Moonquakes - an overview | ScienceDirect Topics

III.C Seismicity
A large number of very small moonquakes have been detected by the Apollo seismic network. The total seismic energy release within the moon appears to be about 80 times less than that in the earth. The moonquakes are concentrated at great depth—between 600 km and 1000 km—which is deeper than earthquakes. The difference in distribution and magnitude of seismic energy relase on the earth and the moon is probably related to fundamental differences in interior dynamics. Correlation of deep moonquakes with lunar tides indicates that tidal stresses must play an important role in triggering them.
3.2 Seismology
The lunar seismic signals have a large degree of wave scattering and a very low attenuation so that during moonquakes the Moon “rings like a bell” owing to the absence of water and the very fractured nature of the upper few hundred meters. Observed moonquakes have been mostly less than 3 on the Richter scale; the largest recorded ones have a magnitude between 5 and 5.7. Many are repetitive and reoccur at fixed phases of the lunar tidal cycle. The Apollo seismometers recorded the impacts of 11 meteorites with masses of more than one ton. The Moon is seismically inert compared to the Earth, and tidal energy is the main driving force for the weak lunar seismic events.
10.03.3.2 External Seismic Activity: Artifical and Natural Impacts
Impacts constitute about one-fifth of the detected events (1742 of the 9315 identified and classified events) on the Apollo seismic network, and the strongest ones have amplitudes comparable to the largest moonquakes. As discussed in Section 10.03.2, these surface seismic sources permit more detailed studies of crustal seismic structure, including lateral variations in thickness (Chenet et al., 2006). A meteorite impact can be modeled as a point force seismic source, with an amplitude equal to the momentum of the impactor (i.e., mass times impact velocity, in kg m s−1 or N s). During the Apollo seismic experiment, controlled impacts of the Saturn V upper stages and the Lunar Modules were performed, and it may be possible to use these to calibrate the energy flux from natural impacts. Signals from the impact of the Apollo 17 upper stage are shown in Figure 15. The momentum of this impact was about 37 000 kg km s−1 (14 487 kg at 2.55 km s−1), and knowledge of its position and occurrence time, provides unique seismic data with well-constrained P- and S-wave arrival times.
Figure 15. Seismic records, from the Apollo seismic network, of the impact of the Apollo 17 Saturn V upper stage (Saturn IVB) on the Moon on 10 Dec 1972 at 4.21 °S, 12.31 °W (at distances of 338, 157, 1032, and 850 km from the Apollo 12, 14, 15, and 16 stations, respectively). X, Y, Z are for the long-period seismometers, z is for the short-period seismometer. Bit sensitivity is 4 × 10−10 m s−2 for the LP (at 0.45 Hz) and 1.3 × 10−7 m s−2 for the SP (at 8 Hz). Amplitudes at the Apollo 14 station, 157 km from impact, reach about 10−5 m s−2 with an amplitude mainly related to S waves trapped in the regolith. The first P arrival is typically 10 times smaller. Note the 10 db gain change at the middle of the fourth and fifth traces. Reprinted from Lognonné P (2005) Planetary seismology. Annual Review of Earth and Planetary Sciences 33: 19.1–19.34. (doi:10.1146annurev.earth.33.092203.122605).
Even if the locations of natural impactors are not known, they are still useful seismic sources. In the case of the Moon and Earth, the frequency and size of impactors can be obtained using statistical properties of the Earth-crossing asteroids (ECAs, e.g., Poveda et al., 1999) and of the collision rates (e.g., Shoemaker et al., 1990). Figure 16 shows the frequency distribution of impactors as a function of seismic impulse, for the lunar cross-section. Even when uncorrected for geometrical spreading of seismic waves, the frequency of the detected impacts with respect to the amplitude of the seismic signal (extracted from Figure 1), matches the distribution of impactors fairly well for the largest impacts (note that only the distribution in impactor diameter is taken into account). The amplitude of the largest impacts was estimated by Gagnepain-Beyneix et al. (2006) and typically releases a seismic momentum of about 3 × 108 N s. This is a factor of 10 larger than the Apollo Saturn V upper-stage impact, and a factor 100 larger than the Lunar Module impact. We have used this value to anchor the observed impact curve on Figure 16.
Figure 16. Impact statistics for Earth-crossing asteroids on the Moon. The distribution of impactors is from Poveda et al. (1999) and the collision probability is assumed to be 4.2 × 10−9 per year (Shoemaker et al., 1990). Lunar case (green dotted line); impactor detection curves at the Apollo 14 station (purple circles and line), with the largest event anchored to a seismic moment of 3 × 108 N s; Mars (red dot-dashed) and Earth (blue solid) curves, with atmospheric effects included. The latter leads to a reduction of both the mass and velocity of the impactor. Pale blue stars are the seismic impulses, in increasing impulse, of the LEM, the Saturn V, and the largest natural impacts, respectively. The natural impact corresponds to typical impacts expected on Near-Earth objects (NEOs) missions. Note that event frequency scales with the squared ratio of the NEO radius (crossing diameter) and the Moon’s radius. At 1 AU, the number of events, compared to the Moon, are reduced by about 120 000 for a 5 km radius NEO. Entry velocity is taken as 20 km s−1, for impactor mass of 2400 kg m−3.
On planets with an atmosphere, both the velocity and the mass of the impactors are reduced during atmospheric entry. This effect can be assessed by integrating the impactor equation in the atmosphere (Chyba, 1993; Poveda et al., 1999):
[6]mdvdt=12CDρv2A−mgdmdt=12CHQρv3A
where m is the mass of the meteoroid, v its velocity, ρ its density, g the gravity, CD the drag coefficient (here taken as 1.7, the cylinder value), CH the heat transfer coefficient (taken as 0.1), Q the heat of ablation of the meteoroid (taken as 107 J kg−1), A its cross section, and t is time. Figure 16 gives the effect of the Mars and Earth atmospheres on the ECAs or Moon-impacting meteoroids, for an asteroid density of 2400 kg m−3, an entry velocity of 20 km s−1, and standard atmospheric models (US standard atmosphere 1976) of the planets. We see that in contrast to Earth, the atmospheric shielding effect on Mars is small. Shoemaker et al. (1990) and Davis (1993) have proposed that the entry flux of meteorites entering the atmosphere of Mars is 2.6 times that on Earth due to the proximity of the asteroid belt. Based on these data, the number of impacts per year detected by a seismic station on Mars is comparable to that on the Moon (Davis, 1993), assuming an instrument detection level similar to that of the Apollo seismometers.
Lastly, we consider small bodies. Currently the experiment SESAME/CASSE is onboard the Rosetta mission to comet 67P/Churyumov-Gerasimenko (Kochan et al., 2000), for a landing expected in 2014. This experiment aims to investigate the outermost surface of the comet by means of active and passive acoustic-wave monitoring in a frequency range from a few hundred to several kilohertz, and may reopen the study of small-bodies seismology after the failure of acoustic sensors onboard the Phobos 1 and 2 landers (Surkov, 1990). Both internal (related to degassing events) and external sources might generate seismic signals on comets, in contrary to asteroids, for which only impacts are expected. Future missions may be targeted toward asteroids, either to study the interior or as a deflection strategy in asteroid mitigation (Ball et al., 2003). Both active seismology (using explosives or impactors) or passive seismology (with natural impacts) have been already proposed (e.g., Walker and Huebner, 2004). Due to their limited size, these bodies have low gravity. As a consequence, minor impacts on asteroids produce ground accelerations higher than the local gravity, and seismic reverberation may have a major effect on the morphology of the asteroid regolith, including in the shape and density of small craters (Richardson et al., 2004). The amplitudes of the body P and S seismic waves generated by impacts can be estimated with simple models. Scattering processes will affect body-wave amplitudes, and the irregular asteroid surface will strongly affect the amplitudes of surface waves.
For large impacts, such as the artificial impacts made in the Deep Impact mission (3.8 × 106 N s by the impact of 370 kg at 10.2 km s−1) (A’Hearn et al., 2005) or those expected in the Don Quichotte mission (4 × 106 N s by 400 kg at 10 km s−1) (e.g., Ball et al., 2003), the displacement amplitude of body waves can be roughly expressed by
[7]ap=Ft−r/cp4πρcp2randas=Ft−r/cs4πρcs2r
where cp and cs are P and S velocities, ρ is bulk density, and r is radial distance (Aki and Richards, 1980) and F are the point force source functions (in N). By scaling the impact probability of Figure 16 by the ratio of the surface body areas, we see that these impacts, on a 5-km radius asteroid, will occur naturally at a rate of about one event every 1000–1500 years for an ECA flux, and will generate accelerations in the range of 0.01–1 m s−2 depending on frequency (Figure 17). As the frequency of impacts is typically a function of the diameter with a power of −5/2 (e.g., Poveda et al., 1999), and therefore of the mass with a power of −5/6, accelerations 106 times smaller (in the range of 1–100 ng) will occur at a rate of 105 events a year at 1 AU, and possibly higher at larger distances from the Sun. This translates into rates of hundreds of events per year that could possibly be recorded by space-qualified short-period seismometers with nanogram sensitivities. Such natural impacts might provide excellent seismic sources for tomography studies of asteroids.
Figure 17. Direct P and S waves amplitude vs distance estimates of the P and S waves for an impact of 400 kg at 10 km s−1 for different frequency bands (a) and diffraction/attenuation effects (b). Velocities are 3670 and 2100 m s−1, respectively and densities 2600 kg m−3. A Q of 1000 is assumed in (a). Figure redrawn after Lognonné 2005.
III.B.2 Other Experiments
Although images are the most spectacular data returned by spacecraft, a whole array of equally valuable experiments are included in each scientific mission. For example, a gamma ray spectrometer aboard the lunar orbiters was able to map the abundance of iron and titanium across the Moon's surface. Seismometers placed on the Moon recorded waves from small moonquakes. Measurements of heat flow from the interior of the Moon enabled scientists to understand something about its composition and present evolutionary state. Onboard the Voyager spacecraft were several experiments which are valuable for satellite investigations, including an infrared spectrometer capable of mapping temperatures; an ultraviolet spectrometer; a photopolarimeter, which simultaneously measured the color, intensity, and polarization of light; and a radio science experiment that was able to measure the pressure of Titan's atmosphere by observing how radio waves passing through it were attenuated. The Galileo magnetometer obtained crucial evidence for subsurface liquid water on Europa and perhaps Ganymede and Callisto.
4.8.3.2 Moon quakes
The stations of the Apollo lunar seismic network, installed on the Moon between 1969 and 1972, recorded the only confirmed seismic events on any extraterrestrial body so far (e.g., Lognonné, 2005; Lognonné and Johnson, 2007). The large number of events recorded on the Moon, which had been considered as tectonically “dead,” came as a big surprise at that time (e.g., Frohlich and Nakamura, 2009). The waveforms recorded on the Moon were readily recognized to be very different from what had been expected based on Earth data (Gold and Soter, 1970; Dainty and Toksöz, 1981). Lunar event seismograms have emergent onsets, a strong body wave coda, and a lack of coherent, dispersive surface-wave trains. In contrast to classical terrestrial signals, the lunar events are characterized by longer event durations of up to 100 min (Fig. 4.24). Source processes of these events probably differ from those on the Earth, where most quakes is related to plate tectonics.
Figure 4.24. Typical lunar seismograms for different event types. The plotted events were used as prototypes in the event classifier. (A) Shallow moonquake. (B) Deep moonquake. (C) Impact. Note the different amplitude scale of the plots as given by the black bars. Black parts of LPY and SPZ components indicate the time intervals used to learn the features of the respective events for the classifier. Traces signed as “LPX (Y, Z)” were recorded with a long-period seismometer having an eigenperiod of 15 s “SPZ” is a vertical short period record (eigenperiod of the sensor 1 s).
From Knapmeyer-Endrun and Hammer (2015).
The Apollo seismic recordings are a valuable test case for any future experiment in planetary seismology and algorithms that could be employed in it. The main lunar event types, listed in the long-period event catalog (Nakamura et al., 1981), are given by deep moonquakes, impacts, and shallow moonquakes. Deep moonquakes, occurring in a depth range between 700 and 1200 km (Nakamura, 2005), make up more than 55% of all identified events and form spatial clusters of periodic activity (e.g., Lammlein et al., 1974; Lammlein, 1977; Nakamura, 2003; Bulow et al., 2007). The established periodicities of 1 month, 7 months, and 6 years correlate with the Moon's position with respect to the Sun–Earth–Moon system, indicating variations in tidal stress as cause of deep moonquakes (Lammlein, 1977; Nakamura, 1978; Minshull and Goulty, 1988; Weber et al., 2010). Frohlich and Nakamura (2009) have argued that fluid phases or partial melts in the lunar mantle could additionally play an important role, similar to the case of intermediate-depth subduction-zone seismicity on the Earth, as deep moonquakes occur at depths where brittle failure is impossible in dry rock. The second largest group of lunar seismic events, around 13% of all identified events, are generated by impacts, both of natural meteoroids and of artificial sources. Artificial sources are spent stages of the Saturn-V launch vehicles and lunar modules that were directed at the lunar surface on purpose to create seismic sources of well-defined timing and location. The natural impacts, especially of meteoroids with masses smaller than 1 kg, show temporal clustering correlating with known meteor showers (Nakamura et al., 1982; Oberst and Nakamura, 1991). In contrast to deep moonquakes, shallow moonquakes with focal depths of less than 200 km are rare events, making up less than 1% of all identified events. They are significantly larger than the deep moonquakes (Mw of up to 4.7), have a low b value (Nakamura, 1977), and are sometimes also associated with exceptionally large stress drops (Oberst, 1987). As tectonic quakes on a stagnant-lid planet, they might be comparable to intraplate earthquakes, which exhibit some similar properties (Nakamura et al., 1982).
Knapmeyer-Endrun and Hammer (2015) identified and classified specific seismic event types in the Apollo 16 lunar seismic data by using an HMM based classification system. The original event catalog was composed in 1981 by Nakamura et al. (1981) and contained 13,058 events. More than a half of the events was classified as deep moonquakes while shallow moonquakes and impacts represent only a small part of detected events. In contrast to classical terrestrial signals, the lunar events are characterized by longer event durations of up to 100 min. However, not all events, listed in the catalog, are assigned to a specific seismic signal class. Ranging from two third in the initial catalog, this number could be reduced to about 25% through intense multiple analyses throughout the last 25 years (Knapmeyer and Weber, 2015). To test the proposed “learning-while-recording” approach also for nonterrestrial signals, Knapmeyer-Endrun and Hammer (2015) first focus on matching the existing event labels. The automatic system classified more than 80% of the events correctly (i.e., according to the manually given label). As a next step, the previously unclassified signals were fed into the classifier and about 60% could be assigned with a reliable class label. Finally, it was possible to find 210 new events in the continuous raw data which were not recognized before. The additional events can be used for a more detailed study of temporal variations in event occurrence or to improve the cluster stacks for phase picking.
6.1 Seismicity
To date, the Moon is the only body besides the Earth on which extensive seismological experiments have been done. The Apollo 12, 14, 15, and 16 astronauts installed a network of four three-component, long-period seismometers on the Moon, which detected more than 13,000 seismic events. Although the Apollo seismometers were highly sensitive state-of-the-art instruments and recorded lunar seismic events for almost 8 years, the relatively small network aperture and their restriction to only the nearside limited both the detection capability and the accuracy with which seismic sources could be located. Unfortunately, such location uncertainties are propagated into uncertainties of the interior structure. Nevertheless, the Apollo data revealed that lunar seismicity is different from that of the Earth due to the absence of tectonic plate movement. Sources of seismic activity on the Moon are related to deep moonquakes, shallow moonquakes, and meteoroid impacts.
Deep moonquakes occur in so-called “nests”, which repeatedly release seismic energy in specific cycles correlated with lunar tides (Figure 23.6). More than 7000 individual deep moonquakes were recorded, originating from more than 250 of these nests, which are situated at depths between 700 and 1100 km. From the limited data available, neither the spatial nor the temporal patterns of deep lunar quakes are fully understood. However, it is apparent that the activity of these nests correlates with the tidal cycles caused by the Earth and the Sun. The shallow lunar quakes are even more enigmatic. Compared to the deep moonquakes, they are relatively rare; during almost 8 years of continuous data collection, only 28 such events were recorded and their epicenters were not correlated with known geologic or geographic features. One interpretation is that these moonquakes were formed by the release of thermoelastic stresses in the lunar crust or the upper mantle, possibly connected to large impact basins.
While early seismic studies favored crustal thicknesses of about 60 km beneath the Apollo network, more recent studies favor thinner crusts of only 30–45 km. The mean thickness of the lunar crust is estimated at 49 ± 15 km if an Airy-type compensation mechanism applies. These crustal thicknesses are in agreement with estimates of 34–43 km on the basis of GRAIL gravity data and assumed mantle densities of 3220 kg/m3 and 3150 kg/m3.
For the upper mantle, the Apollo seismic data indicate almost constant seismic velocities, whereas they are more variable in the underlying middle mantle. At about 560 km below the surface of the Moon, the seismic data suggest a discontinuity in the mantle, which might be related to a compositional change to more aluminous and MgO-rich mafic silicates. If this discontinuity exists globally, it may represent the initial depth of melting during the magma ocean phase of the Moon. If it only exists on the nearside, it can be interpreted as the maximum depth of melting of the mare basalt source region beneath the Procellarum KREEP Terrane (PKT). In the lower mantle below a depth of about 1150 km, seismic waves are strongly attenuated, indicating a partially molten zone extending to the core–mantle boundary.
With few exceptions, almost all recorded seismic events within the Moon occurred on the nearside, a dichotomy that might be either observational bias or related to geology. Because of the depth distribution of the deep moonquakes, P and S waves do not sample depths below 1200 km and thus the core region remains poorly understood. Among the ∼1700 meteoroid impacts detected by Apollo seismometers, only one was at a distance great enough to allow for investigations of the deepest mantle. This observation constrains the lunar core to have a radius between 170 and 360 km.
3.7 Seismology
The Apollo astronauts deployed four seismometers on the lunar surface between 1969 and 1972 (Figure 24.7). These instruments gathered data continuously until 1977, making the Moon the only extraterrestrial body for which extensive seismic data has been gathered.
FIGURE 24.7. (Left) Map of the lunar nearside (15° increments of latitude and longitude) showing the locations of the four seismometers that comprised the Apollo Passive Seismic Experiment (red diamonds) and the nominal epicenter locations of the deep moonquake source regions (white rectangles). Farside source regions (outer black circles) are projected on the nearside. The deep moonquakes are roughly constrained to a wide swath trending northeast–southwest across the entire nearside of the Moon. (Right) Photograph of the Apollo 16 seismometer as installed on the lunar surface. It is covered with a mylar shield that was intended to thermally protect the instrument. Other instruments from the Apollo Lunar Surface Experiments Package can be seen in the background, as well as the central station that telemetered data continuously to Earth.
Photo AS16-113-18,347 from the NASA Apollo Archive.
The Moon is much less seismically active than the Earth, due to its lack of oceans and plate tectonics. Still, the Apollo network recorded several types of both naturally occurring and artificial seismic events, resulting in a total number of approximately 13,000 catalogued events over the 8-year span of the experiment. Because the Moon has no atmosphere to burn off potential impactors, there were a significant number of meteoroid impacts on the surface that registered on the seismic array. The booster rockets and lunar modules from the Apollo spacecraft were also purposely impacted onto the surface, in part to test and calibrate the seismic array. Observed naturally occurring moonquakes include the relatively large but rare shallow moonquakes of unknown origin (similar to intraplate earthquakes), and the relatively small but frequent deep moonquakes, (triggered with monthly periodicity by the lunar tides). Observed deep moonquakes generally had body wave equivalent magnitudes less than three, with most less than magnitude one; shallow moonquakes were larger, with the larger recorded events having moment magnitudes ranging between 2.4 and 4.1. In addition, the network detected many noiselike thermal events that were associated with the large temperature fluctuations between lunar day and night.
Deep moonquakes are the most numerous type of seismic event, comprising approximately half of the event catalog. They are known to originate from distinct source regions located in a wide swath across the nearside, at depths between approximately 700 and 1200 km (Figure 24.7). Events from a single source are periodic at monthly (tidal) periods, and exhibit high degrees of waveform similarity, likely representing repeated failure on existing fault structures at depth.
Compared to terrestrial seismograms, lunar seismic signals exhibit characteristics typical of a large degree of wave scattering and very low attenuation, due in part to the fractured nature of the upper few hundred meters of lunar regolith. During events, seismic energy reverberates within the Moon, resulting in recorded signals of extremely long duration, sometimes an hour or more, with P- and S-wave codas that mask secondary arrivals. The small number of seismic stations and lack of high-quality seismic events limited the types of analyses that could be performed. Despite these limitations, data from the Apollo passive seismic network have been extensively analyzed to reveal details on the Moon's internal structure, confirming the presence of separate crust, mantle, and core layers (Figure 24.8).
FIGURE 24.8. Schematic cross-section of the Moon showing the crust, mantle, and core layers inferred through analyses of the Apollo seismic data. The Apollo station locations and seismically active regions are indicated; arrowheads show a sample of meteoroid impact locations.
Figure reproduced from Nakamura, Latham, and Dorman (1982).
The detailed structure of the upper kilometer of the lunar crust was determined by two additional seismic experiments: the Active Seismic Experiment on Apollo 14 and 16, and the Lunar Seismic Profiling Experiment on Apollo 17. In both experiments, the astronauts detonated a series of small explosives on the lunar surface (Figure 24.9). A network of geophones then recorded the ground motions generated by these explosions. On Apollo 14 and 16, up to 19 explosions were detonated by an astronaut using a “thumper” device along a 90-m-long geophone line. Additionally, on Apollo 16, three mortar shells were used to launch explosive charges to distances of up to 900 m from the ALSEP. On Apollo 17, the astronauts were able to position eight explosive charges at distances of up to 3.5 km from the Lunar Module, with the assistance of the Lunar Roving Vehicle. Both the Apollo 16 mortar shells and the Apollo 17 explosives were detonated by radio control after the astronauts left the lunar surface.
FIGURE 24.9. (Left) The Apollo 16 mortar package mounted on its base. The cable running off to the left connects the experiment to the ALSEP Central Station. The red flag at the top of the mast provided a visual warning for the crew to steer clear when driving the Lunar Rover (photo AS16-113-18,378 from the NASA Apollo Archive). (Right) Apollo Lunar Module Pilot Edgar Mitchell walks along the geophone line during the Apollo 14 mission, operating the “thumper”.
Photo AS14-67-9374 from the NASA Apollo Archive.
These experiments showed that the seismic P-wave velocity is between 0.1 and 0.3 km/s in the upper few hundred meters of the crust at all three landing sites. These velocities are much lower than observed for intact rock on Earth, but are consistent with a highly fractured material produced by the prolonged meteoritic bombardment of the Moon. At the Apollo 17 landing site, the surface basalt layer was determined to have a thickness of 1.4 km.
The lower crust and mantle seismic velocities have been estimated using the classical nonlinear inversion of compressional and shear wave arrival time readings made from the Apollo seismograms (Table 24.2). Shallow structure is constrained largely using surface and near-surface events (impacts and shallow moonquakes), while deep structure is constrained using midmantle and deeper events (deep moonquakes). Seismic velocities increase steadily down to approximately 20 km. At that depth, there is a change in velocities within the crust that probably represents the depth to which extensive fracturing, due to massive impacts, has occurred (see Figure 24.11 and Section 4.1). At an earlier stage of seismic data analysis, this velocity change was thought to represent the base of the mare basalts, but these are now known to be much thinner. The main section of unbroken crust has rather uniform velocities that correspond to the velocities expected from the average anorthositic composition of the lunar samples.
FIGURE 24.11. Schematic cross-section of the upper 25 km of the lunar surface, illustrating the effects of large-scale cratering on the structure of the lunar crust.
Figure reproduced from the Lunar Source Book.
Very few seismic rays detected by Apollo traverse the region below the deep moonquake zone. Evidence for a highly attenuating region in the deep interior such as a layer of partial melt or a fluid lunar core is implied in part by a lack of observation of seismic signals originating from the farside of the Moon (Figure 24.10). Since deep moonquakes are generally small, their energy cannot penetrate the attenuating region to reach the nearside Apollo array. An additional interpretation of the lack of farside signals is that the farside is aseismic, which, given the other global nearside/farside asymmetries (e.g. crustal thickness and mare distribution), is not outside the realm of possibility. Further seismic exploration is needed to resolve this issue.
FIGURE 24.10. Schematic cross-section of the Moon showing the bounding seismic rays for shear wave shadows from the deep attenuating region. The direction to Earth is indicated. Small squares denote the Apollo seismic station locations (Apollo 12 and 14 are approximately colocated, so only one square is shown for clarity). The crosses mark the locations of hypothetical deep moonquake hypocenters for which the seismic rays to a station at a corner of the triangular station network graze the lower mantle, in which shear waves are severely attenuated. For events in Zone A, clear shear wave arrivals may be observed at all three corners of the network, i.e. none of the seismic stations are in the S-wave shadow. For events in Zone B, clear shear wave arrivals may be observable at two corners of the network, i.e. one corner of the array is in the S-wave shadow. For events in Zone C, clear shear wave arrivals may be observable only at one corner of the network, i.e. two corners of the array are in the S-wave shadow. For events in Zone D, no clear S-wave arrivals are observable at all three corners of the network, i.e. all corners of the array are in the S-wave shadow.
5.27.4.1 Trigger of mass movements on the Moon
The three major triggers of terrestrial landslides are water, volcanic and tectonic activity, and saturation of water is the most important one. On the Moon, forces that could cause deformation and failure in slope materials include seismic activity induced by impact cratering and endogenic geodynamic processes (e.g., volcanism, tectonism, tides and body shrinkage) and thermal stresses raised by diurnal temperature changes.
Thousands of seismic events were recorded by the Apollo passive seismic networks, and most of them are sourced from > 700 km in the lunar interior due to tides (Nakamura et al., 1979). The dozens of shallow moonquakes are possibly formed by impact cratering and tectonic activity. The widespread small and fresh-appearing tectonic features on the Moon are believed to be young and possibly active (Fig. 1G). In theory, moonquakes generated by the activity of these small structures can be on the same magnitude with the 1906 San Francisco earthquake (Nahm and Velasco, 2013). Elastic waves have a much slower decaying rate on the Moon due to the lack of pore water, the Copernican-aged tectonic features could theoretically release a large amount of stored energy over the entire Moon. Recently, by relocating the possible epicenter for a shallow moonquake recorded by the Apollo seismometers, the possible triggering fault, which is located beneath young compressional structures is revealed (Watters et al., 2019; Senthil Kumar et al., 2019). Interestingly, clusters of rock falls are observed near the lobate scarp, and the falls were interpreted to be caused by the recent fault movement (Senthil Kumar et al., 2019).
Impact generated seismic activity can affect areas much larger than the source craters. For example, formation of the < 100 km diameter Copernicus crater on the Moon could trigger ground deformation of ∼ 1–10 m at three radii from the final crater rim (Schultz and Gault, 1976). On the basis of the kilometer-scale topographic roughness across the Moon (Fig. 2A and B), Kreslavsky and Head (2012) noticed a sharp difference in the occurrences of steep slopes between the Late Imbrian and Early Imbrian craters, as terrains older than the Early Imbrian are significantly muted in topography. It was argued that basin-forming impacts can cause global resurfacing by strong seismic shaking that induce ground deformations up to hundreds of meters, so intensive mass movements occurred globally due to the basin impact (Kreslavsky and Head, 2012). The Orientale basin-forming event, which separates the Early and Late Imbrian epochs, was the last event that induced global mass movements at kilometer scales (Kreslavsky and Head, 2012). This explanation predicted that numerous such global resurfacing events have occurred on the Moon by impact basins older than Orientale, and observed slopes larger than ∼ 32°–35° on the Moon are all formed by younger craters (Kreslavsky et al., 2013).
The rate of diurnal temperature change is less than 2 K/min on the Moon (Molaro and Byrne, 2012), but grain-scale stresses induced by diurnal temperature changes are on the order of 100 MPa (Molaro et al., 2015). This magnitude of stress is comparable with typical strength of lunar rocks. Thirumalai and Demou (1973) performed physical experiments on lunar analogue samples and found that thermal expansion in vacuum was the same as that in atmosphere, and that the tested samples displayed permanent and irrecoverable strains, which were caused by grain reorientation from the differential expansion. Numerical modelling of thermal cycling on lunar analogue rocks revealed that surface boulders ≤ 30 cm in dimensions exhibit a weak response to thermal forcing, but thermal stresses in boulders larger than 1 m in diameters increase to several tens of MPa, since their surfaces approach that of an infinite halfspace (Molaro et al., 2017). Based on fracture patterns observed in meter-sized boulders, fractures formed by both thermal weathering and micrometeorite impacts were interpreted (Ruesch et al., 2020). It is highly plausible that thermal weathering can break down exposed bedrocks on the Moon, and fast landslides can be triggered by thermal weathering on steep slopes (Molaro et al., 2017). However, it is also notable that numerical modelling showed that a thin cover of regolith on competent boulders, which should be very common on the Moon, acts as a good thermal shield, and the thin regolith cover would prohibit the initiation of thermal fracturing (Molaro et al., 2017).
Besides thermal fracturing, thermal moonquakes may be another mechanism driving lunar landslides. Most shallow moonquakes recorded by the Apollo seismometers events are small and local quakes that may be triggered by diurnal temperature changes. Thermal moonquakes are recognized by the repetition of nearly identical signals at the same time during each lunation, which begin abruptly about 2 Earth days after the lunar sunrise and decrease rapidly after sunset (Duennebier and Sutton, 1974). The source energy of thermal moonquakes is between 7 × 107 and 1010 ergs within a 1 to 4 km range. The surface effect of thermal moonquakes is similar to booming sands on Earth, which only occurs in extremely dry sands that feature smooth surfaces (Lindsay et al., 1976; Criswell and Lindsay, 1974). On the dry lunar surface, the ubiquitous cover of a layer of fine soil and the repeatedly occurring thermal moonquakes may enable mechanical decoupling among grains, causing slow downslope creeping of lunar regolith (Pike, 1971).
10.03.2.1 The Apollo PSE Data
From 1969 to 1977, the Apollo Lunar Surface Experiment Package (ALSEP) recorded the lunar PSE data that now provide a unique and valuable resource for determining the interior structure of the Moon. The Apollo seismic ‘network’ comprised four stations at Apollo sites 12, 14, 15, and 16. Stations 12 and 14 were about 180 km apart, and, together, they formed one corner of an approximately equilateral triangle, with stations 15 and 16 at the other corners, about 1100 km apart (Figure 2). The lunar event catalog (Nakamura et al., 2008) documents 13 058 events recorded by the network. Recorded events exhibit different signal characteristics and originate from five types of sources originally classified as artificial impacts (9), meteoroid impacts (~ 1700), shallow moonquakes (28), deep moonquakes (~ 1360), and unclassified events. Many of the originally unclassified events have since been classified as deep moonquakes, with a total of 7084 deep moonquakes identified (as reported in Nakamura et al., 2008).
Compared with terrestrial quakes, moonquakes are small-seismic-moment events, and, as a result, the moonquakes recorded by the Apollo missions generated small signals, often with amplitudes smaller than the microseismic noise on Earth, even in Earth’s best seismic vaults. At first, events were visually identified in the lunar data, using hardcopy printouts and overlays of the seismograms recorded at Apollo stations 12, 14, 15, and 16. The criteria used to identify and classify seismic sources included the time interval from the start of a signal to its maximum amplitude (rise time), the dominance of low-frequency versus high-frequency content in the records, and the presence of compressional (P) and shear (S) waves (see review in Lammlein, 1977 and details in Lammlein, 1973 and Lammlein et al., 1974). Shallow events are also referred to in the literature as high-frequency teleseismic events, because of their unusually high frequency content and poor depth location (Nakamura, 1977a, 1980; Nakamura et al., 1974a, 1979). Analyses of data during early ALSEP operations suggested that deep moonquakes have two important characteristics (Nakamura et al., 1980) that were subsequently used to help identify such events. These characteristics are similar waveforms and tidal periodicities in moonquake occurrence times (Ewing et al., 1971; Lammlein, 1977). The main properties of the seismic sources, including seismic moment, stress drop, and rupture times, are given in Section 10.03.3. Figure 3 also shows examples of waveforms recorded at the Apollo stations, according to source type. In general, events become apparent in the seismic readings only after data filtering. Filtering is required because the raw waveform data are dominated by large amplitude signals that result from the expansion and contraction of the thermal Mylar shroud at sunrise and sunset (Duennebier and Sutton, 1974; see Figure 3 of Bulow et al., 2005).
Figure 3. From top to bottom, examples of waveforms for natural impacts (a), artificial impacts (b), a shallow event (c), and deep moonquakes (d–f) recorded on LP channels by the Apollo passive seismic network. All records span 55 min, with the x-axis representing time in minutes and the y-axis representing digital units. The following characteristics aided the identification of different seismic sources in the Apollo data: (1) The rise time (defined as the time from event onset to maximum signal amplitude) of moonquakes is typically less than that of meteoroids – the rise time in (c), (d), and (e) is approximately 5 min, while the rise time in (a) is approximately 10 min; (2) Deep moonquakes typically have low amplitude and produce poor data quality, as in (e). However, the waveform repeatability of deep moonquakes from distinct source regions allows the use of stacking to improve the signal-to-noise ratio, as in (f).
During the Apollo era, digital data were extracted from the complete, continuous PSE records for time windows encompassing identified events, and the digital data were then stored as the ‘event’ waveform data set. This derived data set and the original continuous records are now available, with documentation, from the IRIS Data Management Center (http://www.iris.edu/dms/nodes/dmc/) and at the ISAS DARTS web site (http://www.darts.isas.jaxa.jp/planet/seismology/apollo/index.html, Nagihara et al., 2011). Recent analyses of the Apollo 17 gravimeter complement the lunar data set with seismic records from the Apollo 17 landing site (Kawamura et al., 2010a,b), including records for deep moonquakes. Although the generic transfer functions of the Apollo seismometer have been published (see, e.g., on ISAS DARTS) and are generally used in most publications, it is possible that aging led to some change in these transfer functions. For example, when corrected with the generic transfer function, the noise spectrum of the Apollo 12 vertical records in peaked mode shows significant peaks near the peaked frequency of about 0.45 Hz (blue curve of the right box of Figure 1). These peaks decrease significantly, however, when slight changes are made to some of the transfer function parameters, as when the coil-magnet efficiency and the damping constant of the natural suspension are changed by − 6% and − 16%, respectively, from their nominal values for the light blue curve. For this reason, time-related changes in seismometer function will be kept in mind when detailed analyses are done on Apollo seismic data, including on the subsurface resonances. Although accurate time-dependent transfer functions can be obtained by processing calibration signals, which were recorded periodically (usually once a week) throughout the experiment, the coefficients of these functions (i.e., gain, pole, and zero) for each of the Apollo instruments throughout the 7 years of operation are not available in the literature.
Later work (Nakamura, 2003) has classified a significant fraction of identified, but previously unclassified, events (Nakamura et al., 1981) and has led to a nearly fivefold increase in the number of cataloged deep moonquakes. Additionally, work by Bulow et al. (2005) showed that the waveform repeatability of deep moonquakes can be exploited to search for additional unidentified events associated with known deep source regions. Implementation of this algorithm led to an approximately 35% increase in the number of events at the nine most seismically active (at the detection level of the Apollo seismometers) source regions (Bulow et al., 2005, 2007). Despite the seemingly large number of lunar seismic events, the poor signal-to-noise ratio for emergent arrivals in most records means that it is only possible to accurately measure arrival times and locate events for a subset of the lunar catalog data (Figure 4). Small event magnitudes and restricted receiver locations mean that events for which travel times can be clearly identified mainly occur on the near-side of the Moon.
The Apollo passive seismic data set has contributed enormously to our understanding of the internal structure of the Moon. The temporal and spatial distributions of lunar seismicity provide constraints on subsurface structure. Of particular importance is the evidence of brittle failure at depths of 700–1200 km, with failures apparently concentrated in the 800–1000-km depth interval. Unfortunately, the small magnitude of these events, combined with the limited number and geographical distribution of receivers, and associated errors in the locations of deep moonquakes (see, e.g., Hempel et al., 2012 for a recent analysis of the location errors), means that it has not been possible to unambiguously establish a focal mechanism for any of the deep moonquakes (Koyama and Nakamura, 1980). Seismic-velocity models for the crust and mantle (see Section 10.03.2.2) constrain the physical properties of the lunar interior (including temperature) and provide indirect constraints on mineralogy and the thermal evolution of the Moon. Studies of seismic attenuation indicate a dry lunar mantle (Nakamura and Koyama, 1982). Attempts have been made to investigate lunar normal modes (Khan and Mosegaard, 2001) to constrain the core size and state. However, these have largely been unsuccessful because of the limited bandwidth of the Apollo seismometers (Figure 1).
The Moon's shallow structure has received less attention in the literature, but it is of particular interest given the results from the Gravity Recovery and Interior Laboratory (GRAIL) mission (see below). The long coda in the observed waveforms likely result from the combined effects of low attenuation and a highly scattering near-surface layer with low seismic velocity (Latham et al., 1970a; Nakamura, 1976, 1977b). These coda can thus be used to probe the structure of the brecciated upper crustal layer, or megaregolith. Early lunar studies investigated the geological structure in the immediate vicinity of the receivers. For example, Nakamura et al. (1975) estimated regolith thicknesses of a few meters beneath Apollo stations 11, 12, and 15, with an underlying layer displaying P-wave velocities of 250–400 m s− 1, and Horvath et al. (1980) calculated seismic-velocity profiles down to a depth of about 200 m. The coda decay can be characterized by the quality factor Qc (where Qc− 1 represents the fraction of energy dissipated per cycle at a given frequency), and, for a lunar module impact at Apollo station 12, Latham et al. (1970a) found Qc ~ 3600, a value much larger than those recorded for terrestrial crustal rocks. Blanchette-Guertin et al. (2012) have undertaken the first systematic survey of the coda characteristics of impacts and moonquakes recorded at all four Apollo passive seismic stations. The results indicate decay times and Qc increase with depth, and Qc increases with increasing frequency and epicentral distance for shallow events. The results are consistent with scattering in a global, low-velocity and low-attenuation megaregolith layer with a size distribution of scatterers (blocks) that have more small-scale scatterers than large-scale scatterers. In addition, this study finds that scattering can be intensified locally in one or more frequency bands due to local structure.
In the following section, we discuss attempts to establish seismic velocity profiles for the lunar crust and mantle, and we also highlight GRAIL mission results related to lunar crustal structure. Because of their potential to provide significant insights concerning the lunar interior, these seismic-velocity models have been used for many seismological, mineralogical, and thermal studies of the Moon.
10.02.5.5 Lunar Seismology and Mineralogy
The Moon is the only planetary body other than the Earth for which a seismic velocity structure has been derived from analyses of the Apollo seismic data set (see Chapter 10.03). A seismic network of four stations on the lunar near side, installed at the landing sites of Apollo 12, 14, 15, and 16 from 1969 to 1972, continued to operate until it was turned off in September 1977. The seismic activity of the Moon is provided by deep moonquakes located at 850–1000 km depth, shallow moonquakes or high-frequency teleseismic events situated at 50–220 km depth, and meteoroid impacts onto the lunar surface. The deep moonquakes, apparently confined to the lunar near side, are most numerous and triggered by tidal deformation of the Moon. More than 12 000 events associated with 81 identified sources (‘nests’) have been recorded during the lifetime of the Apollo seismic network. It is likely that about 30 nests are located on the lunar far side. This implies that either the Moon's deep interior within about 40 ° from the sub-Earth antipode is nearly aseismic or, alternatively, that the region of the deep lunar interior deflects seismic waves and is highly attenuating for S waves, thereby impeding the localization of far side moonquake source regions (Nakamura, 2005).
Early seismic analyses were based on a limited set of arrival-time readings (Goins et al., 1981), whereas the model of Nakamura (1983) employs the complete data set of 5-year simultaneous operation of four Apollo seismometers including more deep moonquake sources. More detailed lunar velocity and density structure models have been obtained from a new inversion of different subsets of the Apollo seismic data set by applying various inversion techniques to revised arrival-time readings (Khan et al., 2000; Khan and Mosegaard, 2002; Khan et al., 2004; Lognonné et al., 2003; Chenet et al., 2006).
The multilayer model of Goins et al. (1981) consists of a dual-layer crust, an upper mantle, an intermediate transition zone, and a lower mantle. The crust is subdivided into a 20-km-thick upper layer with P-wave velocity VP = 5.1 km s−1 and S-wave velocity VS = 2.96 km s−1 and underlain by a lower-crust layer with VP = 6.8 km s−1 and VS = 3.9 km s−1 varying between 20 and 60 km in thickness. The upper mantle then extends from 60 to 400 km depth with average seismic velocities of VP = 7.7 ± 0.15 km s−1 and VS = 4.45 ± 0.05 km s−1. The moderate velocity decrease with depth is believed to be related to the temperature increase with depth. The corresponding seismic quality factors are QP = 5000 and QS = 3000, respectively. A narrow transition zone is situated between 400 and 480 km depth and reveals a prominent decrease of the shear-wave velocity, thereby creating a seismic shadow zone. The substantial decrease of the shear-wave velocity at the base of the transition zone has been interpreted in terms of a compositional change to a more primitive undifferentiated interior. The lower mantle ranges from 480 to 1100 km depth with average seismic velocities of VP = 7.6 ± 0.6 km s−1 and VS = 4.2 ± 0.1 km s−1 and corresponding seismic quality factors of QP = 1500 and QS = 1000, respectively. However, significant shear-wave attenuation accompanied by a QS drop by one order of magnitude occurs below a depth of 1100 km and suggests that the lowermost lunar mantle is at least partially molten.
The model of Nakamura (1983) includes a four-layer crustal structure that was introduced to satisfy the P-wave traveltime data of artificial impacts. The lunar mantle is subdivided into three layers with discontinuities at 270 and 500 km depth. The average velocities in the upper mantle above 270 km depth decrease from VP = 7.74 km s−1 and VS = 4.49 km s−1 to VP = 7.46 km s−1 and VS = 4.25 km s−1 between 270 and 500 km depth. In contrast to earlier studies based on a limited data set, the average velocities in the middle mantle between depths of 500 and 1000 km are about 8.26 and 4.65 km s−1, respectively. The relatively high velocities in the middle mantle layer may represent Mg-rich olivine rock left over from early melting and differentiation of the deep lunar interior. The multilayer lunar interior model suggested by Nakamura (1983) is an approximation to more realistic continuous models with P and S velocities gradually decreasing with depth.
Contrary to previous studies, mean crustal thickness of only 45 ± 5 km (Khan et al., 2000) or 38 ± 3 km (Khan and Mosegaard, 2002) have been proposed based on a re-analysis of Apollo lunar seismic data (Figure 7). The P-wave velocity increases from the surface to the crust–mantle transition. The upper mantle extends from 45 to 500 km depth with average P- and S-wave velocities of 8.0 ± 0.8 km s−1 and 4.0 ± 0.4 km s−1, respectively. In contrast to the model of Nakamura (1983), the upper mantle is characterized by constant seismic velocities indicating a homogeneous structure and composition. Similar to the model of Nakamura (1983), the transition to the middle mantle at a depth of 560 ± 15 km is marked by a sharp increase to VP = 9.9 ± 1.9 km s−1 and VS = 5.9 ± 0.9 km s−1. It should be noted, however, that such high seismic velocities are difficult to reconcile with any known mineral phase assemblages. An intermediate low-velocity layer may cover the depth range from 620 to 700 km. Below that depth, extends a high-velocity layer with P- and S-wave velocities of 11.0 ± 2.1 and 6.0 ± 0.7 km s−1, respectively, encompassing the deep moonquake source region.
Figure 7. Marginal posterior probability distributions illustrating the range of possible (left) S-wave and (right) P-wave velocity structures of the Moon based on a total number of 50 000 models. The contour lines define nine equally sized probability density intervals for the distributions. Note that maximum probability does not necessarily correspond to maximum likelihood of a calculated velocity structure. Adapted from Khan A and Mosegaard K (2002) An inquiry into the lunar interior: A nonlinear inversion of the Apollo lunar seismic data. Journal of Geophysical Research 107(E6): 50036 (doi:10.1029/2001JE001658).
The most prominent increase in seismic velocity that occurs at a depth of ∼500 km represents the transition from the upper mantle to the middle mantle. This discontinuity has been interpreted in terms of a mineralogic phase transition from the spinel to the garnet stability field. Alternatively, a change in composition to more aluminous and MgO-rich mafic silicates has been invoked thereby increasing the Mg number below this depth (Nakamura, 1983; Khan et al., 2006a). The re-analysis of the Apollo lunar seismic data indicates a homogeneous, constant-velocity upper mantle extending down to 560 ± 15 km depth, whereas the radial velocity distribution suggests more inhomogeneous middle-and lower-mantle layering (Khan et al., 2000; Khan and Mosegaard, 2002). The compositional change across the discontinuity together with the homogeneity of the upper mantle and the inhomogeneity of the middle mantle has been interpreted in terms of the initial depth of melting and differentiation during the magma-ocean phase of the Moon (Hood and Zuber, 2000). Based on mare-basalt petrology and thermal evolution considerations, Elkins-Tanton et al. (2003a) identify the 500-km discontinuity with the maximum depth of melting beneath the Procellarum-KREEP-terrane on the lunar near side. It is also possible that the 500-km discontinuity represents stratified olivine- and orthopyroxene-rich cumulates that were subsequently emplaced at the bottom of the lunar magma ocean. The mineralogical layering is supported by seismic inversions that account for thermodynamic mineral phase equilibria (Kuskov and Fabrichnaya, 1995; Kuskov, 1995, 1997; Kuskov and Kronrod, 1998).
The existence of a small lunar core could not be confirmed or rejected solely on the basis of the seismic data acquired by the Apollo missions. During the limited operational period of the lunar seismic network with stations on the near-Earth hemisphere, only one meteoroid event was detected with seismic rays crossing a central low-velocity zone from which the size of a lunar core was estimated for the first time. However, Sellers (1992) has shown that a relocation of this meteorite impact on day 263 in 1973 can explain the arrival-time and signal-amplitude data without requiring the presence of a lunar core. From tentative arrival-time determinations of two of the largest, most distant impacts it has been found that the corresponding seismic phases could have traveled at a P-wave velocity of about 5 km s−1 through a lunar core about 400–450 km in radius (Sellers, 1992). Considering the ambiguity inherent in the Apollo lunar seismic data, first reliable recordings of seismic phases traveling through the central region of the Moon are therefore expected from future measurements, like those once envisaged for the Japanese Lunar-A mission (Mizutani 1995).