Lunar and Planetary Geology
Summary and Keywords
Lunar and planetary geology can be described using examples such as the geology of Earth (as the reference case) and geologies of the Earth’s satellite the Moon; the planets Mercury, Mars and Venus; the satellite of Saturn Enceladus; the small stony asteroid Eros; and the nucleus of the comet 67P Churyumov-Gerasimenko. Each body considered is illustrated by its global view, with information given as to its position in the solar system, size, surface, environment including gravity acceleration and properties of its atmosphere if it is present, typical landforms and processes forming them, materials composing these landforms, information on internal structure of the body, stages of its geologic evolution in the form of stratigraphic scale, and estimates of the absolute ages of the stratigraphic units. Information about one body may be applied to another body and this, in particular, has led to the discovery of the existence of heavy “meteoritic” bombardment in the early history of the solar system, which should also significantly affect Earth. It has been shown that volcanism and large-scale tectonics may have not only been an internal source of energy in the form of radiogenic decay of potassium, uranium and thorium, but also an external source in the form of gravity tugging caused by attractions of the neighboring bodies. The knowledge gained by lunar and planetary geology is important for planning and managing space missions and for the practical exploration of other bodies of the solar system and establishing manned outposts on them.
What Is Lunar and Planetary Geology
Geology is an earth science which studies the rocks and sediments of which the outer solid shells of our planet are composed, various deformations affecting them, surface landforms, and the processes by which those global features and phenomena formed and changed over time. Similar study approaches can be applied to any solid planet, satellite, or small body of the solar system, so one can consider the geology of Venus, the Moon, a comet nucleus, and so on. The extraterrestrial geologic materials, depending on body size and position in the solar system, range from silicate rocks, ices, and organic materials (in the sense of chemical composition), to products of their surface destruction by a set of processes. Serious differences of planetary geology from the geology of Earth that have been observed may arise from the fact that planets, satellites and small bodies are being studied mostly by remote sensing and by rare landing missions with even rarer sample returns (supported by geophysical and geochemical modeling) that may cause problems of misidentification and misunderstanding. Geologic study of any body and its parts implies investigation of changes and variations in time that demand measurements of time. While on Earth one can study organic fossils and isotopic measurements of absolute age, in the case of planetary geology some surrogates have been invented, including most usable measurements of spatial density (number per unit area) of impact craters superposed on various geologic units. One productive approach in planetary geology is a comparative study of various planets and other objects. In particular, it has been discovered that early in the history of the solar system was a period of intense “meteorite” bombardment which influenced all the bodies including Earth. Below are given very short descriptions of the geologies of Earth (as the reference case), the Moon, Mercury, Mars and Venus, Enceladus (an icy satellite of Saturn), the small asteroid Eros, and the nucleus of the Churumov-Gerasimenko comet; this will be followed by cross-comparisons and resulting conclusions.
Earth is the third planet from the Sun, orbiting it for ~365 days (one Earth’s year) at the mean distance 149,000,000 km, also called an astronomical unit (AU) and often used for measurements of distances within the solar system. The mean radius of Earth is 6,371 km, mean density is 5.5 g/cm3, and surface gravity is 9.8 m/s2. Earth has an essentially iron core having a radius of ~2900 km, above which are essentially silicate mantle and then crust. About 71 percent of Earth surface is covered by the ocean with mean depth ~3.7 km and 29 percent is occupied by continents and islands with a mean height above the ocean level of ~0.8 km. The crust beneath the oceans is on average 5 to 10 km thick and consists mostly of basalt. The continental crust is 30 to 70 km thick and is composed of granites and andesites with an admixture of sedimentary and metamorphic rocks. Earth’s crust and upper ten kilometers of mantle are mechanically rigid and called the lithosphere. Beneath it down to several hundred kilometers the mantle is mechanically weaker and called the asthenosphere.
The Earth’s lithosphere is divided into several rigid plates that move across the surface over periods of many millions of years, diverging or converging with regard to each other with a speed of a few centimeters per year. These movements are caused by the thermal convection in the mantle and this dynamic system is called plate tectonics. The latter leads to the rising of the originally deep-seated hot material in the divergent boundaries and the subducting of the relatively cold near-surface lithosphere material to a hotter depth. This causes generation of magma both at the converging and diverging plate boundaries with its emplacement in the crust and subsequent volcanism as well as the majority of tectonic deformations. In addition, locally there are centers of so-called hot spot volcanism caused by hot plumes which are believed to rise from the core-mantle boundary. When the plume comes to the surface within the moving plate, chains of volcanic islands form. Volcanic and tectonic processes are caused by internal (in relation to Earth) factors and are thus called endogenic geological processes.
Geologic processes of another category, exogenic ones, are caused by external factors, such as changes in surface temperature, water- and wind-driven surface erosion, and material transportation and sedimentation. Earth’s axis of rotation is ~23° tilted to the orbital plane, producing seasonal variations on the planet’s surface. Earth has a relatively thin atmosphere consisting of nitrogen (78 percent) and oxygen (21 percent) with admixtures of other gases—mostly argon, carbon dioxide, and water vapor. The atmospheric pressure near the ocean level is ~100 kPa. The average surface atmosphere temperature on Earth is +14°C, and it varies very widely depending on the locality, season, and time of day. The range is from -90°C (once observed in Antarctica) to +70°C observed in Australia’s Queensland desert (see Universe Today). Changes of atmospheric temperature and pressure are the driving force for the so-called water cycle of water evaporation, condensation, precipitation, infiltration, surface runoff, and subsurface flow. This, in turn, is a driving force of surface erosion, the transportation of fragmented solid and dissolved materials with subsequent sedimentation, and thus is a very powerful factor.
What we observe at Earth’s surface are the results of a continuous interaction of endogenic and exogenic processes. The first is driven by the energy of radiometric decay of potassium, uranium, and thorium, present in the mantle and crust, and leads to the formation of volcanic materials and landforms and tectonic deformations, including formation of large uplifts or deep depressions, while the second ones are driven by the sun’s radiation and destroy the endogenic (and exogenic) materials and landforms by cutting through them with systems of valleys and filling them with sediments. Figure 1 shows a global image of Earth with some traces of those endogenic and exogenic interactions.
On Earth magmatic (divided into volcanic—erupted on the surface, and plutonic—solidified at some depth) materials are mostly represented by various basalts and their plutonic analogs—gabbro and norite consisting of the minerals plagioclase, pyroxene(s), and olivine; plutonic granites (the volcanic analog is rhyolite); and volcanic andesites (the plutonic analog is diorite), consisting mostly of plagioclase, potassium feldspar and, in the case of granites, quartz. Sedimentary materials are represented by sediments—clay, silt, sand, gravel and pebbles—and their lithified analogs,—mudstone, siltstone, sandstone, gravellite, and conglomerate. Locally on Earth there are glaciers whose gravity-driven downslope movement leads to erosion of the underlying materials and formation of so-called moraines (or tills) which usually are unsorted mixtures of materials of different grain size (e.g., clayish sediment with inclusions of pebbles and even boulders).
The movements caused by the plate-tectonic mechanism at the converging boundaries bring volcanic and sedimentary materials to depths with significantly high pressures and temperatures. This process causes changes in mineralogical compositions and their structures or textures or both, so-called metamorphism, so granites and sandstones become gneisses; basalts, gabbro and norites become to be amphibolites; and so on. The materials moving to depth (subducting) are relatively rich in water (both free and chemically bound) and at some depth are dehydrated; part of the freed water supports the partial melting and formation of granitic or andesitic magmas which finds ways to move up and form chains of volcanoes.
In a number of places on Earth there are relatively recent impact craters formed by high-velocity impacts of meteorites, asteroids, and comets. If the craters are significantly destroyed (typically by erosion), their remnants are called astroblemes. Rocks of these landforms are impact breccias and impact melts. These landforms and the rocks they are composed of play negligible roles in the today’s geology of Earth, but sometimes there are large-scale impacts, for example the impact which 65 Ma ago formed the ~200 km in diameter astrobleme Chicxulub on the Yucatán Peninsula in Mexico, which caused significant regional changes and the global mass extinction of many biological species (see, e.g., Schulte et al., 2010). Large high-velocity impacts on Earth can eject some amount of terrestrial material into space and part of it can land on the Moon (Armstrong et al., 2002; Crawford et al., 2008).
One essential part of geologic science is the working out of the geologic history of our planet. Formal presentation of geological history is based on establishing a time sequence of so-called stratigraphic units. The key principle of stratigraphy is a principle of superposition: “In a sequence of undisturbed sedimentary layers the younger layer is on the top and the older layer at the bottom.” (Hendrix & Thompson, 2014, p. 480). It was suggested in 17th century and by subsequent observations that fossil organisms’ remnants were changing over time and this allowed scientists to trace the stratigraphic units laterally and establish global stratigraphic scale (e.g., Cohen et al., 2013; see the updated version at International Chronostratigraphic Chart). So the geologic history of Earth was subdivided into several eras, consisting of periods, which, in turn, are subdivided into epochs. With the progress of radiometric dating, absolute ages of the stratigraphic units’ boundaries have been determined. The beginning of the youngest, the Cenozoic era was found to be at 66 Ma; next, the Mesozoic—250 Ma; Paleozoic—540 Ma; Proterozoic—2500 Ma; Archean—4000 Ma, and the earliest era is called the Hadean, which started ~4.5 Ga ago when Earth was formed. Naturally, details of the subdivisions and understanding of what was happening during each time unit are greater for recent periods.
Earth harbors life and this plays a significant role in the exogenic geology of our planet. It is considered that early terrestrial atmosphere was reducing, consisting of CO or CH4 (Zahnle et al., 2010) and oxygen in it appeared as a product of biological photosynthesis. The appearance of abundant oxygen led to drastic changes in the geochemistry of surface processes and probably even caused a short-term “snowball” climate disaster ~2.3 Ga ago (e.g., Kopp et al., 2005). Later, wide distribution of vegetation would drastically decrease rates and scales of fluvial and eolian surface erosion.
The Moon is a satellite of Earth, orbiting it at the distance of 384,400 km. Its orbital period is 27.3 Earth days and its rotation period is the same. The rotation period of the Moon is not the lunar solar day. Because the Moon together with Earth moves around the Sun, the combination of this movement and the around-axis rotation makes the solar day on the Moon 29.5 Earth days. The long solar day results in a high surface temperature (up to 120°C). But the Moon has no atmosphere so during its long night the temperature may drop down to -150°C.
Because the Moon’s orbital period and its rotation period are the same the Moon always looks at Earth from the same side; thus in descriptions of the Moon the terms near side (visible from Earth) and far side (invisible from Earth) of the Moon are often used. About 1/3 of the near side is covered by relatively dark basaltic plains called maria (single mare—English “sea”), the rest are areas of heavily cratered highlands. On the far side of the Moon mare surfaces are rare and cratered highlands predominate. So in total lunar maria occupy ~16% of lunar surface (Figure 2).
The mean radius of the Moon is 1,738 km, mean density is 3.34 g/cm3, and surface gravity is 1.62 m/s2 (1/6 of Earth’s) (Taylor, 2007). Using the LOLA data Hare et al. (2015) reestimated the mean radius as 1737.4 km. According to these new results the surfaces of most maria are 2 to 3 km below the mean radius. The floor of the lunar South Pole—Aitken (SPA) basin is 5 to 6 km below the datum and locally, at the floors of superposed craters, down to -8 km. Most highland areas are 1 to 3 km above the datum and the highest areas north of SPA basin are up to 9 to 10 km above it.
The Moon is a differentiated body. It has a geochemically distinct crust, mantle, and core. The core with a radius of 300 km is iron-rich and may be partly molten. The lunar mantle and crust are essentially silicate. The mantle is composed of ultramafic rocks, the crust is essentially anothositic with basalts on top within the maria. The GRAIL mission estimated the lunar crust thickness; average value was found to be between 34 and 43 km, with variations from 20 to 30 km in the maria and 35 to 60 km in the highlands. Within large impact basins the crust thickness may be less than 10 km (Wieczorek et al., 2013).
Dominating landforms on the Moon are impact craters whose diameters vary from centimeters to meters, kilometers, and hundreds or thousands of kilometers. Craters smaller than 1 to 2 km in diameter are present practically everywhere. Larger craters are mostly in highlands. Craters smaller than 15 to 20 km are typically bowl-shaped. The larger ones have a more complex morphology, with central peak(s) and concentric rings of rims; craters with rings are called basins. Some craters, especially large ones, are filled with mare basalts. Most large craters formed earlier than 3.9 to 4 Ga ago when the rate of impact bombardment was hundreds time higher than in subsequent time. Due to this heavy bombardment the upper part of the lunar crust is composed not of pristine magmatic rocks but of impact breccias with impact melts. In the highlands they are outcropped on the surface, covered only by relatively thin (meters) of regolith, the layer of ejecta from small craters. In the maria the brecciated crust is covered by flows of basaltic lavas, which, in turn, are covered by regolith.
Plains of lunar maria bear volcanic and tectonic landforms. The first are mostly represented by flows of solidified basaltic lavas, usually seen only when the sun is very low above the horizon, and so-called sinuous rilles, channels cut by the flowing lava. The second are represented by extentional grabens and compressional upthrust faults, usually in the form of wrinkle ridges. In highlands volcanic and tectonic landforms are rare and represented by solidified “lakes” of mare-like lavas as well as grabens and thrust faults usually in the form of scarps (Watters et al., 2010).
The geologic history of the Moon was first deduced from Earth-based telescopic observations and then developed through the results of space missions. The key marker of lunar stratigraphy was suggested to be the large impact basin emplacing Mare Imbrium (Figure 2). Formation of this basin was considered as the beginning of Imbrian period. Most maria were considered to be Imbrian rather than pre-Imbrian. The youngest craters, having bright radial rays of their ejecta (e.g., the crater Copernicus) were classified as formed in the Copernican period. And nonrayed craters superposed on Imbriam maria and subsequently superposed by Copernican rays (e.g., the crater Eratosthenes) were classified as formed in the Eratosthenian period. Later, the pre-Imbrian period was separated into Nectarian and pre-Nectarian periods. Radiometric datings of returned lunar samples allowed a determination of the absolute ages of the stratigraphic units. The pre-Nectarian period began ~4.5 Ga ago, the Nectarian ~4.1 Ga, Imbrian ~3.85 Ga, Eratosthenian ~3.2 Ga, and Copernican ~1 Ga ago.
The major geologic activity during the pre-Nectarian and Nectarian periods was intense impact cratering. This heavy bombardment should have also been happening on neighboring Earth; the Imbrian and part of the Eratosthenian periods were a time dominated by volcanic activity accompanied by some tectonic activity. The youngest part of the geologic history of the Moon was considered mostly to be an arena of non-intense impact cratering. Recent analyses of high-resolution images taken by Lunar Reconnaissance Orbiter narrow angle cameras showed that volcanic and tectonic activity had extended into the geologically current epoch although on a very reduced scale (e.g., Watters et al., 2010; Braden et al., 2014).
An interesting issue of lunar geology is the presence of water on the Moon. Analyses of returned lunar samples in the 1960s and 1970s showed a practical absence of water and until recently the Moon was considered as an anhydrous body except for the polar areas. The axis of rotation of the Moon is tilted only at 1.3° from the ecliptic pole so floors of craters at the poles are permanently shadowed and thus represent cold traps for water vapor and other volatiles mostly brought by the impacts of meteorites or asteroids or both, as well as comets. This should lead to accumulation at the poles of some water ice. Presence of water ice there was recently confirmed by orbital neutron spectrometry (e.g., Mitrofanov et al., 2010) and by the results of the LCROSS experiment (e.g., Colaprete et al., 2012). Orbital IR spectroscopy showed that on the surface of grains of lunar regolith there is a thin (1-2 mm) layer with enriched contents of H2O-OH obviously resulting from the interaction of solar wind protons with oxygen of the regolith grains (Pieter et al., 2009). And finally, study of samples lunar magmatic rocks and minerals using new analytical techniques showed that lunar magmas, at least some of them, contained water in amounts comparable with primary magma of midoceanic basalts of Earth (Greenwood et al., 2011; Saal et al., 2008). While water in the polar regolith was trapped in geologically recent time and H2O-OH on the surface of grains of lunar regolith is being formed at the present time, water in lunar magmatic systems was trapped on the Moon during its formation.
High-velocity impacts on the Moon lead to ejection of some parts of lunar materials in space and part of that comes to Earth as lunar meteorites (e.g., Korotev, 2012; Papike et al., 1998). They are valuable for science because they increase the still limited amount of available lunar samples. On August 2017 the total number of lunar meteorites was 306, probably presenting more than 30 meteorite falls (Meteoritical Bulletin Database, 2017). Their additional value is that they significantly extend areas from which samples have been derived, although in each given case the source location of the lunar meteorite is unknown.
Mercury is the closest to the Sun planet of the Solar system. Its average distance from the Sun is 57,900,000 km (0.38 AU) with closest distance 46,000,000 km (0.3075 AU) and furthest, 69,800,000 km (0.4667 AU). The mean radius of Mercury is 2440 km (0.38 Earth’s), mean density is 5.44 g/cm3 (0.99 Earth’s), and surface gravity is 3.7 m/s2 (0.38 Earth’s). Mercury’s period of rotation is 58.64 Earth days and orbital period 87.97 days, which results in a solar day being 176 Earth days long. The obliquity of the Mercury rotation axis is close to 0°, so the sun in the polar regions is always close to horizon and it is always very cold there. In the shadowed parts of polar regions were found radar-bright deposits considered to be accumulations of water ice with an admixture of organic compounds. These features are probably analogous to the water-ice accumulations in polar areas of the Moon (Chabot et al., 2016). The source of the water and organics is probably also external—namely, the impacts of asteroids and comets. Because of the very long solar day and closeness to the Sun, the surface temperature at the equator can reach +470°C. But Mercury has no atmosphere so during long night the temperature may drop down to about -180°C.
Mercury has an iron core having a radius of ~2000 km (82 percent of the planet’s radius), above which are a silicate mantle (~400 km) and crust (30–50 km) (Margot et al., 2017). A significant part of Mercury’s surface is cratered, thus resembling the lunar highlands. The largest crater is the Caloris basin, with a diameter of ~1500 km. In between craters and locally in distinct regions are seen plains, which, in their relative smoothness, resemble lunar maria, though not as dark (Figure 3). The majority of Mercury’s craters are older than the plains, suggesting a period of heavy bombardment probably contemporaneous with that on the Moon.
Plains on Mercury are mostly volcanic and composed of lavas, which, based on the orbital X-ray spectrometer measurements, are compositionally intermediate between basalts and komatiites (ultramafic volcanic rocks) (Head &Wilson, 2015). Beside the lava fields on Mercury are found surface mantles suggesting explosive pyroclastic activity. The pyroclastic source vents in the form of planimetrically irregular craters are often associated with large impact craters (Goudge et al., 2014). The oldest plains are considered to be ~3.9 Ga old. The age of pyroclastic activity is estimated between 3.5 and 1 Ga ago.
In many areas of Mercury are observed tectonic scarps suggesting compressional environment due to thermal contraction of the planet that led to a decrease of the planet radius by ~7 km (Byrne et al., 2014). Recently through the high-resolution imaging there were found small tectonic scarps whose relationship with small craters suggest a very young age, indicating that the planet thermal shrinking still continues. In some places on Mercury are observed grabens suggesting local extentional environment.
Based on analysis of superpositions of various Mercurian landforms there was suggested a stratigraphic scale of this planet consisting of 5 time periods: the pre-Tolstojan (starting at ~4.5 Ga), Tolstojan at (~3.9 Ga), Calorian (~3.85 Ga), Mansurian (~3.0 Ga), and Kuiperian. The first two are subdivisions of the heavy bombardment era resembling lunar pre-Nectarian and Nectarian periods. The Calorian period started with the formation of the Caloris impact basin and continued with the formation of volcanic plains (analogous to the Imbrian period on the Moon). The Mansurian and Kuiperian periods to some degree are analogs of lunar Eratosthenian and Copernican periods. Absolute dating of the Mercurian stratigraphic units is based on counts of density of superposed craters and thus is model dependent.
Mars is the fourth planet from the Sun orbiting it for ~687 Earth days (1.88 Earth’s year) at the mean distance 228,000,000 km (1.52 AU). The mean radius of Mars is 3,390 km (0.53 Earth’s), mean density is 3.93 g/cm3 (0.71 Earth’s), and surface gravity is 3.7 m/s2 (0.38 Earth’s). Mars rotates around its axis for 24 hours 37 minutes. This Martial solar day is often called sol. The axis of rotation is now inclined in relation to the ecliptic pole by 25°. But modeling (e.g., Laskar et al., 2004) and geologic observations (e.g., Head et al., 2003) show that in the geologic past the axial tilt significantly varied, which could lead to serious climate changes.
Mars has essentially a CO2 atmosphere with surface pressure close to 6 mbar. The mean annual surface temperature is close to -60°C at the equator, reaching at noon about +20°C, and the lowest temperatures are at the poles (down to -120°C). So for most of the planet surface liquid water is not stable and if for some reason it appears it will vaporize or rapidly freeze.
According to results of geophysical modeling Mars has an essentially iron (with addition of sulfur) core with radius ~1800 km, above which are essentially silicate mantle and crust. Mars has a typical north-south global dichotomy which is expressed in three ways: as differences in elevations (the northern hemisphere is on average 5.5 km lower than the southern one), differences in crustal thickness (~30 km in the north and ~60 km in the south), and differences in crater densities (the north is essentially plains and in the south are cratered terrains). This picture is complicated by the so-called Tharsis bulge—a broad (several thousands of kilometers across) region elevated by several kilometers—which contains the giant volcanoes Arsia, Pavonis, Ascreus, Olympus, and Alba, as well as the very large essentially tectonic canyon of Valles Marineris (Figure 4).
Cratered terrains of the southern hemisphere of Mars suggest the presence in early Mars history of the era of heavy bombardment, showing that this phenomenon was not limited to only the Moon and Mercury but was solar-system wide (e.g., Carr & Head, 2010). The largest known impact basins are Argyre (D = 800 km), Isidis (1500 km), and Hellas (D = 2300 km). They have been seriously modified by subsequent geologic activity. Many craters of Mars have specific ejecta features in appearance like mudflows which were probably formed by melting of water ice presenting at some depth as an admixture in Martian crust (e.g., Carr, 1996). Recent analysis of the HIRISE images revealed at medium and high latitudes of Mars some small (4 to 15 miles in diameter) craters of excavated water ice suggesting that it may be present at a very small depth (Dundas et al., 2014)
Cratered terrains probably underlie the essentially volcanic materials of the northern plains and Tharsis bulge. Orbital observations and in situ measurements showed that these volcanic materials are mostly various basalts. Materials of cratered terrains are obviously impact breccias and impact melts. Composition of their source rocks is not well known but growing evidence shows that among them are present felsic rocks: granitoids and anorthosites (e.g., Bandfield, 2006; Sautter et al., 2015; Wray et al., 2013).
Volcanic activity on Mars formed the northern plains and volcanic constructs—giant, medium-sized, and small—all formed by effusions of basaltic lavas (e.g., Carr & Head, 2010). In some cases signatures of explosive basaltic volcanism were found—for example, from the HRSC orbital imaging (Lanz et al., 2010) and in in situ observations by Spirit rover (Squyres et al., 2007). Geologic analysis of HRSC images for the eastern flank of Olympus Mons accompanied by crater counts showed very recent (<25–40 Myr) activity suggesting that this volcano may be not extinct, but dormant (Basilevsky et al., 2006). Tectonic activity is represented by extensional faults and grabens, especially prominent in Valles Marineris and its vicinities and by compressional wrinkle ridges.
Martian cratered terrains are often modified by fluvial valley networks and show evidence of the presence of weathering products such as phyllosilicates, suggesting at least episodic warm conditions with presence of liquid water during the heavy bombardment era and soon after it. Presence of liquid water on the surface of impact-fragmented crust implies existence of subsurface hydrosphere. During that time the surface erosion rates on Mars were two to five orders of magnitude higher than in later times but essentially lower than average terrestrial rates (Carr & Head, 2010).
Fluvial activity continued in later times but in the very specific mode of formation of so-called outflow channels, specific valleys reaching hundreds to a few thousands kilometers in length, tens to hundreds kilometers in width, and 1 to 2 km in depth begun abruptly in areas of chaotic terrain. These were catastrophic floods generated by rapid groundwater evacuation whose nature still is not well understood. Large volumes of water at the termini of outflow channels could form short-lived seas or even oceans, but evidence for that is still equivocal. The outflow channels are typical for the middle part of the geologic history of Mars. During this time most aqueous activity at the surface other than large floods was suppressed. However, presence of discrete sulfate-rich deposits suggests that water surface activity at this time did not decline to zero. The subsurface hydrosphere certainly continued to exist although at the background of the global cooling of Mars it became essentially a cryosphere.
In later periods formation of outflow channels declined to almost zero and the fluvial activity happened mostly in formation of gullies. Those typically consist of an upper alcove that tapers downslope to converge on one or more channels that extend further downslope to terminate in a debris fan. They range in width from meters to tens of meters, are hundreds of meters long, and are common on steep slopes in the 30°–60° latitude belts, particularly in the south, tending to occur on pole-facing slopes (Dickson et al., 2009).
Martian poles host surface accumulations of solid H2O and CO2. There are extended seasonal polar caps formed in winter season of the given hemisphere and there are permanent caps at both poles. The seasonal cap is a relatively thin (~1m) layer of CO2 snow, which extends to the ~50° latitudes. In some sense they are analogs of seasonal H2O-snow covers on Earth. At its formation the seasonal cap captures up to 25 to 30 percent of the atmospheric CO2 and this seasonally captured gas migrates back and forth between the northern and southern hemispheres, causing strong winds. The permanent caps are composed mostly of thinly-layered deposits of, at various degrees, dusty H2O ice and their thickness is a couple of kilometers. The layering is attributed to accumulations of dust and ice modulated by the orbital and rotational motions of Mars.
One serious agent of surface processes on Mars is wind, which erodes and moves the surface material in the saltation regime and may cause dust storms which sometimes are of global scale. The mean wind velocity on Mars is a few meters per second and it could increase up to tens of meters per second during dust storm events. In considering these numbers one should keep in mind that density of Martian near-surface air is ~0.02 kg/m3 , which is only 1.5% to 2% of the Earth’s near-surface air. Despite this small density the Martian winds erode the surface (including carving so-called yardangs—sharp ridges), transport the mobilized material, and deposit it in the form of dust mantles and sandy dune and ripple fields. Frequent eolian landforms on Mars are also wind streaks: erosional or depositional features downwind of craters and other wind obstacles.
Combining observations of relative time relations of geologic units with the results of counts of craters superposed on them, the following stratigraphy of Mars was suggested: Its beginning is Pre-Noachain period which started at the planet accretion time ~4.5 Ga ago and ended before the formation of Hellas impact basin between 4.1 and 3.8 Ga. Most of the geologic record of this interval has been erased by subsequent geologic activity. The crustal dichotomy is thought to have formed during this time. Noachain period activity started with the formation of the Hellas basin and lasted until ~3.7 Ga. Major geologic activity of that time was intense impact cratering and formation of the Tharsis bulge. The Hesperian period (~3.7 to ~3 Ga) is known as the time of the formation of extensive lava plains and catastrophic releases of water-carved outflow channels. The formation of Olympus Mons probably began during this time period. The Amazonian period (~3Ga to present) was marked by low-intensity impact cratering, glacial activity, and minor releases of liquid water. Bibring et al. (2006), based on analysis of the Omega IR spectroscopy, suggested parallel “mineralogic” eras: Phyllocian (after phyllosilicates) before ~4 Ga ago; Theiikian (after “sulphurous” in Greek for the sulfate minerals) 4 to 3.5 Ga ago; and Siderikan (named for “iron” in Greek, for the iron oxides that formed that time) from 3.5 Ga till now.
As previously mentioned, in the geologic past of Mars liquid water was present, at least episodically, on its surface; at present, Mars has subsurface liquid water and sources of thermal and chemical energy under dormant volcanoes. So one may guess that early in its history, and even later, Mars had conditions that allowed the appearance of life and successful establishment of microorganisms. If Martian life developed, it was (and may still be) chemotrophic and anaerobic. This hypothetical similarity with early terrestrial chemotrophic life may suggest a key for understanding the potential distribution of Martian chemotrophs and their fossilized traces (e.g., Westall et al., 2015).
High-velocity impacts on Mars lead to ejection of some part of Martian materials in space and part of that part comes to Earth as Martian meteorites (e.g., McSween & Treiman, 1998; Smith et al., 1984). They are extremely valuable for planetary geology because they could be (and are) studied with all the potential abilities of modern analytical techniques, while the orbital remote sensing and lander contact studies are always limited. There are more than 130 known Martian meteorites (Meteoritical Bulletin Database, 2017).
The supply of Marian meteorites on Earth together with suggestions that early and even present underground Mars could be habitable for primitive organisms revives the old idea of panspermia: the “seeds” of life exist all over the universe and can be propagated through space from one location to another (see Panspermia-Theory.com). So from time to time appear discussions that life could have first appeared on Mars, been brought on Earth, and then flourished here (e.g., Adcock et al., 2013).
Venus is the second planet from the Sun, orbiting it for ~225 Earth days at the mean distance 108,000,000 km (0.72 AU). The mean radius of Venus is 6,052 km (0.95 Earth’s), mean density is 5.24 g/cm3 (0.95 Earth’s), and surface gravity is 8.87 m/s2 (0.904 Earth’s). The internal structure of Venus is not well known. Model calculations show that Venus has an essentially iron core having a radius of ~3000 km, above which are essentially silicate mantle and then crust. Most parts of Venus are occupied by plains formed by effusions of basaltic lavas. About 7.5 percent of Venus’s surface is occupied by so-called tessera terrain—heavily deformed highland regions. The crust thickness within those plains is probably 20 to 50 km. Within the large massifs of tessera the crust may be thicker.
Venus’s axis of rotation is inclined to the ecliptic pole only 3.4°. The period of rotation is 243 Earth days. The combination of periods of orbiting and rotation (which is retrograde) has resulted in a Venusian solar day duration of 117 days.
Venus has a massive essentially CO2 atmosphere which exerts surface pressure on the mean radius level of 93 bar; the density is 65 kg/m3. This dense atmosphere causes a strong greenhouse effect so the mean surface temperature is ~470°C with practically no seasonal and daily variations, and only decreases or increases with altitude change (8.1 K/km within the lower 70 km). At altitudes of 45 to 70 km there is a cloud layer composed of suspended droplets of concentrated sulfuric acid. On the cloud level the atmosphere is in state of superrotation circling the planet in four Earth days; the speed of this planet-wide wind is ~100 m/s. Visual picture of this superrotation can be seen in the left portion of Figure 5. Closer to the surface the windspeed decreases down to a few m/s and smaller.
About 80 percent of Venus’s surface is composed of volcanic plains formed by flows of basaltic lavas. Their composition was determined by several in situ measurements made by Soviet Venera landers (Surkov, 1997) and supported by morphological analysis of the Magellan radar images (e.g., Head et al., 1992). Among the plains are locally observed steep-sided volcanic domes (typically 20 to 30 km in diameter) whose morphology suggests that they may be composed of nonbasaltic, more felsic lavas. Among the plains are observed large and small elevated massifs of previously mentioned tessera terrain, which occupy in total ~8% of the planet surface. Its composition could be nonbasaltic. Plains and tesserae are cut through by prominent geologically young rift zones resembling continental rift zones of Earth. They occupy ~5% of the surface. And ~8% of the surface is occupied by old rift zones. Like on Earth Venusian rift zones are associated with volcanic landforms. Signatures of the present volcanism in the form of transient hot spots were found within the young rift zone of Ganis Chasma (2 in Figure 5 right; see Shalygin et al., 2015).
In radar images of the Venusian surface are seen ~1000 impact craters with diameters from several kilometers to 280 km (the crater Mead). The absence on Venus of craters smaller than a few kilometers is due to atmospheric shielding. Some craters with diameters close to this lower limit are present in clusters formed by impacts of fragments of meteoroids destroyed in the passage through the very dense atmosphere. Spatial density of the observed craters suggests that the mean absolute age of Venusian plains is between 0.5 and 1 Ga. The youngest 10 percent of Venusian craters have radar-dark parabolic haloes (no. 3 in Figure 5 right) which have formed due to expulsion of some part of crater ejecta above the atmosphere, its finer fraction carried downwind by the superrotating atmosphere and deposited as relatively fine material, whose smooth surface is seen as dark in synthetic aperture radar images (e.g., Campbell et al., 1992). Turbulence during the deposition of the parabola materials and superposition of deposits of different parabolas may result in the cm-scale layering observed in the Venera landing sites (Basilevsky et al., 2004).
In several areas of Venus are seen supposedly eolian features: wind streaks, yardangs (ridges carved by wind), and fields of dunes (Greeley et al., 1997). Their orientations do not conform to the global east-west atmosphere rotation. So their formation could be related to some occasional wind-forming events, maybe to winds induced by shock waves in the atmosphere caused by crater-forming impacts (Ivanov et al., 1992).
There is no widely accepted “official” stratigraphc scale of Venus. As a result of photogeologic mapping of Venus Ivanov and Head (2011) suggest distinguishing four periods of the geologic history of this planet: (a) pre-Fortunian (no geologic record is known), (b) Fortunian (time when the tessera terrain was formed), (c) Guineverian (most now observed plains were formed), and (d) Atlian (young plains and young rift zones were formed). The boundary between Guineverian and Atlian periods corresponds to an absolute age between 0.5 and 1 Ga, the boundary between Fortunian and Guineverian periods is about 1.5 older.
Current surface conditions on Venus are too extreme for life as we know it, but it could have had abundant water and favorable conditions for life when the early sun was fainter. Schulze-Makuch and Irwin (2004) suggest the possibility that after conditions on Venus’s surface became untenable for life it could have continued to exist in restricted environmental niches such as the high-pressure subsurface habitats with water at a supercritical liquid state and in the zone of dense cloud cover where thermoacidophilic life might survive.
Enceladus is one of relatively large satellites of the planet Saturn, which is 9.5 times farther from the Sun than Earth. Enceladus orbits the planet for 1.37 Earth days at the mean distance 238,000 km. Its diameter is ~500 km, mean density 1.68 g/cm3, surface gravity is 0.113 m/s2 (0.0113 Earth’s g). It rotates around its axis synchronously, for the same time as it orbits the planet. Because of large distance from the Sun, the minimum surface temperature on Enceladus is -240°C, mean, -198°C and maximum, -128°C. It has practically no atmosphere, only traces of it, composed of water-group molecules and ions being continuously replenished (Dougherty et al., 2006).
Observations using Cassini’s Visual and Infrared Mapping Spectrometer showed that the surface of Enceladus is composed mostly of nearly pure water ice except at its south pole, where there are light organics, CO2, and amorphous and crystalline water ice (Brown et al., 2006). The geometric albedo of Enceladus is 1.41 at 0.55 μm (Verbiscer et al., 2005), the highest of any known solar system body. The reason for this anomalous value is coherent enhancement of backscattering (Shkuratov et al., 2004).
The interior structure of Enceladus is not well known. Based on previously mentioned data, within its bulk density it should have a silicate fraction probably differentiated into a core, although the direct evidence for differentiation is absent. Above the core there should be a global or regional water layer and an ice shell thick enough to support the ~1 km amplitude topography observed on this body (Spencer et al., 2009).
A significant part of Enceladus’ surface is covered with impact craters at various areal densities and levels of degradation (Figure 6 left). Two regions of smooth plains are also observed on this body. They generally have low relief and much fewer craters than in the cratered terrains, indicating a relatively young surface. There are also several types of tectonic features, including troughs, scarps, and belts of grooves and ridges. Results from Cassini suggest that tectonics is the dominant mode of deformation on Enceladus in geologically recent time (Spencer et al., 2009).
The Cassini mission detected on Enceladus chronologically variable active plumes erupting from warm fractures near its south pole (Figure 6 right). This activity is presumably powered by tidal heating maintained by Enceladus’ 2:1 mean-motion resonance with Dione, another satellite of Saturn (Spencer et al., 2009). Southern polar activity is concentrated along the four “tiger stripe” fractures (Figure 6 left), which radiate heat and are the source of multiple plumes ejecting ~200 kg/s of H2O vapor along with significant N2 (or C2H4), CO2, CH4, NH3, and higher-mass hydrocarbons (Spencer et al., 2009). With this plume activity and presence of liquid water in its subsurface Enceladus is considered as the most promising potential habitat for life in the outer Solar system.
Eros is a typical asteroid representing one of the classes of small bodies of the solar system. It has an irregular shape—11 x 11 x 34 km (Figure 7)—and orbits the Sun for 1.76 Earth years at a mean distance from the Sun of 1.46 AU. The closest distance (perihelion) is 1.13 AU and the farthest (aphelion) is 1.78 AU. So it belongs to the so-called near-Earth asteroids, which approach the Sun for less than 1.3 AU. Eros rotates around its axis for 5.27 hours. The daytime temperature on Eros can reach ~100°C at perihelion. Nighttime measurements fall near −150°C. The mean density of Eros is ~2.67 g/cm3 and surface gravity varies as 2 to 5 x 10-4 Earth’s g (Miller et al., 2002; Murdoch et al., 2017).
Eros belongs to S-type asteroids considered to be a source of ordinary chondrite meteorites. But the bulk density of Eros is noticeably lower than that of ordinary chondrites (3.3–3.5 g/cm3, see Wilkinson & Robinson, 2000), which suggests that it is significantly porous or fractured or both (Cheng, 2002). Its geometric albedo is 0.25 ± 0.06. Measurements by the NEAR Shoemaker X-ray/gamma spectrometer did not establish a definite meteorite link (Trombka et al., 2001). It is not clear if Eros’s composition is actually unrelated to any known meteorite type, or if it’s actually chondritic at a depth below surface layers that may have been altered and fractionated by unknown weathering processes (Cheng, 2002).
The surface of Eros is covered by impact craters from meters tens to several hundred meters in diameter and larger. The largest craters of Eros are Himeros (D = 11 km), Shoemaker (D = 7 km), and Psyche (D = 5 km). Material under them should be significantly fractured down to a depth of several kilometers. The number of craters smaller than 100 m in diameter is unexpectedly low (in comparison with the Moon), and the cause of this phenomenon is unknown. Because of Eros’s very small surface gravity, a significant part of ejecta from impact craters of Eros escapes to space and this is the way asteroids supply majority of meteorites which come to Earth. However the combination of the irregular shape and rapid rotation of the asteroid makes distribution of crater ejecta rather complicated (Durda et al., 2012). Eros has a regolith cover, whose thickness in some areas may reach 100 m, and a consolidated but fractured substrate (Cheng, 2002).
Along with Eros’s impact craters are observed structural features including sinuous and linear grooves, ridges, and scarps. Eros is too small to have a source of internal energy to produce these features and the obvious causes of their formation are stresses originating in large impact-cratering events. The large variation in directions, patterns, and relative ages of these features indicate that they were formed during many different unrelated events. The largest ridge of Eros, Rahe Dorsum, cross-cuts the largest crater Himeros and follows across more than a third of Eros’s circumference.
Despite the low gravity, steep slopes on Eros show signatures of downslope material movement. In some gravitational lows are observed extremely level ponded deposits. The fine particles could have been mobilized by electrostatic levitation and seismic shaking from impacts (Cheng, 2002).
Comet 67P Churyumov-Gerasimenko
This comet is now the most studied and this provides a possibility to consider what kinds of geologic processes occur on comet nuclei. Comet 67P was visited by the ESA mission Rosetta, which studied it with 11 instruments of the mission orbiter and 10 instruments of the Philae lander. The 67P comet belongs to the Jupiter family of comets, which are short-period comets whose current orbits are primarily determined by the gravitational influence of Jupiter. The comet orbit aphelion is 5.68 AU, perihelion 1.24 AU, orbital period 6.44 years, period of rotation 12.4 h, and obliquity 52° (JPL Small-Body Database). The specifics of its orbit dictate that summer in the nucleus’s northern “hemisphere” is relatively long and cold while summer at the south is much shorter but hotter. The size of the nucleus is about 2x4x3 km3, its bulk density is 533 kg/m3, and surface gravity in the nucleus’s different localities varies from 1.2 to 2.2 x 10-4 m/s2 (Groussin et al., 2015; Jorda et al., 2016; Pajola et al., 2017). The nucleus has a bilobate shape with a smaller (head) and larger (body) lobe and a narrow neck between them (Figure 8).
Two major types of nucleic material are distinguished: (a) the consolidated nucleus material and (b) the loose material, a kind of cometary regolith, covering the nucleus’s consolidated material. This mostly resulted from the transportation of “fines” from the summer-hot nucleus south to the colder north and thus is more typical for the north (Keller et al., 2017). On the surface of the consolidated material rather long (up to hundreds meters) straight lineaments are seen. They probably correspond to fractures and in some cases to strata. Their presence suggests that the consolidated material is rather compact and lacks voids larger than tens of meters across. Surfaces of consolidated nucleus material typically show knobby appearance at scales from tens of meters and meters to centimeters and millimeters. This suggests that this material is grainy, consisting of more and less resistant (to surface weathering) “particles” on the scale of the visible knobs.
The geometric analysis of the steep slopes is based on the nucleus shape model allowing for an estimate of tensile, shear, and compressive strength of the consolidated material (Basilevsky et al., 2017b; Groussin et al., 2015). It has been shown that the 67P consolidated nucleus material is very fragile, and taking into account the scale effect one can conclude that it is as fragile as terrestrial fresh fallen snow and perhaps even more fragile. In addition, estimates of the compressive strength of the surface material were considered at the sites of the first and the last contacts of the Philae lander with the surface and it was found that it is mechanically even more weak that the consolidated material (Biele et al., 2015). Observations also showed evidence of various downslope and lateral movements of rather large material masses (landslide? avalanche?) as well as boulders and “fines,” which are driven primarily by gravity and then by the acquired inertia, but in some cases a material transport by dust or gas jets outbursts could play a role. The latter could also be responsible for formation of the eolian-type ripples.
In many places of the nucleus are seen pinnacles, local promontories of varied shapes including spires with pointed tops (Basilevsky et al., 2017a). They are typically asymmetric with somewhat different slope angles at different slopes whose maximum values range from 40° to 90°, sometimes with small overhangs. Their heights vary from 10–20 to 100–200 m and the foot diameters from 30–300 m. They have a surface texture similar to that of the consolidated nucleus material and are probably composed of it. The observed characteristics of the pinnacles suggest that they are the erosion remnants formed due to a slowdown of erosion in places where nucleus material is more resistant to sublimation than the material around it. In this case the maximum heights of the pinnacles (100–200 m) define a lower boundary for the amount of surface material lost and their diameters (typically tens of meters) are a measure of the size of the erosion-resistant parts.
Lessons Learned and Potential Perspectives
Solid bodies of the Solar system in general can be described and analyzed using the set of terms, approaches, and techniques of what is worthwhile calling “lunar and planetary geology.” This science despite its rather short history has already showed significant progress in aiding our understanding of the environments in which we live, from areas on Earth to the whole solar system, and in the near future it will certainly be applied to the studies of planets around other stars. Now there are no doubts that early in the solar system’s history there was a period of heavy “meteoritic” bombardment. We may not understand yet details of what caused it, but it is already considered by terrestrial geologists and scientists working in the field of the origin of life as a leading process on the early Earth. We have learned that small bodies, like Eros or comet nuclei, have a rather primitive geology, driven by exogenic processes caused by meteorite impacts and solar radiation. From the example of Enceladus and some other bodies (Jupiter satellites Io and Europa) we have learned that internal processes like volcanism and large-scale tectonics may be driven not by internal sources of radioactive decay energy, but “external” gravity tugging caused by attractions of the neighboring bodies. The approaches and techniques of lunar and planetary geology also have a crucial significance for planning and managing space missions and, in some not very distant future, for the practical exploration of other bodies of the solar system and establishing manned outposts.
Badro, J., & Walter, M. (Eds.). (2015). The early Earth: Accretion and differentation. New York: Wiley.Find this resource:
Burbine, T. H. (2017). Asteroids: Astronomical and geological bodies. New York: Cambridge University Press.Find this resource:
Carr, M. H. (2006). The surface of Mars. New York: Cambridge University Press.Find this resource:
Dougherty, M. K., Esposito, L. W., & Krimigis, S. M. (Eds.). (2009). Saturn from Cassini-Huygens. Dordrecht, the Netherlands: Springer.Find this resource:
Harland, D. M. (2008). Exploring the Moon: The Apollo Expeditions (2nd ed.). Chichester, UK: Springer Praxis Books/Space Exploration.Find this resource:
Heiken, G. H., Vaniman, D. T., & French, B. M. (1991). Lunar Source Book: A Users’ Guide to the Moon. New York: Cambridge University Press.Find this resource:
Mason, J. W. (Ed.). (2008). Exoplanets: Detection, formation, properties, habitability. Chichester, UK: Springer.Find this resource:
McFadden, L.-A., Weissman, P. R., & Johnson, T. V. (Eds.). (2007). Encyclopedia of the Solar System (2nd ed.). Amsterdam: Elsevier Academic Press.Find this resource:
Seargent, D. A. J. (2017). Weird Comets and Asteroids: The Strange Little Worlds of the Sun’s Family. Cham, Switzerland: Springer.Find this resource:
Strom, R. G., & Sprague, A. L. (2003). Exploring Mercury: The Iron Planet. Chichester, UK: Praxis Books.Find this resource:
Taylor, F. W. (2014). The Scientific Exploration of Venus. New York: Cambridge University Press.Find this resource:
Adcock, C. T., Hausrath, E. M., & Forster, P. M. (2013). Readily available phosphate from minerals in early aqueous environments on Mars. Nature Geoscience, 6, 824–827.Find this resource:
Armstrong, J. C., Wells, L. E., & Gonzales, G. (2002). Rummaging through Earth’s attic for remains of ancient life. Icarus, 160(1), 183–196.Find this resource:
Bandfield, J. L. (2006). Extended surface exposures of granitoid compositions in Syrtis Major, Mars. Geophysical Research Letters, 33, L06203, 1–4.Find this resource:
Basilevsky, A. T., Head, J. H., & Abdrakhimov, A. M. (2004). Impact crater air fall deposits on the surface of Venus: Areal distribution, estimated thickness, recognition in surface panoramas, and implications for provenance of sampled surface materials. Journal of Geophysical Research, 109(E12003), 1–18.Find this resource:
Basilevsky, A. T., Krasilnikov, S. S., Mall, U., Hviid, S. F. S., Skorov, Yu. V., & Keller, H. U. (2017a). Pinnacles on the 67P comet nucleus: Evidence for large scale erosion and hierarchical agglomeration of the nucleus. Planetary and Space Science, 140, 80–85.Find this resource:
Basilevsky, A. T., Mall, U., Keller, H. U., Skorov, Yu. V., Hviid S. F., Mottola, S., . . . Dabrowski, B. (2017b). Geologic analysis of the Rosetta NavCam, Osiris, and ROLIS images of the comet 67P/Churyumov-Gerasimenko nucleus. Planetary and Space Science, 137, 1–19.Find this resource:
Basilevsky, A. T., Werner, S. C., Neukum, G., Head, J. W., van Gasselt, S., Gwinner, K., & Ivanov, B. A. (2006). Geologically recent tectonic, volcanic and fluvial activity on the eastern flank of the Olympus Mons volcano, Mars. Geophysical Research Letters, 33(13), L13201, 1–4.Find this resource:
Bibring, J.-P., Langevin, Y., Mustard, J. F., Poulet, F., Arvidson, R., Gendrin, A., . . . Forget, F. (2006). Global mineralogical and aqueous Mars history derived from OMEGA/Mars Express data. Science, 312(5772), 400–404.Find this resource:
Biele, J., Ulamec, S., Maibaum, M., Roll, R., Jurado, E., Arnold, W., . . . Witte, L. (2015). The landing(s) of Philae and inferences about comet surface mechanical properties. Science, 349 (6247).Find this resource:
Braden, S. E., Stopar, J. D., Robinson, M. S., Lawrence, S. J., van der Bogert, C. H., & Hiesinger, H. (2014). Evidence for basaltic volcanism on the Moon within the past 100 million years. Nature Geoscience, 7, 787–791.Find this resource:
Brown, R. H., Clark, R. N., Buratti, B. J., Cruikshank, D. P., Barnes, J. W., Mastrapa, R. M. E., . . . Sotin, C. (2006). Composition and physical properties of Enceladus’ surface. Science, 311, 1425–1428.Find this resource:
Byrne, P. K., Klimczak, C., Şengör, C., Solomon, S. C., Watters, T. R., & Hauck, S. A. (2014). Mercury’s global contraction much greater than earlier estimates. Nature Geoscience, 7, 301–307.Find this resource:
Campbell, D. B., Stacy, N. J. S., Newman, W. I., Arvidson, R. E., Jones, E. M., Musser, G. S., . . . Schaller, C. (1992). Magellan observations of extended impact crater related features on the surface of Venus. Journal of Geophysical Research, 97(E10), 16249–16277.Find this resource:
Carr, M. H. (1996). Water on Mars. New York: Oxford University Press.Find this resource:
Carr, M. H., & Head, J. W. (2010). Geologic history of Mars. Earth and Planetary Science Letters, 294(3–4), 185–203.Find this resource:
Chabot, N. L., Ernst, C. M., Paige, D. A., Nair, H., Denevi, B. W., Blewett, D. T., . . . Murchie, S. L. (2016). Imaging Mercury’s polar deposits during MESSENGER’s low-altitude campaign. Geophysical Research Letters, 43(18), 9461–9468.Find this resource:
Cheng, A. F. (2002). Near Earth asteroid rendezvous: Mission summary. In W.F. Bottke et al. (Eds.), Asteroids III (pp. 351–366). Tucson, AZ: University of Arizona Press.Find this resource:
Cohen, K. M., Finney, S. C., Gibbard, P. L., & Fan, J.-X. (2013). The ICS International Chronostratigraphic Chart. Episodes, 36, 199–204.Find this resource:
Colaprete, A., Elphic, R. C., Heldmann, J., & Ennico, K. (2012). An overview of the Lunar Crater Observation and Sensing Satellite (LCROSS). Space Science Review, 167(1–4), 3–22.Find this resource:
Crawford, I. A., Baldwin, E. C., Taylor, E. A., Bailey, J. A., & Tsembelis, K. (2008). On the survivability and detectability of terrestrial meteorites on the Moon. Astrobiology, 8(2), 242–252.Find this resource:
Dickson, J. L., Fassett, C. I., & Head, J. W. (2009). Amazonian-aged fluvial valley systems in a climate microenvironment on Mars: Melting of ice deposits on the interior of Lyot crater. Geophysical Research Letters, 36(8), L08201, 1–5.Find this resource:
Dougherty, M. K., Khurana, K. K., Neubauer, F. M., Russell, C. T., Saur, J., Leisner, J. S., & Burton, M. E. (2006). Identification of a dynamic atmosphere at Enceladus with the Cassini magnetometer. Science, 311(5766), 1406–1409.Find this resource:
Dundas, C. M., Byrne, S., McEwen, A. S., Mellon, M. T., Kennedy, M. R., Daubar, I. J., & Saper, L. (2014). HiRISE observations of new impact craters exposing Martian ground ice. Journal of Geophysical Research (Planets), 119(1), 109–127.Find this resource:
Durda, D. D., Chapman, C. R., Merline, W. J., & Enke, B. L. (2012). Detecting crater ejecta-blanket boundaries and constraining source crater regions for boulder tracks and elongated secondary craters on Eros. Meteoritics and Planetary Science, 47(6), 1087–1097.Find this resource:
Goudge, T. A., Head, J. W., Kerber, L., Strom, R. G., Xiao, Z., Zuber, M. T., . . . Solomon, S. C. (2014). Global inventory and characterization of pyroclastic deposits on Mercury: New insights into pyroclastic activity from MESSENGER orbital data. Journal of Geophysical Research Planets, 119(3), 635–658.Find this resource:
Greeley, R., Bender, K. C., Saunders, R. S., Schubert, G., & Simpson, R. A. (1997). Aeolian processes and features of Venus. In S. W. Bougher et al. (Eds.), Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment (pp. 547–589). Tucson: University of Arizona Press.Find this resource:
Greenwood, J. P., Itoh, S., Sakamoto, N., Warren, P., Taylor, L., & Yurimoto, H. (2011). Hydrogen isotope ratios in lunar rocks indicate delivery of cometary water to the Moon. Nature Geoscience, 4, 79–82.Find this resource:
Groussin, O., Jorda, L., Auger, A.-T., Gaskell, R., Capanna, C., Scholten, F., . . .Vincent, J.-B. (2015). Gravitational slopes, geomorphology and material strengths of the nucleus of comet 67P/Churyumov-Gerasimenko from OSIRIS observations. Astronomy and Astrophysics, 583(A32), 1–12.Find this resource:
Hare, T. M., Hayward, R. K., Blue, J. S., Archinal, B. A., Robinson, M. S., Speyerer, E. J., . . . Mazarico, E. (2015). Image mosaic and topographic map of the moon. U. S. Geological Survey Scientific Investigations Map 3316, 2 sheets.
Head, J. W., Crumpler, L. S., Aubele, J. C., Guest, J. E., & Saunders, R. S. (1992). Venus volcanism: Classification of volcanic features and structures, associations, and global distribution from Magellan data. Journal of Geophysical Research, 97(E8), 13153–13197.Find this resource:
Head, J. W., Mustard, J. F., Kreslavsky, M. A., Milliken, R. E., & Marchant, D. R. (2003). Recent ice ages on Mars. Nature, 426, 797–802.Find this resource:
Head, J. W., & Wilson, L. (2015). Volcanism on Mercury. In H. Sigurdsson, B. Houghton, H. Rymer, J. Stix, & S. McNutt (Eds.), The Encyclopedia of Volcanoes (pp. 701–716). San Diego, CA: Elsevier Academic Press.Find this resource:
Hendrix, M., & Thompson, G. R. (2014). Earth2. Boston, MA: Cengage Learning.Find this resource:
Ivanov, B. A., Nemchinov, I. V., Svetsov, V. A., Provalov, A. A., Khazins, V. M., & Phillips, R. J. (1992). Impact cratering on Venus: Physical and mechanical models. Journal of Geophysical Research, 97(El0), 16167–16181.Find this resource:
Ivanov, M. A., & Head, J. W. (2011). Global geological map of Venus. Planetary and Space Science, 59(13), 1559–1600.Find this resource:
Jorda, L., Gaskell, R., Capanna, C, Hviid, S., Lamy, P., Durech, J., . . .Wenzel, K. -P. (2016). The global shape, density and rotation of Comet 67P/Churyumov-Gerasimenko from preperihelion Rosetta/OSIRIS observations. Icarus, 277, 257–278.Find this resource:
JPL Small-Body Database Browser: 67P/Churyumov-Gerasimenko. NASA/Jet Propulsion Laboratory. Retrieved from http://ssd.jpl.nasa.gov/sbdb.cgi?sstr=67P.
Keller, H. U., Mottola, S., Hviid, S. F., Agarwal, J., Kührt, E., Skorov, Y., . . . Thomas, N. (2017). Seasonal mass transfer on the nucleus of comet 67P/Chuyumov-Gerasimenko. Monthly Notices of the Royal Astronomical Society, 469(suppl. 2), S357–S371.Find this resource:
Kopp, R. E., Kirschvink, J. L., Hilburn, I. A., & Nash, C. Z. (2005). The Paleoproterozoic snowball Earth: A climate disaster triggered by the evolution of oxygenic photosynthesis. Proceedings of the National Academy of Sciences, 102(32), 11131–11136.Find this resource:
Korotev, R. (2012). Lunar meteorites from Oman. Meteoritics and Planetary Science, 47(8), 1365–1402.Find this resource:
Lanz, J. K., Wagner, R., Wolf, U., Kröchert, J., & Neukum, G. (2010). Rift zone volcanism and associated cinder cone field in Utopia Planitia, Mars. Journal of Geophysical Research, 115(E12019), 1–21.Find this resource:
Laskar, J., Correia, A. C. M., Gastineau, M., Joutel, F., Levrard, B., & Robutel, P. (2004). Long term evolution and chaotic diffusion of the insolation quantities of Mars. Icarus, 170(2), 343–364.Find this resource:
Margot, J.-L., Hauck, S. A., Mazarico, E., Padovan, S., & Peale, S. J. (2017). Mercury’s internal structure. In S. C. Solomon, B. J. Anderson, & L. R. Nittler (Eds.), The View after MESSENGER (pp. 1–36).Find this resource:
McSween, H. Y., & Treiman, A. H. (1998). Martian meteorites. In J. J. Papike (Ed.), Planetary Materials. Reviews in Mineralogy (Vol. 36, pp. 6-1–6-53). Washington, DC: Mineralogical Society of America.Find this resource:
Meteoritical Bulletin Database—Lunar Meteorite search results. (2017). The Meteoritical Society. Retrieved from https://www.lpi.usra.edu/meteor/metbull.php.
Miller, J. K., Konopliv, A. S., Antreasian, P. G., Bordi, J. J., Chesley, S., Helfrich, C. E., . . . Yeomans, D. K. (2002). Determination of shape, gravity, and rotational state of asteroid 433 Eros. Icarus, 155(1), 3–17.Find this resource:
Mitrofanov, I. G., Sanin, A. B., Boynton W. V., Chin, G., Garvin, J. B., Golovin, D., . . . Zuber, M. T. (2010). Hydrogen mapping of the lunar South Pole using the LRO Neutron Detector Experiment LEND. Science, 330(6603), 483–486.Find this resource:
Murdoch, N., Sanches, P., Schwartz, S. R., & Miyamoto, H. (2017). Asteroid surface geophysics. In P. Michel, F. DeMeo, & W. Bottke (Eds.), Asteroids IV (pp. 1–29). Tucson: University of Arizona Press.Find this resource:
Pajola, M., Lucchetti, A., Fulle, M., Mottola, S., Hamm, M., Da Deppo, V., . . . Thomas, N. (2017). The pebbles/boulders size distributions on Sais: Rosetta’s final landing site on comet 67P/Churyumov-Gerasimenko. Monthly Notices of the Royal Astronomical Society.Find this resource:
Papike, J. J., Ryder, G., & Shearer, C. K. (1998). Lunar samples. In J. J. Papike (Ed.), Planetary Materials (Vol. 36, pp. 5-1–5-234). Washington, DC: Mineralogical Society of America.Find this resource:
Pieters, C. M., Goswami, J. N., Clark, R. N., Annadurai, M., Boardman, J., Buratti, B., . . . Varanasi, P. (2009). Character and spatial distribution of OH/H2O on the surface of the Moon seen by M3 on Chandrayaan-1. Science, 326(5952), 568–572.Find this resource:
Robinson, M. S., Thomas, P. C., Veverka, J., Murchie, S. L., & Wilcox, B. B. (2002). The geology of 433 Eros. Meteoritics and Planetary Science, 37(12), 1651–1684.Find this resource:
Saal, A. E., Hauri, E. H., Cascio, M. L., Van Orman, J. A., Rutherford, M. C., & Cooper, R. F. (2008). Volatile content of lunar volcanic glasses and the presence of water in the Moon’s interior. Nature, 454(7201), 192–195.Find this resource:
Sautter, V., Toplis, M. J., Wiens, R. C., Cousin, A., Fabre, C., Gasnault, O., . . . Wray, J. J. (2015). In situ evidence for continental crust on early Mars. Nature Geoscience, 8, 605–609.Find this resource:
Schulte, P., Alegret, L., Arenillas, I., Barton, P. J., Bown, P. R., Bralower, T. J., . . . Whalen, M. T. (2010). The Chicxulub asteroid impact and mass extinction at the Cretaceous-Paleogene boundary. Science, 327(5970), 1214–1218.Find this resource:
Schulze-Makuch, D., & Irwin, L. N. (2004). Reassessing the possibility of life on Venus: Proposal for an astrobiology mission. Astrobiology, 2(2), 197–202.Find this resource:
Shalygin, E. V., Markiewicz, W. J., Basilevsky, A. T., Titov, D. V., Ignatiev, N. I., & Head, J. W. (2015). Active volcanism on Venus in the Ganiki Chasma rift zone. Geophysical Research Letters, 42(12), 4762–4769.Find this resource:
Shkuratov, Yu., Videen, G., Kreslavsky, M. A., Belskaya, I. N., Ovcharenko, A., Kaydash, V. G., . . . Zubko, E. (2004). Scattering properties of planetary regoliths near opposition. In G. Videen, Y. Yatskiv, & M. Mishchenko (Eds.), Photopolarimetry in Remote Sensing (pp. 191–208). NATO Science Series. London: Kluwer Academic.Find this resource:
Smith, M. R., Laul, J. C., Ma, M. S., Huston, T., Verkouteren, R. M., Lipschutz, M. E., & Schmitt, R. A. (1984). Petrogenesis of the SNC (shergottites, nakhlites, chassignites) meteorites: Implications for their origin from a large dynamic planet, possibly Mars. Journal of Geophysical Research, 89(S02), B612–B630.Find this resource:
Squyres, S. W., Aharonson, O., Clark, B. C., Cohen, B. A., Crumpler, L., de Souza, P. A., . . . Yen, A. (2007). Pyroclastic activity at Home Plate in Gusev crater, Mars. Science, 316(5825), 738–742.Find this resource:
Spencer, J. R., Barr, A. C., Esposito, L. W., Helfenstein, P., Ingersoll, A. P., Jaumann, R., . . . Waite, J. H. (2009). Enceladus: An active cryovolcanic satellite. In M.K. Dougherty, L.W. Esposito, & S. M. Krimigis (Eds.), Saturn from Cassini-Huygens (pp. 683–724). New York: Springer Science.Find this resource:
Surkov, Yu. A. (1997). Exploration of terrestrial planets from spacecraft: Instrumentation, investigation, interpretation. New York: Wiley.Find this resource:
Taylor, S. R. (2007). The Moon. In L.-A. McFadden, P. R. Weissmann, & T. V. Johnson (Eds.), Encyclopedia of the Solar System (2nd ed., pp. 227–250). San Diego, CA: Elsevier.Find this resource:
Trombka, J. I., Nittler, L. R., Starr, R. D., Evans, L. G., Mccoy, T. J., Boynton, W. V., . . . Murphy, M. E. (2001). The NEAR-Shoemaker x-ray/gamma-ray spectrometer experiment: Overview and lessons learned. Meteoritics and Planetary Science, 36(12), 1605–1616.Find this resource:
Verbiscer, A. J., French, R. G., McGhee, C. A. (2005). The opposition surge of Enceladus: HST observations 338–1022 nm. Icarus, 173(1), 66–83.Find this resource:
Watters, T. R., Robinson, M. S., Beyer, R. A., Banks, M. E., Bell, J. F., III, Pritchard, M. E., . . . Williams, N. R. (2010). Evidence of recent thrust faulting on the Moon revealed by the Lunar Reconnaissance Orbiter Camera. Science, 329(5994), 936–940.Find this resource:
Westall, F., Foucher, F., Bost, N., Loizeau, D., Vago, J. L., Kminek, G., . . . Cockell, C. S. (2015). Biosignatures on Mars: What, where, and how? Implications for the search for Martian life. Astrobiology, 15(11), 998–1029.Find this resource:
Wieczorek, M. A., Neumann, G. A., Nimmo, F., Kiefer, W. S., Taylor, G. J., Melosh, H. J., . . . Zuber, M. T. (2013). The crust of the Moon as seen by GRAIL. Science, 339, 671–675.Find this resource:
Wilkinson, S., & Robinson, M. S. (2000). Bulk density of ordinary chondrite meteorites and implications for asteroidal internal structure. Meteoritics and PIanetaty Science, 35(6), 1203–1213.Find this resource:
Wray, J. J., Hansen, S. T., Dufek, J., Swayze, G. A., Murchie, S. L., Seelos, F. P., . . . Ghiorso, M. S. (2013). Prolonged magmatic activity on Mars inferred from the detection of felsic rocks. Nature Geoscience, 6, 1013–1017.Find this resource:
Zahnle, K., Schaefer, L., & Fegley, B. (2010). Earth’s earliest atmospheres. Cold Spring Harbor Perspectives in Biology, 2(10), 1–17.Find this resource: