The interior structures of the Earth and Moon are determined from seismic data. The existence and sizes of cores in other planets are inferred from observations of planetary sizes, masses, and shapes, which constrain their uncompressed mean densities and moment of inertia factors. Mantle and crust thicknesses can also be estimated from gravity data obtained by orbiting spacecraft. Successful models of planetary interiors constructed from compositional data must be consistent with observed densities and moments of inertia. High-pressure laboratory experiments can constrain the mineralogy of mantles and cores and the partitioning of elements between silicate and metal in the terrestrial planets. The interiors of the giant planets are not well understood, because of uncertainties in their compositions and internal temperatures and pressures. The states of hydrogen and helium in the interiors of Jupiter and Saturn, and the crystalline forms of ices in Uranus, Neptune, and icy satellites, are inferred from experimentally determined or calculated phase diagrams. The giant planets may have small rocky cores, with successive layers of either metallic hydrogen (Jupiter and Saturn) or ices (Uranus and Neptune), and molecular hydrogen. Planetary mantles and cores evolve over geologic time, through cooling and extraction (or reintroduction, in the case of Earth) of crustal components.

We present a brief overview of the planets, moons, dwarf planets, asteroids, and comets – intended as a primer for those with limited or no familiarity with planetary science. The terrestrial planets (Earth, Mars, Venus, and Mercury) are rocky bodies having mean densities that indicate metal cores; the giant planets are composed mostly of hydrogen and helium and can be divided into gas giants (Jupiter and Saturn) and ice giants (Uranus and Neptune), based on their physical states. Small bodies, composed of rock and ices, are either differentiated or not, depending on their thermal histories. Each section of this chapter is generally organized in the historical order in which the objects have been explored by spacecraft. We will return to these bodies repeatedly in the book, focusing on understanding their geologic characteristics and materials, and the processes that produced them.

Physical weathering of rocks on bodies other than the Earth occurs mostly through impact fragmentation, producing regoliths. The lunar regolith is finer-grained and contains more agglutinates than asteroidal regoliths, indicating its greater maturity. Mars exhibits both physical and chemical weathering, and its sedimentary deposits superficially resemble those on Earth. However, its basalt-derived sediments differ from those formed from felsic protoliths on Earth, and evaporation of its aqueous fluids is dominated by sulfates, distinct from terrestrial evaporites that are mostly carbonates and halides. On the surfaces of airless bodies, recondensation of vapor produced by micrometeorite impacts accounts for spectral changes, known as space weathering. In the interiors of carbonaceous chondrite asteroids, isochemical reactions of rocks with cold aqueous fluids produced by melting of ice have altered their mineralogy. Thermal metamorphism of dry chondritic asteroids has modified all but near-surface rocks. Hydrothermal metamorphism on Mars, likely associated with large impacts, has produced low-grade mineral assemblages in metabasalts and serpentinites. Conditions at Venus’ surface are severe enough to cause thermal metamorphism, and reactions with rocks may control the composition of the atmosphere. Because all bodies have gravity, some sloping topography, and some unconsolidated materials, mass wasting is among the most common processes modifying planetary surfaces.

We explain nucleosynthesis in evolving stars and use this foundation to understand the chemical composition of our own star and of the Solar System. Element abundances are determined from the Sun’s spectrum, and from laboratory measurements of the solar wind and chondritic meteorites. The metal-rich Solar System composition reflects the recycling of elements formed in earlier generations of stars. Condensation models of a cooling nebular gas having this composition produced the minerals found in refractory inclusions in chondrites. The deuterium enrichment in organic matter in chondrites suggests that hydrocarbons formed at low temperatures in molecular clouds and were subsequently processed into complex molecules in the solar nebula and in parent bodies. Ices condensed far from the Sun and were incorporated into the giant planets and comets. Element fractionations in the nebula were largely controlled by element volatility or by the physical sorting of solid grains. Separation of isotopes by mass was common in the nebula, although oxygen shows mass-independent fractionation.

Planetary volatiles occur in gas, liquid, and solid forms. In this chapter, we will see that the terrestrial planets have secondary atmospheres formed by outgassing of their interiors. The chemical compositions of the atmospheres of Venus and Mars are dominated by CO2, but the Earth’s atmosphere is distinct because CO2 is sequestered in the lithosphere and life has added O2 to the mix. Giant planet atmospheres are mostly hydrogen with some helium. Titan has an atmosphere of N2 and reducing gases, along with seas of hydrocarbons. Mars has briny groundwater and had lakes and possibly oceans in the distant past. Some moons of the giant planets have subsurface oceans. Noble gases and stable isotopes hold keys to the origin and evolution of volatiles. Differences in temperature and pressure cause atmospheric circulation, controlled by planetary rotation and energy transport. Frozen volatiles are common as polar deposits and sometimes permafrost, but they are especially abundant in the outer Solar System, where they may comprise the crusts of giant planet satellites. Volatile behaviors can be described in terms of geochemical cycles. Greenhouse warming has important implications for planetary climates.

Planetary exploration typically advances in step with technology. Improvements in spatial and spectral resolution yield discoveries that progress from global to regional scales, and exploration on a planet’s surface provides ground truth for remote sensing data and a level of observation and measurement that geologists crave. Samples that can be analyzed in the laboratory provide geochemical and geochronologic information that complements spacecraft data and enhances its interpretation. We illustrate how data at all these scales have been integrated to characterize the complex geology of Mars and to constrain its geologic history.

Only a quarter of a century ago, the only planets known to exist were those orbiting our Sun. Now we know that other stars host extrasolar planets, or “exoplanets.” The earliest discoveries were made using Earth-based telescopes that detected stellar wobbles caused by nearby orbiting planets. To be detectable, these exoplanets must be large and orbit very close to their stars. Using a different method, NASA’s Kepler orbiting space telescope (2009–2013) observed partial eclipses as planets transited in front of stars (Batalha, 2014). Kepler’s high-precision measurements – only possible in space where stars don’t twinkle – have enabled the discovery of ~4000 candidate planets; of these candidates, perhaps 90 percent are thought to be real planets (Lissauer et al., 2014). The list includes nearly 700 multiple planet systems (Figure E.1), and the large numbers imply that these must be very common. Planetary orbits in multi-planet systems are usually coplanar, consistent with their formation within accretion disks (like our own Solar System’s ecliptic plane).

In addition to spectroscopy, planetary geoscience uses some other tools familiar to most geologists, and some tools that are either unique or involve new twists in how they are employed. We explain how stratigraphic principles are adapted for planets (using strata produced by impacts), how the density of craters can be quantified to derive relative ages of geologic units, and how radioisotope measurements on samples, where available, give absolute ages. We explain how images from orbiting and landed spacecraft are used, along with chronologic and remote-sensing data, to make planetary geologic maps at different scales. We consider various geophysical techniques that are used on spacecraft to obtain information about planetary potential fields, interior structure, and surface topography. We summarize the kinds of extraterrestrial materials that are available for laboratory investigations, and briefly describe the analytical techniques used to characterize their mineralogy, petrology, and geochemistry. We also examine some techniques that are adapted as remote sensing tools for analyses of rocks and soils on planetary surfaces.

We describe the near-surface wind profile, its relation to environmental conditions, and how it can be quantified. The freestream wind speed can be converted to a friction wind speed, which relates to the flow at the atmosphere–surface interface and thus to the entrainment of sediment. The minimum wind speed for entrainment of aeolian sediment depends on gravity and grain size, so that threshold wind speed differs for varying planetary conditions. The difference in transport mechanism for grains leads to different depositional morphologies, which provide clues to the wind speed, wind direction, and sediment availability. Erosional landforms likewise provide information on near-surface atmospheric processes and surface sediments, as well as bedrock lithologies. The study of aeolian landforms thus informs our understanding of the atmosphere, surface geology, and sedimentology on other planets.

The surfaces of terrestrial planets and icy satellites have enjoyed deformation marked by faults and folds. We use these geologic structures not only to characterize the morphology of the surfaces, but also to describe the motions, stresses, and deformation processes that created the structures. Ultimately, we can sum these structural data and interpretations to infer the tectonic deformation for large portions of, or even entire, planets and satellites. As we will see, understanding deformation at this large tectonic scale enables us to investigate what is driving overall planet or satellite development. We will also learn that while the expected will happen, conundrums exist too. For example, the rocks of terrestrial planets deform quite differently from the icy shells of the satellites of the gas giants, yet the magnitude of these differences and their causes can surprise us. On the other hand, Venus and Earth are quite similar in many planetary characteristics but have strikingly different tectonic histories, which challenges us to understand why.

Heavily cratered surfaces emphasize the important geologic role of impacts on almost all planetary bodies. Large impacts (referred to as “hypervelocity” to indicate velocities of tens of kilometers per second) produce micro-, meso-, and macroscale deformations that can influence the structures of planetary crusts. Crater morphologies are described as simple, complex, and multi-ring, and correlate with crater size and inversely with gravity of the target body. Crater formation is envisioned in three stages: contact/compression, excavation, and modification, each characterized by different processes and geologic features. Shock metamorphism has affected all Solar System bodies, producing breccias containing planar deformation features, high-pressure polymorphs, and melts. Craters provide the basis for planetary stratigraphy and chronology. Massive impacts on the Earth have potential consequences for damaging the planetary ecosystem and biosphere.

With our preliminary survey of planetary bodies in the Solar System complete, we next turn our attention to developing an understanding of the tools that are used by planetary geologists to study these bodies. In this chapter, we will learn about some of the remote sensing techniques that are most commonly used in planetary geologic exploration. We will focus greatest attention on methods that employ light as the carrier of information about the remote target, but we will also consider some methods that employ the detection of other types of carriers. We will describe techniques that can be used from aircraft flying above the Earth, from spacecraft orbiting planets, and from landed and roving platforms on other worlds. We will examine both active and passive remote sensing techniques, provide guidance on how to select the right type of remote sensing method for the desired scientific outcome, and discuss the role of a successful ground campaign in the analysis of remote-sensing data when that option is available.

We explain how heat is produced by radioactive decay, segregation and exothermic crystallization of metallic cores, impacts, and tidal forces. Planetesimals in the early Solar System were most affected by the decay of short-lived radionuclides. Larger, rocky planets were heated primarily by large impacts and core segregation. Because rocks are poor conductors, heat retention in rocky bodies is a function of planet size. Large-scale melting to produce magma oceans was likely a common process facilitating differentiation to form cores, mantles, and crusts. Metallic liquids are probably necessary for core segregation. Primary crusts, formed during planetary differentiation, are rarely preserved. Mantles are residues from the extraction of silicate crustal melts and core materials. Differentiation of the giant planets was driven by density variations in high-pressure forms of gases, ices, and rock more than by heating and melting. The importance of the various planetary heat sources changes over time; in modern planets the effective heat sources are decay of long-lived radioisotopes and, for the Earth, exothermal crystallization of the liquid outer core.