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date: 06 June 2020

The Surface Composition of Terrestrial Planets

Summary and Keywords

Planetary surface compositions are fundamental to an understanding of both the interior activity through differentiation processes and volcanic activity and the external evolution through alteration processes and accumulations of volatiles. While the Moon has been studied since early on using ground-based instruments and returned samples, observing the surface composition of the terrestrial planets did not become practical until after the development of orbital and in situ missions with instruments tracking mineralogical and elemental variations. The poorly evolved, atmosphere-free bodies like the Moon and Mercury enable the study of the formation of the most primitive crusts, through processes such as the crystallization of a magma ocean, and their volcanic evolution. Nevertheless, recent studies have shown more diversity than initially expected, including the presence of ice in high latitude regions. Because of its heavy atmosphere, Venus remains the most difficult planetary body to study and the most poorly known in regards to its composition, triggering some interest for future missions. In contrast, Mars exploration has generated a huge amount of data in the last two decades, revealing a planet with a mineralogical diversity close to that of the Earth. While Mars crust is dominated by basaltic material, recent studies concluded for significant contributions of more felsic and alkali-rich igneous material, especially in the ancient highlands. These ancient terrains also display widespread outcrops of hydrous minerals, especially phyllosilicates, which are key in the understanding of past climate conditions and suggest a volatile-rich early evolution with implications for exobiology. Recent terrains exhibit a cryosphere with ice-rich landforms at, or close to the surface, of mid- and high latitudes, generating a strong interest for recent climatic variability and resources for future manned missions. While Mars is certainly the planetary body the most similar to Earth, the observation of specific processes such as those linked to interactions with solar wind on atmosphere-free bodies, or with a thick acidic atmosphere on Venus, improve our understanding of the differences in evolution of terrestrial bodies. Future exploration is still necessary to increase humankind’s knowledge and further build a global picture of the formation and evolution of planetary surfaces.

Keywords: chemistry, mineralogy, infrared spectrometry, Mars, Moon, Mercury, Venus

The composition of terrestrial planet surfaces is a fundamental parameter for understanding their geologic and climatic evolution. Surface composition is also often the most accessible information regarding a planetary body’s internal structure. Tools used to determine planetary surface composition have improved throughout time. Initially limited to earth-based telescope observations (e.g., Moroz, 1964), orbital remote sensing, in situ analyses, and returned samples progressively enabled more widespread and precise mapping of planetary surfaces. Scientists also gained knowledge from meteorites found on Earth, some of which are natural samples from the Moon and Mars ejected by impacts. This article focuses predominantly on results from space missions, starting with a summary of the various techniques. These techniques can be used to identify either volatiles deposited as ices or rock composition. Instead of reporting each planetary body individually, this review summarizes first the distribution of volatiles on each body, then describes the igneous composition to help explain the formation of planetary crusts and their volcanic evolution from surface observations; finally it focuses on the alteration of igneous rocks, as far as it can be understood from surface data. As the level of knowledge of the surface composition of terrestrial planets remains variable from one body to the other, this review concludes with unresolved questions and looks ahead to future missions planned to find these answers.

Determining Composition on Planetary Surfaces

This section describes the main techniques used by space missions to determine surface composition useful for the understanding of the later discussion (Table 1). This summary is not exhaustive, neither in terms of instrument names nor in terms of techniques.

Mineralogy by Orbital and In Situ Measurements

Determination of minerals is the most fundamental observation for geologists trying to classify rocks. Remote sensing techniques emerged with the achievement of the first space probes in Earth orbit in the 1960s, using spectroscopy as a main remote sensing technique for mineral determination (in a broad sense, including ices). Spectroscopy uses the properties of matter to modify light depending on the fundamental vibrations of various molecules present in rocks or ices. While the remote sensing of the Earth’s surface remains technically challenging due to Earth’s atmosphere and the presence of water at the surface, these techniques have been powerful on planetary surfaces lacking or presenting a tenuous atmosphere (Mars) or an exosphere only (Moon, Mercury).

Visible and near infrared (VNIR) spectroscopy uses the reflectance of sunlight from surfaces at wavelengths of 0.4 to 3 μ‎m. This technique is powerful for the Moon from ground-based observations thanks to its proximity (e.g., McCord et al., 1981). Spectrometers were sent in orbit only relatively recently to improve spatial resolution, such as the near infrared (NIR) camera on board the Clementine spacecraft (e.g., Le Mouélic et al., 2000), the SMART1 infrared spectrometer (SIR and SIR2) on board the Indian mission Chandrayaan (Mall et al., 2009; Mall, Nathues, & Keller, 2003), and the Moon Mineralogy Mapper (M3; Mustard et al., 2011; Pieters, Boardman, et al., 2009). After first hints at Mercury by Mariner 10, the Mercury Dual Imaging System (MDIS; Hawkins et al., 2007) and Mercury Atmosphere and Surface Composition Spectrometer (MASCS; McClintock & Lankton, 2007) on board the Messenger probe have shown a diversity of terrain on this ancient, moon-like surface. For Mars, after first tries with limited success during the Mariner 6/7 missions (e.g., Herr, Forney, & Pimentel, 1972; Martin, 1985) and the Mars Phobos mission (using Infrared Spectrometer for Mars; e.g., Bibring et al., 1989; Mustard et al., 1993), results from the Observatoire pour la Minéralogie, l’Eau, les Glaces et l’Activité (OMEGA) on board Mars Express (e.g., Bibring, Soufflot, et al., 2004; Bibring et al., 2005) and Compact Reconnaissance Infrared Spectrometer for Mars (CRISM) on board Mars Reconnaissance Orbiter (e.g., Murchie et al., 2007, 2009) demonstrated that this technique can change the face of a planet.

The main limitation of the VNIR spectroscopy technique is the lack of penetration, usually considered to be <100 μ‎m, thus strongly dependent on mantling/coating such as dust accumulation or space weathering. Thermal infrared (TIR) spectroscopy is more penetrative than VNIR (to a few millimeters or centimeters, depending on the substratum), enabling the determination of the mineralogy through a veneer of dust, as is the case in many locations on Mars. Nevertheless, it is a more difficult technique in the sense that it acquires a weaker signal emitted by rocks at longer wavelength (3–50 μ‎m; Kahle, Palluconi, & Christensen, 1993), and, consequently, it does not provide the same spatial resolution (one or two orders of magnitude lower than the 18 m/pixel reached by CRISM). TIR was first used to determine measurements of the temperature during Viking missions on Mars (using the Infrared Thermal Mapper of Mars instrument; Kieffer, Neugebauer, Munch, Chase, & Miner, 1972) with limited mineralogical application but was later used for widespread mineralogical application onboard the Mars Global Surveyor with the Thermal Emission Spectrometer (TES; Christensen et al., 1992, 2001) and on board the Mars Odyssey with the Thermal Emission Imaging System (THEMIS; Christensen et al., 2004). TIR orbital observations of the Moon are fewer than images from ground-based telescopes, but the Diviner Lunar Radiometer Experiment on board the Lunar Reconnaissance Orbiter (LRO) has recently given some interesting results (Paige et al., 2009; Song, Bandfield, Lucey, Greenhagen, & Paige, 2013).

In situ instruments based on these techniques have not been as numerous, especially because landers and rovers are more oriented toward local analyses than remote sensing. Nevertheless, mini-TES present on the Mars exploration rovers (MERs) Spirit and Opportunity (Christensen et al., 2003) has been an in situ tool comparable in its technique to the orbital TES instrument On the Moon, the Chinese missions Chang’E-3 and Chang’E-4, which landed in 2013 and 2019, respectively. The latter used a visible and near-infrared imaging spectrometer (He et al., 2014). Apart from these techniques, in situ mineralogy can be achieved by techniques used in the laboratory on Earth and adapted to landed missions. The “Chemistry and Mineralogy” (CheMin) instrument on board the Curiosity rover uses X-ray diffraction, a classic technique used on Earth, which was utilized for the first time on another planet one century after its discovery (Blake et al., 2013).

Elemental Chemistry by Orbital or In Situ Measurements

Studying the elemental chemistry of the surface of planetary bodies enables observations complementary to mineralogy. It also requires distinct techniques through the recording of higher energy photons emitted by surfaces in the X-ray and gamma ray wavelengths. Photons at these wavelengths are emitted through various processes. When exposed to cosmic rays (particles that come from the Sun and other stars), chemical elements in soil and rocks emit unique signatures at specific wavelengths in the form of gamma rays due to collisions in the first centimeter of the ground. In addition, some elements are naturally radioactive (U, Th, K), enabling a direct measurement of their presence. Various versions of gamma-ray spectrometers (GRS) have been sent to the Moon, Mars, and Mercury with successful global mappings (Boynton et al., 2002, 2004; Feldman et al., 1999, 2004; Goldsten et al., 2007; Lawrence et al., 1998). Neutron spectrometers (NS) on board space probes to the Moon, Mars, and Mercury are based on the same type of phenomena, but they measure neutrons instead of photons. Neutrons can form by interactions of cosmic rays with the ground. Neutrons are captured by atoms of hydrogen present in the ground, enabling the identification and quantification of water (Feldman et al., 1998, 1999, 2002). Radar sounding can be used in addition to NS for deeper assessment of water ice. GRS and NS instruments enable estimations of the elemental chemistry over several tens of centimeters with a coarse spatial resolution of several hundreds of kilometers.

In situ instruments have been built using similar techniques. On Venus landers (Venera 13 and 14 and Vega 2, 1982 and 1986, respectively), the in situ instruments were X-ray fluorescence spectrometers for determining the major rock-forming elements and GRS for radioactive element abundances (U, Th, K). On Mars, the alpha-particle X-ray spectrometer (APXS) instrument exposes rocks or soils to alpha (α‎) and X-rays emitted during the radioactive decay of the element curium. When X-rays and α‎ particles interact with atoms in the surface of rocks, they produce photons in the X-rays’ wavelengths that can be measured by spectrometers. Three versions of APXS were present onboard, respectively: the Mars Pathfinder Sojourner rover, the MER Spirit and Opportunity rovers, and the Curiosity rover (e.g., Gellert et al., 2006). An APXS instrument was also present on the Yutu rover of the Chang’E-3 mission to the Moon (Fu et al., 2014). APXS can identify all major elements in the uppermost tens of micrometers of soils and rocks. A new technique of laser ablation (laser induced breakdown spectroscopy) has successfully been used on the Mars Curiosity rover with the ChemCam instrument. ChemCam enables the identification and quantification of most major elements, some minor elements, and volatiles (including hydrogen), by analyzing the light emitted by a plasma initiated by a laser focused on the rock (Maurice et al., 2012; Wiens et al., 2012).

Table 1. List of Instruments Sent to Terrestrial Planetary Bodies to Analyze Their Composition

Instrument Name or Capability




Gamma-Ray Spectrometer, X-Ray Spectrometer


Apollo 15–16


UV and VIS Cameras




NIR Camera




Gamma Ray Spectrometer


Lunar Prospector


Neutron Spectrometer


Lunar Prospector


Spectrometer Infrared




X-Ray Fluorescence Spectrometer




Spectrometer Infrared




X-Ray Fluorescence Spectrometer




Moon Mineralogy Mapper (VNIR)




Hyper-Spectral Imaging Spectrometer




Spectral Profiler




Multi-Band Imager




Gamma Ray Spectrometer




X-Ray Fluorescence Spectrometer




Diviner Lunar Radiometer Experiment


Lunar Reconnaissance Orbiter


Lyman-Alpha Mapping Project


Lunar Reconnaissance Orbiter


Imaging Interferometer Spectrometer




Gamma Ray Spectrometer


Chang’E1 and 2


X-Ray Spectrometer


Chang’E1 and 2


Alpha Particle X-Ray Spectrometer



Moon (in situ)

VIS-NIR Imaging Spectrometer


Chang’E3 and Chang’E4

Moon (in situ)

Gamma-Ray and Neutron Spectrometer




X-Ray Fluorescence Spectrometer




Mercury Dual Imaging System




Mercury Atmospheric and Surface Composition Spectrometer




Infrared Spectrometer for Mars


Phobos 2


Thermal Emission Spectrometer


Mars Global Surveyor


Thermal Emission Imaging System


Mars Odyssey


Gamma Ray Spectrometer


Mars Odyssey


Observatoire pour la Minéralogie, l’Eau, les Glaces et l’Activité


Mars Express


Compact Reconnaissance Imaging Spectrometer for Mars


Mars Reconnaissance Orbiter


Miniature Thermal Emission Spectrometer


Mars Exploration Rovers

Mars (in situ)

Alpha Particle X-Ray Spectrometer


Mars Exploration Rovers

Mars (in situ)

Alpha Particle X-Ray Spectrometer


Mars Science Laboratory

Mars (in situ)

Chemistry & Camera


Mars Science Laboratory

Mars (in situ)

X-Ray Diffraction Mineralogy


Mars Science Laboratory

Mars (in situ)

Sample Analysis on Mars (Organic chemistry and isotopes)


Mars Science Laboratory

Mars (in situ)

X-Ray Fluorescence

Venera 13 and 14

Venus (in situ)

Ices at the Surface and Close Subsurface of Terrestrial Planets

Ices on Atmosphere-Free Bodies

Ices at the surface of atmosphere-free bodies of the inner Solar System are unstable and tend to sublime toward space. Although the Moon and Mercury are generally dry bodies at their surface, the presence of water ice in the subsurface of the polar regions of these bodies was, nevertheless, suggested by Earth-based radar analyses (Arnold, 1979; Slade, Butler, & Muhleman, 1992; Watson, Brown, & Murray, 1961). Measurements by the MESSENGER NS show decreases in the flux of epithermal and fast neutrons from Mercury’s north polar region that are consistent with the presence of water ice in permanently shadowed regions (Lawrence et al., 2013). The bistatic radar reflection on board the Clementine spacecraft has confirmed evidence for water ice at the lunar poles (Nozette et al., 1996). The NS on board Lunar Prospector has detected hydrogen present as water ice at both the north and south lunar poles (Figure 1; Feldman et al., 1998). Water ice may be found as buried bodies or pore fillings beneath a dry regolith. Water ice at a level of 1% has been confirmed by the LCROSS mission that impacted Moon polar regions and analyzed the escaped material in which hydroxyl bounds were observed by spectrometers (Colaprete et al., 2010). It has also been suggested that water ice can be found preferentially on the floor of permanently shadowed craters near both poles, which act as cold traps (Paige et al., 2010; Slade, Butler, & Muhleman, 1992), explaining the fact that water ice can be stable over geological timescales without sublimation toward space (e.g., Paige, Wood, & Vasavada, 1992; Paige et al., 2010).

The Surface Composition of Terrestrial Planets

Figure 1. Neutron map of the Moon polar regions above 70° north and south. A low neutron flux indicates their capture by hydrogen (Feldmann et al., 1998).

Image credit: NASA/Lunar Prospector/NS/ASU.

Ices on Mars Surface and Shallow Subsurface

Mars possesses two polar caps whose surface compositions are different. The northern cap is a ~1000 km × 1000 km, ~3 km thick body of water ice mixed with up to 10% dust (Figure 2; Phillips et al., 2008). The smaller southern cap (~400 km in diameter) has long been considered as dominated by dry ice (CO2 ice; Kieffer, 1979). However, high resolution images, thermal imagery and NIR spectroscopy have progressively shown that the layer of CO2 ice is only 10 to 20 m thick and displays concentric shapes typical of sublimation patterns, suggesting that this layer is currently under sublimation (Bibring, Langevin, et al., 2004; Byrne & Ingersoll, 2003; Titus, Kieffer, & Christensen, 2003). Below this 10 m thick layer, the southern polar cap is dominated by water ice mixed with dust, like the northern cap (Figure 2). However, significant amounts of buried CO2 ice have recently been detected from radar data. This buried dry ice could represent twice the current CO2 pressure if sublimed in the atmosphere and is bounded by layers with dust and water ice probably mixed with CO2 under clathrate hydrates (Bierson et al., 2016; Phillips et al., 2011).

The Surface Composition of Terrestrial Planets

Figure 2. (a) Mars water ice northern polar cap (~1000 km across). Spiral-shaped bands are deep troughs that are dustier. The image synthesizes altimetric data from Mars Orbiter Laser Altimeter (MOLA) and images from the Mars Orbiter Camera (MOC) wide-angle channel. Image Credit: NASA/JPL-Caltech/MSSS. (b) False color map of the southern polar cap (~400 km across) done by the OMEGA spectrometer. The pink color indicates dry ice (CO2) covering most of the surface; the blue color indicates water ice present around and below the CO2 ice (Bibring, Langevin, et al., 2004).

Credit: ESA/Mars Express/OMEGA/IAS.

Martian seasonal caps are due to the condensation of CO2 as frost over thicknesses reaching ~2 m. H2O frost in these seasonal deposits is minor but can form a thin frost layer (<100 μ‎m) down to equatorial regions at night when temperature drops below the frost point. CO2 frost forms during the cold temperatures of the winter seasons, down to –130°C, poleward of 50° to 60° of latitude north and south (Langevin et al., 2007). The defrosting of the seasonal caps produces a large diversity of landforms and textures, including dark spots due to quick punctual defrosting and spiders formed by local geysers due to translucent ice, contributing to the huge diversity of Martian landforms and processes (e.g., Piqueux, Byrne, & Richardson, 2003). Martian seasonal caps are also fundamental in the processes controlling the atmospheric circulation (e.g., Armstrong, Leovy, & Quinn, 2004).

Apart of polar caps, water ice is also present in the polar regions as near surface ice (<1 m depth). The NS on board the Mars Odyssey has detected a huge amount of hydrogen corresponding to a volumetric abundance of water ice from 10% to 60% (Figure 3; Feldman et al., 2002). This water ice is correlated to the presence of polygonal landforms formed by the cyclic effect of seasonal temperatures variations, as on Earth in arctic regions from Canada or Siberia (Mangold, Maurice, Feldman, Costard, & Forget, 2004). The Phoenix lander has provided the ground truth of the presence of water ice at 5 to 10 cm depth at its landing location 70°N of the latitude where polygonal cracks are ubiquitous (Mellon, Arvidson, Marlow, Phillips, & Asphaug, 2008). Further to the equator, glacial tongues initially referred to as lobate debris aprons have been identified as ice-bearing landforms from images and topographic data and were later confirmed as water ice rich from Mars Reconnaissance Orbiter radar data (Head et al., 2005; Holt et al., 2008; Mangold & Allemand, 2001; Squyres, 1989). Mars is the only terrestrial planet (apart from Earth) for which the water cycle plays such an important role in its past and present surface evolution. This is a well-studied topic in constant progress.

The Surface Composition of Terrestrial Planets

Figure 3. Neutron spectrometer map of the polar regions of Mars observed by the Mars Odyssey space probe. Hydrogen is assumed to be mainly found as water in the ground (Feldmann et al., 2002).

Image credit: NASA/Mars Odyssey/NS/ASU.

The Igneous Mineralogy at the Surface of Terrestrial Planets

Composition of the Ancient Crust of the Moon, Mercury, and Mars

The surface of planetary bodies is dominated by igneous rocks, with a large variety of composition and landforms related to various parameters including the size of the body, the initial bulk composition, and the presence of volatiles in its interior (Figures 47, Table 2). Existing models agree that all large, rocky, planetary bodies were partially molten and underwent differentiation early after their accretion (i.e., magma ocean stage), resulting in the formation of compositionally distinct layers in their interiors (Elkins-Tanton, 2012; Ringwood, 1979; Solomatov & Stevenson, 1993a, 1993b, 1993c; Urey, 1955). The mechanism for producing a primary crust depends largely on its size: in the case of small bodies like the Moon, the crystallization of additional buoyant minerals such as plagioclase may have resulted in a floatation crust (Ryder & Wood, 1977; Wood, Dickey, Marvin, & Powell, 1970), whereas for larger bodies (e.g., the Earth), the difference in liquidus and adiabatic gradients led to a bottom-up crystallization of the magma ocean with possible mantle overturn (Elkins-Tanton, 2012). Soon after differentiation, subsequent volcanism intruded the primary crust and erupted lavas at the surface, producing a secondary crust (e.g., Taylor & McLennan, 2009). On Earth, the original crust was destroyed billion of years ago and has been recycled by plate tectonics, forming a tertiary crust and leaving few, if any, clues about its formation. In contrast, Mars, Mercury, and the Moon are small enough to cool down by conduction and may therefore have preserved evidence of their primordial crusts.

The Surface Composition of Terrestrial Planets

Figure 4. New views of the lunar surface composition (a–f: nearside, orthographic projection centered at longitude 0°; g–i: farside, orthographic projection centered at longitude 180°). (a) Color composite of M3 low-resolution mode data for the lunar nearside. Red: pyroxene absorption at 2.0 µm. Green: reflectance at 2.4 µm. Blue: hydroxyl molecules at 3 µm (Credit: ISRO/NASA/JPL-Caltech/BrownUniv/USGS; Pieters, Goswami, et al., 2009). (b) M3 color composite designed to illustrate major mineral absorptions. Red: band depth at ~1000 nm. Green: band depth at ~2000 nm. Blue: reflectance at 1580 nm. In this composite, the feldspathic highlands are largely blue (with few mafic minerals) whereas the basaltic mare are variations of red and yellow, illustrating the presence and diversity of mafic minerals (Pieters, Goswami, et al., 2009; Staid et al., 2011). (c) Clementine UVVIS warped color-ratio mineral map at 200 m/pixel. Red: 750/415 nm. Green: 750/1000 nm. Blue: 415/750 nm. The red channel represents areas that are low in titanium or high in glass content (e.g., pyroclastics), or mature (e.g., pristine highlands). The green channel is sensitive to the amount of iron in the surface. The blue channel reflects the surfaces with high titanium or bright slopes and albedos (see Lucey, Blewett, Taylor, & Hawke, 2000). (d) Lunar Prospector GRS Th abundance (rainbow scale, 0.4–12.9 ppm, at 0.5 pixel/degree) are projected in transparency over the LRO WAC mosaic of the nearside. The red line outlines the limits of the Th-rich PKT. Stars represent return sampling sites from previous Apollo and Luna missions. Remaining terrains belong to the outer Felsdpathic Highlands Terrane as defined by Joliff et al. (2000). (e) Clementine Fe abundance (calculated after Lucey, Blewett, & Jolliff, 2000; rainbow scale, 0–20 wt%, at 0.5 pixel/degree) projected in transparency over the LRO/WAC mosaic of the nearside. The black contours outline the mare units (Image credit: NASA/LROC). (f) M3 and SP high-resolution observations allowed multiple, recent detections of Mg-spinel (pink triangles; Pieters et al., 2014), olivine (green dots; Yamamoto et al., 2010), and pure anorthosite (blue dots; Donaldson Hanna et al., 2014). Nearside red spots spectral anomalies (red squares; Hagerty et al., 2006) and occurrences of Mg-plutons (yellow squares; Tompkins & Pieters, 1999) are also represented. Background = LROC WAC mosaic at 100 m/pixel. (g) Lunar Prospector GRS Th abundance (rainbow scale, 0.4–12.9 ppm, at 0.5 pixel/degree) are projected in transparency over the LRO WAC mosaic of the farside. The yellow line outlines the limits of the Th-enriched South Pole Aitken Terrane. Remaining terrains belong to the central and outer Feldspathic Highland Terrane (FHT). (h) Clementine Fe abundances (calculated after Lucey, Blewett, et al., 2000; rainbow scale, 0–20 wt%, at 0.5 pixel/degree) are projected in transparency over the LRO WAC mosaic of the farside. The black contours outline the mare units (Source: LROC team). (i) The same detections as in (f) are shown here for the lunar farside, background = LROC WAC mosaic at 100m/pixel.

The Surface Composition of Terrestrial Planets

Figure 5. New views of Mercury’s surface composition from MESSENGER (orthographic projection centered at longitude 90°). (a) MDIS enhanced color global mosaic, which uses the 430, 750, and 1000 nm bands and places the second principal component, the first principal component, and the 430/1000 ratio in the red, green, and blue channels, respectively. Volcanic plains appear in orange tones, whereas the most ancient, opaque-rich crustal material appears in dark blue tones (Hawkins et al., 2007; Denevi et al., 2016). (b) Interpolated colour mosaic of MASCS 575 nm (red), 415/750 nm (green), 310/390 nm (blue) bands (Izenberg et al., 2014). (c) XRS map of the weight ratio of Al to Si (rainbow scale, 0–0.35) is overlain in transparency over the MDIS, NAC, and WAC 750 nm mosaic. (d) XRS map of the weight ratio of Mg to Si (rainbow scale, 0–0.75) is overlain in transparency over the MDIS, NAC, and WAC 750 nm mosaic (Weider et al., 2015). The Caloris basin, which is characterized by low Mg/Si and high Ca/Si, is clearly distinguishable on the upper right of each map. High Mg/Si and low Al/Si (and Ca/Si) ratios in volcanic plains suggest compositions comparable to terrestrial komatiites and boninites (e.g., Nittler et al., 2011).

Credit: NASA/Johns Hopkins University Applied Physics Laboratory/Carnegie Institution of Washington.

The Surface Composition of Terrestrial Planets

Figure 6. Comparison of some of the major elemental chemistry of planetary surfaces. Lunar highlands high Al/Si ratio is linked to predominance of anorthosite feldspar. Mercury has lower Al/Si and higher Mg/Si than typical lunar surface materials and terrestrial basalts, indicating a lower fraction of the common mineral plagioclase feldspar. Mars has low Al/Si and Mg/Si ratios related to the high proportion of iron-rich minerals.

Credit: NASA/Johns Hopkins University Applied Physics Laboratory/Carnegie Institution of Washington.

The Surface Composition of Terrestrial Planets

Figure 7. Distinctive surface features on planetary surfaces. (a), (b, and (c): Lunar volcanics as seen by LROC. (All image credits: NASA/GSFC/Arizona State University, source: LROC website). (a) Irregular mare patches near the crater Maskelyne (LROC image M1123370138R). (b) Low reflectance pyroclastics near a fracture on the western floor of Lavoisier crater (LROC image M105055584LR). (c) An oblique view of the northern portion of the Gruithuisen Gamma volcanic dome (LROC NAC M1106087898LR). (d), (e), and (f): Mercury volcanics and hollows as seen by MDIS. (All image credits: NASA/Johns Hopkins University Applied Physics Laboratory/Carnegie Institution of Washington. Source: NASA JPL photojournal). (d) Dome-like feature (white arrow) inside the rim of the Caloris impact basin (see Head et al., 2008, for interpretations, MDIS MET:108826877). (e) Enhanced-color MDIS mosaic showing a candidate site for an explosive volcanic vent in yellow and a double-ring basin with a smooth interior that may be the result of effusive volcanism. (f) Field of hollows in the western portion of the floor of Zeami impact basin (MDIS NAC 8072780). (g), (h), and (i): Mars volcanic features as seen from orbit and in situ. (g) Viking 1 mosaic of Olympus Mons (Source: NSSDCA image catalogue; Images credits: NASA/JPL). (h) Example of Hesperian olivine-bearing flat plain and crater floors in the southern highlands (Image: THEMIS IR day mosaic; see Figure 7 in Ody et al., 2013). (i) Mastcam image of the alkaline rock Bindi (width of about 10 cm) analyzed by the MSL rover Curiosity on sol 544 (Mastcam MR2192001000E1, Credits: NASA/JPL-Caltech/MSSS). (j) and (k) Venus volcanics as seen by Magellan radar (Images credits: NASA/JPL). (j) Pancake domes in Tinatin Planitia (Portion of Magellan C1-MIDR 15N009;1, framelet 36, source: NSSDCA image catalogue). (k) Computer-generated view of the Sapas Mons volcano (width is about 400 km for scale). The simulated hues are based on color images recorded by the Soviet Venera 13 and 14 spacecraft (image produced at the JPL Multimission Image Processing Laboratory).

The presence of a primary anorthositic, upper crust on the Moon was a major finding from the Apollo missions and is the foundation for all terrestrial planets differentiation models (Ringwood, 1979; Wood et al., 1970). Models predict that the lunar lower crust and mantle should be more enriched in denser, mafic minerals such as pyroxenes and olivine (e.g., Lin, Tronche, Steenstra, & van Westrenen, 2017; Snyder, Taylor, & Neal, 1992; Warren, 1985). Remote sensing observations from the lunar highlands and central peaks of impact craters indicate both lateral and vertical compositional heterogeneities contrasting with this schematic view. On the basis of Lunar Prospector gamma-ray data, the lunar crust and underlying mantle has been found to be divided into distinct terranes that possess unique geochemical and geological characteristics (Figure 4d, 4e, 4g, 4h; Joliff et al., 2000; Lawrence et al., 1998). The Procellarum KREEP Terrane (PKT) region indicates a high concentration in thorium and potassium, consistent with chemistry from returned samples. These Apollo samples include KREEP basalts, named after their enrichment in potassium (K), rare earth elements (REE), and phosphorus (P) that may be inherited from the last liquid layer to have crystallized from the magma ocean (Warren & Wasson, 1979). The concentration of heat-producing elements in the PKT (Figure 4d), possibly inherited from an asymmetric magma ocean crystallization, may have led to the nearside being more volcanically active than the farside. Other anomalies include the enrichment of iron of the South Pole Aitken basin (southern farside), thought to display deep crustal or upper mantle rock exposures or a differentiated impact melt sheet (Figure 4g; Nakamura et al., 2009; Pieters, Tompkins, Head, & Hess, 1997). Local anomalies also include detections of olivine and magnesium (Mg)-spinel and rock samples belonging to the Mg- and alkali-rock suites (e.g., Martinot, Besse, Flahaut, Quantin-Nataf, & van Westrenen, 2018; Song et al., 2013; Tompkins & Pieters, 1999; Yamamoto et al., 2010). These anomalies may be explained by the ubiquity of plutons that intruded the lunar pristine crust soon after its formation (Figure 4f, 4i; e.g., Papike, Ryder, & Shearer, 1998; Pieters et al., 2011; Pieters et al., 2014; Shearer & Papike, 2005; Shearer et al., 2006; Snyder, Neal, Taylor, & Halliday, 1995; Tompkins & Pieters, 1999). A large proportion of the current crust volume may consist of intrusions, making the identification of the primary crust fragments more difficult (e.g., Head & Wilson, 1992; Spudis & Davis, 1986).

Because of its comparable size, Mercury was suspected to have a similar primary, anorthositic crust (e.g., Brown & Elkins-Tanton, 2009). This hypothesis is consistent with the general lack of mineralogical diagnostic absorptions in the VNIR at Mercury’s surface (e.g., Lucey & Bell, 1989, Izenberg et al., 2014). However, data from MESSENGER’s orbital X-ray and GRS contrast with the Moon-related hypothesis. Mercury’s surface is characterized by relatively low iron (<2%) and iron oxides (<3%) but high Mg/silicone (Si) and low aluminium (Al)/Si and calcium (Ca)/Si ratios, comparable to komatiites and boninites on Earth (Figures 5c, 5d, and 6; Evans et al., 2012; Lawrence et al., 2010; Nittler et al., 2011; Peplowski et al., 2011; Robinson et al., 2008; Weider et al., 2012). Furthermore, these data have revealed a relatively high abundance of sulphur (up to 4 wt%) and potassium (up to 2400 ppm) for a presumed volatile-poor body. These results suggest that Mercury might not have formed at high temperatures and could have been formed from a different precursor such as enstatite chondrites (e.g., Evans et al., 2012; Nittler et al., 2011; Peplowski et al., 2012). The unexpected enrichment in volatiles could provide a mechanism for the surface hollows that point to recent outgassing (Blewett et al., 2011, 2013, 2016). Most hollows are located within Mercury’s ancient and heavily cratered highlands of unknown nature, referred to as low reflectance material (LRM; Figure 7e, 7f, and 5a; e.g., Head et al., 2008; Head et al., 2011; Nittler et al., 2011; Robinson et al., 2008; Vander Kaaden et al., 2017). The surface albedo of the LRM is abnormally low, suggesting the concentration of a darkening agent. Oxides are excluded, given the very low iron and titanium abundances measured by XRS (Murchie et al., 2015; Nittler et al., 2011; Robinson et al., 2008). Graphite offers an attractive new explanation for the highlands low albedo values as it has also been proposed to be the main component of Mercury’s primary (floatation) crust (Peplowski et al., 2016; Vander Kaaden & McCubbin, 2015). As LRM is generally found in large impact craters and ejecta, it could represent ancient (or perhaps primary) portions of Mercury’s crust, partly buried by subsequent volcanism (e.g., Peplowski et al., 2016).

Larger planetary bodies are not expected to retain float minerals but should rather develop basaltic crusts from melting during cumulate mantle overturn. Remote-sensing data and in situ analyses show that the martian ancient highlands are broadly basaltic in composition, that is, dominated by mafic minerals (olivine and pyroxene) with comparable amounts of low calcium pyroxenes (LCP) and high-calcium pyroxenes (HCP; e.g., Bandfield, 2002; McSween, Grove, & Wyatt, 2003; Mustard et al., 2005; Ody et al., 2012). Mars’s crust also displays a high iron abundance (Taylor & McLennan, 2009), explaining the predominantly low Mg/Si compared to other terrestrial bodies (Figure 6). Most highlands outcrops may correspond to volcanic terrains rather than to a primary crust (e.g., Grott et al., 2013). If preserved, a primary crust would be buried at depth, but impacts may have excavated it locally. Massive, LCP-rich cumulate outcrops at the bottom walls of the Valles Marineris and in surrounding craters’ central peaks could represent such remnants of Mars’s primary crust (Flahaut et al., 2013; Quantin, Flahaut, Clenet, Allemand, & Thomas, 2013). Nevertheless, these rocks may also represent early volcanic products. Indeed, early cumulate from a Mars magma ocean or moderate degrees of partial melting of the martian primitive mantle would lead to the formation of LCP-dominated igneous rocks (Baratoux, Toplis, Monnereau, & Sautter, 2013; Elkins-Tanton, Hess, & Parmentier, 2005). Anorthosite detections (>90% plagioclase) in the surroundings of the Hellas and Argyre impacts have also been interpreted as residue of an early, lunar-like floatation crust (Carter & Poulet, 2013). The findings of granitoïds from orbital data (Wray et al., 2013) and several felsic rocks, such as granodiorites, by the Curiosity rover at Gale crater (Figure 7i; Sautter et al., 2014, 2015) may suggest the presence of a felsic and more differentiated component in Mars’s crust. These observations are in agreement with the presence of a lower density material buried at depths below the southern hemisphere (Baratoux et al., 2014).

Table 2. Main Mineral Groups Detected on Terrestrial Planets

Primary Mineral Names














Plagioclase feldspar





Alkali feldspar













Fe- Ti- oxides (ilmenite, magnetite)



Phosphates (apatite)



Sulphides (pyrite, pyrrothite)





Sulphates (anhydrite)




Secondary Mineral Names





Ca- or Mg-sulfates



Fe-sulfates incl. jarosite



Ca- Mg- or Fe- carbonates


Chlorides or perchlorates


Phosphates (incl. apatite)



Fe-oxides (hematite, goethite)


Smectites (nontronite, saponite)










Opaline silica (incl. chalcedony)


Prehnite or Pumpellyte


Note. Y = Detected; Z = Suspected at detection limit or deduced from chemical analyses; ?: Debated from theoretical arguments). All of these mineral classes exist on Earth.

Composition of Volcanic Plains of the Moon and Mercury

Both the Moon and Mercury’s surfaces are dominated by mare-producing basaltic extrusions and show evidence of widespread effusive volcanism with a peak of activity early in their history (4.0 to 3.6 Gy ago; Figures 4, 5, and 7; Denevi et al., 2013; Head, 1976; Head & Wilson, 1992; Head et al., 2008; Head et al., 2009; Marchi et al., 2013). Basalts sampled by the Apollo and Luna missions range from 3.9 to 3.1 Ga in age; however, crater densities suggest that mare basalts were placed mostly between 4.0 and 2.0 Ga ago and up to ~1.0 Ga ago (Hiesinger, Head, Wolf, Jaumann, & Neukum, 2011; Hiesinger, Jaumann, Neukum, & Head, 2000). The last major phases of lunar volcanism produced spectrally distinct, high-titanium basaltic plains on the western nearside of the Moon, which display strong 1 μ‎m and weak 2 μ‎m absorptions on M3 VNIR data, consistent with olivine-rich basaltic compositions (Figures 5b, 5c, and 5e; Charette, McCord, Pieters, & Adams, 1974; Elphic et al., 2002; Pieters et al., 1980; Staid et al., 2011).

Mercury’s northern plains are formed by smooth, flow features that could have been placed on Mercury up to 1 Ga ago (Head et al., 2011; Marchi et al., 2011; Prockter et al., 2010). Mercury VNIR surface spectra generally lack diagnostic absorptions, but GRS and XRS data allow the characterization of the surface lava composition, which varies between basalt, boninites (effusive rocks with relatively high Mg, >8% wt., magnesium oxide [MgO] and intermediate dioxide form of silicon [SiO2] abundance), and komatiites (Figures 5b5d; e.g., Nitler et al., 2011; Vander Kaaden & McCubbin, 2016; Weider et al., 2015). In addition, high-resolution optical imagery obtained by the recent missions MESSENGER and LRO allowed the discovery of previously unidentified, potentially more recent (<100s My) volcanic eruptive events on both bodies (Braden, Stopar, Robinson, Lawrence, van der Bogert, & Hiesinger, 2014; Thomas, Rothery, Conway, & Anand, 2014). However, the relatively young “irregular mare patches” are located atop ancient lunar shield volcanoes and could originate from associated, ancient magmatic foams (Wilson & Head, 2017) or other processes such as explosive degassing (Figure 7a; Elder et al., 2017). If proven correct, the existence of late-stages volcanism on the Moon and Mercury would have major implications for models of interior dynamics and thermal evolution.

Besides lava flows, pyroclastics deposited by fire fountains represent another product of basaltic volcanism on the Moon and Mercury. Their spatial extent is highly variable, ranging from ~1 km2 to up to ~50 km2 on both the Moon and Mercury (Gaddis, Staid, Tyburczy, Hawke, & Petro, 2003; Kerber et al., 2011). Pyroclastic deposits are often distinguishable due to their low albedo and local to regional extent; their smoothness is comparable to rough lava flows with lower crater density at their surface (Figure 7b and 7e; e.g., Head, 1974; Head et al., 2008; Head et al., 2009). While the knowledge of their composition on Mercury from orbital data only remains limited, the complement between orbital and in situ data on the Moon enables a deeper understanding. Lunar pyroclastic deposits may have consisted largely of fragmented basalt, with substantial components of iron-bearing mafic minerals (pyroxenes, olivine) and smaller amounts of volcanic glass (Gaddis, Hawke, Robinson, & Coombs, 2000; Gaddis, Pieters, & Hawke, 1985; Gaddis et al., 2003). Laboratory analyses of picritic glasses samples from the Apollo 14 and 15 sites from several of these deposits show that the glasses had a greater depth of origin and lesser fractional crystallization than lunar mare basalts (Delano, 1986; Papike et al., 1998; Shearer & Papike, 1993). Recent studies suggest the presence of indigenous water in lunar deposits with abundances of up to 150 ppm within large deposits (Milliken & Li, 2017). Whereas lunar pyroclastic deposits tend to cluster near mare deposits, mercurian pyroclastic deposits are more widely distributed, often located on the floor of impact craters, with the exception of a cluster at the edges of the Caloris basin, which is partially filled with volcanic material (Figure 7d; e.g., Head et al., 2008; Kerber et al., 2009, 2011). Measurements of surface reflectance by MESSENGER indicate that the pyroclastic deposits are spectrally distinct from their surrounding terrains, with redder spectral slope in the VIS and NIR and a downturn at wavelengths shorter than ~400 nm, consistent with a lower iron content of the pyroclastic deposits with respect to the average surface of Mercury (Besse, Doressoundiram, & Benkhoff, 2015; Goudge et al., 2014).

In addition, lunar surface volcanic processes include landform morphology, spectral anomalies (“red spots”; Figures 4f and 7c), and granitic or rhyolitic components found in the Apollo sample suites, which suggest (limited) occurrences of non-mare, geochemically evolved (Si-enriched) volcanic deposits. Recent VNIR and thermal infrared spectroscopy, high-resolution imagery, and topographic data from LRO show steep lunar domes that could have been formed by silicic rocks (Glotch et al., 2010; Jolliff et al., 2011). The existence of evolved volcanic lithologies on the Moon implies either crustal assimilation, extreme fractional crystallization, or silicates/liquid immiscibility processes on the Moon.

Composition of the Volcanic Plains of Mars

Recent spectral image observations provide tantalizing clues about Mars’s volcanism. Global mapping suggests that the volcanic activity peaked in the Hesperian, although the earlier, pre-Noachian and Noachian record is partly—possible largely—missing (e.g., Greeley & Schneid, 1991; Grott et al., 2013; Werner, 2009).The presence of the Tharsis bulge, a 10 km high, 4000 km across plateau, suggests continuous volcanic activity from the Noachian (4.0–3.7 Gy) to recently (<100 My ;Neukum et al., 2004; Phillips et al., 2001). The Hesperian volcanism (3.0–3.7 Gy) is manifested predominantly by a mare-style, flood volcanism (Figure 7h; Syrtis Major Planum, Hesperia Planum, etc.) dominated by basaltic compositions with pyroxene and plagioclase compositions (Figures 8 and 9; Bandfield, 2002; Ody et al., 2012). Isolated volcanic edifices and corresponding mafic mineral exposures are found at Syrtis Major and around the Hellas impact basin (Williams et al., 2009). HCP is by far the predominant pyroxene type in Hesperian volcanic plains (Mustard et al., 2005; Poulet et al., 2007, 2009). This compositional difference in pyroxene type has been interpreted as a progressive cooling of the mantle (Baratoux et al., 2013). Olivine is associated with pyroxenes in many volcanic plains, in agreement with their effusive, low viscosity style (Figure 9; Ody et al., 2013). Olivine is also a major component of the Nili Fossae region, where it represents the presence of an ultramafic igneous region overlapping the LCP-rich ancient crust (Hamilton & Christensen, 2005; Hoefen et al., 2003). This region is located around the Isidis Planitia impact basin and thus may have been formed by impact melt with excavation of olivine from the mantle (Mustard et al., 2007). In contrast, isolated detections of dacite and quartz-rich products in the Syrtis Major area show that magmatic differentiation occurred locally well after the highland crust formation (Bandfield, Hamilton, Christensen, & McSween, 2004; Christensen et al., 2005).

The Surface Composition of Terrestrial Planets

Figure 8. TES maps of Mars plagioclase detections (top), pyroxene detections (middle), and dust (bottom). (Bandfield, 2002).


The Surface Composition of Terrestrial Planets

Figure 9. OMEGA maps of Mars pyroxene detections (top) and olivine detections (bottom). (Ody et al., 2012).

Credit: ESA/MarsExpress/OMEGA/IAS.

The Amazonian period (<3.0 Gy) is marked by larger volcanic edifices, including giant shield volcanoes (Olympus Mons, Elysium Mons, Arsia Mons) reaching 25 km in height (Figure 7g). Effusive plains in the Amazonia Planitia and South Elysium-Cerberus Fossae regions display activity as young as ~10 My according to crater counts (e.g., Grott et al., 2013; Plescia, 1990; Vaucher et al., 2009). Unfortunately, most Amazonian volcanic landforms on Mars are located in dusty regions. VNIR spectroscopic observations are thus limited to few dust-free windows. OMEGA data have shown that platy lava flows at Noctis Labyrinthus and Echus Chasma are basaltic with predominant detection of HCP (Mangold et al., 2010), similar to Hesperian plains. CRISM data on rims of fresh craters probing below dust have recently been used to provide more input into the Tharsis and Elysium volcanic areas. Results show a basaltic composition with no major changes during the Amazonian period but a significant contribution of LCP (Viviano-Beck, Murchie, Beck, & Dohm, 2017). Overall, a gamma-ray data set over Amazonian volcanism indicates a trend toward a more Si-poor content consistent with its apparent effusive style (Baratoux, Toplis, Monnereau, & Gasnault, 2011). The NS has also enabled the detection of excess chlorine in the recent volcanic plains of South Elysium and the nearby Medusae Fossae Formation, which could be the result of ash deposits or aerosols linked to the relatively recent volcanic episodes in these regions (Diez et al., 2009).

Several igneous rocks have been analyzed in situ by rovers. Bounce rock analyzed at Meridiani Planum displays a basaltic composition similar to the martian meteorite class called shergottites (Zipfel et al., 2011). A picro-basalt class of rocks (Adirondack class) was observed at Gusev crater. Modeling of mini-TES spectra indicate an ultra-mafic composition with predominance of olivine (36%) compared to pyroxene (22%), plagioclase (18%), and silica glass (24%; Hamilton & Ruff, 2012; Ruff et al., 2006). Husband Hills at Gusev provide evidence for more evolved rocks such as tephrites and trachybasalt that may be related to the same (but more fractionated) magmatic system as the Adirondack class (Gellert et al., 2004; McSween et al., 2008; Squyres et al., 2006). Alkali-rich igneous rocks (Na2O + K2O >4%–5%) have also been analyzed at Gale craters in rocks such as Jake-M, a mugearite analyzed by APXS (Stolper et al., 2013), and a series of trachyte and trachy-andesite rocks analyzed with ChemCam data (Figure 7i; Cousin, Sautter, et al., 2017; Sautter et al., 2015).

Apart from bedrock exposures and their mechanical grinding into a regolith, the martian surface is covered by many dark dune fields, which cover many crater interiors and part of the northern lowlands (Figure 9). Orbital data helped to determine a mineralogy that is consistent with nearby volcanic plains, usually dominated by mafic minerals (e.g., Ody et al., 2013). Local analyses of sand dunes in Gale crater by Chemistry and Mineralogy (CheMin) onboard Curiosity show a slight predominance of plagioclase (24%) over pyroxene (22%) and a large amount of glass/amorphous phases (35%; Achilles et al., 2017). Enrichments in sulphur and chlorine in sand dunes suggest a contamination by aerosols (Cousin, Dehouck, et al., 2017). Glasses have also been detected from orbital data, but it is unclear whether they come from ash deposits, impact craters, or recent weathering (e.g., Farrand, Wright, Rogers, & Glotch, 2016; Horgan & Bell, 2012). Actually, a large part of the northern plains, initially interpreted to be andesitic due to the high plagioclase content (Figure 8; Christensen et al., 2001), are suspected to be covered by glass-rich sand grains (e.g., Horgan & Bell, 2012; Wyatt & McSween, 2002).

Surface Composition of Venus

The solid surface of Venus is difficult to observe directly because of the thick, optically opaque atmosphere. Between the 1970s and 1990s, the Soviet and American radar orbiters scanned the surface (e.g., Moroz, 1983; Pettengill et al., 1980; Rzihiga, 1987; Saunders & Pettengill, 1991). Radar images show a strong diversity of volcanic landforms all over the planet, ranging from conic volcanoes associated with >100 km long lava flows to convex, pancake domes (Figure 7j7k; e.g., Barsukov et al., 1986; Guest et al., 1992; Head, Crumpler, Aubele, Guest, & Saunders, 1992; Masursky et al., 1980; McKenzie, Ford, Liu, & Pettengill, 1992; Pavri, Klose, & Wilson, 1992; Senske, Schaber, & Stofan, 1992). The Soviet lander Venera 7 first landed on Venus in 1970 and resisted the extreme physical surface conditions (~460°C, 95 bar; e.g., Fimmel, Colin, & Burgess, 1983). Some of these landers acquired data over three hours, taking pictures of the Venusian surface at the toe of volcanic-rifted Beta-Phoebe Region. Images obtained show a surface that is relatively flat, covered by gray-toned, polygonal blocks of rocks similar to terrestrial lava flows fractured mechanically by cooling and weathering (Figure 10). Numerous subangular particles of various sizes (dust to pebbles) fill voids between the rocks (Florensky et al., 1977; Greeley, 1987; Moroz, 1983). Sand-size particles may be transported by wind, forming dunes, as has been observed by the Magellan orbiter radar imager (Avduevskii et al., 1976; Greeley et al., 1992; Weitz, Plaut, Greeley, & Saunders, 1994).

The Surface Composition of Terrestrial Planets

Figure 10. Color-corrected view of Venus surface from Venera 13.

Credit: Russian Academy of Sciences.

The composition of Venus is poorly known, except in three equatorial lowland zones (Phoebe Regio and Atla Regio) over which Venera 13 and 14 and Vega 2 landers analyzed the surface. Analyses of major elements show abundances typical of basaltic composition (~45-47 wt.% SiO2, 9 wt.% FeO, 7 wt.% MgO, and 16 wt.% Al2O3; Surkov et al., 1982; Surkov, Moskaleva, et al., 1986). This chemistry suggests that the Venusian rocks would be dominated by mafic minerals and plagioclases, with rocks ranging from olivine gabbro-norite to tholeiitic basalt (Surkov et al., 1982; Surkov, Moskaleva, et al., 1986). The high concentration of elements such as Th, U, and K (up to 5 wt.% K2O) measured by the GRS of the Venera 8 lander suggests that more differentiated rocks also exist, perhaps in relation to the volcanic domes observed by the Magellan radar (Basilevsky, Nikolaeva, & Weitz, 1992; Surkov, 1983; Surkov, Kirnozov, et al., 1986; Vinogradov, Surkov, & Kirnozov, 1973).

Alteration Mineralogy at the Surface of Terrestrial Planets

Aqueous Alteration

Secondary minerals formed by the alteration of primary (igneous) minerals include a vast range of phases, from hydrated silicates—including the huge family of clay minerals, various form of hydrous silica, a series of salts formed by precipitation (sulphates, carbonates, chlorides, phosphates, etc.), and all kinds of oxides such as ferric or manganese oxides (Table 2). On Earth, these minerals are fundamental in the understanding of aqueous processes, at the surface under ambient temperatures or below the surface down to deep regions of the crust. Venus is too warm for liquid water to exist, and most of the alteration is present as gas-rock interaction. On Mercury and the Moon, these processes are limited due to the lack of surface fluids and atmosphere and relatively dry mantles. Deep alteration of the crust may be possible, as has been deduced for asteroids. No widespread, definitive detection of secondary minerals has ever been made from remote sensing or in situ instruments on these two bodies, but the VNIR spectral imager M3 detected absorption features at 2.8 to 3.0 μ‎m at the surface of the Moon that are typically attributed to hydroxyl- and/or H2O- bearing minerals (Figure 4a; Pieters, Goswami, et al., 2009).

On Mars, first hint of the presence of hydrous minerals date from the analyses of the first martian meteorites (e.g., Karlsson, Clayton, Gibson, & Mayeda, 1992), but the lack of hydrous signatures in global thermal mapping by Mars Odyssey suggested an alteration localized or limited in intensity (Bandfield, 2002; Christensen et al., 2001). However, VNIR instruments OMEGA and CRISM have led to unprecedented, widespread detections of aqueous minerals at the surface of Mars (Figure 11; e.g., Bibring et al., 2005, 2006; Carter, Poulet, Bibring, Mangold, & Murchie, 2013; Gendrin et al., 2005; Murchie et al., 2009; Mustard et al., 2008; Poulet et al., 2005). The lack of detections of aqueous minerals in the TIR range remains poorly understood, although the coarse resolution of these data and the surface texture could both have degraded the signal enough to limit the detection of clay minerals (Michalski, Poulet, Bibring, & Mangold, 2010).

The Surface Composition of Terrestrial Planets

Figure 11. Global map of hydrated mineral sites at the surface of Mars detected using the reflectance data acquired by the Mars Express OMEGA and from the MRO. CRISM imaging spectrometers.


Clay minerals, also referred to as phyllosilicates, or sheet silicates, constitute a group of minerals containing water as OH- and/or H2O- molecular bonds. Formation of clay minerals by aqueous alteration can be due to various processes such as hydrothermal alteration subsequent to impact or volcanism, diagenetic alteration by burial of sediments, and weathering at the surface. Recent studies suggest that these three processes have occurred on Mars (e.g., Carter et al., 2013; Ehlmann et al., 2013; Milliken & Bish, 2010). Hydrothermal alteration is recognized from the proximity of phyllosilicate detections with volcanic and/or impact landforms, or from crustal rocks that were submitted to high thermal gradient before excavation by impact craters (Ehlmann et al., 2011; Marzo et al., 2010; Michalski et al., 2013; Poulet et al., 2005; Viviano, Moersch, & McSween, 2013). Minerals in these locations define complex assemblages of various hydrated silicates, with large proportion of high-T phases (chlorite, prehnite, serpentine, etc.), as observed in the central peaks of impact craters (Bultel, Quantin, Andreani, & Clenet, 2015; Ehlmann et al., 2011; Loizeau et al., 2012; Quantin et al., 2013).

Weathering tends to produce a more limited number of minerals compared to hydrothermal alteration, with predominance of smectite clays and lack of high-T phases. Weathering was first proposed to explain the clay minerals found in the region of Mawrth Vallis (Figure 12; Bishop et al., 2013; Loizeau et al., 2007, 2010; McKeown et al., 2009; Wray, Ehlmann, Squyres, Mustard, & Kirk, 2008). In this huge region (300 km × 300 km), Al-rich clays (dominated by kaolinite) are superimposed over Fe/Mg-rich smectites (dominated by nontronite), in a similar way as observed in terrestrial soil profiles. Several other locations on Mars display a similar stratigraphy in ancient terrains, suggesting that weathering was a planet-wide process at their time of formation, in the late Noachian epoch (~3.7–3.8 Gy; Carter, Loizeau, Mangold, Poulet, & Bibring, 2015). Clay minerals are also detected in sedimentary rocks, such as presumed lacustrine sediments (e.g., Ismenius Cavus, Jezero, Eberswalde, and Terby craters; Ansan et al., 2011; Dehouck, Mangold, Le Mouélic, Ansan, & Poulet, 2010; Ehlman et al., 2008; Rice et al., 2013), but it is often difficult to distinguish between an origin by detrital deposition and one by authigenic processes such as diagenesis (e.g., Milliken & Bish, 2010). For instance, the Fe2+-saponite observed by the X-ray diffraction instrument CheMin of the Curiosity rover has been interpreted as diagenetic (Figure 13; McLennan et al., 2014; Vaniman et al., 2014).

The Surface Composition of Terrestrial Planets

Figure 12. Composite 3D view of HRSC mosaic of the Mawrth Vallis plateau with superimposed colors indicating phyllosilicates (with higher concentration in yellow to red spots).

Credit: ESA/Mars Express/OMEGA/HRSC/DLR/IAS.

Clay minerals are mainly detected in outcrops of the Noachian crust and sedimentary layers within (Figures 11 and 12), while Hesperian and Amazonian volcanic regions do not display significant clay detection (Ehlmann et al., 2011; Mangold et al., 2007; Poulet et al., 2005). This observation led to the reference of the “phyllosian” for the early period of clay formation (Bibring et al., 2006). Nevertheless, limited occurrences of clay minerals related to regional activities have been reported in Hesperian and Amazonian terrains (Carter et al., 2013; Ehlmann et al., 2013; Murchie et al., 2009; Thollot et al., 2012; Sun & Milliken, 2014; Weitz, Bishop, Thollot, Mangold, & Roach, 2011). While clay detections imply surface conditions drastically different from current ones, the exact formation conditions remain debated (acidic or neutral pH, periglacial or more temperate climate, etc.), with strong implications on the timescales considered.

Sulphate minerals (Mg-, Ca-, and Fe- sulphates such as kieserite or gypsum) were found in the bedrock almost simultaneously by the Opportunity rover and by OMEGA in the winter of 2004 (Gendrin et al., 2005; Squyres et al., 2004). Sulphate minerals are present in equatorial layered deposits from Valles Marineris to Meridiani Planum locally in association with iron hydroxides (Chojnacki & Hynek, 2008; Flahaut et al., 2015; Flahaut, Quantin, Allemand, Thomas, & Le Deit, 2010; Gendrin et al., 2005; Mangold et al., 2008; Massé et al., 2008), in a few isolated basins (such as Gale crater; Milliken, Grotzinger, & Thomson, 2010), and as eolian deposits surrounding the northern polar cap (Fishbaugh, Poulet, Chevrier, Langevin, & Bibring, 2007; Langevin, Poulet, Bibring, & Gondet, 2005; Massé et al., 2010). The formation of sulphate deposits requires the presence of sulphur from either volcanic outgassing or sulphide alteration (e.g., pyrite) or both (e.g., Dehouck et al., 2012; King & McLennan, 2010). Meridiani Planum sediments observed by the Opportunity rover contain ~30% of Mg/Fe sulphates (Figure 13; Squyres et al., 2004). These sediments also include local observations of jarosite, a K/Na-Fe-sulphate that forms only in acidic conditions (pH <4; Tosca & McLennan, 2006). These sediments are present over the 30 km traverse of Opportunity and over a 500 km × 500 km region. Their relatively late stage deposition (mostly in the Hesperian) led to the conclusion that they formed during a drier, more acidic period following the phyllosian period (Bibring et al., 2006).

The Surface Composition of Terrestrial Planets

Figure 13. Mars as seen from the surface by the Opportunity and Curiosity rovers. (a) Meridiani Planum imaged by Opportunity shows half-cm small spherules accumulated on the ground after their weathering from the bedrock. Those concretions are hematite-rich (b) hematite being the mineral initially detected from orbit by thermal imagery. (c) PanCam/Opportunity mosaic of the sulphate-rich sediments on the rim of the Endurance impact crater. (Credits: NASA/JPL/MER/MI). (d) Mastcam/Curiosity mosaic of the Mont Sharp showing in the foreground phyllosilicate-bearing mudstones capped by darker eolian sandstones to the right (with basaltic composition). In the background, sediments from the upper Mount Sharp are too dusty for any remote detection of minerals. (e) View of the phyllosilicate-bearing mudstones of the Yellowknife Bay region that are fractured and less resistant that the sandstones to the top left of the image. At this location, brushed mudstones (f) display a lighter-toned, greyish color typical of dust-covered rocks. White nodules and veins correspond to calcium sulfates formed by precipitation of fluids after burial of sediments.


Assuming the presence of a relatively warmer early Mars with a CO2 atmosphere thicker than the current one, carbonates should be present to compensate for the loss of CO2. Carbonates have been detected by CRISM as Mg-rich carbonates (hydromagnesite) in relation to the alteration of an olivine-rich bedrock at Nili Fossae (Ehlmann et al., 2008), while magnesite and/or calcite have been reported in central peaks of impact craters (Michalski & Niles, 2010; Wray et al., 2016), and siderite (Fe-carbonate) has been detected in situ by the Spirit rover (Morris et al., 2010). Overall, the number of detections of carbonates remains limited and dominated by exhumed outcrops likely due to subsurface fluid circulation, leaving unsolved the link with atmospheric processes (Edwards & Ehlmann, 2015; Wray et al., 2016).

At the rover scale, diagenetic processes are highlighted by the presence of concretions, veins, raised fracture fills, or halos of a different composition than the host rock (Figures 13b and 13f). These features observed by rovers display various mineralogy: Ca-sulphate veins observed by Opportunity at Meridiani Planum and Curiosity at Gale crater (Nachon et al., 2014; Squyres et al., 2012) and Fe-rich concretions (the so-called blueberries made of hematite) observed by the Opportunity rover and initially identified by the TES orbital spectrometer (e.g., Christensen et al., 2000; Squyres & Knoll, 2005). In addition, Mg- and Fe-sulphates concretions (including jarosite), Mn-rich fracture fills, and Si-rich halos (including opal-A and -CT) were recently observed by Curiosity (Frydenvang et al., 2017; Lanza et al., 2016; Nachon et al., 2017). These minerals are formed by precipitation due to fluid circulation at various pH and oxidation states; they provide precious information on the postdepositional history of sedimentary layers.

Weathering of Rocks by Solid-Gas Interactions

Atmosphere-bearing planetary bodies (Venus and Mars) present a type of weathering with no or a minor amount of liquid water due to various processes, such as solid-gas interaction, ultraviolet (UV) photo-dissociation, or radiolysis by cosmic rays. The Venus atmosphere is so dense and hot that it can generate specific interactions with the surface. Venus’s atmosphere is composed of a massive compound of CO2 (~96.5%) and N2 (~3.5%), with many trace gases such as CO, SO2, OCS, HCl, and HF (e.g., Fimmel et al., 1983; Gel’man et al., 1979; Hoffman, Hodges, McElroy, Donahue, & Kolpin, 1979; Moroz et al., 1979, 1980; Oyama, Carle, Woeller, & Pollack, 1979; Oyama et al., 1980; Vinogradov et al., 1970), which are concentrated in the ~50 km–in-elevation tropopause (e.g., Pollack, Toon, & Boese, 1980). Trace gases may derive from two processes: Volcanic degassing may explain the presence of S2, H2S, COS, HCl, HF, and H2 while photochemical dissociation and recombination may explain the presence of SO3, Cl2, and O2 (e.g., Hoffmann, Oyama, & von Zahn, 1980; Khodakovsky, 1982; Prinn, 1973). In this context, atmosphere-surface interactions play a significant role. The near-surface troposphere may be in quasi-equilibrium with the surface rocks (Kerzhanovich et al., 1979). Models and experiments developed after Venus exploration tried to determine the chemical reactions between the surface and the atmosphere of Venus, leading to the discovery of chemically weathered rocks (e.g., Fegley & Prinn, 1989; Khodakovsky, 1982). For instance, Ca-bearing pyroxenes and feldspars could react with sulphur-bearing molecules to form sulphates and serve as a sink for sulphur (Fegley& Prinn, 1989).

In the early 1980s, the sulphur cycle was proposed to explain spacecrafts’ and Earth-based observations of the sulfuric acid clouds, sulphur-bearing gases in the lower atmosphere, and sulphur-bearing phases on the Venusian surface (Khodakovsky, 1982; Prinn, 1985; Von Zahn, Kumar, Niemann, & Prinn, 1983). Pyrite (FeS2) and pyrrhotite (Fe1–xS with x = 0 to 0.2) have been invoked to explain the low radar emissivity in the highlands (Pettengill, Ford, & Simpson, 1996). But these minerals would be unstable at the surface of Venus because the atmosphere would be too oxidizing, leading to their decomposition into iron oxides and sulphates (Fegley, Lodders, Treiman, & Klingelhöfer, 1995, Fegley, Zolotov, & Lodders, 1997). Ferro-electric minerals (Shepard, Arvidson, Brackett, & Fegley, 1994) or a coating of volatile metallic vapours, halides, and sulphides emitted from volcanoes (Brackett, Fegley, & Arvidson, 1995; Pettengill et al., 1996), would be more plausible candidates for the radar properties of the Venusian highlands. The oxidation (redox) state of the surface of Venus and its lower atmosphere are crucial to understand the chemical weathering. The oxygen fugacity (effective partial pressure) of the near-surface atmosphere would control the iron oxidation state and the type of trace gas at the surface (CO2, SO2, and H2O) compared to reduced compounds (CO, OCS, H2S, and H2; Fegley et al., 1997; Johnson & Fegley, 2002). For instance, magnetite would be stable only in the lowest and hottest plains because the CO-CO2 conversion is quenched above an elevation of –0.7 km where cooler temperatures exist. Conversely, hematite would form at a higher elevation due to higher atmospheric oxygen fugacity (Fegley et al., 1997) in agreement with spectral data at the Venera 9 and 10 landing sites indicating the presence of hematite (Pieters et al., 1986).

With regard to oxidation processes, Mars has been the “red planet”—a color suggestive of iron oxidation—for Earth observers since the early stages of astronomy. However, it has been debated for decades whether the reddish oxidized surface is a by-product of the aqueous history of Mars or a result of the following 3 billion years long dry period. The first in situ observations made by Viking landers enabled interpretations of the red color of the regolith as a consequence of an enrichment in iron oxides (e.g., hematite) with an oxidation present over a thickness of ~10 cm (Toulmin et al., 1977). Mossbauer spectrometer and magnets present on board MER rovers have confirmed the presence of ferric phases (hematite, maghemite, and goethite) in soils (Madsen et al., 2009; Morris et al., 2006). Although identified thanks to the ferric oxide red color, the bulk dust composition on Mars is basaltic, close to the average Mars composition without significant excess iron but with relative high abundances of sulphur and chlorine (Berger et al., 2016). Dust may thus correspond to a mechanical grinding of Mars surface rocks with additional components, such as poorly crystalline Fe3+ phase(s), often referred to as nanophase oxides, and chlorine- and sulphur-bearing phases (Morris et al., 2006). The global mapping of the Fe3+ absorption bands by the OMEGA spectrometer fits well with the region dominated by the homogeneous red color in visible light and the low thermal inertia typical of loose material (Figures 8c and 9c; Poulet et al., 2007; Ruff, 2004).

The nature of Mars’s surface oxidation remains only partly understood, but the photochemistry of atmospheric molecules seems to be a strong source of oxidation (Huguenin, 1973, 1982). It was observed experimentally that UV radiations through the photo-dissociation of water vapor produce O2, and then H2O2 by recombination of the dissociated molecules (Huguenin, 1982; Hunten, 1979; Zent, 1998), which in turn oxidizes magnetite to ferric oxides (hematite and maghemite). These molecules can diffuse into a porous regolith down to several meters and interact at the surface of grains/rocks (Zent, 1998). Rocks brushed by rovers reveal, however, that this oxidation does not penetrate deep (<0.1 mm) into rocks (Figure 13f). Measurements from Mössbauer on board the Spirit rover on the undisturbed and brushed surfaces of the rock named Mazatzal suggest the presence of a dark iron oxide likely formed by oxidative weathering (Morris et al., 2004). A thin oxidation coating (of unknown thickness but <1 mm) has also been observed from analyses of meteorites found along the traverse of the Opportunity rover (Ashley et al., 2011; Schröder, Bland, Golombek, Ashley, Warner, & Grant, 2016). Soil-atmosphere interactions through UV photo-dissociation on Mars may not be very efficient due to the low abundance of water vapor, but, at a timescale of few billion years, these processes are fair explanations for the presence of the red, oxidised surface material on Mars without the involvement of liquid water.

Other oxidants are found in martian soils, especially as a group of chlorine-bearing minerals called perchlorate. Perchlorates have been detected by the Phoenix lander at 70°N latitude (Hecht et al., 2009), in locations where ice is present close to the surface (Feldman et al., 2002). The Curiosity Sample Analysis for Mars instrument has found perchlorates at low latitudes at Gale crater (Glavin et al., 2013). A reanalysis of Viking data concluded that perchlorates were present in the soils analyzed by Viking (Navarro-Gonzales, Vargas, de la Rosa, Raga, & McKay, 2010). The processes of formation of perchlorates remain undetermined and could involve UV photo-dissociation, as for hydrogen peroxide (Schuttlefield, Sambur, Gelwicks, Eggleston, & Parkinson, 2011); radiolysis from the interaction with cosmic rays (Wilson, Atreya, Kaiser, & Mahaffy, 2016); or thin films of liquid water (Hecht et al., 2009). In the latter case, a similar process has been observed on Earth in Antarctica and in the Atacama Desert. Hydrogen peroxide H2O2 and perchlorates are both superoxidants with strong reactive effects, especially on organics. Enhanced oxidation due to UV photo-dissociation processes of surface water has also been proposed for Meridiani Planum sediments, questioning the role of similar process in the past (Hurowitz, Fisher, Tosca, & Milliken, 2010). While Mars is currently oxidized despite a globally reducing CO2 atmosphere, the degree of oxidation of Early Mars (with a potentially denser CO2 atmosphere) is an important question. Indeed, alteration pathways are strongly influenced by the level of oxidation, as demonstrated by laboratory experiments showing a decrease in clay formation under wet oxidizing conditions (Dehouck, Gaudin, Chevrier, & Mangold, 2016).

Weathering of Atmosphere-Free Bodies

Atmosphere-free planetary bodies (the Moon, Mercury, and asteroids) display a modification of their surface related to interactions with space that is generically referred to as space weathering. Space weathering, also referred to as optical maturation, has been discovered in lunar soils (lunar “agglutinates”) after Apollo exploration (e.g.,Pieters et al., 2000, and references therein), although it had already been suggested by Gold (1955) based on telescope observations. Space weathering is particularly identifiable from visible and infrared spectroscopy by a darkening of surfaces. Apollo regolith samples were found to be darker, with redder continuum spectra, than powders of lunar rocks of similar composition (e.g., Hapke, 1973; Hapke, Cassidy, & Wells, 1975). An illustration of this process can be observed on the Moon and on Mercury with the presence of bright rays formed around young craters that expose unweathered material above a weathered regolith exposed at the surface for longer times.

Space weathering can be classified under weathering processes (as such forming coatings or varnishes) because the darkening occurs at the surface of grains (on a micrometer scale) through the presence of metallic iron (Fe0), also reported as nanophase-reduced Fe, npFe0, or submicroscopic metallic iron, depending on the authors (e.g., Hapke, 2001; Pieters et al., 2000). Various processes have been proposed to explain this Fe-rich coating, such as interactions of grains with micrometeorites, solar wind, and/or cosmic-ray bombardments. The most popular scenario is vaporization of grains by solar wind and subsequent condensation of iron (Keller & McKay, 1993, 1997), but the other processes may contribute as well. Space weathering exists on most atmosphere-free bodies at various levels. Mercury is the most affected as it is the closest to the sun and so is hit by stronger solar winds than the Moon (Lucey & Riner, 2011; Noble & Pieters, 2003; Slavin et al., 2010). The space weathering identified on the asteroids has different spectral properties than on the Moon, suggesting that the target composition is an important parameter (Clark, Hapke, Pieters, & Britt, 2002). The study of this effect on various bodies will help in understanding the Moon and Mercury as well.

Future Missions

The improvement in the detection, mapping, and in situ analyses of planetary surface rocks enable scientists to progressively fill gaps in the knowledge of the composition of these planetary bodies (Table 2). These observations also provide fundamental inputs for the understanding of Earth’s evolution, from primordial magma ocean to specific alteration processes. While the Moon and Mars have been studied extensively, Mercury and Venus have long been looked over. While this gap has started to be reduced for Mercury, the difficulty to access the Venus surface has been challenging and will require more technological improvement to allow data acquisition during the descent of landers or by balloon or to enable landers a longer lifetime at the huge surface pressure. This section provides a nonexhaustive list of future missions and instruments that will continue to improve humankind’s knowledge of planetary surfaces.

The Moon

The Moon is recognized as a natural and accessible laboratory to study planetary processes such as differentiation or impact processes and presents a unique record of the Solar System bombardment history. Still, despite more than 40 years of exploration missions, the Moon’s countless mysteries have barely begun to be solved. VNIR and TIR spectroscopy has shown a diversity of lunar rocks, that are not present in the current sample collection. In addition, recent developments in analytical techniques have allowed for a better characterization of returned samples and have raised new questions. Determining the surface composition more precisely, from orbit and in situ, is critical to future lunar missions aiming to address (a) the vertical and lateral composition of the lunar crust and mantle as well as the extent and asymmetry of the magma ocean (e.g., Chang’E-4, Chang’E-5, SELENE-2, Moonrise) and (b) the nature, origin, and distribution of lunar volatiles (e.g., Luna 25, 26, 27, Resource Prospector). Upcoming missions will focus, for the first time, on in situ exploration of both lunar farside (e.g., Chang’E-4) and polar regions, with the aim to combine science exploration and in situ resources utilization. A wide range of techniques are considered to allow a better remote and in situ characterization of the volatile nature and abundance and the polar regolith properties (e.g., low orbit and in situ neutron spectrometer, VNIR spectroscopy, mass spectrometry, ground-penetrating radar) and will reinforce knowledge of the lunar surface composition.


Before the MESSENGER mission in 2011, data related to Mercury were scarce. Early studies based on telescopic observations and the Mariner 10 flybys in 1974 and 1975 already demonstrated that Mercury was unique among the terrestrial planets, with an unusually high metal to silicate ratio and a lack of diagnostic absorptions in the VNIR. Providing the first complete photographic and spectroscopic mapping of the planet, the MESSENGER mission has revolutionized scientists’ knowledge of Mercury. Still, much remains to be discovered regarding Mercury’s surface and bulk composition and its surprising enrichment of volatiles. The European Space Agency (ESA)–JAXA BepiColombo joint mission to Mercury was launched in October 2018. It includes two VNIR and TIR spectrometers: Spectrometers and Imagers for MPO BepiColombo Integrated Observatory (SIMBIO-SYS, 0.4–2µm) and Mercury Radiometer and Thermal Imaging Spectrometer (MERTIS, 7–14 µm); a Mercury Imaging X-ray Spectrometer (MIXS); and a Mercury Gamma-Ray and Neutron Spectrometer (MGNS) that will help address these questions (e.g., Benkhoff et al., 2010; Rothery et al., 2010).


Mars exploration will continue with one lander (Insight) landed on November 26, 2018, and three rovers (National Aeronautics and Space Administration [NASA] 2020 mission, ExoMars, China HX-1) planned for launch in 2020. The Insight mission is a lander that is not devoted to analyzing surface composition, but it will provide geophysical data from a seismometer (SEIS), an instrument measuring thermal flux (HP3), and a radioscience experiment (RISE) that will improve the understanding of the subsurface and core size and state structure and thus will be helpful for placing the volcanic evolution and the related surface composition of Mars in its context.

On board future rovers, instruments devoted to rock analysis will include VNIR spectrometers, which are fundamental in orbital data measurements, on both ExoMars (MicrOmega) and the NASA 2020 rover (SuperCam). Raman spectrometry will complement infrared spectroscopy in mineral identification for the ExoMars rover with the RLS (Raman Laser Spectrometer) instrument and for Mars 2020 with both SuperCam and Scanning Habitable Environments with Raman & Luminescence for Organics & Chemicals (SHERLOC) instruments. On this rover, the SuperCam suite of instruments will include laser ablation spectrometry inherited from ChemCam as well and will be associated with a new X-ray fluorescence instrument (Planetary Instrument for X-ray Lithochemistry [PIXL]), which will enable a microscopic mapping of the chemistry close to what can be done in terrestrial laboratories. The 2020 NASA rover will therefore be able to provide the composition of rocks at multiple scales. The NASA 2020 rover also aims to cache samples that may be returned to Earth in the next decade. Although the return mission is not budgeted yet, it would be the first to provide samples from a major planetary body since the lunar missions of the 1970s. Analyses of returned samples would open a new era in the understanding of Mars geology and climate evolution by helping to calibrate and investigate surface ages and geological time scales, describe mineralogical assemblages at the microscopic scale, and look for traces of extinct (and potential extant) life at a higher level of precision than in situ instruments.


The surface composition of Venus remains largely unknown despite relatively high-resolution (~100s m.pixel-1) global coverage with radar imagery mapping and topography. What is the crustal composition of the equatorial and boreal highlands? Are these regions composed of lighter material similar to terrestrial continental crust? Are they relics of differentiated material concentrating radiogenic elements? What is the role of atmosphere-solid surface interaction? All of these questions remain without answers, mainly because of the thick Venusian atmosphere. Exploration programs have been in stand-by since the end of the ESA’s Venus Express mission (2014). A new European mission to Venus, Envision, could be scheduled for launch in 2029, if selected among three M5 missions, in order to determine the nature and current state of geological activity on Venus. It will use interferometric radar, sounder radar, and thermal mapping on an orbiter to analyze its relationship with the atmosphere and to understand why Venus and Earth have evolved so differently. A new collaborative Russian-American mission (IKI Venera-D) is also currently being assembled for launch around 2025, including both an orbiter and a lander.

The in situ exploration of the Venusian solid surface remains quite difficult technologically, due to the large surface pressure, high temperature, and corrosive atmosphere. The key challenge consists of the implementation of relatively long-lived electronic systems onboard landers. In the 2000s, several projects were proposed both by NASA (Venus Entry probe and DAVINCI with atmospheric probes, SAGE and VISE with landers) and by ESA (Venus entry probe with balloon). None have been selected, but the project Venus In Situ Composition Investigations (VICI) remains supported by NASA, although not selected for the New Frontiers program yet. This mission has the goal to analyze the surface composition both remotely during the descent and at the surface using laser ablation.


The authors are thankful to Carle Pieters and Melissa Martinot for providing material for the figures. The author are granted by the Centre National d’Etudes Spatiales (CNES) for their work on OMEGA/Mars Express.


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