of clay minerals are described in detail in previous chapters. Transmittance spectra of fine powders depend on the imaginary component of the complex index of refraction k, while reflectance and emission spectra depend on both the real n and imaginary k components of the index of refraction, and are thus susceptible to the effects of surface scattering. Because MIR remote sensing data are measured as TIR, that is the best technique for comparison of lab data to remote sensing data. Lab emissivity spectra have been acquired of numerous samples including clay minerals for comparison with the Thermal Emission Spectrometer (TES) data of Mars (Christensen et al., 2000a). Detailed lab studies of clay minerals (Michalski et al., 2005, 2006a,b) and related materials such as zeolites (Ruff, 2004), glass (Minitti et al., 2002; Byrnes et al., 2007) and amorphous phases (Kraft et al., 2003; Rampe et al., 2012) were performed, although the majority of TIR lab studies focus on other minerals. Hemispherical reflectance spectra can often be used to calculate emission spectra using Kirchhoff’s Law: E = 1 – R; however, this does not always hold true for small particle sizes or if the surface is not a Lambertian scatterer (e.g. Salisbury et al., 1994). Reflectance spectra collected using a biconical system can provide only an approximation of emission spectra. For finely particulate samples it is often easier to obtain high-quality reflectance spectra in the lab than emissivity spectra. Clay minerals are typically pressed into powders in order to improve the band strength in emission spectra (Michalski et al., 2005). Pressing finely grained materials into powders removes pore space and reduces internal scatter, which strengthens the Reststrahlen bands (reflectance maximum or emission minimum). Most importantly, clay minerals occurring naturally within rocks (e.g. mesostatic replacement in volcanic rocks or compacted clays in clastic rocks) behave spectroscopically like coarse materials or pressed pellets whereas clays in dust behave more like fine particulates. TIR spectral studies of clay-bearing rocks were also performed (e.g. Wyatt et al., 2001; Michalski et al., 2004; Hamilton et al., 2008b; Ehlmann et al., 2012; Rampe et al., 2013) in order to support identification of clays on planetary surfaces.

14.2.2.1. Si–O Stretching Vibrations

The strongest bands for identification of clay minerals in thermal remote sensing typically result from vibrations of the tetrahedral sheet Si(Al,Fe)O₄ (Michalski et al., 2006b). However, related SiO₄ vibrations are observed in spectra of other silicates as well. Salisbury (1993) describes the Christiansen Feature (CF) as a reflectance minimum or emission maximum located at the short-wavelength edge of the v(SiO) Reststrahlen feature. The CF occurs where the imaginary component of the index of refraction approaches 0 and the real component approaches 1. Salisbury (1993) noted shifts in the wavelength of this feature with mineral composition. Typically the CF occurs at longer wavelengths for clay minerals compared to mafic silicates (Salisbury, 1993). Michalski et al. (2005) observed a shift in the v(SiO) in spectra of clay minerals, aluminosilicate glasses and altered phases, such that the band near 1000 cm⁻¹ transitions towards lower wavenumbers (longer wavelengths) with increasing abundance of tetrahedral Al and Fe³⁺ and decreasing abundance of Si. This is consistent with the trends observed by Salisbury et al. (1991b) for a large collection of silicate minerals and with observed transmittance spectra of clay minerals (Farmer, 1974b). For example, Michalski et al. (2005) noted emissivity minima of 8.8 μm (1135 cm⁻¹) for SWy-1 montmorillonite with an Si/O ratio of 0.399, and of 9.5 μm (1056 cm⁻¹) for NAu-1 nontronite with an Si/O ratio of 0.349.

The Si(Al,Fe)O₄ vibrations observed in emission spectra of several clay minerals are shown in Fig. 14.2. The v(SiO) band centre exhibits a trend of increasing wavelength across the smectites as the octahedral cation changes from Al to Fe³⁺ to Mg. The Si–O stretching band position for Al-rich smectite is similar to that observed for kaolinite and aluminosilicate gel. The Si–O stretching band occurs at longer wavelengths for the poorly crystalline allophane and imogolite and also for clay minerals rich in Fe³⁺ and Mg (Fig. 14.2).

14.2.2.2. Si–O Bending Vibrations

The bending vibrations for tetrahedral SiO₄ groups as well as lattice deformation modes such as Si–O-M_oct absorptions, occur near 18–25 μm (550–400 cm⁻¹) in emission spectra when little or no tetrahedral substitution occurs (Michalski et al., 2005; Salisbury et al., 1991b). Spectra of serpentines, chlorites and biotite contain a strong SiO₄ bending vibration centred near 20.6–21.3 μm (485–470 cm⁻¹). In TOT (tetrahedral-octahedral-tetrahedral configuration) clay minerals, bands associated with Si–O–M_oct absorptions (δ(SiOMg_oct)) partially overlap with the Si–O bending (δ(SiO)) absorption. In trioctahedral clay minerals, the overlap between δ(SiO) and δ(SiOMg_oct) is significant, resulting in a single, strong band located near 20.8 μm (480 cm⁻¹), the components of which can only be resolved through modeling. However, for dioctahedral clay minerals, the δ(SiOAl_oct) and δ(SiOFe³⁺_oct) bands shift to shorter wavelengths (18.6–19.6 μm or 540–510 cm⁻¹), separating from the δ(SiO) bands. These absorptions are key because they can be used to distinguish different clay mineral compositions using long-wavelength thermal emission remote sensing data (Michalski et al., 2010a). Coordinated bending also occurs along the Si–O–M_oct bonds from the tetrahedral Si–O into octahedral O–M bonds and is responsible for the weaker bands that often occur near 15–18 μm (670–550 cm⁻¹) for Al and Fe³⁺ cations and near 22–24 μm (450–400 cm⁻¹) for Mg cations. Si–O–Mg_oct vibrations are particularly well-resolved in the reflectance spectra of the Mg-rich minerals chrysotile and clinochlore (Fig. 14.2).

14.3. CHARACTERISATION OF CLAY MINERALS ON EARTH

It is out of necessity that remote sensing (VNIR and MIR) is used to study clay minerals beyond Earth. However, remote sensing is also a useful technique for clay mineral identification and mapping on our home planet as well. In fact, the ability to identify clay minerals and interpret their compositions remotely on Earth lends confidence in the capability to do so on other planets. Just as is the case in planetary exploration, remote sensing provides a tool for mapping distant, rugged terrains that are difficult or impossible to access in the field. Remote sensing data, when calibrated by careful sample analysis and field measurements, can allow for extrapolation of detailed information from a local area to mapping of vast study sites. In addition, remote sensing provides opportunities to identify and map alteration mineralogy, which might be obvious in IR data, but difficult to identify in the field.

14.3.1. Challenges of Remote Detection of Clay Minerals on Earth

IR spectroscopy is a powerful tool for clay mineral analyses in the laboratory; in part because significant amounts of radiation can be supplied using an active source in either reflectance or transmission. Having large numbers of photons means that the measured spectrum can be split into many channels, achieving the high spectral resolution required to resolve the precise position of specific absorptions without sacrificing much in terms of signal-to-noiseratio (SNR). This is generally not the case in a remote sensing setting where a sensor can measure only reflected sunlight or emitted heat radiation. Therefore, remote sensing instruments are designed to maximise the science return through judicious use of available photons.

Remote sensing instruments are therefore designed as either ‘multispectral’ (a few to a few tens of channels) or ‘hyperspectral’ (hundreds of channels). Hyperspectral instruments typically offer spectral channels comparable to the spectral sampling of laboratory instruments, but with much lower SNR than what is achieved in the lab setting. Multispectral instruments have fewer channels, but they are strategically located to collect as much relevant information as possible. The photon budget also governs tradeoffs in spatial sampling (pixel size); all else being equal, smaller pixels result in lower SNR.

Spatial sampling is an important factor in remote sensing. In the laboratory, the spot size of an IR measurement can be from <1 mm² to 1 cm² or more. While significantly larger than the clay mineral grains of interest, this ‘pixel size’ is small enough to provide confidence in exactly what part of the sample is being measured. In the case of remote sensing a single pixel can contain several surfaces as different as bedrock, regolith, soil, vegetation, water and man-made materials. An advantage of smaller pixel size is to minimise this effect; the best-case scenario in IR remote sensing of Earth is typically ~10 m².

IR spectral analysis of clay mineral-bearing samples is also affected by spectral sampling as a function of depth, in addition to spatial resolution. Just as in the laboratory, reflected NIR or emitted TIR radiation samples only the uppermost 1–100 μm in most cases. Because of this effect, thin coatings can obscure the dominant mineralogy of rocks in remote sensing studies. Clay minerals compose a significant fraction of some siliceous coatings present on Earth (e.g. desert varnish) and could be important in some remote sensing settings beyond Earth. Clay mineral-bearing rocks are also in some cases masked by coatings and not detected via remote sensing.

The second major difference between laboratory and remote sensing IR studies is that environmental conditions are controlled in the laboratory setting to minimise or eliminate absorptions and scattering from gases and aerosols. The Earth’s atmosphere contains abundant water vapour, most of which occurs in the troposphere and is unavoidable in any remote IR measurement. The most significant absorptions are due to H₂O and occur at 0.718, 0.810, 0.935, 1.13, 1.38, 1.88 and 2.68 μm (e.g. Gao et al., 1993; Farrand et al., 1994; Kruse, 2004). Other major absorptions affecting IR remote sensing are CO₂ absorptions at ~1.4, 1.6 and 2.0 μm, the O₃ absorption at 9.6 μm (e.g. Gao et al., 1993), and aerosol scattering throughout. Complex physical models exist for separation of atmospheric components from remote sensing data in order to retrieve spectra of the surface. Due to the complexity of Earth’s atmosphere, remote spectral observations of other planetary bodies can be more straightforward to interpret. A hyperspectral image of Meteor Crater, Arizona collected by the Airborne VIR Imaging Spectrometer (AVIRIS) (e.g. Porter and Enmark, 1987; Green et al., 1990) illustrates the power of spectral imaging over a planetary analogue site (Fig. 14.3). The atmospherically corrected data show the limitations imposed by the atmospheric windows.

FIG. 14.3 — An example of clay mineral detection in a terrestrial impact crater (Meteor Crater) using hyperspectral AVIRIS data. (A) Approximate true colour image of the 1.2 km diameter crater. (B) A principal component image where the fourth, third and second principal components are displayed as RGB, respectively. The colours demonstrate real compositional differences, though they do not have literal meaning. Pink colours in the uppermost crater rim correspond to the Moenkopi (M) Formation. Yellow-red colours underlying the Moenkopi Formation in the mid-crater wall correspond to the Kaibab (K) formation. Green colours refer to the Coconino (Co) Sandstone. Blue colours in the crater correspond to alluvium and lacustrine sediments. (C) Atmospherically corrected spectra of several sites around the crater compared with a lab spectrum of kaolinite. Black arrows mark the positions of atmospheric H₂O bands that were removed. A grey line marks the position of the (*v + δ*)Al₂OH kaolinite band at 2.21 μm.

Many other important differences between laboratory and remote sensing analyses exist, but they are all secondary to the issues of SNR, spectral sampling, spatial sampling and atmospheric effects. Despite all of the challenges and realities outlined above, remote sensing provides a tremendous opportunity for studying clay mineralogy on Earth. Remotely collected IR data reveal the presence of clay minerals in sites where they were previously unrecognised in the field, they show patterns of the occurrence of clay minerals that are surprising, and that challenge geological assumptions. The data provide a powerful tool for characterising clay mineralogy and constraining crystal chemistry over large areas. In addition, remote sensing of clay minerals on Earth lends tremendous confidence to the remote detection of clay minerals on Mars and other planets—environments where atmospheric effects are lacking or minimal, and where surface vegetative cover, water bodies, and the presence of man-made materials are not concerns.

14.3.2. Instruments and Datasets Available for IR Remote Sensing of Clay Minerals on Earth

Terrestrial remote sensing can be carried out from orbital or airborne platforms (Table 14.1). While an orbital platform provides significant advantages in terms of global coverage, there are substantial tradeoffs with respect to spatial resolution and SNR. Most orbital IR instruments to date have been multispectral, with the exception of Hyperion, a hyperspectral imager aboard the EO–1 satellite (e.g. Kruse et al., 2003; Pearlman et al., 2003). Hyperion was able to image nearly anywhere on the globe at 30 m/pixel spatial resolution, but at the expense of very low SNR. These data can be used for qualitative maps showing compositional diversity, but they are not useful for clay mineral identification. ASTER and LANDSAT satellites have provided global multispectral NIR/short-wave IR (SWIR) coverage of the Earth’s surface at 30 m/pixel resolution. Due to the multispectral nature of these data, spectral bands associated with specific minerals cannot be identified. However, they can provide useful information through broad spectral patterns and these data can be used for mapping alteration, and potentially outcrops of clay minerals through coordination with hyperspectral sensors or ground testing. Such techniques have been developed (Viviano and Moersch, 2012) for detection of clay minerals on Mars using the Mars Odyssey thermal emission imaging system (THEMIS) multispectral sensor (Christensen et al., 2003).

TABLE 14.1.

Instruments and Datasets Available and Planned for Remote Analysis of Clay Minerals on Earth

Sensor Name	# Bands	Wavelength Range (μm)	Spectral Resolution	Spatial Resolution (m/pixel)

Orbital
Landsat 8	9	0.43–2.3	Multi	15–30
ASTER	14	0.52–11.65	Multi	15–90
Hyperion	220	0.4–2.5	Hyper	30
HyspIRI	212	0.38–2.5	Hyper	60
HyspIRI	8	3–12	Multi	60
Airborne
AVIRIS	224	0.4–2.5	Hyper	5–50
NS001	8	0.46–12.3	Multi	5–50
TIMS	6	8.2–12.2	Multi	5–50
MASTER	50	0.4–13	Multi	5–50
HyMap	128	0.45–2.5	Hyper	3–10
HyTES	256	7.5–12	Hyper	2–10

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Note: only the most recent Landsat imager is included in this table; the spatial resolution of airborne sensors depends on the height at which they are flown; HyspIRI is in the planning stage at the time of this writing.

Sensors flown on airborne platforms can generally achieve higher SNR and smaller pixel size by flying at low altitude. The major tradeoff is that the data need to be deliberately collected, which involves planning flights. In other words, the data are only available when targeted for a specific site. Over the last 25 years, AVIRIS has been the primary remote sensing instrument for IR clay mineral mapping (e.g. Kruse et al., 1993; Boardman et al., 1995; Clark et al., 1995; Clark et al., 2002; Crowley and Zimbleman, 1997; Chabrillat et al., 2002; Guinness et al., 2007). A more recently developed instrument, HyMap provides similar data (e.g. Cocks et al., 1998) and has been used as well for detection of clays (e.g. Brown et al., 2005; Nunez et al., 2015).

14.3.3. Remote Characterisation of Clay Minerals on Earth

A number of authors have used remote sensing to study clay minerals on Earth (e.g. Crowley and Zimbleman, 1997; Chabrillat et al., 2002; Clark et al., 2003; Guinness et al., 2007; Seelos et al., 2009; Swayze et al., 2014). A driving factor in remote sensing of clay minerals has been to identify alteration patterns linked to economic mineralisation in rugged, complex terrains (e.g. van der Meer et al., 2012). The global reach of multispectral datasets such as ASTER has proven fruitful for the identification and basic characterisation of hydrothermal alteration associated with ore deposits (e.g. Bakardjiev and Popov, 2015), but hyperspectral mapping by airborne sensors has revealed a tremendous amount of diversity in hydrothermally altered, mineralised terrains (e.g. Swayze et al., 2014; van Ruitenbeek et al., 2012). Clay mineral-bearing soils were mapped along the Front Range Urban Corridor in Colorado using AVIRIS and HyMap imagery (Chabrillat et al., 2002). This study illustrated the utility of hyperspectral imaging for identification of distinct clay minerals such as smectite, inner stratifications of smectite with illite and kaolinite-bearing soils. A set of band shape algorithms termed Tetracorder have also been developed for identification of clays and clay minerals and other minerals using band centre and asymmetry (Clark et al., 2003).

While hydrothermal deposits associated with economic mineral outcrops might be a driver for remote exploration, studies of alteration relevant to environmental issues are also carried out via remote sensing. For example, Crowley and Zimbleman (1997) used AVIRIS data to map various clay minerals on Mount Rainier, in areas that are extremely difficult to access through field work alone. These highly altered terrains are most susceptible to slope failure, leading to landslide hazards.

Remote sensing studies of clay minerals have also been carried out in order to better understand clay minerals detections beyond Earth. For example, Mustard and Pieters (1987) used data from an airborne hyperspectral sensor to identify and map alteration units within serpentinised terrain on the Colorado Plateau. Studies of weathering patterns and occurrences of clay minerals in igneous rocks on Earth has helped to understand how the break-down of primary minerals impacts interpretations of bedrock lithology on the Earth and Mars (Michalski et al., 2004). Spectral analyses of alteration associated with Hawaiian cinder cones and calderas shed light on the detection of silica and clay minerals in analogous environments on Mars (Guinness et al., 2007; Seelos et al., 2009).

14.3.3.1. Clays and Clay Minerals Detected in Hyaloclastites in Askja, Iceland

The central Icelandic desert contains youthful, well-preserved examples of landforms that occur when volcanoes interact with glacial ice. These include tuyas (table mountains) and tindars or moberg ridges that form beneath thick ice (Allen, 1979; Chapman and Smellie, 2007). The products of these subice eruptions contain intensely fragmented volcanic glass that has largely been altered to smectite, oxides and oxyhydroxides, and minor carbonates and sulphates (Honnorez, 1981; Furnes, 1978; Jakobsson, 1978). Examples of palagonitised hyalotuffs were investigated in situ at Hlö∂ufell tuya and Thórólfsfell ridge (Bishop et al., 2002d) and examples of clay mineral-bearing altered basalts were characterised from flows near Hvalfjordur and Berufjordur (Ehlmann et al., 2012).

The remoteness of the Askja region make it difficult to reach with airborne sensors, but ASTER is well suited for clay mineral mapping in the vast (>5000 km²), largely unvegetated, rugged landscape (Graettinger et al., 2013). An ASTER SWIR principal component image of the Askja region (Fig. 14.4) shows strongly altered components of the tuya and moberg in bright blue and slightly less altered parts (most of the structure) in purple. Extracted reflectance spectra illustrate strong absorptions in ASTER bands 7 and 8, which correspond approximately to 2.24–2.37 μm, where trioctahedral clays and Fe³⁺-rich dioctahedral clay minerals absorb. TIR spectra (in the MIR region) of the same surfaces (Fig. 14.4d) show a strong absorption at 9.3 μm but little absorption from 10 to 11 μm. This observation suggests that much of the mafic glass with absorptions from 10 to 11 μm has been converted to smectite, which has strong Si–O absorptions at shorter wavelengths (9–9.5 μm). Taken together, the SWIR and TIR results provide a strong case for the formation of nontronite and/or Fe³⁺-rich trioctahedral smectite in the Askja hyaloclastites.

FIG. 14.4 — Remote sensing of clay minerals in the Askja region of central Iceland (A). The landscape includes lavas, reworked volcanics, moberg ridges and tuyas (B). ASTER SWIR (C) and TIR (D) principal component images show colour variations related to volcanism and hydrothermal alteration. Extracted reflectance (E) and thermal emission (F) spectra show features related to octahedral cation–OH and Si–O vibrational absorptions in clays.

Analyses of materials sampled from these areas confirm the above hypothesis and reveal more information about the altered hyaloclastites (Michalski and Bleacher, 2014). NIR reflectance spectra collected in the lab contain 2v (OH) absorptions near 1.4 μm, 2v(H₂O) absorptions at 1.41 μm, (v + δ) (H₂O) absorptions at 1.91 μm, and structural (v + δ)MOH bands at 2.305 consistent with the presence of Fe/Mg-smectite.

14.3.3.2. Clays and Clay Minerals Detected at Swansea, Arizona

Swansea, Arizona contains extremely, structurally complex Micocene volcanic and sedimentary rocks juxtaposed against Precambrian basement along the Buckskin-Rawhide detachment fault. The Miocene was a period of extreme and active extension in southwestern North America, where movement along the fault is estimated at 65 km or more (Reynolds and Spencer, 1985). As the crust extended, a number of shallow basins containing alkaline lakes formed. These surface and subsurface basin brines pervasively metaso-matised volcanic ash and lava and debris flows composed of fragments of both (Chapin and Lindley, 1986). The K-metasomatism is remarkable because it has resulted in 100% chemical and mineralogical replacement of the protoliths, which preserved the primary textures. Despite the fact that some of the altered rocks have K/Na ratios >500, this alteration is very difficult to identify in the field.

Remote sensing data provide valuable insight into the alteration patterns in the Swansea area. While it was conceivable that the alteration might have been limited to fluid conduits along the Buckskin-Rawhide detachment fault or associated normal faults, ASTER clay mineral maps (Fig. 14.5) show that this is not the case. In fact, analysis of ASTER imagery reveals that the volcanic units (mostly basaltic lava protolith) are altered everywhere they are observed. All of the rock units were exposed to basin brines as the basins were buried during rapid extension. However, the alteration only affected the mafic lithologies; it was driven entirely by susceptibility of the protolith to alteration by the K-rich brines.

ASTER reflectance data suggest the presence of Al-rich clays based on the 2.2 μm absorption band (Fig. 14.5). TIR ASTER data show narrow Si–O absorptions shifted to shorter wavelengths, consistent with the occurrence of clay minerals rather than mafic minerals in basalt exposures. Laboratory analyses of the bulk rock, as well as the <2 mm fraction of the crushed and sieved rock particles, show that the altered basalt contains illite. Micro-FTIR mapping of the alteration at the scale of hundreds of microns reveals how the plagioclase has been replaced by K-feldspar and the pyroxene and amphibole have been completely replaced by illite (Michalski et al., 2007).

14.3.3.3. Clays and Clay Minerals Detected at the Painted Desert, Arizona

The Painted Desert in Arizona is another region rich in clay minerals and low in vegetation with coverage by the HyMap imager. HyMap has a spectral range of 0.45–2.48 μm with a sampling of 13–17 nm per channel, and a spatial resolution of 4 m per pixel. Spectra are shown in Fig. 14.6 from sites 1–1 and 1–6 (McKeown et al., 2009b) that include mixtures of montmorillonite, calcite, dolomite, analcime, quartz and plagioclase. The montmorillonite bands are present in the lab, field and aerial data of these sites, although the relative intensity of the montmorillonite and carbonate bands varies. XRD indicates a higher abundance of montmorillonite at site 1–6, which corresponds well with the stronger montmorillonite signatures in those data.

FIG. 14.6 — Example remote sensing analyses from the Painted Desert, AZ (modified from McKeown et al., 2009b). (A) Aerial photograph of the region showing clay mineral-bearing units in white/grey or red where the topsoil has been eroded away, (B) lab, field and HyMap spectra of site 1–1 from the study, (C) lab, field and HyMap spectra of site 1–6 from the study, (D) visible HyMap image highlighting variations in iron content in the lower-left corner in red/purple tones (colours mapped as R: 0.65 μm, G: 0.55 μm, B: 0.45 μm), (E) NIR HyMap image highlighting variations in mineralogy in the upper half of the image in blue and purple tones (colours mapped as R: 2.48 μm, G: 1.50 μm, B: 1.01 μm) and (F) lab spectra of minerals compared to lab spectra of Painted Desert samples.

14.4. CHARACTERISATION OF CLAYS AND CLAY MINERALS ON MARS

The primary source of data for martian clay minerals are the OMEGA (Bibring et al., 2005) and CRISM (Murchie et al., 2009) orbital imaging spectrometers that collect reflected sunlight of the surface in the VNIR spectral region. The near entirety of Mars has been mapped at VNIR wavelengths at the km to 200 m resolution, while selected spot observations have been carried out at tens of metres resolution. TES orbital imaging provides context for investigations of clay minerals with near global coverage at 1–5 km resolution (Christensen et al., 2001). Clay minerals frequently occur in small outcrops on Mars that are less than 1 km across. Such occurrences are not possible to detect with TES. However, modelling developed by the TES team enables estimations of major and minor surface components including clay minerals and related altered aluminosilicate-bearing materials (Ramsey and Christensen, 1998; Rogers and Christensen, 2007; McDowell and Hamilton, 2009). Band parameters have also been developed for detection of Fe/Mg-clay minerals using THEMIS imagery (Viviano and Moersch, 2012), which has better spatial resolution but only a few channels (Christensen et al., 2003). Detecting clay minerals on Mars with TES has been challenging due to spot size, surface morphology and the nature of the MIR clay mineral bands in relation to other silicate features. Despite these challenges, investigations of selected regions have identified components due to altered basalt, glass, zeolite, high-Si phases and poorly crystalline aluminosilicates (Michalski et al., 2013b; Michalski and Fergason, 2009; Rogers and Christensen, 2007; Rampe et al., 2012; Ruff, 2004). Recently, coordinated studies of CRISM and TES analyses have found that including poorly crystalline aluminosilicates in the model enables detection of clay minerals in TES data as well for Mawrth Vallis, one of the most clay mineral-rich sites of Mars (Bishop and Rampe, 2016).

An early assessment of clay mineral detections on Mars using OMEGA data that indicated they are largely found in the ancient Noachian terrains (Bibring et al., 2006) continues to hold for most of the planet. More recently formed clay-rich materials have been observed in isolated regions such as Noctis Labyrinthus (e.g. Weitz et al., 2011), Coprates Chasma (e.g. Weitz et al., 2014), and some impact craters. Early investigations of outcrops bearing clays and clay minerals at Mawrth Vallis revealed that Al-rich clays and clay minerals always occurred in strata above Fe/Mg-rich smectite (Bishop et al., 2008c). A more recent study found that this trend is common across most of the planet where clay mineral exposures are observed (Carter et al., 2015). Detections of clay minerals from orbit (Wray et al., 2009; Milliken et al., 2010) have even guided detections of clay minerals by surface rovers at Endeavour crater (Arvidson et al., 2014; Fox et al., 2016) and Gale crater (Ming et al., 2014; Vaniman et al., 2014), where more detailed analyses of these ancient aqueous outcrops were performed. Characterisation of clay minerals by instruments on the Curiosity rover at Gale crater support formation of smectites in Hesperian rocks in low water to rock ratio environments (e.g. McLennan et al., 2014; Grotzinger et al., 2014; Grotzinger et al., 2015; Bristow et al., 2015).

14.4.1. Global Mapping of Clays and Clay Minerals and Aqueous Alteration on Mars

Mars exhibits several thousands to tens of thousands of clay mineral exposures at the coarse kilometric scale across its surface (Figs. 14.7 and 14.8). The distribution of these clay mineral-bearing outcrops is not homogeneous, with a majority of detections originating from the southern, Noachian-aged highlands (>3.8 Gyrs old). Terrains younger than the Late Hesperian (~3.5–3.8 Gyrs) are nearly devoid of these signatures, indicating that the bulk of the clay mineral-forming environments occurred during and before that time. Little is known about approximately one third of the martian surface as it is masked to VNIR instruments, either by coatings of anhydrous dust or by the cryosphere (ice, recurring ice clouds and seasonal frosts). Clay mineral exposures are typically small in size (a few 100 m to a few km) and often take the form of discontinuous patches, owing to the degraded state of the ancient surface. Their geological context is often too disrupted to be uniquely inferred from orbit. Recent reviews of orbital remote sensing at Mars describe the current understanding of the martian surface composition including clay minerals (Murchie et al., 2018; Hamilton et al., 2018).

FIG. 14.7 — Global map of aqueous minerals on Mars, based on data from the OMEGA and CRISM VNIR spectrometers. The background is Mars topography from MOLA and colours indicate mineral composition: red for Fe/Mg-clay minerals (80% smectites, 20% chlorites), green for hydrated sulphates and cyan for Al clay minerals. Yellow indicates mixtures of Fe/Mg-smectites and sulphates. Most aqueous mineral exposures, especially non-Fe/Mg-clay minerals detections are not visible at this scale. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

FIG. 14.8 — TES global abundances of clay minerals on Mars. Modelled abundances correspond to clay minerals and other poorly crystalline aluminosilicates (Michalski et al., 2006b). (*modified from Bandfield*, 2002.)

14.4.1.1. VNIR Observations of Clays and Clay Minerals and Aqueous Alteration on Mars

About half of the clay mineral exposures are observed in the peaks, walls and ejecta of impact craters of all sizes (e.g. Carter et al., 2013; Ehlmann et al., 2013, indicating that clay minerals are often buried at shallow depths, up to a few km, usually less). The ubiquity of impact craters and their broad range of excavation depths are used to probe the mineralogy of the subsurface. Impacts over Hesperian and Amazonian aged terrains rarely excavate clay minerals, with the exception of larger craters (typically >20–200 km), which are thought to have penetrated the deeply buried and altered Noachian crust. This applies in particular to the northern lowland hemisphere of Mars, where the few clay mineral signatures are restricted to remnant Noachian units and the largest craters. In the south, consistent trends have been found wherein clay minerals are replaced at depth by higher temperature and lower water activity species that include chlorites, zeolites and prehnite (e.g. Ehlmann et al., 2009; Marzo et al., 2010). Low-grade metamorphic facies have been recognised for some sites. A few craters are also thought to have formed clay minerals in situ, as a result of an impact-driven hydrothermal system.

Non impact-related clay mineral exposures are numerous and include several deposits of vast sizes (up to hundreds of km across) spread across the southern highlands (e.g. Poulet et al., 2005; Bishop et al., 2008c, 2013c; Mustard et al., 2008; Murchie et al., 2009; McKeown et al., 2009a; Milliken et al., 2010; Loizeau et al., 2012a; Carter et al., 2015; Weitz et al., 2012; Weitz et al., 2015; Weitz and Bishop, 2016). Those studied in detail have uncovered over 10 types of clay minerals (smectites including nontronite, saponite, beidellite and montmorillonite, vermiculites, corrensites, chlorites, K/Al-rich micas, Fe²⁺-smectite/mica mixtures, kaolinite, mixed layered kaolinite-smectite, allophane, talc and serpentines). These are summarised by Ehlmann and Edwards (2014). Correlations between the clay mineral and related mineralogy and morphological context across the planet have been interpreted as originating from a number of surface to near-surface formation and depositional environments (e.g. Ehlmann et al., 2011). These include detrital and authigenic lacustrine clay minerals, climate-mediated surface weathering (involving rain or snow fall), evaporitic playa environments and hydrothermal systems (deep crustal, epithermal and possibly spring deposits). Geochemical conditions proposed for these environments range from highly acidic (pH ~ 2) to mildly alkaline, near freezing to >100°C, water-dominated to rock-dominated, mostly oxidising, and with varying ion activities (e.g. Ehlmann et al., 2013; Carter et al., 2015; Murchie et al., 2018).

Beyond this preserved diversity, global mineralogical trends exist. The bulk of Mars alteration (>80% of the deposits) appears to have formed Fe/Mg-rich clay minerals with a strong smectitic contribution (Ehlmann et al., 2011; Carter et al., 2013; Michalski et al., 2015). For some of these, there is evidence for mixed-layering with chlorites and micas, while chlorites are found in a spectrally pure form less than 20% of the time. Hydrated sulphates and silica are the second and third most abundant aqueously altered species (accounting for ~40% of the aqueous deposits on the surface), while aluminium clay minerals (both smectites and kaolinites) are found in about a third of the mineral deposits. It should be noted that these relative abundances may be significantly biased by a number of factors: (1) the selective detection sensitivity to the diagnostic spectral absorptions of different clay minerals, (2) the selective survivability of some species to later chemical transformations (diagenesis) or to disruptive processes such as impact gardening or aeolian erosion, (3) the lack of access to a third of the surface of Mars, (4) the poor penetration depth of VNIR spectroscopy (typically hundreds of microns or less), which largely restrict the study to surficial layers and erosional windows.

The dominance of Fe/Mg-rich clay minerals at Mars is consistent with water-limited alteration of basaltic and ultramafic protoliths, but their associated geological context and alteration process remains elusive in some cases. It is likely that a large fraction of the alteration occurred in the martian subsurface, but it is also clear that surface alteration has occurred throughout ancient Mars. The biggest mystery over climate remains: was Mars episodically warm and wet or did most of the alteration occur in cold, snow or ice-dominated environments?

14.4.1.2. TIR Observations of Clays and Clay Minerals and Aqueous Alteration on Mars

TIR emission spectral provide a global view of martian surface mineralogy, albeit at coarse spatial resolution (hundreds of km²) (Christensen et al., 2001; Hamilton et al., 2018a). Martian dark regions are spectrally dominated by basaltic materials in the thermal IR (Christensen et al., 2000b). Linear spectral unmixing of data from the TES spectrometer provides estimates of mineral abundances in the dark regions. This approach suggests that dark volcanic surfaces are composed of approximately 25–40% plagioclase feldspar, 25–40% pyroxene and 5–10% olivine (Rogers and Christensen, 2007). The data also show, however, that martian dark regions contain 15–35% clay minerals or high-silica phases (Fig. 14.8) (Michalski et al., 2006b), with 10–15% occurring at low latitudes and 20–30% occurring at higher northern latitudes (Wyatt et al., 2004).

TES spectra provide an interesting perspective on the occurrence and abundance of clay minerals on Mars. The spectral unmixing approach that has often been applied to TES data depends heavily on the spectral shape in the 900–1300 cm⁻¹ region, where Si–O absorptions occur in all silicate minerals and amorphous phases. In this spectral range, clays and clay minerals exhibit nonunique absorptions related to Si–O vibrations. However, these somewhat broad unremarkable features are shared by other poorly crystalline materials with similar Si/O molar ratios (Michalski et al., 2005; Rampe et al., 2012).

Early results from TES showing 10–15% of clay minerals or similar materials in the dark, low-latitude regions were difficult to interpret in the pre-OMEGA era, but now these results seem reasonable as an average composition of Noachian martian crust. Such detections might also include allophane, imogolite and other poorly ordered aluminosilicates, possibly associated with weathered volcanic ash in the ancient crust (Rampe et al., 2012; Bishop and Rampe, 2016).

Increased abundances of modelled clay-like, high-silica materials at high latitudes are more complicated to interpret because, while NIR data suggest the presence of mafic glass in some parts of the high latitudes (Horgan and Bell, 2012), neither OMEGA nor CRISM has detected such widespread deposits consistent with high-silica materials. It is possible that the high-latitude deposits detected by TES correspond to thin, but highly absorbing coatings and rinds (Kraft et al., 2003; Michalski et al., 2006a) similar to those that form as immature weathering products in low water/rock ratio settings on Earth.

14.4.2. Regional Mapping of Clays and Clay Minerals and Aqueous Alteration on Mars

At the regional scale, Mars exhibits variations in the clay mineralogy that can sometimes be linked to a particular morphological context or at least a broad geological unit. Detailed investigations often yield additional morphological and mineralogical diversity and allow refining interpretations of aqueous environments. The main limitation is often the lack of deep erosional windows to assess the stratigraphy, obscuration by dust or other mantles (e.g. volcanic capping) or lack of submetre resolution context imagery. Another potential complicating factor is the lack of a direct analogue on Earth for some of the alteration processes and geologic contexts observed on Mars, which provides a great challenge to interpretation of martian geology. A few examples of mineral provinces at Mars are described in detail below that are representative of the clay context globally and/or have been identified as targets of particular interest, including for future exploration.

14.4.2.1. The Clay-Laden Eastern Margin of Chryse Planitia

The Chryse Planitia basin provides a diversity of terrains with ages ranging from the middle Noachian to the Amazonian (Tanaka et al., 2005). Since its early emplacement, the basin has been in-filled by mass wasting, catastrophic flood deposits, and aeolian and volcanic resurfacing processes. The margin of Chryse Planitia retains, however, the signatures of several aqueous environments present during the Noachian to early Hesperian time period. The near entirety of the preserved margin exhibits clay-rich units (Fig. 14.7), the most prominent of which are located toward the east in the Mawrth Vallis plateau and Oxia Planum deposits (Fig. 14.9). The former has been recognised as the best example of past pedogenesis on Mars based on its stratigraphy and composition (Michalski et al., 2010b; Bishop et al., 2013b). The lower altered strata consist of hundreds of meters-thick layered Fe³⁺-rich smectite (mapped in red) that may have formed by the syn-depositional alteration of volcanic sediments involving meteoritic water (e.g. Michalski and Noe Dobrea, 2007; Wray et al., 2008; Bishop et al., 2008c; McKeown et al., 2009a; Loizeau et al., 2010; Loizeau et al., 2012b; Bishop and Rampe, 2016). Upper, highly leached strata consist of mixtures dominated by Al-smectite and kaolinite (mapped in cyan), sulphates and poorly crystallised phases such as allophane and hydrated silica (Bishop et al., 2016). Changes in climate conditions from near neutral to mildly acidic are proposed to have caused this chemical gradient in the Mawrth Vallis region. Although the Oxia Planum deposits appear to pertain to the same circum-Chryse clay mineral unit as Mawrth Vallis, their composition and stratigraphy differ. The aqueous deposits at Oxia Planum consist of thinner strata (tens of metres) of layered clay minerals with a remarkably homogeneous composition, at least as viewed from orbit, and the clay minerals identified at Oxia Planum are either ferroan saponite or vermiculite-bearing. An alluvial fan or delta postdates this Late Noachian unit, and contains hydrated silica and kaolinite. These are interpreted as authigenic phases formed during a shallow lake phase overlying the clay minerals (Carter et al., 2016; Quantin et al., 2016). Located in between Mawrth Vallis and Oxia Planum is McLaughlin crater, where mixtures of clay minerals with carbonate and serpentines have been found (Michalski et al., 2013a). These may originate from a combination of inherited, deeply buried material excavated by a nearby 30 km-large crater, while the carbonates may have formed authigenically within an alkaline, ground-water fed crater lake (Michalski et al., 2013a). The high scientific potential of these three deposits at the margin of the Chryse Plateau has led to multiple landing site proposals for current and upcoming rover missions.

FIG. 14.9 — Regional scale, clay mineral-rich units on the eastern margin of Chryse Planitia. Background is THEMIS IR imagery colour coded (green-brown tones) with altimetry from MOLA. Darkened areas correspond to a dust-mantle partly shrouding the surface to VNIR spectroscopy (mapped with OMEGA). Fe/Mg-clay minerals are overlain in red and vary from Fe³⁺-nontronite at the Mawrth Vallis Plateau, Fe^3+/2+-saponite/vermiculite at Oxia Planum, to serpentine and carbonate bearing at McLaughlin crater. Aluminous clay minerals are mapped in blue, and are comprised of kaolinites, smectites and allophane. Spectrally pure deposits of sulphates, carbonates and minor phases are not visible at this scale. Map is centred approximately at −22°E, 22°N.

Layered outcrops in the Mawrth Vallis region of Mars contain the greatest diversity of aqueous alteration materials that provide an opportunity to infer past aqueous environments. Numerous orbital investigations have documented aluminous and siliceous clay-bearing units overlying an Fe/Mg-smectite-rich unit below (e.g. Bishop et al., 2008c; Wray et al., 2008; McKeown et al., 2009a; Michalski et al., 2010b; Noe Dobrea et al., 2010; Loizeau et al., 2012b). Many different secondary minerals have been identified in the upper aluminous and siliceous clay-rich units (e.g. Bishop et al., 2013b), but the presence of poorly crystalline phases was reported only recently (Bishop and Rampe, 2016). Further, allophane and imogolite comprise a significant portion of the uppermost stratum of the aluminous and siliceous clay-rich units, on top of the montmorillonite- and opal-rich unit. These results signify a change in climate on Mars from a warm and wet environment to one where water was sporadic and likely depleted rapidly (Bishop and Rampe, 2016). A more complete stratigraphy including five units was developed recently for the Mawrth Vallis region using newly created CRISM parameters, HRSC DTMs across hundreds of kms and HiRISE DTMs across hundreds of metres (Bishop et al., 2016). This stratigraphy is illustrated in Fig. 14.10. The distinct mineralogic units were characterised using spectral properties from CRISM and surface morphology from HiRISE. Bishop et al. (2016) describe an ancient wet and warm geologic record that formed the thick nontronite unit, a period of wet/dry cycling to create acid alteration, followed by leaching or pedogenesis to result in aluminous clay minerals, and finally a drier, colder climate that left the altered ash in the form of nanophase clay minerals, rather than crystalline clay minerals.

FIG. 14.10 — Views of clay-bearing outcrops at Mawrth Vallis. (A) Mawrth Vallis region from Fig. 14.9 with a white box showing the location of the region shown in B. (B) Close-up view of northwestern region in CTX imagery over an HRSC DTM with colours from CRISM overlain. Nontronite (Nt) is mapped in red, ferroan clay minerals (FeSm mix) in purple, a sulphate-bearing and/or acidic altered unit in yellow, Al-clay minerals (AlSm) and opal in blue, and poorly crystalline aluminosilicates (allophane, Allo) in green. The white box shows the location of C. (C) HiRISE DTM with colours from CRISM, view towards N. (D) Example spectra from the five stratigraphic units identified recently in the region. (*modified from Bishop et al., 2016*.)

14.4.2.2. The Clay-Bearing Region West and South of Isidis Planitia

After the Chryse Planitia basin the regions surrounding the Isidis basin have the next greatest abundance of clay and clay mineral-rich material (Figs. 14.7 and 14.11). The Nili Fossae region northwest of Isidis hosts pristine olivine and pyroxene outcrops (Mustard et al., 2009) and abundant and varied clay mineral-bearing units in several provinces with distinct mineral assemblages (Mustard et al., 2008; Ehlmann et al., 2009). Fe/Mg-smectite is found throughout the Nili Fossae region in the ancient Noachian rocks. This Fe/Mg-smectite unit occurs in eroded terrains in Eastern Nili Fossae where it is stratigraphically covered by either kaolinite or magnesite (Ehlmann et al., 2009, Fairén et al., 2010). In other outcrops the Fe/Mg-smectite is mixed with mica, illite or chlorite. The presence of zeolites and prehnite together with Fe/Mg-smectite in and around impact craters in the west and south are indicators of low-grade metamorphic or aqueous hydrothermal alteration (Ehlmann et al., 2009; Fairén et al., 2010; Marzo et al., 2010). Sulphates are observed together with Fe/Mg-smectite, magnesite and other clay minerals along the southwestern part of Nili Fossae between Jezero crater and the Syrtis volcanics region (Ehlmann and Mustard, 2012). Jarosite ridges were observed on layered material containing polyhydrated sulphates and Fe/Mg-smectite. The diversity in clay minerals observed across the Nili Fossae region, as well as the geologic context of these clay mineral-bearing outcrops indicates that several episodes of aqueous alteration took place (Ehlmann et al., 2009).

FIG. 14.11 — Regional scale clay mineral-rich units near Isidis Planitia. The background is THEMIS IR as in Fig. 14.9. Red indicates ferroan-magnesian clay minerals (80% smectites, 20% chlorites), green hydrated sulphates, cyan for Al-rich clay minerals, and yellow mixtures of ferroan-magnesian smectites and sulphates. (A) Nili Fossae region northwest of Isidis Planitia centred near 75°E, 20°N, (B) Libya Montes region south of Isidis Planitia centred near 84°E, 2°N and (C) Tyrrhena Terra region centred near 88°E, 10°S between Isidis Planitia, and Hellas Basin.

The Libya Montes region, located at the southern rim of the Isidis impact basin, hosts a variety of geological processes that have shaped the martian surface. Evidence for fluvial, lacustrine, aeolian, volcanic and hydrothermal processes have been documented, resulting in a variety of landforms and ample evidence for chemical alteration of the local rocks (Crumpler and Tanaka, 2003; Tornabene et al., 2008; Jaumann et al., 2010; Erkeling et al., 2012). The Libya Montes region contains widespread pyroxene- and olivine-bearing units (Tornabene et al., 2008) and a variety of smectites and Fe/Mg-carbonate similar to dolomite (Bishop et al., 2013c). The ancient Noachian basaltic crustal materials experienced extensive aqueous alteration at the time of the Isidis impact, during which the montes were also formed, followed by emplacement of a rough olivine-rich lava or melt and finally the smooth pyroxene-bearing caprock unit. The Fe/Mg-smectite in this region vary from nontronite with an OH band at 2.29 μm to Fe/Mg-smectite with a band at 2.30 μm and saponite with a band at 2.31 μm (Fig. 14.12). Beidellite is observed here due to an (v+δ)Al₂OH band at 2.19 μm rather than montmorillonite, which would have a band at 2.21 μm. Beidellite is a higher temperature smectite and this beidellite likely formed via low temperature hydrothermal alteration from the Isidis impact. Abundant fluvial features dating from the Noachian to Amazonian time periods traverse the Libya Montes region; however, clay minerals are primarily observed in the ancient Noachian rocks.

FIG. 14.12 — Views of clay mineral-bearing outcrops at Libya Montes. (A) A portion of the Libya Montes region from Fig. 14.11 with a white box showing the location of B. (B) Close-up view of the central region in CTX imagery over an HRSC DTM with colours from CRISM overlain. Ferroan-magnesian smectite is mapped in red/pink, olivine in green and pyroxene in blue. The white box shows the location of C. (C) CRISM parameter map showing outcrops of multiple units and the locations of example spectra shown in D: 1-beidellite (Bd), 2-saponite (Sap), 3-Fe/Mg -carbonate mixed with smectite (Carb), 4-olivine (Olv) and 5-nontronite (Nt). (D) Example spectra illustrating the features observed for the geologic units containing smectites, carbonate and olivine. Grey lines at 1.92, 2.30 and 2.53 μm mark the spectral features. (*modified from Bishop et al., 2013c*.)

The Tyrrhena Terra region lies south of Libya Montes and north of Hellas basin and exhibits some of the highest and most ancient crust of Mars, ranging from the Early to Late Noachian. It consists of highland plateaus, at times dissected, but rarely resurfaced by the accumulation of later sediments or volcanic deposits. It displays a large number of impact craters of all sizes (up to ~100 km in diameter), which provide natural probes to the upper crust from the near surface to several kms deep. Older craters have experienced erosion so that their depth-to-diameter ratio has decreased, or because they have been locally infilled by younger material. Fresh craters however have a well identifiable rim, central peak (or rings) and ejecta blanket, which are necessary to explore the excavated vertical stratification and disentangle contributions from postimpact, clay mineral-forming hydrothermal environments. As shown in Fig. 14.11C, most fresh craters, regardless of their size, display significant clay mineral abundances in their ejecta blanket, while simultaneously exhibiting clay minerals in intracrater deposits (Loizeau et al., 2012a; Carter et al., 2013). Most of these signatures are interpreted as previously altered excavated material, while postimpact formation is considered minor in this region (Loizeau et al., 2012a). A trend is found such that craters typically smaller than about 10 km excavate smectitic clay minerals, while larger craters have stronger contributions from chlorites. Zeolites and prehnite are observed locally as excavated material within the largest craters (except in a few small craters where they are thought to have been excavated from an underlying ejecta blanket). The vertical stratigraphy of smectites to chlorites ± zeolites to chlorites ± prehnite is interpreted as the signature of burial diagenesis and low-grade meta-morphism, while direct precipitation in hot and deep aquifers can also contribute to the observed mineralogy (Ehlmann et al., 2011; Carter et al., 2013). The metamorphic mineral epidote is found at a few sites, but its original context/stratigraphy has not been inferred yet. Carbonates and talc have also been detected in some regions of Tyrrhena Terra associated with clay minerals, especially near the Hellas rim (Viviano-Beck et al., 2016).

14.5. CHARACTERISATION OF CLAYS IN METEORITES

Meteorites serve as delivery systems for extraterrestrial clay minerals to Earth. They are classified by petrologic type and origination, with some coming from the Moon and Mars and many others coming from the asteroid belt or further sources. The majority of meteorites are stony type (chondrites and achondrites), while fewer are iron meteorites or stony-iron meteorites (Wasson, 1985; Lauretta and McSween Jr., 2006; Papike, 1989). The lunar and martian meteorites fall into the achondrite category. Attempts have been made to connect meteorites to asteroid types through their spectral properties (Chapman and Salisbury, 1973; Gaffey, 1976; Sandford, 1984; Salisbury and Hunt, 1974; Salisbury et al., 1991a), but these have not been entirely successful, likely due to weathering of meteorites and because the meteorites are not well represented on the surfaces of their parent bodies. Clay minerals are rare in meteorites, but are present in some martian meteorites and some chondrites (Zolensky and McSween, 1988).

Early detections of clay minerals and altered materials in carbonaceous chondrites were reported for Allende (Dominik et al., 1978), Plainview (Nozette and Wilkening, 1982), and Mokoia (Cohen et al., 1983), although it was greatly debated in the community whether these clay minerals were formed by terrestrial alteration or had an extraterrestrial origin. Carbonaceous chondrites of the Ivuna and Orgueil types contain clay minerals intimately mixed with a carbon-rich matrix attributed to preterrestrial alteration by aqueous solutions (Morlok et al., 2006). The C2 Tagish Lake meteorite includes mixed clay minerals, carbon, carbonate and sulphides, which are also thought to be preterrestrial due to the brief environmental exposure of this meteorite (Brown et al., 2000). The spectral properties of both Fe^3+/2+-rich (Calvin and King, 1997) and Mg²⁺-rich (King and Clark, 1989) clay minerals were investigated in comparison with the spectral properties of meteorites containing Ivuna and Mighei type chondrites. Mixtures including Fe-rich serpentines and chamosite were found to be most consistent with these meteorites (Calvin and King, 1997).

The majority of martian meteorites follow three categories: shergottites, nakhlites and chassignites, named after meteorites from these locations (McSween and Treiman, 1998). Whether or not clay minerals detected in these meteorites were formed by terrestrial weathering or preterrestrial processes was hotly debated for years, and only became widely accepted after abundant clay minerals were observed on Mars by the orbiters. Smectites together with carbonate and sulphate were detected in interior regions of the Nakhla meteorite and attributed to preterrestrial alteration because of their location in the rock and the morphology of the grains (Gooding et al., 1991). Further investigation of additional shergottite, nakhlite and chassignite meteorites revealed preterrestrial clay minerals and associated aqueous alteration minerals (Gooding, 1992). Continuing investigation of clay minerals in these meteorites discovered Fe-rich smectite (Treiman et al., 1993). Model-ling of alteration pathways for formation of these clays, carbonates and sulphates suggest evaporation from low temperature brines in contact with parent igneous rocks (Bridges et al., 2001). One important martian meteorite that does not fall in one of these categories is the famous orthopyroxenite ALH 84001 (Mittlefehldt, 1994), which received great attention following reports of tiny globules containing preterrestrial organics, carbonates, magnetite and amorphous materials (McKay et al., 1996). Clays in this meteorite are found exclusively in the carbonates and not in the glass, thus strengthening support for their preterrestrial origin (Brearley, 2000). A recent review of clay mineral detections in martian meteorites suggests that broadly similar clay minerals are observed in the meteorites and on the surface of Mars (Velbel, 2012). Unfortunately, the clay minerals present in meteorites tend to occur in extremely tiny aliquots on the scale of nm or smaller, thus challenging researchers to identify and characterise them. The nakhlite meteorite group contains the largest quantity of clay minerals and for this reason they have been characterised in more detail. The most abundant clay minerals present include Fe^3+/2+-enriched trioctahedral saponite and serpentine (Hicks et al., 2014). Modelling of the geochemical conditions required for formation of the observed carbonate and Fe/Mg-smectite in nakhlites indicates there was a CO₂-rich hydrothermal fluid at 150–200°C with pH ~ 6–8 and a water/rock ratio ≤300 (Bridges and Schwenzer, 2012).

14.6. CHARACTERISATION OF CLAY MINERALS AT ASTEROID 1-CERES

The asteroid 1-Ceres is the largest body in the asteroid belt and is composed of silicate minerals (McCord and Gaffey, 1974; McCord and Sotin, 2005). Spectral features near 3 μm in telescopic data first led to the idea of clay minerals on the Dwarf planet Ceres (Lebofsky et al., 1981). Ammoniated clay minerals were detected on Ceres through telescopic measurements using the spectral bands near 3.07 μm (King et al., 1992). Their data were most consistent with NH₄-saponite. Further comparison of ammoniated saponite with emission spectra of Ceres found consistencies near 6.2 and 6.9 μm, but a difference near 9.5 μm (Cohen et al., 1998), suggesting additional components are present as well. A lab study of NH₄-treated smectite showed shifts in the 3.07 μm band depending on how the NH₄ was bound to the smectite (Bishop et al., 2002b). For tightly bound NH₄ this band was observed at 3.06 μm, while for smectite washed in NH₄ solution this band occurred at 3.07–3.1 μm. Additional analyses of improved telescopic data of Ceres reported the presence of carbonate and Fe-smectite (Rivkin et al., 2006). Further analyses of this new data led to the interpretation that Ceres contains brucite, carbonate and serpentine (Milliken and Rivkin, 2009).

More recent spectral data were acquired for Ceres by the Dawn spacecraft from 0.4 to 5 μm (De Sanctis et al., 2015). These spectra are consistent with widespread NH₄-bearing minerals across the surface of Ceres. De Sanctis et al. (2015) support formation of these NH₄-clay minerals during differentiation, perhaps through incorporation of material from the outer Solar System. The type of ammoniated clay or clay mineral present on Ceres has not yet been determined, but it is thought to be mixed with magnetite, antigorite and carbonate (De Sanctis et al., 2015). Brucite was ruled out as contributing to the 3.05–3.1 μm feature in Ceres’ spectrum through laboratory mixture experiments (De Angelis et al., 2016). A recent review of telescopic spectra of asteroids and NH₄-bearing minerals determined that 1-Ceres, 10-Hygiea and 324-Bamberga all exhibit spectral features near 3.05–3.07 μm attributed to ammoniated clays (Berg et al., 2016). Several other spectral bands are present for NH₄-treated clay minerals, but these are more difficult to resolve in spectra of asteroids. Thus, the current understanding of Ceres supports the presence of ammoniated clay minerals, but the specific type of clay minerals cannot yet be uniquely assigned.

14.7. CHARACTERISATION OF CLAY MINERALS IN COMETS

The Spitzer Space Telescope has collected TIR spectra (5–40 μm) of several Jupiter family comets including 9B-Tempel 1 (Werner et al., 2004). Spitzer spectra were collected of dust ejected from the coma of 9B-Tempel 1 before and after the collision of the Deep Impact mission with its nucleus. Features due to clay minerals, carbonate, water ice and organics were detected in the Spitzer spectra of the dust after impact (Lisse et al., 2006). Fine grains of crystalline and amorphous olivine and pyroxene were observed in the dust around 9B-Tempel 1 before and after impact and were modelled together with the clay minerals, carbonate, water ice and organics (Kelley and Wooden, 2009). Lisse et al. (2006) found nontronite to correspond best to the clay mineral signatures and estimated its presence as 5–10% of the silicate material in the dust. Lisse et al. (2007) compared the Spitzer data for comet 9B-Tempel 1 with International Space Observatory (ISO) spectra of comet Hale-Bopp following its 1995–1996 apparition resulting in strong outflow of dust (Williams et al., 1997) and an ISO spectrum of HD100546, a young star often compared to comets (Waelkens et al., 1997). Coordinated spectral analyses of these bodies indicate similar components including nontronite (Lisse et al., 2007). They calculated weighted surface areas for nontronite, carbonate, water ice, organics, sulphides, olivine and pyroxene for these objects. Their results included weighted (unit-less) surface areas of 0.14, 0.08, 0.05 and 0.04 for nontronite, carbonate, water ice and polyaromatic hydrocarbons, respectively, for 9B-Tempel 1, compared to 0.18, 0.33, 0.34 and 0.14 for Hale-Bopp, and 0.14, 0.08, 0.46 and 0.51 for HD 100546 (Lisse et al., 2007). The variations in clay minerals and volatile components suggest different formation histories for these planetary bodies.

14.8. SUΜMARY OF REMOTE OBSERVATIONS OF PLANETARY CLAY MINERALS

Spectral remote sensing in the VNIR and MIR regions has enabled detection and characterisation of multiple clay minerals on Earth and elsewhere in the Solar System. Remote sensing on Earth has the greatest challenge due to atmospheric absorptions that interfere with detection of surface minerals. Still, a greater variety of clay minerals has been observed on Earth than other planets due to the maturity of this planet. Clay minerals have arguably been mapped in more detail on Mars because they are not masked by vegetation and the atmosphere is thinner and primarily composed of CO₂ and thus easier to process. Fe/Mg-smectite is the most abundant clay mineral identified on the surface of Mars, is present (with Fe²⁺/Mg-enriched clay minerals) in some meteorites, and is also the most common clay mineral in comets where detected. The Fe/Mg-smectite on Mars is almost always found in ancient rocks 4 Gy old. Could this clay mineral have been a component of the material formed early in the life of the Solar System? Perhaps this question could be addressed as more spectrometers are sent on missions to other planetary bodies.

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PERMALINK

Remote Detection of Clay Minerals

Janice L Bishop

Joseph R Michalski

John Carter

Abstract

14.1. PRESENCE OF CLAY MINERALS IN OUR SOLAR SYSTEM

14.2. REMOTE DETECTION OF CLAY MINERALS

14.2.1. VNIR Bands Used for Detection of Clay Minerals

FIG. 14.1.

14.2.1.1. VNIR Characterisation of Smectites

14.2.1.2. VNIR Characterisation of Kaolinites-Serpentines

14.2.1.3. VNIR Characterisation of Chlorites

14.2.1.4. VNIR Characterisation of Micas, Illites and Other Clay Minerals

14.2.1.5. VNIR Characterisation of Poorly Crystalline Clay Minerals

14.2.2. MIR Bands Used for Detection of Clay Minerals in TIR Spectra

14.2.2.1. Si–O Stretching Vibrations

FIG. 14.2.

14.2.2.2. Si–O Bending Vibrations

14.3. CHARACTERISATION OF CLAY MINERALS ON EARTH

14.3.1. Challenges of Remote Detection of Clay Minerals on Earth

FIG. 14.3.

14.3.2. Instruments and Datasets Available for IR Remote Sensing of Clay Minerals on Earth

TABLE 14.1.

14.3.3. Remote Characterisation of Clay Minerals on Earth

14.3.3.1. Clays and Clay Minerals Detected in Hyaloclastites in Askja, Iceland

FIG. 14.4.

14.3.3.2. Clays and Clay Minerals Detected at Swansea, Arizona

FIG. 14.5.

14.3.3.3. Clays and Clay Minerals Detected at the Painted Desert, Arizona

FIG. 14.6.

14.4. CHARACTERISATION OF CLAYS AND CLAY MINERALS ON MARS

14.4.1. Global Mapping of Clays and Clay Minerals and Aqueous Alteration on Mars

FIG. 14.7.

FIG. 14.8.

14.4.1.1. VNIR Observations of Clays and Clay Minerals and Aqueous Alteration on Mars

14.4.1.2. TIR Observations of Clays and Clay Minerals and Aqueous Alteration on Mars

14.4.2. Regional Mapping of Clays and Clay Minerals and Aqueous Alteration on Mars

14.4.2.1. The Clay-Laden Eastern Margin of Chryse Planitia

FIG. 14.9.

FIG. 14.10.

14.4.2.2. The Clay-Bearing Region West and South of Isidis Planitia

FIG. 14.11.

FIG. 14.12.

14.5. CHARACTERISATION OF CLAYS IN METEORITES

14.6. CHARACTERISATION OF CLAY MINERALS AT ASTEROID 1-CERES

14.7. CHARACTERISATION OF CLAY MINERALS IN COMETS

14.8. SUΜMARY OF REMOTE OBSERVATIONS OF PLANETARY CLAY MINERALS

References

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