Internal Structure/Mantle Motions of the Moon

Over the years since the Apollo era, the current lunar interior structure has been investigated using seismic, gravity, and magnetic field data. The Apollo seismic network recorded about 1,800 meteoroid impacts, 28 energetic shallow moonquakes (with body wave magnitudes up to five and hypocenters about 100 km below the surface), and about 7,000 extremely weak deep moonquakes that were located about halfway to the center of the Moon (e.g., Wieczorek 2009). The deep moonquakes are very enigmatic in that their occurrences are correlated with the tides raised by the Earth, they involve very low stress drops (less than 1 bar), they appear to originate from about 300 “nests” that are repeatedly activated, and almost all of these nests are located on the Moon’s nearside hemisphere (Nakamura 2003; Bulow et al. 2007; Qin et al. 2012). The nearside distribution of the deep moonquakes could perhaps indicate that the farside hemisphere is seismically inactive. Such a hypothesis is possible, especially if one considers that almost all of the Moon’s mare basalts erupted on its nearside hemisphere. Alternatively, it is possible that the deepest interior of the Moon attenuates seismic signals originating from farside events. In support of this hypothesis, seismic arrivals along ray paths passing through regions deeper than 700 km radius are not detected, suggesting mantle in this region hot enough to be seismically attenuating and possibly partially molten (Nakamura 2005). One scenario that could be consistent with the above observations is that numerous magma-filled fractures currently located in the deep mantle are capable of relieving the small stresses induced by Earth-raised tides. The Lunar Laser Ranging mission (Williams et al. 2001) and recent Gravity Recovery and Interior Laboratory (GRAIL) mission (Zuber et al. 2012) provide us the gravity measurements. Especially the Love numbers from the GRAIL mission are unprecedentedly accurate. These Love numbers have been used to invert the bulk solid-body “dissipation factor” Q (Williams et al. 2001) as well as the tidal tomography (Zhong et al. 2012). The Q number is around 30, which is compatible to that of the Earth andMars. This Q number infers a core–mantle boundary (CMB) temperature 1,400 C (Nimmo et al. 2012). This temperature is below the harzburgite mantle solidus while higher than ilmenite-mixed harzburgite solidus (Zhang et al. 2013a, see Fig. 1). Joint inversion of seismic and gravitational data (Khan et al. 2006) suggests a CMB temperature range from 1,170 C to 1,500 C, consistent with that inferred from the Q facter. To gain a partial melt, the lunar deep mantle has to be chemically heterogeneous (either ilmenite rich or water rich (Evans et al. 2014)). The surface and orbit magnetic records have also been used to explore the present-day mantle temperature (Sonett 1982). The electrical conductivity profile of the Moon has been estimated by using electromagnetic-sounding techniques, whereby the relationship between simultaneous time variations in the ambient magnetic field as measured from orbit and the resulting field measured on the surface was analyzed (Wieczorek 2009; Hood et al. 1982). This electrical conductivity profile suggests the lunar mantle temperature from 1,100 C to 1,250 C (see Fig. 1). The lunar mantle is energetically interacting with its core. A number of seismic and gravity studies identify the lunar core with a radius of about 250–430 km (Wieczorek 2009; Weber et al. 2011; Garcia et al. 2011). The core is believed to have liquid outer and solid inner regions based on seismic