Geological Characteristics of the Moon
Summary and Keywords
The geological characteristics of the Moon provide the fundamental data that permit the study of the geological processes that have formed and modified the crust, that record the state and evolution of the lunar interior, and that identify the external processes that have been important in lunar evolution. Careful documentation of the stratigraphic relationships among these features can then be used to reconstruct the sequence of events and the geological history of the Moon. These results can then be placed in the context of the geological evolution of the terrestrial planets, including Earth. The Moon’s global topography and internal structures include landforms and features that comprise the geological characteristics of its surface. The Moon is dominated by the ancient cratered highlands and the relatively younger flat and smooth volcanic maria. Unlike the current geological characteristics of Earth, the major geological features of the Moon (impact craters and basins, lava flows and related features, and tectonic scarps and ridges) all formed predominantly in the first half of the solar system’s history. In contrast to the plate-tectonic dominated Earth, the Moon is composed of a single global lithospheric plate (a one-plate planet) that has preserved the record of planetary geological features from the earliest phases of planetary evolution. Exciting fundamental outstanding questions form the basis for the future international robotic and human exploration of the Moon.
Prior to the advent of the space age in 1957, geologists were occupied with the study of Earth and with the reconstruction of Earth’s history; the planets were largely the domain of astronomers. Soon after the space age began, the field of geology was revolutionized by the recognition that Earth’s history was not a series of independent uncorrelated events, but rather the planet as a whole operated as an integrated system in a global process that came to be known as plate tectonics, encompassing seafloor spreading at divergent plate boundaries, subduction at convergent plate boundaries, and continental drift. Seafloor spreading and subduction, combined with atmospheric weathering processes, were found to operate at such rapid rates that the vast majority of Earth’s most ancient geological record had been destroyed. Most of Earth’s surface (>60%; the ocean floor) is less than 200 Ma, only a few percent of the total history of the Earth (Figure 1). Thus, the formative years of Earth’s history are no longer readily available for study on the Earth itself.
Fortunately, there was a second parallel revolution in the understanding of Earth brought about by the space age. Exploration of the planets by spacecraft flybys, orbiters, and landers by the United States and the Soviet Union caused a major change in perspective: The Moon and planets, originally seen as astronomical objects, began to be viewed as objects of geological interest. Furthermore, during the NASA Apollo Lunar Exploration Program, six scientific expeditions were sent to the Moon, and human explorers undertook field trips to different sites chosen on the basis of their geological characteristics and scientific importance.
Following the Luna Program and the Apollo Lunar Exploration Program, there have been numerous successful lunar exploration missions from different countries such as Clementine (1994, NASA), Lunar Prospector (1998, NASA), SMART-1 (2003, ESA), SELENE (i.e., Kaguya, 2007, Japan), Chang’e-1 (2007, China), Chandrayaan-1 (2008, India), LRO/LCROSS (2009, NASA), Chang’e-2 (2010, China), GRAIL (2011, NASA), Chang’e-3 (2013, China), LADEE (2013, NASA), and Chang’e-4 (2018, China). From a geological point of view, Clementine acquired remote sensing composition data covering the entire Moon (e.g., Lucey, Blewett, & Hawke, 1998; Lucey, Blewett, & Jolliff, 2000), SELENE obtained high-resolution imagery and high-quality elemental data (e.g., Kato, Sasaki, & Takizawa, 2010; Otake, Ohtake, & Hirata, 2012), Chang’e-1 and Chang’e-2 obtained high-resolution imagery, elemental and brightness temperature data (e.g., Fa & Jin, 2007; Wu, 2012; Wu et al., 2012; Zuo, Li, & Zhang, 2014), Chang’e-3 disclosed the subsurface structure of Mare Imbrium for the first time (e.g., Xiao et al., 2015; Zhang et al., 2015), Chang’e-4 became the first mission landed on the farside and explored the Von Karman crater (CLEP, 2019; Xiao et al., 2019; see Figure 2), Chandrayaan-1 conducted mineral mapping (e.g., Pieters et al., 2009a) and discovered signs of water in polar regions (e.g., Pieters et al., 2009b), LRO/LCROSS acquired high-resolution topography (Lunar Orbiter Laser Altimeter, LOLA; e.g., Smith et al., 2010, 2017) and imagery (Narrow Angle Camera, NAC; e.g., Robinson et al., 2010), conducted volatile detection experiments of the polar regions (e.g., Sanin et al., 2017), and GRAIL provided high-resolution gravity data and made many new findings (e.g., Zuber et al., 2016).
The fruits of these explorations led to unlocking the secrets of the Moon, the history of the early solar system, and the geological processes that were operating at that time (Figure 3).
The origin of the Moon, the formation of primary crust, the interior structure of the Moon, the nature and importance of impact cratering, the role of vertical tectonics on a planetary body without plate tectonics, and the nature of volcanism in early planetary history were all themes that emerged during the geological exploration of the Moon. Furthermore, the proximity of the Moon to the Earth permitted multiple international missions to investigate a wide array of the Moon’s geology, mineralogy, petrology, and geophysics. The Moon became a fundamental cornerstone for the understanding of early planetary evolution, including that of Earth. This article outlines the nature of the geological characteristics of the Moon, the geological processes that operate there, and what these processes tell us about the geological history and thermal evolution of the Moon.
The Formation, Internal Structure, and Topography of the Moon
The Moon is approximately one-fourth the diameter of Earth and its density is considerably less than that of Earth, implying a much less dense interior and a relatively small core (Table 1; Figure 4).
Table 1. Main Properties of the Moon and Comparison to Earth
Escape Velocity/km s−1
Surface Temperature/ °C
Lunar interior structure is revealed by geophysical experiments and model-based geophysical inversions. Geophysical experiments include in situ seismic experiments, observations of lunar gravity, electromagnetic sounding, and lunar laser ranging (LLR). In the Apollo era, the information about the lunar interior was derived mainly from in situ seismic experiments at several Apollo landing sites and the Lunar Laser Retroflector experiment. With plausible assumptions for the lunar interior, the internal structures of the Moon were reconstructed by interpreting the distribution of seismic wave velocities from Apollo seismic data (e.g., Nakamura, 1983). Similar to the Earth’s interior, the interior of the Moon is stratified into three parts: a crustal layer, a mantle layer, and a core, which all differ in density and chemical composition. The lunar crust, which is composed mainly of plagioclase, occupies the uppermost 40–60 km, with a density of about 2,700–2,800 kg m−3. The lunar mantle, which is composed primarily of olivine and pyroxene and serves as the source of mare basalt, has a density of about 3,300–3,400 kg m−3. Furthermore, the lunar mantle can be subdivided into two sub-layers by the discontinuity of seismic velocities at a depth of about 500 km. This discontinuity boundary could mark an abrupt variation of Mg# or the spinel–garnet phase transition in the lunar mantle (Nakamura, 1983; Hood & Jones, 1987; Snyder, Taylor, & Neal, 1992). The lunar core is much denser than the overlying silicate portion of the Moon due to the metallic iron therein. Assuming the core is Fe-FeS eutectic, the estimated lunar core radius varies between 200 and 530 km (Hood & Jones, 1987; Kuskov, Kronrod, & Hood, 2002). Weber, Lin, Garnero, Williams, and Lognonne (2011) also reinterpreted the Apollo seismic data using a seismic array technique to estimate the structures of the deep lunar interior. The results indicate that the lunar core comprises a solid inner core and a liquid outer core, with a total radius of about 330 km. A partially molten layer, with a thickness of about 160 km, is interpreted to overlie the core–mantle boundary, with a viscosity estimated at about 2 × 1016 Pa s (Harada et al., 2014).
The prevailing theory for the origin of the Moon is the accretion of impact ejecta orbiting the Earth in the aftermath of one (perhaps Mars-sized) or more planetary bodies impacting into the proto-Earth (e.g., Canup & Asphaug, 2001; Canup, 2012; Cuk & Stewart, 2012; Rufu, Aharonson, & Perets, 2017). To a first order, the low density of the Moon relative to the Earth is attributed to the ejecta from the impacts being largely composed of crust and upper mantle rather than denser core material of the Earth. Following the accretion of the Moon, continued impact of the solar system projectiles at a very high rate contributed to the heating and melting of the outer part of the Moon, forming a thick exterior veneer of molten rock known as a magma ocean. Sequential cooling and crystallization of this magma ocean led to the flotation of the less dense mineral plagioclase and the formation of a global low-density plagioclase-dominated crust, averaging several tens of kilometers in thickness (Figure 4). The crystallization of the magma ocean led to stratification of the crust (Figure 5) and upper mantle and to basal denser residual layers (e.g., ilmenite-bearing cumulate, IBC) at the base of crust that are likely to have foundered during or in the immediate aftermath of the initial solidification process.
A residual layer relatively richer in potassium, rare earth elements, and phosphorus (designated by the acronym KREEP) formed in the vicinity of the base of the crust. This global lunar crustal chemical/mineralogical layer (dominantly anorthositic upper crustal layer and underlying more noritic layer) (e.g., Warren & Wasson, 1979) thus formed in a fundamentally different manner than the continental crust on Earth. Results from the Gravity Recovery and Interior Laboratory (GRAIL) mission have shown that the average thickness of the anorthositic lunar crust is between 34 and 43 km (Wieczorek et al., 2013), and that there is a major difference in the average thickness of the lunar crust between the nearside and the farside of the Moon (Figure 6). This nearside–farside crustal thickness difference is one of the fundamental characteristics of the Moon and is interpreted to explain its center of figure–center of mass offset. The origin of the nearside–farside crustal thickness difference has been attributed to (a) broad-scale convection patterns in the magma ocean, (b) an excess of impact basins on the lunar nearside, and (c) one or more extremely large impact basins (Pieters et al., 1993; Neumann, Zuber, Smith, & Lemoine, 1996; Loper, 2002; Bierson et al., 2016).
As the magma ocean cooled and solidified, it also formed a mechanical layer, the lunar lithosphere, representing the outer thermal boundary layer between the hotter interior and the loss of heat to deep space. From very early in lunar history, this layer was globally continuous, and in contrast to the lithosphere of Earth, was unsegmented. Thus, the Moon is known as a “one-plate planet” (Solomon, 1978), in contrast to the multiple segmented laterally moving plates of Earth. Due to the small size of the Moon and the high surface area–volume ratio, the lunar lithosphere cooled conductively and relatively rapidly with time, precluding segmentation and subduction. This relatively rigid lithosphere served as a platform on which the geological history of the first half of the solar system history was preserved and could be decoded (Figure 1).
During and following the formation of the lunar highlands crust and global lithosphere, abundant relatively smaller bodies of debris were circulating in the solar system (e.g., asteroids, comets), eventually colliding with larger planetary bodies at various angles and velocities and forming a host of impact craters of a wide range of sizes. Overlapping with formation of the large impact basins, such as Imbrium and Orientale, internal heating of the mantle of the Moon generated partial melts of the mantle that propagated to the surface in “magma-filled” cracks generated by melt overpressure and erupted effusively or explosively to form the lunar maria. Mare basalts cover about 17% of the global lunar surface (Head, 1976; Head & Wilson, 1992) and are concentrated on the lunar nearside and in ancient impact basins and craters.
Together, these elements make up the first-order characteristics of global lunar topography (Figure 7).
Reflected in the unimodal lunar surface elevation hypsogram (Figure 8) is the breadth of the unimodal peak, owing to crustal thickness differences and the skewness to lower elevations, caused by the presence of large, unrelaxed impact basins and the lunar maria.
The lunar surface elevation hypsogram, dominated by a globally continuous crust and thick rigid lithosphere, lies in contrast to the bimodal hypsogram of Earth owing primarily to oceanic and continental crustal thickness differences, and secondarily to contrasting crustal composition and density and the evolving oceanic thermal boundary layer. Second-order lunar topographic characteristics are dominated by the rims, steep walls, and peaks of impact craters and basins and volcanic resurfacing that forms relatively smooth plains typically at lower elevations, and the large farside South Pole–Aitken impact basin (Figure 9). Crustal characteristics can be classified into several broad geochemical terrains on the basis of FeO and Th content (Figure 10): (a) a Procellarum KREEP Terrane (PKT), (b) a South Pole–Aitken Basin Terrane (SPAT), and (c) a Feldspathic Highlands Terrane (FHT) (Jolliff, Gillis, Haskin, Korotev, & Wieczorek, 2000).
Variation in the vertical composition of the lunar crust has been proposed and accepted for many years. Ryder and Wood (1977) suggested that the lunar crust becomes more mafic with depth because the impact melts associated with the large Imbrium and Serenitatis basins are more mafic than the surface composition of the Moon. However, Lemelin, Lucey, Song, and Taylor (2015) analyzed the composition of crater central peaks by using recent remote sensing data and combining the best practices of previous studies. They found that there is no increase in mafic mineral abundances with proximity to the crust and mantle boundary or with depth from the current lunar surface and, therefore, that the crust apparently does not become more mafic with depth.
The Main Geological Features of the Lunar Surface
The most prominent visible topographic features on the surface of the Moon are craters and basins originating from impact-cratering processes. Impact cratering is a fundamental geological process for all solid planets, including the Moon (e.g., Melosh, 1996; Osinski & Pierazzo, 2012), and the presence of a stable lunar crust and lithosphere since early lunar history has provided a template on which craters can be described and studied in detail. Lunar craters form a very important baseline in the understanding of craters on all planets.
Several groups have compiled morphologic and morphometric classifications and catalogs of lunar craters.
1. In an early catalog (Wood & Andersson, 1978), craters were classified according to their morphology (Figure 11). In the LPL (Lunar and Planetary Laboratory, United States) Catalog, fresh craters were classified into five different morphologies. An updated version of this catalog was published as the NASA Catalog of Lunar Nomenclature (Andersson & Whitaker, 1982). It contains the positions and diameters of 8,497 craters.2
2. Russian planetary scientists under the leadership of V. V. Shevchenko created The Morphologic Catalog of Lunar Craters in 1987. It includes a compendium of 14,923 lunar craters, all larger than 10 km in diameter. Craters in this database are referred to by their “SAI index” (SAI meaning Sternberg Astronomical Institute where the work was done). The Catalog has also served as a Russian (Cyrillic) Gazetteer of Lunar craters.
3. In 2010, a global catalog of 5,185 lunar craters ≥20 km in diameter was derived using the precise measurements of the Lunar Reconnaissance Orbiter Lunar Orbiter Laser Altimeter Instrument (Head et al., 2010; Kadish et al., 2011).3
4. A Lunar Impact Crater Database was created partly during the Lunar and Planetary Institute’s Lunar Exploration Intern Program in 2008 (Losiak et al., 2009) and revised by Teemu Öhman in 2011. It contains calculated morphologic, geological, and also stratigraphic (age) information for 8,862 lunar craters.4
5. Robbins (2019) published a new global database of lunar impact craters >1–2 km (the database itself is available through the NASA Planetary Data System). In this database, he identified and measured >2 million lunar craters; 1.3 million are ≥1 km in diameter. There are more craters ≲30 km than all other published catalogs due to the use of multiple data sets and a lower diameter bound, and including subdued and secondary craters.
Other lunar crater databases include that of Salamunićcar, Lončarić, Grumpe, and Wöhler (2014), which has 19,396 craters, and Wang, Cheng, and Zhou (2015), who calculated 106,016 craters with diameters larger than 500 m using Chang’e-1 imagery.
There are more than one million lunar craters with diameters larger than 1 km and countless smaller ones that have been seen at micrometer scale on Apollo samples. Fresh lunar impact craters vary systematically in their shapes and depth–diameter relationships (Figure 12) (Pike, 1980; see review in Hiesinger, 2006). Simple, circular bowl-shaped craters are typically less than about 15 km in diameter. Scalloped-walled craters represent a transition from bowl-shaped to flat-floored, as the circular rim crest becomes more irregular due to material slumping toward the floor. These craters occur in the diameter range of 15–25 km and have lower depth to diameter ratios due to the scalloped wall material sliding toward the crater floor. In the diameter range from ~25 km to several hundred kilometers, craters are characterized by polygonal shapes, scalloped wall terraces, flat floors, and central peaks and are known as complex craters. In fresh craters, the flat floors are covered with abundant rough crenulated material (primary crater-floor roughness) interpreted as impact melt, such as the well-preserved floor of Tycho crater (Krüger, van der Bogert, & Hiesinger, 2016). Central peaks in larger complex craters are often characterized by clusters, which transition to rings of central peaks, generally concentric with the crater rim crest. The advent of peak rings marks the transition from impact craters to impact basins (Baker, Head, Collins, & Potter, 2016; Baker et al., 2017).
Fresh lunar impact craters are characterized by bright rays and clusters of craters extending multiple crater radii away from the crater rim crest; these rays and secondary crater clusters are formed from the impact of fragments ejected from the primary crater (Xiao et al., 2014). Because of their low-angle impact, they typically produce asymmetric crater shapes and ejecta and a distinct V-shaped herringbone pattern that can be used to link the secondary crater and cluster back to the parent crater (the V-shape points back to the crater). The projectiles that form the secondary craters impact at the velocity with which they leave the crater and thus excavate many multiples of their mass, a relationship that changes with increasing radial distance from the impact point (Oberbeck & Morrison, 1976). As the bright rays degrade and darken due to micrometeoroid bombardment and space weathering, the secondary craters and clusters remain until ultimately they are degraded sufficiently so that they cannot be distinguished from the background of primary impact craters. Secondary craters admixed with primary craters can cause errors in the use of small craters in impact crater size-frequency distribution (CSFD) age dating of surfaces (Stöffler et al., 2006), but this situation is typically avoided by excluding irregularly shaped craters and craters lying in the typical secondary crater size range.
Between the fresh crater rim crest and the rays and secondary crater clusters lies a more braided textured deposit known as the continuous ejecta deposit, and is formed on top of the uplifted crater rim (Xiao et al., 2014). There is a progression of the surface texture and morphology as a function of increasing size, with smaller craters typified by boulders and dune-like surface features, and larger craters by secondary crater chains and braided texture that can be traced in toward the crater rim crest.
In the interior of simple lunar impact craters, the shape is circular, the walls are not terraced and appear as talus slopes, and the floor is convex downward or flat and hummocky (Pike, 1980). In complex lunar craters, the polygonal rim crest walls are generally directly correlated with the development of multiple wall terraces formed from the slumping of the rim of the crater down into the collapsing cavity along what appear to be listric faults. Central peaks appear in the middle of the flat floor of the crater, with impact melt deposits coating the crater floor and forming the first-order flatness of the crater floor (e.g., Dhingra, Head, & Pieters, 2017). The rough impact melt deposits appear to have settled onto the crater floor in the modification stage of the cratering event as the uplift of the central peaks removed support from the walls and formed the wall terraces. Asymmetric collapse of the crater in the modification stage has been shown to preferentially propel impact melt up over the opposite crater wall and onto the crater rim, forming flows and puddles of impact melt (Hawke & Head, 1977).
Subsequent to their formation, fresh impact craters undergo modification and degradation due to a variety of processes. Subsequent crater superposition can modify or even obliterate the underlying crater. Formation of a nearby crater can cause “proximity weathering,” degradation from the ejecta, and secondary craters of the subsequent crater (Head, 1975), and this, combined with the general background projectile flux, can degrade craters with time, first the rays, then the secondary craters and clusters, and finally the crater rim and interior. Impact basins, with their huge size and far-reaching ejecta, can modify preexisting crater populations over more than a lunar hemisphere. Furthermore, the seismic effects of such large events can cause instantaneous mass wasting and degradation as well (Kreslavsky & Head, 2012; Schultz & Gault, 1975).
Another source of fresh lunar crater modification and degradation is effusive volcanic activity, most importantly the flooding and filling of craters and their ejecta by lava flows (Whitten & Head, 2013). Typically, this process preferentially modifies topographic lows and leaves the crater rim and rim crest preserved unless the lava exceeds a kilometer or two in thickness. Lunar pyroclastic activity can also modify fresh craters, owing to the widespread emplacement of pyroclastics as the gas-rich magmas expand into a vacuum to produce regional dark mantle deposits. Fresh impact craters can even be modified by intrusions, specifically the sills that are interpreted to underlie the floors of the floor-fractured craters (FFC) (Schultz, 1976). In this case, the crater floors have been uplifted and tectonically disrupted by the intrusive activity. If the crater was large enough and the lithosphere thin enough, viscous relaxation could also in principle modify a crater, relaxing the floor and lowering the crater rim crest; however, confident identification of viscously relaxed lunar impact craters has not been made (Dombard & Gillis, 2001; Hall, Solomon, & Head, 1981), but large, older impact basins were certainly topographically degraded by this wavelength-dependent process. Tectonic activity can also modify fresh lunar craters, but the low level of tectonic activity on the Moon means that degradation is limited to a few normal faults, wrinkle ridges, and arches, and to floor fractures in FFC (e.g., Jozwiak, Head, Zuber, Smith, & Neumann, 2012).
Baker et al. (2017) analyzed the transition with increasing size from simple, to complex, to peak-ring basins, and finally to multiring basins (Figures 11 and 13). They showed that what is important to understanding the relationship between complex craters with central peaks and multiring basins is the analysis of proto-basins (these exhibit a rim crest and interior ring plus a central peak) and peak-ring basins (these exhibit a rim crest and an interior ring). Baker, Head, Fassett, and Kadish (2010) utilized topographic data from the Lunar Orbiter Laser Altimeter (LOLA) instrument onboard the Lunar Reconnaissance Orbiter and from the Lunar Reconnaissance Orbiter Camera (LROC) mosaics to describe and classify these transitional features on the Moon, updating the existing catalogs of lunar peak-ring basins and proto-basins. They found 17 peak-ring basins (rim-crest diameters ranging from 207 km to 582 km) and 3 proto-basins (137–170 km). They also documented 23 craters exhibiting small ring-like clusters of peaks (50–205 km), one of which (Humboldt) exhibits a rim-crest diameter and an interior morphology that may be uniquely transitional to the process of forming peak rings. A power-law fit to ring diameters and rim-crest diameters of peak-ring basins on the Moon revealed a trend that is very similar to a power-law fit to peak-ring basin diameters on Mercury. Baker et al. (2010) developed plots of ring and rim-crest ratios versus rim-crest diameters for peak-ring basins and proto-basins on the Moon and showed that they also reveal a continuous, nonlinear trend that is similar to trends observed for Mercury and Venus. This result suggests that proto-basins and peak-ring basins are parts of a continuum of basin morphologies. Comparisons of the predictions of models for the formation of peak-ring basins with the characteristics of the Baker et al. (2010) basin catalog for the Moon suggest that formation and modification of an interior melt cavity and nonlinear scaling of impact melt volume with crater diameter provide important controls on the development of peak rings. Specifically, a power-law model of the growth of an interior melt cavity with increasing crater diameter is consistent with power-law fits to the peak-ring basin data for the Moon and Mercury. They concluded that the relationship between the depth of melting and depth of the transient cavity offers a plausible control on the onset diameter and subsequent development of peak-ring basins and also multiring basins, consistent with both planetary gravitational acceleration and mean impact velocity being important in determining the onset of basin morphological forms on the terrestrial planets (Baker et al., 2010).
Using high-resolution gravity data from the Gravity Recovery and Interior Laboratory (GRAIL) mission, Baker et al. (2017) analyzed the detailed gravity and crustal structure of lunar impact features in the morphological transition from complex craters to peak-ring basins. They showed that complex craters and proto-basins are characterized by free-air anomalies that are positively correlated with surface topography, unlike the prominent lunar mascons associated with large basins (positive free-air anomalies in areas of low elevation). The complex crater Bouguer gravity anomaly profiles are highly irregular, with central positive anomalies generally absent or not clearly tied to interior morphology. Gravity profiles for peak- ring basins (∼200 km to 580 km), in contrast, are much more regular and are highly correlated with surface morphology; a central positive Bouguer anomaly is confined within the peak ring. Baker et al. (2017) interpret the gravity anomalies within basins to be due to uplift of the mantle confined within the peak ring. They hypothesized that mantle uplift is influenced by interaction between the transient cavity and the mantle. They also found that mascon formation occurred over a wide range of basin sizes and was generally disconnected from the number of basin rings formed and the transition from peak-ring basins to multiringed basins.
Peak-ring basins (rim-crest diameters of 207–582 km) transition to multiring basins at larger diameters. As exemplified by the relatively fresh Orientale basin (Head, 1974), these multiple rings are characterized by an outer ring (the Cordillera Ring at 930 km diameter) separating the radially textured exterior ejecta from the basin interior, the Outer Rook Ring (620 km diameter), the next innermost Inner Rook Ring (480 km diameter), which is morphologically similar to peak rings in peak-ring basins, and finally, an inner depression whose outer edge defines an inward-facing scarp about 320 km in diameter. Neumann et al. (2015) used Lunar Reconnaissance Orbiter (LRO) topography and GRAIL gravity data to identify the global distribution of peak-ring and multiring basins and found that multiringed basins started at about 500 km diameter. The largest basin is South Pole–Aitken basin, located mostly on the farside, which has a diameter greater than 2,500 km. Oceanus Procellarum has been proposed to be the site of a huge (Gargantuan) basin (Wilhelms, 1982), but GRAIL data suggest that this region is not formed by several large impact basins, and alternate explanations have been proposed (Andrews-Hanna et al., 2014, Figure 5). The GRAIL inventory (Neumann et al., 2015) of lunar basins reveals 20 basins larger than 500 km, improving on earlier lists by more than a factor of 2. The nearside and farside hemisphere basin size-frequency distributions differ substantially; more basins larger than 350 km in diameter are seen on the nearside, whereas the farside has more smaller basins. Miljkovićć et al. (2013) attribute these differences to hemispherical differences in target properties (temperature and porosity).
Basin formation is still inadequately understood. This is because laboratory experiments and nuclear explosion tests are unsuitable because of the greater effect of gravity on the much larger-scale basin-forming impacts and size and energy differences. Currently, formation models differ in their interpretation of the modification stage of basin formation. Three hypotheses have been proposed:
(a) Computer models (e.g., Collins, Melosh, & Ivanov, 2004) suggest features within the basin rim are a result of the uplift of the transient crater floor above the pre-impact target surface creating a central peak, which subsequently collapses back into the target. This has been tested and supported by the IODP-364 expedition, which drilled the inner ring of Chicxulub crater and confirmed that the inner ring is made of basement rocks (Morgan et al., 2016).
(b) Other theories suggest features within the basin rim are a result of listric (curved normal) faulting along the boundary of material displaced by the transient crater, which uplifts the crater floor (e.g., Cintala & Grieve, 1998). These theories, however, do not directly treat the formation of multiple outer rings outside of the basin rim, a process that was elaborated on by Head (2010).
(c) Outer ring formation has been proposed to be dependent on the thickness and strength of the target lithosphere. If the transient crater penetrates through the lithosphere into the asthenosphere below, asthenospheric material is predicted to flow toward the basin center exerting a drag force on the lithosphere above. A sufficiently weak lithosphere will fracture, forming the ring structures (Holsapple, 1989; McKinnon & Melosh, 1980).
Recent very high-resolution GRAIL gravity data (Zuber et al., 2016), combined with iSALE hydrocode models (Johnson et al., 2016), provide new insights into the Orientale multiring basin formation and permit development of an updated model for mascon formation (Melosh et al., 2016). Zuber et al. (2016) used 3- to 5-km horizontal resolution GRAIL data to show that a volume of at least 3.4 ± 0.2 × 106 km3 of crustal material was removed by the event and redistributed during basin formation and that the transient crater diameter could be inferred to be between 320 and 460 km. The high-resolution gravity field resolved distinctive structures of the outer three rings of Orientale; the presence of faults that penetrate to the mantle are interpreted to be associated with the outer two rings. Johnson et al. (2016) used a hydrocode to simulate the formation of the Orientale multiring basin. They produced a subsurface structure consistent with high-resolution GRAIL gravity data. The simulated impact produced a transient crater, ~390 km in diameter, that was not maintained because of subsequent gravitational collapse. The simulations indicated that the flow of warm, weak material at depth was crucial to the formation of the outer rings of the basin, which are large normal faults that formed at different times during the collapse stage. The key parameters controlling ring location and spacing are interpreted to be (a) impactor diameter and (b) lunar thermal gradient. On the basis of these and related data, Melosh et al. (2016) noted that GRAIL gravity showed free-air gravity anomalies over lunar impact basins that reveal a bullseye pattern consisting of a central positive (mascon) anomaly, a positive outer annulus, and an intermediate negative collar. Melosh et al. (2016) interpreted this pattern to result from impact basin excavation and collapse, followed by isostatic adjustment and cooling and contraction of a voluminous melt pool. They used (a) a hydrocode to simulate the impact, and (b) a self-consistent finite-element model to simulate the subsequent viscoelastic relaxation and cooling. They concluded that the main factors controlling mascon-basin modeled gravity signatures are (a) the impactor energy, (b) the lunar thermal gradient at the time of impact, (c) the crustal thickness, and (d) the extent of volcanic fill.
Impact Craters and Lunar Chronology
Impact craters can also be used as an indication of the relative and absolute ages of geological units by counting craters and determining the CSFD. Crater density is closely related to accumulation time; thus different geological units formed at different times have different crater densities. The lowest crater density is found in the young impact basins, especially Orientale, and in the mare regions. The highest crater density occurs in the highlands of the southern nearside and north-central farside that reached a state of saturation equilibrium (Head et al., 2010; Kadish et al., 2011).
CSFD analysis, compiled on the basis of different crater densities on different planetary surfaces, has been the major technique used in remote age determination for planetary surfaces. While crater density comparisons reveal relative ages for geological units on a given planetary body, absolute model ages derived from measured crater density are determined by a known crater production rate, which in turn is calibrated by the observed crater density on surfaces of known age. Since the Moon is the only planetary body besides the Earth for which samples were collected from known locations, the lunar crater chronology was first established after the Luna and Apollo missions, and it has been extrapolated to the other planetary bodies based on impact-flux ratios estimated from orbital dynamics (Stöffler et al., 2006). The core part of the lunar chronology contains two constraints: (a) a crater production function that describes the size–frequency relationship of different-sized craters formed during a given time interval; and (b) a crater chronology that describes the relationship between surface age and crater density.
The lunar crater chronology has been demonstrated to be robust and efficient at least to the first order via many successful applications. For example, lunar mare ages estimated by crater counts are broadly consistent with radiometric ages of samples. However, an area of active research focus is on the problems and potential caveats that exist in the current understanding of lunar crater chronology. For example: (a) A robust link between geological units and the Apollo and Luna samples has been questioned, so that perhaps many of the calibration points on the lunar crater chronology function may not be correct (e.g., the ages of Copernicus, Tycho, and Imbrium). (b) The observed crater size–frequency distribution on the sample return areas may not represent that of the production population, since the effects of secondaries, target properties, saturation equilibrium, and boundaries of geological units are not fully resolved. (c) The suitable stratigraphic ranges that could be dated by the crater chronology function are highly debated because the lunar crater production functions may have changed with time.
Stöffler and Ryder (2001) re-evaluated the lunar flux curve using isotopic ages of lunar samples and the latest views on the lunar stratigraphy as well as the principles of relative and absolute age dating of geological surface units of the Moon. They derived the following best estimates for the ages of the multiring basins: 3.92 ± 0.03 Gyr for Nectaris; 3.89 ± 0.02 Gyr for Crisium; 3.89 ± 0.01 Gyr for Serenitatis; and 3.85 ± 0.02 Gyr (Ar-Ar age) and 3.91–3.92 Gyr (Pb-Pb age; Gnos et al., 2004; Liu et al., 2012; Merle, Nemchin, Grange, Whitehouse, & Pidgeon, 2014; Snape et al., 2016) for Imbrium. The best estimates for the ages of the mare landing areas are: 3.80 ± 0.02 Gyr for Apollo 11 (old surface); 3.75 ± 0.01 Gyr for Apollo 17; 3.58 ± 0.01 Gyr for Apollo 11 (young surface); 3.41 ± 0.04 Gyr for Luna 16; 3.30 ± 0.02 Gyr for Apollo 15; 3.22 ± 0.02 Gyr for Luna 24; and 3.15 ± 0.04 Gyr for Apollo 12. These data result in a revised lunar impact flux curve that differs from the previously used flux curve. The impact flux curve for this pre-Nectarian time remains unknown. The new calibration curve for lunar crater retention ages less than about 3.9 Gyr provides an updated standard reference for the inner solar system bodies, including Mars. Obviously, available sample and chronology data are insufficient to precisely define the lunar impact-flux curve. Reconciling ages recorded by the different chronometers is an ongoing issue. Samples are needed from the oldest pre-Nectarian and youngest Eratosthenian terranes, where samples of known provenance have not been collected. The coming new lunar exploration missions, especially the future sample return missions, together with the available comprehensive remote sensing data sets, will undoubtedly solve many of the potential problems and caveats in the lunar chronology function.
Volcanic Features and Volcanism
Prior to Lunar Reconnaissance missions and the Apollo Lunar Exploration Program, there was an intense debate about the age of the lunar surface (young or old), the origin of craters (volcanic or impact), and whether the Moon accreted and evolved cold or hot (Taylor, 1982). Early Reconnaissance missions, such as the Lunar Orbiters, confirmed that there was a wide diversity of volcanic features on the Moon (Figure 14), including lava flows, lava channels, pit craters, cones, domes, volcanic complexes, and mantling pyroclastic blankets that are also commonly seen on Earth, that the craters are of impact origin, and that the samples collected by the Apollo astronauts Armstrong and Aldrin demonstrated the basaltic and ancient nature of the lunar maria. Using lunar penetrating radar (LPR) data from Chang’e-3 Yutu Rover, Xiao et al. (2015) reported multilayered subsurface structures of shallow crust within the mare (Figure 15), providing valuable information to reveal the lava eruption extents, style, and filling history within the Imbrium basin.
Also important is the fact that there are some volcanic features observed on Earth and Mars, such as major shield volcanoes (e.g., Hawaii, Olympus Mons) and large calderas, that were not observed on the Moon, and others, such as sinuous valleys (sinuous rilles) tens to over a hundred kilometers long, that are not observed on Earth (Head & Wilson, 1992). Furthermore, the volcanic features observed on the Moon appeared to be overwhelmingly basaltic in nature, and little evidence was found for the petrologic diversity observed on Earth’s continental crust (e.g., dacites, granites, and rhyolites). Initial investigations emphasized the similarity in the lunar volcanic features to those on Earth, and much emphasis was placed on terrestrial analogs in the interpretation of basaltic volcanism on the Moon (Greeley & King, 1977). These excellent studies were instrumental in helping to choose the landing sites for the Apollo Lunar Exploration Program.
Over time, however, the fundamental physical and geological differences between the Earth and Moon (e.g., one-sixth gravity, lack of an atmosphere, thick anorthositic crust, no plate tectonics, rapid heat loss, thick lithosphere, and emerging detailed knowledge of lunar petrology) led planetary geoscientists to consider the generation, ascent, and eruption of magma in a lunar context (Wilson & Head, 1981). In the early 21st century, the ascent and eruption of lunar mare basalt magmas has been modeled with new data on crustal thickness and density (from GRAIL), magma properties, and surface topography, morphology, and structure (from Lunar Reconnaissance Orbiter) (Head & Wilson, 2017; Wilson & Head, 2017b). Comparing the spatial variation of the bulk density structure of the crust of the Moon with the densities of lunar basaltic magmas shows that essentially all lunar magmas were negatively buoyant everywhere within the lunar crust. This situation requires positive excess pressures in melts at or below the crust–mantle interface in order to enable them to erupt. In mantle partial-melt regions, probably at least a few percent of melting must have taken place, likely over a vertical extent of up to 150 km. If melt percolates upward from a partial-melt zone and accumulates as a magma reservoir, either at the density trap at the base of the crust or at the rheological trap at the base of the elastic lithosphere, the excess pressure at the top of the magma body will exert an elastic stress on the overlying rocks. This stress will eventually cause the overlying rocks to fail in tension, allowing a dike to propagate upward from this point. Magma accumulations at the base of the crust would have been able to intrude dikes partway through the crust, but not able to feed eruptions to the surface. In order to be erupted, magma must have been extracted from deeper mantle sources, consistent with petrologic evidence.
Buoyant dikes propagating upward from deep mantle partial-melt sources can disconnect from their source regions and travel through the mantle as isolated bodies of melt that encounter and penetrate the crust–mantle density boundary. For larger source region extents, the dike can reach the surface and erupt on the lunar nearside but still cannot reach the surface on the farside; for even larger source extents, eruptions could occur on both the nearside and the farside. Wilson and Head (2017b) concluded that the paucity of farside eruptions implies a restricted range of vertical extents of partial-melt source–region sizes, between ~16 and ∼36 km. They found that when eruptions can occur, the available pressure in excess of what is needed to support a static magma column to the surface provides information on the pressure gradient driving magma flow. The resulting typical turbulent-magma rise speeds are ∼10 to a few tens of m s−1, dike widths are of the order of 100 m, and eruption rates from 1 to 10 km long fissure vents are of the order of 105 to 106 m3 s−1. These values are huge compared to typical terrestrial eruptions and help to explain many of the unusual characteristics of lunar basaltic eruptions.
Volume fluxes in lunar eruptions (derived from lava flow thicknesses and surface slopes or rille lengths and depths) are found to be of the order of 105 to 106 m3 s−1 for volume-limited lava flows and >104 to 105 m3 s−1 for sinuous rilles, with dike widths of ∼50 m. The lower end of the volume-flux range for sinuous rilles corresponds to magma rise speeds near the limit set by the fact that excessive cooling would occur during flow through a 30 km-long dike kept open by a very low excess pressure. Therefore, these eruptions were probably fed by partial-melt zones deep in the mantle. Longer eruption durations appear to be the key to the ability of these flows to erode sinuous rille channels.
High flux effusive eruption can produce lave tubes. Lava tubes are common on terrestrial basaltic lava fields, such as Hawaii island, Juju Island, Iceland, western Arizona of the United States, and northeastern China. Lava tubes were also predicated on the Moon in the middle of the 20th century (Oberbeck, Quaide, & Greeley, 1969; Cruikshank & Wood, 1972). The SELENE (Kaguya) cameras have detected the collapsed lava tube in the Marius Hill area (Haruyama et al., 2009; Kaku et al., 2017), and the LRO narrow-angle camera has further confirmed and found new skylights of buried lava tubes (Wagner & Robinson, 2014) (Figure 16). The width and length of the lava tubes were estimated as several tens of meters wide and up to hundreds of kilometers long. These new findings attracted wide attention and many groups have proposed to use the lava tubes as candidate sites for lunar base and in situ resource utilizations of the Moon (Coombs & Hawke, 1992; Walden, Billings, York, Gillett, & Herbert, 1998; Haruyama et al., 2012a, 2012b; Wagner et al., 2014; Xiao, Huang, Zhao, & Zhao, 2018).
Differences in the overall abundance of volatiles, the volatile species generated, the depth of their production, and their behavior on reaching the surface also correspond to significant differences in eruption behavior between the Earth and Moon. The great vertical extent of dikes means that the major lunar magmatic volatile, likely CO, was produced in large amounts (up to 2,000 ppm by mass) over a wide range of depths in the dike (potentially extending down to at least 50 km), and that up to at least 1,000 ppm of water and sulfur compounds could be released in the upper few hundred meters of dikes (Wilson & Head, 2018). Furthermore, because basaltic volcanic eruptions on the Moon took place in conditions of low acceleration due to gravity and negligible atmospheric pressure, essentially all lunar eruptions were characterized by an explosive component. Lunar versions of Hawaiian and strombolian explosive activity differ greatly from those on Earth, however (e.g., there is no lunar analog of a convecting plinian eruption cloud). The extreme expansion of even the smallest amounts of gas produced pyroclasts predominantly of sub-millimeter size. Following an acceleration phase as magmatic gas expanded away from the vent into the vacuum, pyroclasts were always dispersed ballistically. For example, steady eruption of relatively volatile-rich lunar magmas would lead to pyroclast speeds of up to 180 m s−1 and maximum ranges of ~20 km (Head & Wilson, 2017). Even greater speeds and ranges are possible in the initial stages of eruptions as gas concentrated in the upper tips of dikes was released, a candidate explanation for producing regional pyroclastic blankets.
In summary, theoretical treatment of the generation, ascent, and eruption of magma on the Moon (Wilson & Head, 2017b) predicts that the following major features are factors in the nature of volcanic deposits and styles, and how they change with lunar thermal evolution.
Magma Generation, Ascent, and Eruption
Contrasts in density between the bulk mantle and regions with greater heat sources causes larger heated regions to rise buoyantly as melt-rich diapirs; these generate partial melts that undergo collection into magma source regions. These diapirs can rise to the crustal density trap at the base of the anorthositic crust (when the crust is thicker than the elastic lithosphere). Later in history, when the thickening lithosphere exceeds the thickness of the crust, the diapirs rise to the base of the lithospheric rheological trap. Upon arrival and stalling, sufficient stress to cause brittle deformation of the elastic part of the overlying lithosphere is caused by (a) residual diapiric buoyancy and (b) continued production and arrival of diapiric material, enhancing melt volume and over-pressurizing the source regions. The direct result is initiation and propagation toward the surface of a magma-filled crack as a convex upward, blade-shaped dike.
These theoretical considerations also predict that since all eruptions began with the arrival of a dike at the surface, the initial vents are always fissures at the tip of the dike. Magma volume released in a single dike event is likely to lie in the range of 102–103 km3, corresponding to dikes with widths of 40–100 m and 60–100 km vertical and horizontal extents. These properties favor eruption on the lunar nearside. In contrast to the small dike widths and volumes and the low propagation velocities typical on Earth, lunar dike propagation velocities are typically sufficiently high so that shallow sill formation is not favored.
Lunar Thermal Evolution and Magmatic History
As the Moon cooled with time, the lithosphere thickened, source regions were less abundant, and magma rising diapirs encountered rheological traps at ever increasing depths. In parallel, the global lithospheric state of stress became increasingly contractional. These two factors served to inhibit dike emplacement and surface eruptions in the last third of lunar history.
Using these basic principles of generation, ascent, and eruption, and the implied predictions for eruption characteristics, Head and Wilson (2017) synthesized and documented the array of observed lunar volcanic features and styles and compared them to the theoretical predictions.
Dikes emplaced into the shallow crust (Figure 17) are predicted to stall and produce crater chains due to active and passive gas venting (e.g., Mendeleev Crater Chain). If the dikes are sufficiently shallow, they can create a near-surface stress field that forms linear and arcuate graben, with accompanying pyroclastic and small-scale effusive eruptions (e.g., Rima Parry V). Local low-density breccia zones beneath impact crater floors, however, can cause lateral magma migration to form laccoliths (e.g., Vitello Crater) and sills (e.g., Humboldt Crater) in floor-fractured craters.
These are modulated by effusion rates, eruption durations, cooling and supply limitations to flow length, and preexisting topography (Wilson & Head, 2018). Small-shield volcanoes (e.g., Tobias Mayer, Milicius) are predicted to occur from relatively low effusion rate cooling-limited flows, whereas higher effusion rate cooling-limited flows lead to compound flow fields of the type that occur in most mare basins. Sinuous rilles (Hurwitz, Fassett, Head, & Wilson, 2010; Hurwitz et al., 2013) (e.g., Rimae Prinz) are predicted to form from even higher effusion rate longer-duration flows that lead to thermal erosion of the vent, enhancement of effusion rate, and thermal erosion of the substrate. Volume-limited flows with lengths of many hundreds of kilometers (e.g., the young Imbrium basin flows) are predicted to occur during extremely high effusion rate flows on slopes.
Explosive Pyroclastic Eruptions
Pyroclastic eruptions are common on the Moon because of the low-pressure environment in propagating dike crack-tips, which can cause gas formation at great depths and throughout dike ascent. Furthermore, at shallow crustal depths, both the smelting reaction and the recently documented abundant magmatic volatiles in mare basalt magmas contribute to significant shallow degassing (Rutherford, Head III, Saal, Hauri, & Wilson, 2017) and pyroclastic activity associated with the dike as it erupts at the surface (Wilson & Head, 2017a). As the gas-rich dike penetrates to the surface, a wide range of explosive eruption types can be produced (Gaddis et al., 1998; Gaddis, Hawke, Robinson, & Coombs, 2000; Gaddis, Staid, Tyburczy, Hawke, & Petro, 2003), whose manifestations are modulated by lunar environmental conditions (e.g., Head & Wilson, 2017): Terrestrial strombolian-style eruptions map to cinder and spatter cone-like constructs (e.g., Isis and Osiris); Hawaiian-style eruptions map to broad flat pyroclastic blankets (e.g., Taurus-Littrow Apollo 17 dark mantle deposits); vulcanian-like eruptions caused by solidification of magma in the dike tip, buildup of gas pressure, and explosive disruption can form dark halo craters with mixed country rock (e.g., Alphonsus Crater floor); Ionian-like eruptions can be caused by artificial gas buildup in wide dikes, energetic explosive eruption, and formation of a dark pyroclastic ring (e.g., Orientale dark ring); and regional dark mantle deposits (e.g., Rima Bode, Sinus Aestuum) can be caused by multiple eruptions from many gas-rich fissures. Long-duration, relatively high effusion-rate eruptions accompanied by continuing pyroclastic activity are predicted to cause a central thermally eroded lava pond and channel, a broader pyroclastic “spatter” edifice, an even broader pyroclastic glass deposit, and, if the eruption lasts sufficiently long, an associated inner thermally eroded vent and sinuous rille channel. The Cobra Head, Schröteri Valley, and the Aristarchus Plateau dark mantle are interpreted to be examples of this type of eruption.
Global Distribution of Mare Basalts
The nearside–farside asymmetric distribution of mare basalt deposits is most plausibly explained by crustal thickness differences; intrusion is favored in the thicker farside crust and extrusion is favored on the thinner nearside crust. Several second-order effects on the global distribution include a regional and global thermal structure (areal and temporal variations in lithospheric thickness) and broad geochemical anomalies (the Procellarum KREEP Terrane).
Chronological Distribution of Mare Basalts
The rapidly decreasing integrated flux of mare basalts is a predicted result of the thermal evolution of the Moon. The following factors progressively inhibited the generation, ascent, and eruption of basaltic magma with time: Continued cooling and crystallization of the Moon’s core decreased diapiric rise and mantle melting, thickened the lithosphere, and caused the global state of stress to be increasingly contractional.
The vast majority of lunar volcanic landforms and samples are related to basaltic compositions. A few exceptions were noted in the early stages of lunar exploration, most specifically the 20 km-diameter 1,200 m-high Gruithuisen Gamma dome, characterized by a distinctive reflectance spectrum with a downturn in the UV (thus, these features are known as “red spots”) (Head, Hess, & McCord, 1978). Some red spots are also associated with thorium anomalies (Hagerty et al., 2006). The morphology, geology, and spectral characteristics of these deposits argued for more evolved compositions (such as dacites or rhyolites) (Wilson & Head, 2003), an interpretation supported by recent LRO DIVINER experimental data (Glotch et al., 2011). A variation on this theme has been reported in the Compton–Belkovich area (Jolliff et al., 2000). Dating of the Gruithuisen domes (Ivanov, Head, & Bystrov, 2016; Wagner, Head, Wolf, & Neukum, 2002, 2010; Shirley, Zanetti, Jolliff, van Der Bogert, & Hiesinger, 2016) places them generally coincident with the early post-Imbrium basin, near the beginning of mare basalt volcanism. The age of the Compton–Belkovich structure has been interpreted to range from Imbrian (>3 Gyr) (Shirley et al., 2016) to perhaps as young as Copernican (Jolliff et al., 2000). Four hypotheses have been proposed for their origin: (a) liquid immiscibility (Longhi, 1990); (b) differentiation, fractional crystallization; (c) assimilation (Wilson & Head, 2017b); and (d) bimodal volcanism (Hagerty et al., 2006). The role of these features in the integrated petrogenetic history of the Moon has yet to be firmly established.
Late-Stage Volcanic Eruptions
Late-stage volcanic eruptions are predicted to be widely separated in time and characterized by high-volume, high-effusion rate eruptions producing extensive volume-limited flows; these are the characteristics of the young Eratosthenian-aged Imbrium basin flows (e.g., Schaber, 1973; Bugiolacchi & Guest, 2008), a predicted characteristic of deep source regions below a thick lithosphere later in lunar history. Recently, Braden et al. (2014) proposed that a series of dozens of Irregular Mare Patches (IMPs), including the enigmatic feature Ina (characterized by unusual mounds and rough, hummocky floors) (Figure 17), represented surface volcanic activity that occurred, on the basis of impact CSFD ages, in the past 100 Ma years of lunar history. Such mantle-derived surface volcanic activity so late in lunar history presents a major challenge to the current paradigm of lunar thermal evolution (the evidence for waning volcanism in middle lunar history, the evolving global state of stress in the lithosphere, and lunar thermal evolution models). Models for the young formation of Ina included recent gas emissions (Schultz et al., 2010) and recent lava extrusions (Braden et al., 2014). An additional possible interpretation was proposed by Qiao et al. (2017) and Wilson and Head (2017a), who examined Ina (Figure 18), one of the key IMP examples that was dated by Braden et al. (2014) at ~33 Ma. The IMP associated with Ina lies within a summit pit crater atop a small-shield volcano that dates from about 3.5 billion years ago, during the peak phase of lunar mare basalt volcanism. They examined the ascent and eruption of magma in the waning stages of the eruptive process in small-shield summit pit crater floors. They showed that a candidate alternate interpretation of many IMP characteristics might be by basaltic magma behavior as the rise rate of the ascending magma in dikes below pit craters slows to zero, volatiles exsolved in the dike and lava lake, a very vesicular foam forms, and the dike begins to close. They found that stresses in the very vesicular and porous lava lake crust could produce fractures through which the foam could extrude to produce convex mounds whose physical properties inhibit typical impact crater formation and regolith development. Such properties would create an artificially young crater retention age. This alternative mechanism for the production and extrusion of very vesicular magmatic foams is also applicable to waning-stage dike closure associated with pit craters atop dikes and fissure eruptions in the lunar maria, providing a potential explanation for many irregular mare patches. This late-stage foam extrusion mechanism implies that IMPs and associated mare structures (small shields, pit craters, and fissure flows) formed synchronously billions of years ago, and provides an alternate explanation for the very young ages (less than ~100 Ma years) proposed for IMPs by others.
Another unique mare volcanic feature is Ring-Moat Dome Structure (RMDS) (Zhang et al., 2017; Figure 19). These low domes (a few meters to ~20 m height, with slopes <5°) are typically surrounded by narrow annular depressions or moats. There are more than 4,000 RMDSs in the lunar maria with diameters ranging from tens to hundreds of meters. Zhang et al. (2017) proposed a mechanism for the formation of the RMDS related to modification of the initial lava flows through inflated flow squeeze-ups and extrusion of magmatic foams below a cooling lava flow surface. These newly discovered features provide new insights into the nature of emplacement of lunar lava flows, suggesting that in the waning stages of a dike emplacement event, magmatic foams can be produced, extrude to the surface as the dike closes, and break through the upper lava flow thermal boundary layer (crust) to form foam mounds and surrounding moats.
Distribution, Duration, and Flux of Lunar Volcanism
Volcanism is one of the most important geological processes on the Moon and lasted for several billion years. Following the formation and solidification of the lunar crust over 4.4 billion years ago, magmatism occurred and magma ascended to the surface to form ancient cryptomare basalts (Schultz & Spudis, 1983; Head & Wilson, 1992), later mare basalts and pyroclastic deposits, and produced distinct volcanic landforms, such as lava plains, rilles, lava tubes, volcanic complexes, domes, and unique IMPs (irregular mare patches) (e.g., Head & Wilson, 2017).
The presence of ancient basalt clasts in the Apollo sample records, some as old as 4.23 Gyr (Taylor et al., 1983), contrasts sharply with the age range of observed mare deposits, which are only 3.9–2.5 Gyr (Head, 1976). These hidden mafic deposits, termed “cryptomaria” by Head and Wilson (1992), are mare basalt deposits whose low albedo signature has been hidden or obscured by superposed high albedo material, emplaced prior to ~3.9 Gyr during the era of large-impact basin formation (Antonenko, Head, Mustard, & Hawke, 1995). Criteria for the detection of cryptomare deposits include the presence of dark-haloed craters (craters that penetrate through the overlying ejecta to excavate underlying mafic material) and the identification of spectral or geochemical anomalies. Cryptomare deposits cover as much of the lunar surface as half of the known maria and have a volume that is equal to 30% of the known mare volume (Head & Wilson, 1992) (Figure 20). It is clear that they represent a significant contribution to global lunar volcanism (Jolliff, Wieczorek, Shearer, & Neal, 2006; Hiesinger & Head, 2006).
Recent research (Whitten & Head, 2015a, 2015b) shows that the total area covered by cryptomare volcanism is ~2% of the lunar surface, which increases the total area covered by mare volcanism. Crustal thickness variations are likely to play an important role in mare basalt emplacement. Most cryptomaria are located on the lunar nearside, in association with younger mare deposits. They are mainly located within ancient impact basins and they are more frequent in the eastern hemisphere than the western hemisphere. Sori, Zuber, Head, and Kiefer (2016) used GRAIL gravity and Lunar Orbiter Laser Altimeter (LOLA) topography data to construct maps of the Moon’s positive Bouguer and isostatic gravity anomalies in order to explore crypto-volcanism. They interpreted several anomalies as cryptomare deposits; these candidate deposits would increase the lunar surface area containing volcanic deposits from 16.6% to between 17.9% and 19.5%, underlining the fact that early (pre-3.8 Gyr) lunar volcanism is an important element of lunar thermal evolution (Figure 20).
The volcanic flux on the Moon has varied as a function of time. Higher fluxes occurred early in lunar history, then decreased steadily over time (Kirk & Stevenson, 1989). Hiesinger, Head, Wolf, Jaumann, and Neukum (2011) dated the major nearside mare basalts using CSFD analysis and found that (a) lunar volcanism was active for almost 3 Gyr, starting at ~3.9–4.0 Gyr ago and ceasing at ~1.2 Gyr ago, (b) most basalts erupted during the late Imbrian period at ~3.6–3.8 Gyr ago; (c) significantly fewer basalts were emplaced during the Eratosthenian period; and (d) basalts of possible Copernican age have been found only in limited areas in Oceanus Procellarum (Figure 20). Hiesinger et al. (2011) further underscored the predominance of older mare basalt ages in the eastern and southern nearside and in patches of maria peripheral to the larger maria, in contrast to the younger basalt ages on the western nearside (i.e., in Oceanus Procellarum).
Unlike Earth, with multiple, segmented, laterally moving and colliding plates, the Moon is a one-plate planet (Solomon, 1978) with a single global lithospheric plate that has thickened with time. In contrast to the divergent plate boundaries, orogenic belts, and subduction zones of Earth, the Moon is characterized by six major types of tectonism and tectonic style: (a) cooling and contraction associated with the formation of major impact basins (transfer of kinetic energy from the impactor to the target, and uplift of deep isotherms); (b) regional and global seismic effects associated with crater and basin formation; (c) surface and near-surface deformation caused by the emplacement of dikes and near-surface extensional stresses to form graben (e.g., Klimczak, 2014); (d) emplacement of thick volcanic deposits on the lunar surface, resulting in lithospheric loading and flexural deformation (e.g., Solomon & Head, 1981); (e) tidal deformation and moonquake formation caused by the slight eccentricity of the lunar orbit (e.g., Watters et al., 2015); and (f) long-term changes in lunar thermal evolution causing changes in the global state of stress in the lithosphere to change from extensional to contractional (e.g., Solomon & Head, 1981). The vast majority of the visible tectonic features are due to (a) loading of mare basalts and resulting flexural deformation, and (b) the global thermal evolutionary change in the net state of stress in the lithosphere. As such, the major tectonic features are generally classified as compressional (e.g., lobate scarps and wrinkle ridges) and extensional (e.g., graben).
Lunar wrinkle ridges are linear to sinuous landforms on the lunar surface, and most of them occur in the mare basins (Figure 21). They typically consist of a broad arch and a superposed sharper ridge, although the detailed morphologies can vary to a large extent. Although wrinkle ridges are generally large in scale and their morphology is distinct, their formation mechanism (tectonic or volcanic) and formation time have been extensively debated. The debate on the lunar ridges being of tectonic or volcanic origin gradually reached the consensus (e.g., Sharpton & Head, 1988) that they result from horizontal shortening (Ono et al., 2009; Watters & Johnson, 2010), with lunar thrust faulting being the main mechanism (e.g., Watters et al., 2015). This consensus has been supported by the interpretation of Lunar Sounder experiment data from Apollo 17 Command/Service Module, the study of candidate terrestrial analogs, and the interpretation of lunar radar sounder data from the Kaguya spacecraft (SELENE) (Ono et al., 2009).
Global mapping by Yue, Michael, Di, and Liu (2017) has identified nearly 3,000 segments with a total length of over 25,000 km. These wrinkle ridge segments were classified into three categories: concentric, parallel, and isolated ridges, and most of them are located within the maria regions, over basalts. A global stress field is required to produce these wrinkle ridges (Yue, Li, Di, Liu, & Liu, 2015). Recent study suggests that the ridge groups formed with average ages between 3.5 and 3.1 Gyr ago, or 100–650 Ma after the oldest observable erupted basalts where they are located, except for the ridges in mare Tranquillitatis. These have been interpreted as a result of local stresses from loading by basalt fill as the principal agent responsible for the formation of lunar wrinkle ridges. Wrinkle ridge formation in Tranquillitatis is likely to indicate a different mechanism of stress accumulation at this site.
Lobate scarps are another expression of lunar shortening deformation. These features are interpreted as the surface expression of thrust faults where the horizontal stress component is larger than the vertical component (Binder, 1982; Watters et al., 2010). They commonly appear as asymmetrical ridges with steeply dipping scarp faces and gently sloping back limbs. These scarps are also indicators of the geological and thermal history of the Moon. Two differing lunar initial thermal models have been proposed: the initially totally molten Moon (ITM) (Binder & Lange, 1980; Runcorn, 1977) and the lunar magma ocean (LMO) (Solomon & Chaiken, 1976). Each model has different implications for the timing and duration of faulting on the Moon. The ITM model suggests that the Moon was in a hot initial state, either near or above basalt solidus, and the lunar surface should have a young age (~3 Gyr) (Solomon & Chaiken, 1976). In contrast, the LMO model predicts that the highlands would be devoid of young lobate scarps or other shortening landforms, since more stress should be required for faults to slip in the highlands versus the maria, as the highly brecciated material in the highlands would need to first compact before thrust faulting can occur (Solomon & Chaiken, 1976). The fact that lobate scarps are observed in the highlands as well as in the maria seems to disagree with the LMO model. Studies of scarps observed in Apollo imagery and the use of thermoelastic stress models have revealed that horizontal compressional stresses resulting from cooling a completely molten Moon would be much higher than those predicted by the LMO model, and sufficient to allow thrust faulting in both the highlands and mare deposits. Recently, these fresh landforms have been discovered throughout the lunar surface in Lunar Reconnaissance Orbiter Camera (LROC) images, which implies a stress state where horizontal compressive stresses globally greater than vertical compressive stresses have persisted to the present day (Watters et al., 2014), although highland faults are still possible in the LMO model (Watters, Robinson, Banks, Tran, & Denevi, 2012; Williams, Watters, Pritchard, Banks, & Bell, 2013).
Another crucial question is the timing of scarp formation; that is, how ancient or recent are these tectonic landforms? In past studies, the timing of contraction (and extension) is poorly constrained. Binder and Gunga (1985) used the method of Trask (1971), where craters were classified according to their state of degradation and calibrated data of Moore, Boyce, and Hahn (1980) to compute crater ages for those partially cross cut by faults, and revealed that the majority of lobate scarps were formed in the past 700 Ma. Initial morphological evaluations by Watters et al. (2010) revealed that these scarps are relatively young because of their small sizes, crisp appearances, and cross-cutting relationships of small craters. Likewise, van der Bogert et al. (2012) derived absolute model ages (AMAs) for the well-known Lee-Lincoln and Mandel’shtam scarps and affirmed that movement along the faults has occurred as recently as 75 Ma ago. Furthermore, the timing of scarp formation from thrust faults could help distinguish between the validity of the two lunar thermal models (Watters et al., 2015). Using crater size-frequency distribution (CSFD) measurements, Clark et al. (2017) derived absolute model ages for the scarp surfaces proximal to the trace of the fault and found that the last slip event occurred in the past ~132 Ma. Along with young model ages, lunar lobate scarps exhibit a youthful appearance with their crisp morphologies, interpreted to be indicative of late-stage horizontal shortening.
Graben are characterized by narrow, flat-floored depressions with lengths that exceed widths and are bounded by two steeply dipping antithetic normal faults (Golombek, 1979; McGill & Stromquist, 1979) (Figure 22). Typical dimensions of small-scale graben range from tens to hundreds of meters wide and up to a couple of kilometers in length, making these a distinct population from the basin-related graben. Small-scale grabens were noted by Watters et al. (2010) in association with the Lee–Lincoln scarp and are oriented subparallel and perpendicular to the strike of the scarp. The formation of these graben was interpreted to be the result of regolith and bedrock extension due to flexural bending during formation of the Lee–Lincoln scarp (Watters et al., 2012). Additional small-scale grabens have been discovered, with estimated ages on the order of 50 Ma (Watters et al., 2012) based on crosscutting relationships with small-diameter impact craters, lack of superposed craters, and infilling rates of shallow depressions, indicating recent extension on a body that is globally contracting.
Using LROC high-resolution image data, French, Bina, Robinson, and Watters (2015) completed global mapping of small-scale lunar graben, assessed their distribution and dimensions, and discussed their formation processes. They found that small-scale grabens are globally distributed and are commonly found near lobate scarps and wrinkle ridges. Some graben can be related to localized tension from flexural bending or dilation associated with lobate scarp and wrinkle ridge formation. The ages of these graben are very young. Maximum graben ages range from late Eratosthenian to early Copernican based on superposing and crosscutting relationships, and crater ages, with a group of graben deforming ejecta from the Copernicus crater. Globally distributed, young, small-scale graben populations were formed as a result of localized extension either from flexural bending or dilation due to contractional faulting or volcanic uplift, indicating a significant level of recent geological activity.
Recently, dozens of linear grabens that are about 10–400 m wide and less than 1 km long are recognized in the southeastern continuous ejecta deposits of Copernicus, supporting the idea that Copernican-aged tectonism and possibly magmatism have occurred on the Moon (Xiao, Huang, Zeng, & Xiao, 2017).
Surface Composition and Regolith
Major elements of the Moon (e.g., Fe, Ti, Mg, Al, Ca, and Si) play very important roles in understanding its composition, origin, and evolution. Until this new era of lunar exploration, global high spatial resolution maps of lunar elements were only derived from Clementine data and for FeO and TiO2 only (Blewett, Lucey, Hawke, & Jolliff, 1997; Lucey et al., 1998; Gillis, Jolliff, & Elphic, 2003). Chang’e-1 Interference Imaging Spectrometer (IIM), a hyperspectral imaging interferometer with 32 continuous channels within the wavelength range of 480–946 nm, acquired first global high spatial resolution maps for all the six major elements and Mg# (Wu, 2012), which contributed to the geological research of local areas. The 200-m resolution maps of elemental abundances revealed a number of implications for the understanding of the Moon (Wu et al., 2012; Wu, 2012).
Rock Type and Distribution
So far, lunar rock types and distribution were mainly investigated by examination of lunar rock samples and by orbital data. The results suggested that lunar rocks include three varieties: igneous, metamorphic, and sedimentary (Lucey et al., 2006). There are mainly five types of igneous rock suites: (a) ferroan anorthosite suite (FAN); (b) magnesian suite; (c) alkali suite; (d) KREEP basalt; and (e) mare basalt and related pyroclastic deposits (Lucey et al., 2006; Wieczorek et al, 2006). Lunar breccias were generated from the impact mixing of these primary rock types (Wieczorek et al., 2006). In addition, lunar meteorites represent random rock samples from the Moon and were produced from the impact mixing of various rock types (Korotev, 2015).
Recent studies disclosed some new rock types. A new olivine-rich type of mare basalt was disclosed by Chang’e-3, which contains abundant olivine and high-Ca ferropyroxene (Ling et al., 2015; Zhang et al., 2015). Mg-spinel rocks, a previously unsampled member of the magnesian suite, were discovered by Chandrayaan-1 M3 data (Pieters et al., 2011, 2014; Prissel et al., 2014). This new rock type, dominated by Mg-rich spinel, contains < 5% pyroxene and olivine and may originate from magmatic intrusion into the lower crust (Pieters et al., 2011).
Recent spacecraft mission instruments have provided new perspectives on lunar rock and mineral distribution. A new global map of the purest anorthosites (PAN) was derived from Kaguya Spectral Profiler data (Yamamoto et al., 2012). According to lunar feldspathic meteorites and orbital data, Gross, Treiman, and Mercer (2014) indicated that the highland crust may be dominated by magnesian anorthosites and may be composed of less ferroan anorthosites. Nine potential candidates for magnesian suite exposures were suggested by some previous studies (Klima et al., 2011; Dhingra, Pieters, Boardman, Isaacson, & Taylor, 2011; Pieters et al., 2011; Shearer, Elardo, Petro, Borg, & McCubbin, 2015). Prissel et al. (2014) proposed that magnesian suite may be excavated on both the nearside and farside. Alkali suite may dominate the central area of the South Pole–Aitken Terrane, the periphery of the Procellarum KREEP Terrane, and some isolated regions (Wang & Zhao, 2017). The global distribution of cryptomaria (buried mare basalts) was documented by M3 data (Whitten et al., 2015).
There are still many unsolved problems (e.g., how to distinguish mare basalts from other mafic materials from the orbital data). Were some of Mg-rich rocks exposed by volcanism, as indicated by Prissel et al. (2016)? With regard to the Low-K Fra Mauro (LKFM) impact-melt breccias, a special and important rock type, what is their distribution on the Moon?
Lunar regolith is the fragmental soil layer covering almost the entire lunar surface and consisting of unconsolidated material. This layer was formed mainly by impact gardening, micrometeorite impact, and solar wind and cosmic ray-caused space weathering on preexisting bedrock (e.g., lava flows). The lunar regolith is important because it is the actual boundary layer between the solid Moon and its space environment and contains critical information about both of these regions. Based on in situ exploration (e.g., drilling by Apollo astronauts), grain-size distribution studies (i.e., Apollo and Luna samples), and crater shape models, the thickness of the regolith layer was estimated generally about 4–5 m in the mare areas but average about 10–15 m in older highland regions (Oberbeck & Quaide, 1968; Taylor, 1982; McKay, Fruland, & Heiken, 1974; McKay et al., 1991; Fa & Jin, 2010).
Returned lunar soil and drill core analyses and remote reflectance spectra are essential to study the properties of regolith. Reflectance spectra records many of the Moon’s characteristics such as composition, grain size, surface roughness, and state of maturity. The reflectance spectra of the Moon have been investigated with laboratory measurements, Earth-based observations, and spacecraft observations (Pieters et al., 2000; Kieffer & Stone, 2005; Velikodsky et al., 2011; Hillier et al., 1999; Besse et al., 2013; Wu et al., 2012). The Chang’e-3 mission accomplished the first in situ measurement of the Moon’s spectra using the Visible-Near Infrared Spectrometer (VNIS) onboard the Yutu Rover. These measurements extended the range of existing photometric geometry with large emission and phase angles. Compared to previous measurements, the absolute reflectance of the Moon was accurately characterized by the equipment of the onboard calibration panel. The Chang’e-3 landing site was suggested as a new calibration site (Wu, Wang, Cai, & Lu, 2018). Also, for the first time, the VNIS detected the microscale thermal characteristics of the lunar regolith, although with much shorter wavelength range than a typical thermal radiometer. The measured temperatures are 10 K higher than expected from theoretical models, indicating low thermal inertia of the lunar soil and the effects of grain facets on soil temperature at the submillimeter scale. Although the reflectance increases with wavelength, the emitted fraction also increases with wavelength (Wu & Hapke, 2018).
Space weathering is an important surface process occurring on the Moon and airless bodies. The optical effects of the Moon’s space weathering have been largely investigated in the laboratory for lunar samples and simulants (Noble et al., 2001, 2007). However, duplication of the real space environment and pristine regolith here on Earth is not possible. The Chang’e-3 in situ spectra provide the unique opportunity of investigating space weathering of the real lunar surface by measuring the regolith in its pristine state as well as comparison to the regolith naturally disturbed by rocket exhaust from the spacecraft. It revealed that the brightness increases measured after the spacecraft landed for all landing sites (i.e., Apollo, Luna, Surveyor) is due to the removal of the finest highly weathered particles by the lander’s rocket exhaust and not due to smoothing of the surface; this settled the long-standing debate. A new model of space weathering for the real lunar surface was derived from the Chang’e-3 in situ spectra: (a) The uppermost surficial regolith, perhaps several millimeters to tens of centimeters, is much more weathered than the regolith immediately below. (b) The finest fraction is much more mature than the coarser fraction (Wang et al., 2017).
In addition to extensive mineralogical and geochemical studies of Apollo lunar regolith samples, great advances have been made in the beginning of the 21st century for the distribution, thickness, and formation rate of lunar regolith. The Microwave Radiometers onboard Chang’e-1 and Chang’e-2 lunar missions have detected the global distribution and thickness of lunar regolith. The results indicate that the thickness in the maria varies from about 0.5 m to 12 m, and the mean is about 6.52 m, while the thickness in the highlands is a bit thicker than the previous estimation, where the thickness varies widely from 10 m to 31.5 m, and the mean thickness is about 16.8 m (Meng et al., 2014). The Chang’e-3 soft landing mission provides another chance to better understand the regolith properties of the northern Imbrium basin. Specifically, the ground-penetrating radar onboard the Yutu Rover has recovered the subsurface structures along its track. Four shallow interfaces were detected by the high-frequency radar and they are (a) top thin layer regolith (<1 m); (b) ejecta blanket materials from Ziwei crater (2–6 m); (c) paleoregolith (4–11 m); and (d) Eratosthenian bed rock (Fa, Zhu, Liu, & Plescia, 2015; Xiao et al., 2015). Up to six deep interfaces represented by paleoregolith formed between the eruption of Imbrian basalts and Eratosthenian basalts, interbedded thin basalts (or pyroclastic rocks), basalts, paleoregolith, and Imbrian basalts (Figure 15; Xiao et al., 2015; Yuan et al., 2017). It suggests that there were multiple volcanic eruption events in the late Imbrian and Eratosthenian periods. The thickness estimated by these studies also indicate that the production rate of regolith is faster than previously thought (Fa et al., 2015). The young volcanism studied by Chang’e-3 in situ exploration also was confirmed by others (Zhao et al., 2014; Xiao et al., 2015; Zhang et al., 2015). Mineral abundances were inferred from Chang’e-3 spectral reflectance (Zhang et al., 2015) and a new type of lunar basalts was proposed (Ling et al., 2015). On the basis of comparative studies with the previous Luna 17 landing site in the same unit several hundred kilometers away, a generally consistent understanding of regolith properties and geological features within the Imbrium basin was presented (Basilevisky et al., 2015).
Helium-3 (3He) was implanted by solar wind in the lunar regolith and is a potentially valuable resource because of its potential as a fusion fuel. The global inventory of 3He was estimated as being 6.6 × 108 kg, in which 3.7 × 108 kg was estimated for the lunar nearside and 2.9 × 108 kg was estimated for the lunar farside (Fa & Jin, 2010).
Volatiles and Water Ice
During the Apollo era, sample studies and Moon-forming models all suggested that there was no sign of water (H2O) in lunar rocks. The vesicles of mare basalts indicate abundant gas phase within the magma, and these vesicles were formed by degassing when they were erupted. Different from terrestrial lavas where the dominant volcanic gases are H2O and CO2, there is no water in lunar rocks and CO was the dominant phase (Taylor, 1991) in mare basalts. Studies on returned lunar samples did not find H2O-bearing minerals. Even for phosphates, the most volatile-bearing minerals (e.g., apatite and whitlockite), there is little hydrogen or H2O (Jolliff, Haskin, Colson, & Wadhwa, 1993). Thus, before the year of 2010, it was generally accepted that the Moon was very “dry.”
Because of a combination of new instruments flown on a number of orbiting spacecraft missions (Chandrayaan-1 of ISRO, LCROSS/LRO of NASA) and advances in analytical technology used in studies of lunar samples, traces of H2O were reported by several groups (Sunshine et al., 2009; Pieters et al., 2009; Boyce et al., 2010; Schultz et al., 2010; Paige et al., 2010), and the detection of water (either as OH or H2O) or water ice on the Moon became one of the most exciting recent developments in the field of lunar science (Arand, 2010). The water-bearing pyroclastic glasses provided the most fundamental sample evidence (e.g., Saal et al., 2008) and water-bearing minerals were found not only in phosphates, but also in highland plagioclase (Hui, Peslier, Zhang, & Clive, 2013).
Evidence from Diviner instrument thermal data, Lunar Prospector investigations, LCROSS impact results, and the Moon Mineralogy Mapper (Li et al., 2018) has been found for the presence of current water ice in polar permanently shadowed regions, thought to be the result of solar wind, comets, or water from the interior, erupted from explosive (pyroclastic) volcanic eruptions (Mitrofanov et al., 2010; Spudis et al., 2013). However, there is still a debate about how wet the mantle of the Moon is (Sharp, Shearer, McKeegan, Barnes, & Wang, 2010). The presence of preserved polar volatiles is both a fundamental scientific record of the evolution of planetary volatiles and a potential resource for future human exploration.
The Geological History of the Moon
The lunar geological time scale has been divided into five periods (pre-Nectarian, Nectarian, Imbrian, Eratosthenian, and Copernican), with one of these (the Imbrian) being subdivided into two epochs.
Table 2. A Summary of Lunar Stratigraphy and Geological History
Age (Billion Years)
1.0 to present
Includes fresh ray craters beginning with the formation of Copernicus.
From the youngest lavas to craters without visible rays (e.g., Eratosthenes).
From the formation of Imbrium Basin to the youngest dated mare lavas from returned Apollo samples. Included are Imbrium Basin deposits, Orientale and Schrödinger multiring basins, most visible maria, etc.
Extends from the formation of Nectaris Basin to the Imbrium Basin, consisting of 12 large basins and some buried maria.
Defined from the point at which the lunar crust formed to the time of the Nectaris impact event. Nectaris is a multiring impact basin that formed on the nearside of the Moon, and its ejecta blanket serves as a useful stratigraphic marker. Includes 30 identified multiring basins.
These divisions of geological time are based on the recognition of convenient geomorphological markers and as such, they should not be taken to imply that any fundamental changes in geological processes have occurred at these boundaries. The Moon is unique in the solar system in that it is the only body (other than Earth) for which we possess rock samples with a known geological context. By correlating the ages of samples obtained from the Apollo and Luna missions to known geological units, it has been possible to assign absolute ages to some of these geological periods. The timeline in Table 2 represents one such attempt, but it is important to note (see “The Moon as a Cornerstone in Solar System History”) that some of the ages are either uncertain or disputed. In many lunar highland regions, it is not possible to distinguish between Nectarian and pre-Nectarian materials, and these deposits are sometimes simply labeled pre-Imbrian.
The Moon as a Cornerstone in Solar System History
In summary, the nature of geological processes operating on the Earth’s Moon provide important insight into Earth’s early history and constitute a framework and an interpretative cornerstone for understanding the early history of the terrestrial planets, Mercury, Venus, and Mars (Figure 1). It has been learned that the Moon originated from an impact of a Mars-sized object into early Earth and accreted from the ejecta from that massive event (Canup & Asphaug, 2001). In the terminal stages of this accretion, the outer part of the Moon was melted to create a molten rock ocean (“magma ocean”) that differentiated into plagioclase, which was buoyant and concentrated in the crust to create the primary lunar crust, and denser residual material that became the lower crust and mantle. This global anorthositic crust set the stage for the further evolution of the Moon. Material below the crust was denser than the underlying mantle and foundered into the deeper interior, carrying radioactive elements, and setting the stage for the generation of mare basalts.
The newly created global crust was immediately subjected to impact bombardment at all scales, ranging up to the huge multiringed basins, in excess of 1,000 km. Impact processes caused excavation of deeper crustal and mantle material, redistribution over areas often in excess of a lunar hemisphere, and the mixing of rock and soil material to create a megaregolith 1 km thick or perhaps multi-kilograms thick, a fragmental debris layer. Rock was instantaneously melted to create impact-melts on the floor of impact basins, and huge circumferential mountain ranges were created. Early basins formed in a thin lithosphere and the major topographic features relaxed; later basins such as Imbrium and Orientale formed in thicker lithospheres and retained their topography.
Following the period of late heavy bombardment, melting in the mantle generated magma that would rise to erupt on the surface to form the mare basalt deposits, preferentially on the lunar nearside. This melting began prior to the last of the major impact basins, creating cryptomaria, the early mare deposits covered by crater and basin ejecta. The period of mare basalt emplacement extended from about 4 billion years ago to perhaps as recently as 1 billion years ago, with the mare emplacement peaking during the period of about 3.0–3.8 billion years ago. The exact end of mare volcanism is uncertain, with some saying it has extended to near the present time. The youngest basalts, as indicated by crater size-frequency distribution (CSFD) analyses, are currently a primary target for proposed sample return missions.
Because the Moon is characterized by a large surface area-to-volume ratio, the Moon cooled down relatively rapidly, thickening the outer thermal boundary layer (the lithosphere) such that it could not subduct and produce plate tectonics. The Moon is thus known as a one-plate planet, characterized by a single rigid global lithospheric plate that is stable over geological time, thickens as the Moon loses heat, and preserves the record of events in early planetary history. Tectonics on the Moon are largely vertical, not lateral as with Earth plate tectonics. The geologically recent tectonics are largely driven by radial contractional stresses owing to interior cooling and tidal stresses (Watters et al., 2015), allowing that many of the older features resulted from loading of thick mare basaltic basin-filling deposits. Loading of the lithosphere by mare basalts, particularly in the impact basin interiors, caused flexing of the lithosphere and formation of extensional tectonic features, such as concentric arcuate and linear graben (rilles). In early lunar history, the Moon was characterized by a global extensional state of stress in the lithosphere, favoring formation of tectonic rilles and graben. As the Moon cooled and the lithosphere thickened, the global state of stress in the lithosphere became increasingly contractional, favoring crustal shortening and the formation of tectonic wrinkle ridges and lobate scarps in mare areas.
Outstanding Problems and Future Exploration
There are still many unanswered fundamental questions about lunar geology. The nature and origin of dichotomy reflected in the surface and interior is still unknown. Lunar impact history has not been well constrained, particularly in the first 500 Ma (Morbidelli et al, 2018) and the Eratosthenian and Copernican periods. Dating of impact melts of major impact events could calibrate the crater geochronology model throughout lunar history and assist immensely in determining broader planetary chronology. Recent observations and interpretations suggested that there may be young mare basalts and volcanic features (e.g., irregular mare patches) that might be less than 100 Ma in age. Determination of these ages is critically important for understanding lunar volcanism and thermal history. The nature of the lunar magnetic field and a dynamo is poorly understood. Layering subsurface structure in the mare area suggested multiple volcanism and complex geological history need more studies. New rock types and mantle composition require more investigation to confirm their properties. Little new seismic experiment data have been acquired for decades; this limits our understanding of the lunar deep interior structure. GRAIL data revealed many questions but also raised more questions about the formation mechanism of the heterogeneous lunar crust. Composition, distribution, and origin of volatiles, including polar region deposits and pyroclastic rocks, will be one of the major lunar science questions for the coming decades. To answer these remaining questions, future robotic and human missions, particularly sample return, are critical. As a cornerstone, the advance of lunar geology will greatly promote knowledge about all terrestrial planets.
There is currently intense interest in both national and commercial efforts to undertake science-focused lunar missions. For example, the Chinese Chang’e-5 mission plans to collect samples from young basalts in the Rümker area and return them to Earth for analysis (Wang & Xiao, 2017; Zhao, Xiao, Qiao, Glotch, & Huang, 2017; Jolliff, Wang, Ling, & Zhang, 2017; Xiao, 2018; Qian et al., 2018; CNR, 2019). Once isotopic chronological data are obtained, the current lunar chronology curve can be refined and a better understanding can be obtained of the volcanic history and thermal evolution of the Moon. Chang’e-4 spacecraft has landed and is exploring the Von Karman crater (Figure 23), which is located within the SPA basin on the farside of the Moon. On-site geological surveys by imagery, composition, and subsurface radar detection would open a new window to look into the unique aspects of the SPA impact basin and the geology of the lunar farside (Wu et al., 2017; Huang et al., 2018; Xiao et al., 2019). The polar regions are rich in science interest and volatile resources and have attracted much attention from many countries and space agencies (e.g., Europe, United States, Japan, Russia, Korea, China, and India). Several missions are planned for the near future to explore ice deposits to determine the presence of water and other volatiles within the permanently shadowed areas. Detailed studies of the composition and origin of these volatiles will greatly help in understanding fundamental lunar and planetary processes.
The authors thank review comments from the editor and anonymous reviewers. These greatly helped them to improve the quality of this article. They also thank Z. Y. Xiao, Y. Z. Wu, Q. Huang, S. Y. Zhao, Z. Q. Xue, and S. R. Yu for helping to collect data. This study was supported by the Natural Science Foundation of China (Grants 41830214, 41772050), the Key Research Program of the Chinese Academy of Sciences (Grant No. XDPB11), and the Science and Technology Development Fund (FDCT) of Macau (Grant 121/2017/A3).
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