Preservation potential of implanted solar wind volatiles in lunar paleoregolith deposits buried by lava flows

Icarus 207 (2010) 595–604

Contents lists available at ScienceDirect

Icarus journal homepage: www.elsevier.com/locate/icarus

Preservation potential of implanted solar wind volatiles in lunar paleoregolith deposits buried by lava flows Sarah A. Fagents a,*, M. Elise Rumpf a, Ian A. Crawford b, Katherine H. Joy b a b

Hawaii Institute of Geophysics and Planetology, University of Hawaii, 1680 East–West Road, Honolulu, HI 96822, United States Centre for Planetary Sciences at UCL/Birkbeck, Department of Earth and Planetary Sciences, Birkbeck College London, Malet Street, London WC1E 7HX, UK

a r t i c l e

i n f o

Article history: Received 16 June 2009 Revised 19 November 2009 Accepted 25 November 2009 Available online 4 January 2010 Keywords: Moon, Surface Regoliths Solar wind Cosmic rays Volcanism

a b s t r a c t The lunar surface is bathed in a variety of impacting particles originating from the solar wind, solar flares, and galactic cosmic rays. These particles can become embedded in the regolith and/or produce a range of other molecules as they pass through the target material. The Moon therefore contains a record of the variability of the solar and galactic particle fluxes through time. To obtain useful temporal snapshots of these processes, discrete regolith units must be shielded from continued bombardment that would rewrite the record over time. One mechanism for achieving this preservation is the burial of a regolith deposit by a later lava flow. The archival value of such deposits sandwiched between lava layers is enhanced by the fact that both the under- and over-lying lava can be dated by radiometric techniques, thereby precisely defining the age of the regolith layer and the geologic record contained therein. The implanted volatile species would be vulnerable to outgassing by the heat of the over-lying flow, at temperatures exceeding 300–700 °C. However, the insulating properties of the finely particulate regolith would restrict significant heating to shallow depths. We have therefore modeled the heat transfer between lunar mare basalt lavas and the regolith in order to establish the range of depths below which implanted volatiles would be preserved. We find that the full suite of solar wind volatiles, consisting predominantly of H and He, would survive at depths of 13–290 cm (for 1–10 m thick lava flows, respectively). A substantial amount of CO, CO2, N2 and Xe would be preserved at depths as shallow as 3.7 cm beneath meter-thick flows. Given typical regolith accumulation rates during mare volcanism, the optimal localities for collecting viable solar wind samples would involve stacks of thin mare lava flows emplaced a few tens to a few hundred Ma apart, in order for sufficient regolith to develop between burial events. Obtaining useful archives of Solar System processes would therefore require extraction of regolith deposits buried at quite shallow depths beneath radiometrically-dated mare lava flows. These results provide a basis for possible lunar exploration activities. Ó 2009 Elsevier Inc. All rights reserved.

1. Introduction Among the many compelling reasons for returning a human presence to the Moon (Taylor, 1985; Spudis, 1996, 2001; Crawford, 2004, 2006; NRC, 2007) is the potential for greatly enhancing our understanding of early Solar System history by accessing the record contained in the lunar regolith. The continuous influx of particles from the solar wind, galactic cosmic rays and micrometeorites, unimpeded by an atmosphere or significant magnetic field, is recorded as a host of resultant changes within the target regolith. Extraction of these exogenously-implanted materials from regolith layers therefore has the potential to yield a record of critical details of the inner Solar System environment throughout the Moon’s history (e.g., Spudis, 1996). * Corresponding author. Fax: +1 808 956 6322. E-mail address: [email protected] (S.A. Fagents). 0019-1035/$ - see front matter Ó 2009 Elsevier Inc. All rights reserved. doi:10.1016/j.icarus.2009.11.033

In order to be preserved, discrete deposits of ancient regolith must be shielded from continued bombardment and irradiation, so as to avoid becoming disrupted and mixed with material implanted at later times. One mechanism for this is the burial of ancient regolith layers (‘paleoregoliths’, by analogy to terrestrial paleosols; Retallack, 2005) by subsequent lava flows. However, the implanted materials must survive heating by the over-lying lava flow at temperatures potentially well in excess of 1100 °C (e.g., Taylor et al., 1991). In this paper we focus on the fate of implanted solar wind volatiles, which represent by far the greatest flux of incoming particles. We establish the conditions under which solar wind volatiles implanted long ago in the lunar regolith would survive heating by an over-lying lava flow. We model numerically the heat transfer from lunar lava flows to the underlying particulate regolith to derive the depths below which various implanted volatiles would be preserved. Finally, we discuss the implications for retrieval of such deposits during future robotic or human lunar exploration.

596

S.A. Fagents et al. / Icarus 207 (2010) 595–604

2. Background 2.1. Regolith deposits as archives of Solar System history Upon exposure of fresh bedrock on airless bodies to the space environment, a host of processes act to fragment and modify the target surface. These processes include bombardment by meteoroids (down to microscales), solar wind particles and cosmic rays, which induce sputtering, vaporization, agglutinate formation and comminution. The result over time is a finely fragmented regolith layer—a complex mixture of rock and mineral fragments, impact glasses, agglutinates, and breccias (McKay et al., 1991; Korotev, 1997; Lucey et al., 2006). Regolith formation rates are rapid after initial exposure of the fresh surface, but decline with time as the growing regolith layer shields the underlying bedrock from impactors. Contemporary regolith formation rates are estimated at 1 mm/Ma (Hörz et al., 1991), but may have been as high as 20 mm/Ma prior to 4 Ga due to the greater impactor population early in Solar System history (Crozaz et al., 1970; Duraud et al., 1975). The Apollo 11 basalts emplaced at 3.6–3.8 Ga are estimated to have accumulated regolith at a rate of 5 mm/Ma (Hörz et al., 1991). Typical present-day regolith thickness estimates range from 4–5 m in the maria to 10–15 m in the highlands (Oberbeck and Quaide, 1968; Shoemaker et al., 1968; McKay et al., 1974; Langevin and Arnold, 1977). The variety of processes acting to create and modify the regolith, and its propensity for retaining implanted materials, means that it potentially holds critical records of the environment of the inner Solar System throughout the Moon’s history. This record likely includes the history of the solar wind (e.g., Wieler et al., 1996), solar particle events produced by solar flares, and galactic cosmic rays produced by distant supernova explosions. The regolith may also preserve evidence of the early Earth’s atmosphere (Ozima et al., 2005, 2008), and samples of the early Earth ejected by meteoroid impacts (Armstrong et al., 2002; Crawford et al., 2008). The regolith therefore represents a unique repository of geochemical signatures of these processes, and of inner Solar System history more generally, that is unlikely to be preserved on more dynamic planetary bodies (Wieler et al., 1996). Of the impacting charged particles, cosmic rays, consisting of high-energy electrons, protons and neutrons, penetrate the surface to depths of centimeters to meters (McKay et al., 1991; Vaniman et al., 1991; Lucey et al., 2006). This influx produces a cascade of nuclear reactions in the target materials, leading to linear tracks of crystal damage. The products of these reactions (mainly CO, CO2, H2O, N2) can be used to measure particle influx, surface exposure age, and burial depth of sample in the regolith (Crozaz et al., 1970; Fleischer et al., 1970; Drozd et al., 1974, 1977; Reedy et al., 1983; McKay et al., 1991; Lorenzetti et al., 2005; Nishiizumi et al., 2009). The solar wind, consisting of a plasma of ionized atoms from the Sun’s atmosphere, represents a substantially greater particle influx (3  108 protons cm2 s1) and is the main source of volatiles in the regolith (Haskin and Warren, 1991; Vaniman et al., 1991). It consists of 95% H, 4% He, and less than 0.5% of C, N, O, Ne, Mg, Si, Fe, Ar, Kr, Xe (Haskin and Warren, 1991; McKay et al., 1991; Vaniman et al., 1991). Solar wind particles penetrate to depths of microns to millimeters in target materials, gradually accumulating in the outer layers of exposed grains (Dran et al., 1970; McKay et al., 1991). It is generally thought that the Moon accreted or retained few volatiles when it formed because of the high temperatures generated by ejection and subsequent reaccretion of material resulting from the impact of a planet-sized body into the early Earth (Taylor, 1982; Lucey et al., 2006). This notion has been challenged recently for H (in the form of H2O or OH) by detections of

significant abundances in the lunar surface (Clark, 2009; Pieters et al., 2009; Sunshine et al., 2009). Although the source is equivocal (extra-lunar vs. endogenic), recent work by Saal et al. (2008) indicates an indigenous source for some H2O, as well as CO2, F, S, and Cl, which further calls into question the notion that the Moon is essentially anhydrous (Chaussidon, 2008). However, Haskin and Warren (1991) summarize a number of studies indicating that abundances of solar wind volatiles H, He, C, N, Ne, Ar, Kr, and Xe are typically 1–4 orders of magnitude greater in lunar soils and regolith breccias than in ‘pristine’ lunar rocks. Furthermore, isotopic compositions of regolith samples cluster tightly around solar wind values, and the solar wind deuterium/hydrogen ratio should be extremely low (Haskin and Warren, 1991). These factors allow solar wind volatiles to be readily distinguished from indigenous sources (Haskin and Warren, 1991; McKay et al., 1991). Studies of the solar wind have implied some variability in intensity and composition over the lifetime of the Solar System (Geiss and Bochsler, 1991; Kerridge et al., 1991). For example, solar wind Xe is suggested to have been a factor of 2–3 greater in the past (Geiss, 1973; Kerridge, 1989); this is supported by measurements on N (Clayton and Thiemans, 1980; Kerridge, 1980). The 15N/14N ratio may have increased by 15% per Ga, suggesting an increase in solar activity (Kerridge, 1975, 1989). Other changes have been reported with lesser confidence (Eberhardt et al., 1972; Geiss, 1973; Kerridge, 1980; Pepin, 1980; Becker and Pepin, 1989). More recently, Wieler et al. (1996) found temporal changes in the solar wind Kr/Ar and Xe/Ar ratios. The picture is far from complete, and additional studies of caches of solar wind particles from throughout lunar history would greatly further our knowledge of solar variability (Geiss and Bochsler, 1991). One particularly interesting aspect of the lunar record of ancient solar activity concerns the strength of the ancient solar wind, rather than just its composition. The standard solar model predicts that the Sun’s luminosity was about 70% of its present value 4 Ga ago (e.g., Gough, 1981), and this poses problems for explaining geomorphological evidence for liquid water on the early Earth and Mars (the so-called ‘faint young Sun paradox’, Sagan and Mullen, 1972). Identifying plausible atmospheric greenhouse models able to account for the ‘paradox’ remains problematical (e.g., Haqq-Misra et al., 2008). An alternative possibility for explaining apparently higher temperatures on early Earth and Mars is that the young Sun was a few percent more massive, and therefore significantly more luminous than estimates based on models that assume its present mass (Whitmire et al., 1995; Sackmann and Boothroyd, 2003). There is some evidence for this hypothesis from astronomical studies of other young solar type stars (e.g., Wood et al., 2002), but these models predict a very intense solar wind early in the Sun’s evolution (i.e., sufficient to reduce the Sun’s mass from perhaps 1.03–1.07 solar masses 4.5 Ga ago to 1 solar mass today). If this theory is correct, evidence for this greatly enhanced solar wind should be preserved on the Moon if sufficiently ancient regolith deposits could be identified. Alternatively, a lack of such evidence could disprove the hypothesis of a more massive early Sun. Either way, the ancient lunar solar wind record has the potential to help solve a major outstanding question in Solar System science (and one of clear astrobiological significance), as well as to advance our understanding of the early evolution of solar type stars more generally (Wood et al., 2002). In order to exploit these volatile archives we need to define the conditions under which they might be preserved. One possibility is that a layer of regolith becomes blanketed by ejecta from nearby impact craters. However, this simple stratigraphy will be destroyed as a complex structure develops due to repeated burial, exhumation and mixing of regolith and ejecta packets by subse-

S.A. Fagents et al. / Icarus 207 (2010) 595–604

quent impacts. Indeed, Apollo core samples were too complex to unravel the detailed record of solar activity (McKay et al., 1991). A potentially more useful scenario for preserving paleoregolith layers is to bury them by later lava flows (Fig. 1) (McKay et al., 1989; Spudis, 1996; Crawford et al., 2007, 2009; McKay, 2009; Spudis and Taylor, 2009). In order to access the record of solar activity contained within a paleoregolith deposit, both the timing and duration of exposure must be established (McKay et al., 1991). The archival value of deposits sandwiched between lava layers is therefore enhanced by the fact that both the under- and over-lying basalt can be dated by isotopic radiometric techniques, thereby precisely defining the age of the paleoregolith and the geologic record contained therein. Samples retrieved from between multiple flow units at different depths would host solar wind materials over a range of ages, therefore allowing a long baseline record to be acquired. However, for useful information to be retrieved from such deposits, the implanted volatiles must withstand the thermal consequences of the emplacement of the over-lying lava flow. Heating experiments on Apollo soil samples showed that the bulk of CO, CO2, N2, and Xe is released at temperatures in excess of 700 °C, although significant amounts of other gases are lost down to temperatures of 500 °C (Ne, CH4, Ar) and 300 °C (H2, He) (see Fig. 2 and Table 1; Gibson and Johnson, 1971; Simoneit et al., 1973; Haskin and Warren, 1991; Fegley and Swindle, 1993). Eruption temperatures of lunar basalts are likely to have been greater than terrestrial basalts, with mare basalt liquidus temperatures in the range 1170– 1440 °C (Murase and McBirney, 1970; Taylor et al., 1991; Williams et al., 2000), and solidus temperatures >1050–1100 °C (Taylor et al., 1991). Thus, volatiles would be baked out of the regolith for some depth beneath the base of the flow. It is therefore important to determine the range of depths below which implanted volatiles will be retained in a regolith deposit capped by lava. Once the volatiles have been outgassed from regolith grains they may become temporarily trapped and concentrated beneath the flow. Upon cooling and solidification, however, the flow will likely develop extensive fracturing which, together with void spaces in surface or basal clinker or crustal plates, would produce a highly permeable medium (USGS, 1999), ultimately promoting loss of the liberated gases to space. micrometeorite bombardment

regolith formation vaporization, sputtering, comminution, agglutinate formation

solar wind particles and cosmic rays

lava flow 2

2.2. Lava–substrate heat transfer A number of classical studies have investigated the heat transfer of igneous bodies in contact with substrate or country rock (e.g., Jaeger, 1961; Bruce and Huppert, 1990; Turcotte and Schubert, 2002). Interest in understanding lava–substrate heat transfer on the Moon was motivated by the hypothesis that lunar sinuous rilles were formed as a result of lava flows entrenching into the surface by thermomechanical erosion, i.e., melting, abrasion, and assimilation of the substrate (Hulme, 1973, 1982; Carr, 1974; Coombs et al., 1988; Bussey et al., 1997; Williams et al., 2000; Fagents and Greeley, 2001). However, in most cases, these models assume a priori that melting at the lava–substrate interface does occur, and are aimed primarily at determining the extent and rate of thermal erosion. Furthermore, an assumption commonly adopted to facilitate solution of the governing equations is that the lava and substrate possess similar thermophysical properties. This approach does not account for the possibility that the lava and substrate are inherently different in nature, as in the case of a high-temperature, high-density lunar lava over-lying a porous, fragmental regolith. The low-density, particulate regolith has a low thermal conductivity and specific heat capacity due to the presence of interstitial voids and vacuum conditions (Fountain and West, 1970; Hemingway et al., 1973; Langseth et al., 1976). The low conductivity imbues the regolith with significant insulating properties that restrict the extent to which the substrate would be heated. The lava and regolith properties also vary with temperature. What is lacking to date, therefore, is a treatment that specifically seeks to determine the temperature distribution within a regolith having contrasting properties to that of the lava. This is the purpose of the present paper. 3. Modeling lava–regolith heat transfer 3.1. Governing equation We seek to determine the depth to which implanted volatiles would be degassed from a regolith layer heated by an over-lying, initially molten lava flow. The scenario we model is that of a lava flow of thickness d and initial temperature Tl0 over-lying a regolith substrate at initial temperature Tr0 (Fig. 3; Table 2 lists all notation used). We treat the lava as instantaneously emplaced, i.e., we are not treating the flow dynamics, but instead model the heat transfer of the flow at rest. The lava surface cools by radiative heat loss to space, whereas the base cools conductively to the regolith below. The thermal pulse from the lava will propagate down into the regolith, reaching a maximum depth of penetration determined by the initial thermal energy of the flow. The general equation for energy conservation in a fluid flow is given by

@ ðqhÞ þ r  ðquhÞ ¼ r  ðkrTÞ þ Sh ; @t

buried paleoregolith containing implanted volatiles

lava flow 1 Fig. 1. Mechanism for preserving volatile-rich regolith deposits by burial by lava flow. Upon emplacement, lava flow 1 will be subjected to an influx of impactors that will generate a surficial regolith. Incoming solar wind and cosmic wind particles will become embedded in the regolith until such time as lava flow 2 is emplaced. This flow will both shield the underlying regolith from further implantation and bake out implanted volatiles to a certain depth. The process of regolith formation commences again at the surface of lava flow 2. Repeated lava emplacement, regolith formation and burial yields a time series archive of solar wind and cosmic ray samples.

597

ð1Þ

in which h, q, u, T and k are the specific enthalpy, density, velocity, temperature, and thermal conductivity of the medium, respectively. The source term Sh describes the volumetric rate of generation of additional heat (e.g., via viscous dissipation, crystallization, etc.). The temperature and enthalpy are related as crT = rh, with c denoting the specific heat capacity. Using this relation and adopting constant c, Eq. (1) can be written with T as the dependent variable:

  @ k S rT þ h : ðqTÞ þ r  ðquTÞ ¼ r  @t c c

ð2Þ

598

S.A. Fagents et al. / Icarus 207 (2010) 595–604

Temperature (˚C)

(a)

0

200

400

600

800

1000

1200

1400

100

Ion Intensity (arbitrary units)

CO, N2 N2

H2 O2

10

H2O

CO2 He H2S SO2

1

400

600

800

1000

1200

1400

1600

Temperature (K) 800

1000

1200

1400

1000

1200

1400

(b)

(c)

% Gas Released

% Gas Released

800 100

100

H He

50

50

Ar

C

Xe N

0 400

0

600

800

1200 400

1000

600

800

1000

1200

Temperature (˚C) Fig. 2. Gas release patterns for regolith volatiles. Release patterns for (a) Apollo 11 soil sample 10086,16 (Gibson and Johnson, 1971); (b) He, Ar, Xe, and (c) H, C, N in lunar soils. Data in (b) and (c) are summarized in Fegley and Swindle (1993) from Hohenberg et al. (1970), Pepin et al. (1970), Srinivasan et al. (1972), Basford et al. (1973), Becker and Clayton (1977, 1978), Becker (1980), Thiemans and Clayton (1980) and Norris et al. (1983). Paired curves for each element indicate range of observations.

Table 1 Temperature ranges over which the main solar wind derived volatiles evolve upon heating of lunar regolith samples (Simoneit et al., 1973). Species

Temperature range

H2, He CH4, Ne, Ar CO, CO2, N2, Xe

300–700 °C (573–973 K) 500–700 °C (773–973 K) >700 °C (>973 K)

For a stationary lava the convective term on the left hand side disappears. Considering the vertical (y-) dimension only, neglecting the source term and assuming incompressibility, Eq. (2) reduces to the well-known one-dimensional heat conduction equation:

@T @2T ¼j 2; @t @y

ð3Þ

in which j is the thermal diffusivity, given by j = k/qc. In this scenario we consider the heat transfer from a fluid (lava) to a particulate solid (regolith substrate) having distinct thermophysical properties, which necessitates the use of numerical solution techniques. We account for these contrasts by applying different values for thermal conductivity k, density q, and heat capacity c in the parts of the computational domain representing the lava (denoted by subscripted l) and the regolith substrate (denoted by subscripted r); see Fig. 3 and Table 3. At the interface between the lava and the regolith, the conductive heat flux supplied from the lava to the interface, qcl, must be

Fig. 3. Schematic illustration of model formulation. Lava of thickness d, initial temperature Tl0, density ql, specific heat capacity cl, and thermal conductivity kl heats the underlying regolith of initial temperature Tr0, density qr, specific heat cr and thermal conductivity kr. The lava surface cools by radiation. The computational mesh is refined at the interfaces to ensure accurate calculation of the temperature distribution in the regions of high thermal gradients.

balanced by the conductive heat flux conducted away from the interface by the regolith, qcr. We thus apply the boundary condition qcl = qcr at the interface, according to Fourier’s law:

S.A. Fagents et al. / Icarus 207 (2010) 595–604 Table 2 Notation. Symbol

Meaning

Units

c d h k L qcl qcr qrad Sh T Tamb Tl0, Tr0 Tliq Tsol u

Specific heat capacity Flow thickness Specific enthalpy Thermal conductivity Latent heat of crystallization Conductive heat flux supplied to interface by lava Heat flux conducted away from interface by regolith Radiative heat flux from lava surface Source term in energy equation Temperature Temperature of the space environment Initial temperature of lava, regolith Lava liquidus temperature Lava solidus temperature Velocity vector Emissivity of lava, 0.99 Thermal diffusivity = k/qc Density Stefan–Boltzmann constant, 5.7  108

J kg1 K1 m J kg1 W m1 K1 J kg1 W m2 W m2 W m2 W m3 K K K K K m s1 – m2 s1 kg m3 W m2 K4

e j q r

 kl

@T @y

 ¼ kr l

  @T ; @y r

ð4Þ

where kl and kr are the thermal conductivities of the lava and regolith, respectively. In the absence of an atmosphere, the upper surface of the lava cools by radiation to the space environment. We therefore apply a radiative flux boundary condition, given by

  qrad ¼ re T 4y¼d  T 4amb ;

ð5Þ

in which r is the Stefan–Boltzmann constant, e is the emissivity of the lava surface, and Ty=d and Tamb are, respectively, the temperatures of the lava surface at y = d and the ambient lunar environment. 3.2. Model parameters Table 3 lists the characteristics of the lava and regolith that were chosen for the modeling study based on values reported in the lunar literature for laboratory studies of returned lunar core samples and regolith simulants, as well as for in situ measurements conducted during the Apollo missions. For the regolith, we adopted a value for thermal conductivity of kr = 0.011 W m1 K1, as representative of the range (0.009–0.013 W m1 K1) derived from in situ heat flow experiments during the Apollo 15 and 17 missions (Langseth et al., 1976; Vaniman et al., 1991). Data on specific heat for a range of regolith materials show little variability between samples (Robie et al., 1970; Hemingway and Robie, 1973; Hemingway et al., 1973), and we thus adopt a value of 760 J kg1 K1. Table 3 Model parameters. Parameter 1

Thermal conductivity, k (W m

1

K

)

Lavaa

Regolith

(i) 1.5 (ii) 0.75

0.011

Density, q (kg m3)

2980

1660

Specific heat, c (J kg1 K1)

(i) 1500 (ii) 3200

760

Initial temperature, T0 (K)

1500

200

Flow depth, d (m)

1 10

–

a Two entries for a given lava property in table denote values for the end member lava types under consideration; see text for details.

599

Density data from in situ and returned regolith core measurements lie in the range 800 to >2000 kg m3 (Carrier et al., 1991). The best estimates of typical values for the density of the lunar regolith are 1550 ± 50 kg m3 for 0–15 cm depth, and 1660 ± 50 kg m3 for the top 0–60 cm (Carrier et al., 1991). The data can be fitted by the expression q = 1.92(z + 12.2)/(z + 18) (where z is depth in centimeters), which asymptotes to a value of 1920 kg m3 beyond 3 m depth (Carrier et al., 1991). However, there is considerable variability in the data; the deepest measurement (down to 3 m), made during Apollo 17, is 1730 kg m3. Given the uncertainties involved, and the shallow depths to which the most intense heating is confined (see Section 5), we adopt 1660 kg m3 as a reasonable estimate of regolith density. Upon emplacement at t = 0 the lava would be predominantly molten but then crystallize and solidify as cooling progresses. Properties of lunar basalts in the melting range have been derived in laboratory studies of synthetic and natural mare basalts (summarized in Basaltic Volcanism Study Project (1981) and Taylor et al. (1991)), as well as through thermodynamic modeling (e.g., MELTS; Ghiorso and Sack, 1995; Williams et al., 2000). In comparison to terrestrial basalt compositions, lunar basalts have low SiO2, high FeO and MgO, very low alkalis, generally elevated TiO2, and negligible dissolved water component. They are expected to have had low viscosities, relatively high densities and high eruption temperatures. The available data indicate liquidus temperatures for mare basalts in the range 1170–1400 °C (1443–1673 K) (Taylor et al., 1991), and solidi typically P1050–1100 °C (1323–1373 K). We adopt an initial value of 1227 °C (1500 K) as representative of lava erupting at a temperature in the mid- to upper end of the melting interval of most compositions measured (Taylor et al., 1991), but note that this temperature is greater than the liquidus of some compositions. We discuss in Section 4 the consequences of varying the initial lava temperature. Specific heat capacity and thermal conductivity of mafic rocks can vary by a factor of two or more from solid to molten states over the temperature range of interest. Values for specific heat at subsolidus temperatures increase nonlinearly from 550 J kg1 K1 at 200 K to 1100 J kg1 K1 at 1200 K (Hemingway et al., 1973; Touloukian et al., 1981). In the melting range, values of 1460– 1700 J kg1 K1 have been reported (Lange and Navrotsky, 1992; Büttner et al., 1998; Williams et al., 2000). For our application, we chose a value of cl = 1500 J kg1 K1 but recognize that this is applicable only at the upper end of the temperature range of interest. This choice largely compensates for the fact that we do not explicitly treat the release of latent heat due to crystallization in the solidification interval, given that the additional latent heat loss is much less than the heat lost over the entire cooling history of the flow. A method of accounting for the resulting buffered cooling due to crystallization is to elevate the specific heat in the melting interval (DT = Tliq  Tsol) according to Carslaw and Jaeger (1986) thus: cl0 = cl + L/DT. Though few direct measurements of L have been made, a variety of sources suggest that a value of 5–6  105 J kg1 is appropriate for basaltic compositions (Lange et al., 1994; Navrotsky, 1995; Williams et al., 2000). A reasonable value for cl0 in a 300 K solidification interval is therefore 3200 J kg1 K1. However, use of this value applied over the entire cooling duration of the flow will significantly overestimate the heat capacity of the lava at subsolidus temperatures, such that the total heat available to heat the substrate will exceed (by a factor of two) that lost from a stationary solidifying flow. With this caveat, we use this value in some of our simulations to illustrate the effect of additional heat sources on regolith heating, such as would be the case for continued advection of lava into the domain. The 1500 and 3200 J kg1 K1 end members therefore cover the range of thermal influence of a lava flow of a given thickness and initial temperature.

600

S.A. Fagents et al. / Icarus 207 (2010) 595–604

Thermal conductivity of mafic rocks varies in a complex way with temperature. It may or may not show an initial decrease with increasing temperature (Cˇermàk et al., 1982), followed by an increase on entering the melting interval (Murase and McBirney, 1973; Touloukian et al., 1981), and sometimes a subsequent decrease at higher temperatures (Büttner et al., 1998). Given this complexity, we chose thermal conductivities kl = 0.75 and 1.5 W m1 K1 to represent the factor of two variability expected over the temperature range of interest. Measurements of the density of returned lunar basalt samples ˇ ermàk et range from 2716 to 3370 kg m3 (data summarized in C al. (1982)). Calculations of densities at liquidus temperatures, based on the method of Bottinga and Weill (1970), lie in the range 2900–2990 kg m3 for lunar compositions (Williams et al., 2000). Laboratory measurements on a synthetic lunar sample gave a density of 2980 kg m3 (Murase and McBirney, 1973), which we adopt here as our preferred value, noting that modeled substrate heating depths are insensitive to quite large variations in the choice of lava density (Fagents and Greeley, 2001). For the surface radiative boundary condition, we use Tamb = 4 K for the far-field temperature of the environment into

(a) 1 m lava, low cl 0

200

400

which lunar lava flows would radiate. Lava emissivities exceeding 0.9 have been used in thermal remote sensing studies of lava flows (Crisp et al., 1990; Pieri et al., 1990; Flynn et al., 1993), with experimentally determined values lying close to 1 (Pinkerton et al., 2002). We therefore adopt a value of e = 0.99 for this study. In considering the range of lava thicknesses to model, we note that outcrops of individual lava units as thin as 1 m have been identified in Apollo 15 images of the wall of Hadley Rille (e.g., AS-15-89-12116; Vaniman et al., 1991). Other considerations suggest that low-viscosity lavas corresponding to Apollo 11, 12, and 15 basalt samples should produce flows 8–10 m thick (Hiesinger and Head, 2006). Observation and measurements of flow units made from orbit yield thicknesses from one to several tens of meters (Gifford and El-Baz, 1981; Wilhelms, 1987; Hiesinger et al., 2002). It is likely that mare lavas, for which flow margins are often indistinct, consist of multiple overlapping individual flow lobes of lesser thickness. We have chosen flow thicknesses of 1 and 10 m to explore the range of depths to which the regolith, initially at T0r = 200 K, would be heated by typical lunar lava flows initially at T0l = 1500 K.

(b) 1 m lava, high cl

Temperature, T (˚C)

600

800

1000

0

1200

200

400

600

800

1000

lava

Depth, z (m)

1 day

0.4

40 d

20 d

40 d

20 d

0.4

10 d

0.2

5d 10 d

0.6

5d

0.6

0.2 100 d

100 d

-0.2 -0.4 200

lava

2d

0.8

0.8

0.0

1200

1.0

1.0

400

600

800

1000

1200

regolith

-0.2

1400

-0.4 200

1600

regolith

400

600

800

1000

600

800

1000

1200

1400

1600

Temperature, T (K)

(c) 10 m lava, low cl 200 10

400

600

800

(d) 10 m lava, high cl 1000

1200

1400

1600

200 10

400

1200

1400

1 day

1 day

Depth, z (m)

10 d 50 d

50 d 200 d

100 d 200 d

5

1600

500 d 5

500 d

1000 d

1000 d 1500 d 2000 d

0

10000 d

0

6000 d

4000 d

2000 d

regolith

regolith

-5

-5 0

200

400

600

800

1000

1200

0

200

400

600

800

1000

1200

Temperature, T (˚C) Fig. 4. Temporal evolution of temperature profiles through the lava and regolith. (a) 1 m thick lava flow with cl = 1500 J kg1 K1, (b) 1 m thick lava flow with cl = 3200 J kg1 K1, (c) 10 m thick lava flow with cl = 1500 J kg1 K1, (d) 10 m thick lava flow with cl = 3200 J kg1 K1. Note difference in y-axis scale between 1 and 10 m lava cases. Dashed lines indicate temperatures and above which different volatiles would be largely released from the regolith: CO, CO2, N2, Xe by 700 °C; CH4, Ne, Ar by 500 °C; H2, He by 300 °C (Table 1).

601

S.A. Fagents et al. / Icarus 207 (2010) 595–604

0

Depth to Isotherm, z (m)

Having defined the problem, we use PHOENICS, a finite-volume computational fluid dynamics code (http://www.cham.co.uk), to solve numerically the governing equation (Eq. (3)) and determine the temperature distribution within the two subdomains of contrasting material properties (lava and regolith) as a function of time. A computational mesh was constructed within the simulation domain (shown schematically in Fig. 3), which was refined near the surface and interface boundaries to ensure the accuracy of the calculations in regions of high temperature gradient. Mesh density, time-step size, and solution parameters (relaxation, etc.) were optimized, and numerical residuals were monitored throughout the solution process, to ensure stability, convergence and accuracy of the solutions. A total of 120 computational cells for simulations of 1 m of lava over-lying 1 m of regolith, and 600 cells for the 10 m lava over 5 m of regolith, were sufficient to achieve meaningful solutions. A time-step of 30 min was adopted, and the simulations were run until the 300, 500, and 700 °C isotherms, representing key volatile release thresholds (Table 1), had reached their maximum depths in the regolith.

10

20

30

40

50

60

70

80

0

1 m lava low cl high cl

700˚C

0.05

700˚C 500˚C

0.10

300˚C 500˚C

0.15 0.20 0.25

300˚C

0.30

Time After Emplacement, t (days)

(b) 0

4. Results Fig. 4 shows the temperature distribution within the lava and regolith as a function of time for 1 m thick (Fig. 4a and b) and 10 m thick lavas (Fig. 4c and d), and for each of the two adopted lava heat capacities (1500 and 3200 J kg1 K1). In each case, the vertical dashed lines denote the threshold temperatures above which key volatile species would be released (300, 500, and 700 °C; Table 1). It can be seen that the surface of the lava cools rapidly as the base of the lava heats the substrate, leaving a warmer core. For any given temperature (vertical dashed lines), the T–z curves propagate downward into the regolith to some maximum depth, after which they move upwards (and to the left) due to the cooling of the lava core. The curves become kinked at the lava–substrate interface due to the differing thermophysical properties (primarily thermal conductivity). The cases for which cl = 3200 J kg1 K1 (Fig. 4b and d) show much greater depths of heating compared with the cl = 1500 J kg1 K1 cases (Fig. 4a and c), as would be expected for a lava releasing a factor of >2 greater heat per unit mass. Similarly, the 10 m thick lava (Fig. 4c and d) heats the substrate to significantly greater depths than for a 1 m lava (Fig. 4a and b), because of the greater thermal energy contained within the factor of 10 greater lava mass. To determine the maximum depths reached by the key volatilerelease temperatures, Fig. 5 shows the penetration depths of the 300, 500, and 700 °C isotherms as a function of time after emplacement. The difference in heating depths for the high- and low-cl case is quite apparent in both the 1 m (Fig. 5a) and 10 m thick lavas (Fig. 5b). The maximum isotherm penetration depths are summarized in Table 4. For a 1 m thick lava with cl = 1500 J kg1 K1, it can be seen that the bulk of CO, CO2, N2, and Xe (700 °C) would be protected from being baked out of the regolith at depths exceeding just 3.7 cm beneath the lava. CH4, Ne, and Ar (500 °C) would be fully preserved at depths exceeding 7 cm, and H2 and He (300 °C) beyond 13 cm. These depths increase by roughly a factor of two for the high-cl case, and a factor of 10 for a 10 m thick lava. In the extreme case (high-cl, 10 m lava) the release depths are 78, 143, and 289 cm for the 700, 500, and 300 °C isotherms, respectively. In exploring the sensitivity of the results to plausible variations in other model parameters, we find that variations in initial lava temperature, Tl0, of ±100 °C lead to variations in the maximum penetration depth of ±9% for the 300 °C isotherm to ±20% for the 700 °C isotherm. These translate to just a few centimeters variation

Time After Emplacement, t (days)

(a)

Depth to Isotherm, z (m)

3.3. Solution procedure

0 0.5

2000

4000

700˚C 500˚C

1

6000

700˚C

300˚C

1.5

500˚C

8000

10000

10 m lava low cl high cl

.

2 2.5 3

300˚C

3.5

Fig. 5. Modeled depths of penetration of the key volatile-release isotherms (700, 500, and 300 °C) into the regolith as a function of time: (a) 1 m thick lava flow and (b) 10 m thick lava flow. In each case two sets of curves are shown, one for cl = 1500 J kg1 K1 and one for cl = 3200 J kg1 K1.

Table 4 Summary of the maximum depths reached by isotherms at which different volatile species would be baked from the regolith. Flow thickness (m)

Lava type

Depth to isotherm (cm) 700 °C (973 K)

500 °C (773 K)

1

Low-cl High-cl

3.7 7.8

7.0 14

300 °K (573 K) 13 28

10

Low-cl High-cl

38 78

70 143

124 289

for the 1 m lava to ±16–26 cm for the 10 m lava. Finally, we note that Fagents and Greeley (2001) found that the thermophysical properties of the regolith dominate over those of the lava in controlling the temperatures reached at depth in the regolith. Thus we infer that, despite the greater range of variability in lava properties (as a result of the simplicity of the model formulation or uncertainties in composition and temperature), our efforts to constrain substrate properties as closely as possible should provide reliable first-order results. 5. Discussion One limitation of our current model is that we have treated the lava as instantaneously emplaced, i.e., we have neglected to account for additional heat advected into the computational domain during a period of flow prior to coming to rest. In our approximation of latent heat release, applying a modified heat capacity (cl = 3200 J kg1 K1) over the entire simulation significantly overestimates the depths of heating of the substrate by a stationary flow, but goes some way towards compensating for the fact that we do not treat advective heat transfer. Applying a heat capacity cl = 1500 J kg1 K1 may compensate for latent heat release, as discussed in Section 4, but not for heat advected by continued flow.

602

S.A. Fagents et al. / Icarus 207 (2010) 595–604

The cl = 1500 J kg1 K1 simulations therefore provide adequate estimates of the baking depths at the distal margins of flow lobes, i.e., the low-cl baking depths given for each flow thickness (Table 4), for which little heat advection would have taken place prior to solidification. Other simplifications include neglecting the variation of thermophysical properties with temperature, which can be significant over the range between the lava eruption temperature and the lunar ambient surface temperature. However, we have adopted values to represent averages over the temperature range of interest. Our model thus provides bounds on the depths to which the lunar regolith would be baked free of implanted solar wind volatiles. Lava flow thickness has the greatest control on heating depth, but critically, the low thermal conductivity of the regolith restricts the depth of heating, such that pristine solar wind samples are expected to be found at depths exceeding 13–28 cm and 1.2–2.9 m in paleoregolith deposits buried beneath 1 and 10 m thick lava flows, respectively. At shallower depths, some but not all implanted volatiles will remain. The composition and abundance of these remaining volatiles are controlled by chemical fractionation when species are liberated at different temperatures and depths. Thus, we predict that abundance ratios of easily detectable volatiles i.e. H2/CO or He/N2 (where CO and N2 are released at high temperatures and He and H2 at low temperatures; Table 1) could be used in paleoregolith cores to act as markers of the extent of subsurface heating and degassing within the regolith. Adopting an average regolith accumulation rate of 5 mm/Ma (Hörz et al., 1991), the minimum time required to form regolith deposits that would preserve the entire suite of solar wind volatiles (Table 1) below a 1 m lava flow, ranges between 26 and 56 Ma (depending on cl). These time estimates scale in direct proportion to lava thickness. Thus, successive mare lavas would have to be emplaced at intervals exceeding these timeframes in order to entrap a viable solar wind sample. Mapping and age dating of flow units in mare basins (e.g., Oceanus Procellarum) suggests that this should have been a common occurrence (Hiesinger et al., 2000, 2003). Location of stacks of multiple flow units emplaced tens of Ma apart intercalated with paleoregolith deposits would therefore provide a time series of implanted volatile samples, suitable for analysis for temporal variations in the strength and composition of the solar wind. The results of our simulations allow us to provide some recommendations for future lunar exploration strategies with respect to locating and extracting an archive of the solar wind, galactic cosmic rays and, possibly, the Earth’s early atmosphere. Although the current lunar exploration architecture favors a polar locality for a manned outpost, we suggest that valuable science can be conducted at mare sites at lower latitudes (Crawford et al., 2007). Existing (Hiesinger et al., 2000, 2003) and future analyses of orbital image data (e.g., the Lunar Reconnaissance Orbiter Camera; Crawford et al., 2009; McKay, 2009) will provide suggestions for sampling locations based on the age range of mapped lava units. These localities may then be targeted with robotic in situ analyses, sample return missions, and human exploration efforts. Sampling of buried regolith deposits will be facilitated if stacks of lavas can be found in outcrop, as at Hadley Rille (Vaniman et al., 1991). Otherwise, detection and sampling of buried paleoregolith deposits would require technology such as ground-penetrating radar (e.g., Sharpton and Head, 1982; Ono et al., 2009; Heggy et al., 2009) and drilling equipment with a depth capability of several tens of meters. In order to accurately bracket the age of formation of sampled paleoregolith deposits, samples of the under- and over-lying lavas would also need to be collected for radiometric dating. Although robotic sampling missions might in principle be able to achieve these objectives, we suggest that they would most effectively be accomplished by future human expeditions (e.g. Spudis,

2001; Crawford, 2006). We have elaborated elsewhere on the operational requirements of a human lunar exploration architecture that would be able to meet these objectives (Crawford et al., 2007).

6. Conclusions The principal conclusions of this work are: (1) Extraction of exogenously-implanted materials from precisely-dated lunar regolith layers has the potential to yield a record of the evolution and variability of the solar wind and galactic cosmic rays throughout the more than 3 Ga represented by the Moon’s volcanic history. (2) The low thermal conductivity of particulate regolith restricts heating of the substrate to shallow layers, thus ensuring that samples of solar wind derived volatiles will remain protected from outgassing at depths as shallow as 3.7–38 cm (beneath the distal margins of 1–10 m thick lava flows; low-cl, Table 4). In locations closer to the vent, continued advective heat transfer could cause volatile outgassing for tens to hundreds of centimeters beneath the flow, as suggested by our high-cl results (Table 4). (3) Mare basalt flow units emplaced tens to hundreds of Ma apart have the potential to generate and entrap paleoregolith deposits suitable for extraction, dating and analysis of the evolution of the solar wind. Similar analyses can be conducted for the preservation of volatiles produced by solar flare events and galactic cosmic rays. Lunar paleoregolith deposits therefore provide a potentially very valuable archive of early Solar System history. We argue that they should be a high priority objective of future lunar exploration efforts, and have provided some recommendations for future exploration strategies that would facilitate their detection and recovery.

Acknowledgments This work was funded in part by NASA Lunar Advanced Science and Exploration Research Program Grant NNX08AY75G. I.A.C. and K.H.J. thank the Leverhulme Trust for financial support. I.A.C. wishes to acknowledge Lionel Wilson for helpful discussions and encouragement. This is HIGP Publication 1827 and SOEST Publication 7851.

References Armstrong, J.C., Wells, L.E., Gonzales, G., 2002. Rummaging through the Earth’s attic for remains of ancient life. Icarus 160, 183–196. Basaltic Volcanism Study Project, 1981. Basaltic Volcanism on the Terrestrial Planets. Pergamon Press, New York. 1286 pp. Basford, J.R., Dragon, J.C., Pepin, R.O., Coscio, M.R., Murthy, V.R., 1973. Krypton and xenon in lunar fines. Proc. Lunar Sci. Conf. 4, 1915–1955. Becker, R.H., 1980. Light elements in lunar soils revisited: Carbon, nitrogen, hydrogen, and helium. Proc. Lunar Sci. Conf. 11, 1743–1761. Becker, R.H., Clayton, R.N., 1977. Nitrogen isotopes in lunar soils as a measure of cosmic-ray exposure and regolith history. Proc. Lunar Sci. Conf. 17, 3685–3704. Becker, R.H., Clayton, R.N., 1978. Nitrogen isotope systematics of two Apollo 12 soils. Proc. Lunar Sci. Conf. 9, 1619–1627. Becker, R.H., Pepin, R.O., 1989. Long-term changes in solar wind elemental and isotopic ratios: A comparison of two lunar ilmenites of different antiquities. Geochim. Cosmochim. Acta 53, 1135–1146. Bottinga, Y., Weill, D.F., 1970. Densities of liquid silicate systems calculated from partial molar volumes of oxide components. Am. J. Sci. 269, 169–182. Bruce, P.M., Huppert, H.E., 1990. Solidification and melting along dikes by the laminar flow of basaltic magma. In: Ryan, M.P. (Ed.), Magma Transport and Storage. John Wiley and Sons Ltd., pp. 87–101. Bussey, D.B.J., Sørenson, S.-A., Guest, J.E., 1997. On the role of thermal conductivity on thermal erosion by lava. J. Geophys. Res. 102, 10905–10908.

S.A. Fagents et al. / Icarus 207 (2010) 595–604 Büttner, R., Zimanowski, B., Blumm, J., Hagemann, L., 1998. Thermal conductivity of a volcanic rock material (olivine–melilitite) in the temperature range between 288 and 1470 K. J. Volcanol. Geotherm. Res. 80, 293–302. Carr, M.H., 1974. The role of lava erosion in the formation of lunar rilles and martian channels. Icarus 22, 1–23. Carrier, W.D., Olhoeft, G.R., Mendell, W.W., 1991. Physical properties of the lunar surface. In: Heiken, G.H., Vaniman, D.T., French, B.M. (Eds.), The Lunar Sourcebook: A User’s Guide to the Moon. Cambridge University Press, Cambridge, pp. 475–594. Carslaw, H.S., Jaeger, J.C., 1986. Conduction of Heat in Solids. Oxford University Press, Oxford. 510 pp. ˇ ermàk, V., Huckenholz, H.-G., Rybach, L., Schmid, R., Schopper, J.R., Schuch, M., C Stöffler, D., Wohlenberg, J. (Eds.), 1982. Physical Properties of Rocks. Numerical Data and Functional Relationships in Science and Technology, vol. 1a. SpringerVerlag, New York. 604 pp. Chaussidon, M., 2008. The early Moon was rich in water. Nature 454, 170–172. Clark, R.N., 2009. Detection of adsorbed water and hydroxyl on the Moon. Science 326, 562–564. doi:10.1126/science.1178105. Clayton, R.N., Thiemans, M.H., 1980. Lunar nitrogen: Evidence for secular change in the solar wind. In: Pepin, R.O., Eddy, J.A., Merrill, R.B. (Eds.), The Ancient Sun: Fossil Record in the Earth, Moon, and Meteorites. Pergamon, New York, pp. 162– 164. Coombs, C.R., Hawke, B.R., Wilson, L., 1988. Geologic and remote sensing studies of Rima Mozart. Proc. Lunar Planet. Sci. Conf. 27, 339–353. Crawford, I.A., 2004. The scientific case for renewed human activities on the Moon. Space Policy 20, 91–97. Crawford, I.A., 2006. The astrobiological case for renewed robotic and human exploration of the Moon. Int. J. Astrobiol. 5, 191–199. Crawford, I.A., Fagents, S.A., Joy, K.H., 2007. Full Moon exploration. Astron. Geophys. 48, 3.18–3.21. Crawford, I.A., Baldwin, E.C., Taylor, E.A., Bailey, J.A., Tsembelis, K., 2008. On the survivability and detectability of terrestrial meteorites on the Moon. Astrobiology 8, 242–252. Crawford, I.A., Joy, K.H., Fagents, S.A., Rumpf, M.E., 2009. The importance of lunar palaeoregolith deposits and the role of Lunar Reconnaissance Orbiter. Lunar Reconnaissance Orbiter Science Targeting Meeting, June 9–11, 2009. Tempe, Arizona, p. 29 (LPI Contribution No. 1483). Crisp, J., Kahle, A.B., Abbott, E.A., 1990. Thermal infrared spectral character of Hawaiian basaltic glasses. J. Geophys. Res. 95, 21657–21669. Crozaz, G., Haack, U., Hair, M., Maurette, M., Walker, R., Woolum, D., 1970. Nuclear track studies of ancient solar radiations and dynamic lunar surface processes. Proc. Apollo 11 Lunar Sci. Conf. 3, 2051–2080. Dran, J.C., Durrieu, L., Jouret, C., Maurette, M., 1970. Habit and texture studies of lunar and meteoritic materials with 1 MeV electron microscope. Earth Planet. Sci. Lett. 9, 391–400. Drozd, R.J., Hohenberg, C.M., Morgan, C.J., Podosek, F.A., Wroge, M.L., 1977. Cosmicray exposure history of Taurus-Littrow. Proc. Lunar Sci. Conf. 8, 3027–3043. Drozd, R.J., Hohenberg, C.M., Morgan, C.J., Ralston, C.E., 1974. Cosmic-ray exposure history at the Apollo 16 and at other lunar sites: Lunar surface dynamics. Geochim. Cosmochim. Acta 38, 1625–1642. Duraud, J.P., Langevin, Y., Maurette, M., Comstock, G., Burlingame, A.L., 1975. The simulated depth history of dust grains in the lunar regolith. Proc. Lunar Sci. Conf. 6, 2397–2415. Eberhardt, P., Geiss, J., Graf, H., Grögler, N., Mendia, M.D., Mörgelli, M., Schwaller, H., Stettler, A., Krähenbühl, U., von Gunten, H.R., 1972. Trapped solar wind noble gases in Apollo 12 lunar fines 12002, and Apollo 11 breccia 10046. Proc. Lunar Sci. Conf. 3, 1821–1856. Fagents, S.A., Greeley, R., 2001. Factors influencing lava–substrate heat transfer and implications for thermomechanical erosion. Bull. Volcanol. 62, 519–532. Fegley, B., Swindle, T.D., 1993. Lunar volatiles: Implications for lunar resource utilization. In: Lewis, J., Matthews, M.S., Guerrieri, M.L. (Eds.), Resources of Near-Earth Space. University of Arizona Press, Tucson, pp. 367–426. Fleischer, R.L., Haines, E.L., Hanneman, R.E., Hart Jr., H.R., Kasper, J.S., Lifshin, E., Woods, R.T., 1970. Particle track, X-ray, thermal, and mass spectrometric studies of lunar material. Science 167, 568–571. Flynn, L.P., Mouginis-Mark, P.J., Gradie, J.C., Lucey, P.G., 1993. Radiative temperature measurements at Kupaianaha Lava Lake, Kilauea Volcano, Hawaii. J. Geophys. Res. 98, 6461–6476. Fountain, J.A., West, E.A., 1970. Thermal conductivity of particulate basalt as a function of density in simulated lunar and martian environments. J. Geophys. Res. 75, 4063–4069. Geiss, J., 1973. Solar wind composition and implications about the history of the Solar System. Papers Presented to the 13th Int. Cosmic Ray Conf., pp. 3375– 3398. Geiss, J., Bochsler, P., 1991. Long time variations in solar wind properties: Possible causes versus observations. In: Sonett, C.P., Giampapa, M.S., Matthews, M.S. (Eds.), The Sun in Time. University of Arizona Press, Tucson, pp. 98–117. Ghiorso, M.S., Sack, R.O., 1995. Chemical mass transfer in magmatic processes: IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid equilibria in magmatic systems at elevated temperatures and pressures. Contrib. Mineral. Petrol. 119, 197–212. Gibson, E.K., Johnson, S.M., 1971. Thermal analysis – Inorganic gas release studies of lunar samples. Proc. Lunar Sci. Conf. 2, 1351–1366. Gifford, A.N., El-Baz, F., 1981. Thicknesses of mare flow fronts. Moon Planets 24, 391–398.

603

Gough, D.O., 1981. Solar interior structure and luminosity variations. Solar Phys. 74, 21–34. Haqq-Misra, J.D., Domagal-Goldman, S.D., Kasting, P.J., Kasting, J.F., 2008. A revised hazy methane greenhouse for the Archean Earth. Astrobiology 8, 1127–1137. Haskin, L., Warren, P.H., 1991. Lunar chemistry. In: Heiken, G.H., Vaniman, D.T., French, B.M. (Eds.), The Lunar Sourcebook: A User’s Guide to the Moon. Cambridge University Press, Cambridge, pp. 357–474. Heggy, E., Fong, T., Kring, D., Deans, M., Anglade, A., Mahiouz, K., Bualat, M., Lee, P., Horz, F., Bluethmann, W., 2009. Potential of probing the lunar regolith using rover-mounted ground penetrating radar: Moses Lake dune field analog study. Lunar Planet. Sci. Conf. 40. Abstract #2183 (CD-ROM). Hemingway, B.S., Robie, R.A., 1973. Specific heats of lunar basalt, 15555, and soils 15301 and 60601 from 90 to 350 K. Proc. Lunar Sci. Conf. 3, 355–356. Hemingway, B.S., Robie, R.A., Wilson, W.H., 1973. Specific heats of lunar soils, basalt, and breccias from the Apollo 14, 15, and 16 landing sites, between 90 and 350 K. Proc. Lunar Planet. Sci. Conf. 4, Geochim. Cosmochim. Acta (Suppl. 3), 2481–2487. Hiesinger, H., Head, J.W., 2006. New views of lunar geoscience: An introduction and overview. In: New Views of the Moon. Rev. Min. Geochem. 60, 1–81. Hiesinger, H., Head, J.W., Wolf, U., Jaumann, R., Neukum, G., 2002. Lunar mare basalt flow units: Thicknesses determined from crater size-frequency distributions. Geophys. Res. Lett. 29. doi:10.1029/2002Gl014847. Hiesinger, H., Head, J.W., Wolf, U., Jaumann, R., Neukum, G., 2003. Ages and stratigraphy of mare basalts in Oceanus Procellarum, Mare Nubium, Mare Cognitum, and Mare Insularum. J. Geophys. Res. 108 (5065). doi:10.1029/ 2002JE001985. Hiesinger, H., Jaumann, R., Neukum, G., Head, J.W., 2000. Age of mare basalts on the lunar nearside. J. Geophys. Res. 105, 29239–29275. Hohenberg, C.M., Davis, P.K., Kaiser, W.A., Lewis, R.S., Reynolds, J.H., 1970. Trapped and cosmogenic rare gases from stepwise heating of Apollo 11 samples. Proc. Apollo 11 Lunar Sci. Conf., 1283–1309. Hörz, F., Grieve, R., Heiken, G.H., Spudis, P.D., Binder, A., 1991. Lunar surface processes. In: Heiken, G.H., Vaniman, D.T., French, B.M. (Eds.), The Lunar Sourcebook: A User’s Guide to the Moon. Cambridge University Press, Cambridge, pp. 61–120. Hulme, G., 1973. Turbulent lava flow and the formation of lunar sinuous rilles. Mod. Geol. 4, 107–117. Hulme, G., 1982. A review of lava flow processes related to the formation of lunar sinuous rilles. Geophys. Surv. 5, 107–117. Jaeger, J.C., 1961. The cooling of irregularly shaped igneous bodies. Am. J. Sci. 259, 721–734. Kerridge, J.F., 1975. Solar nitrogen: Evidence for a secular change in the ratio of nitrogen-15 to nitrogen-14. Science 188, 162–164. Kerridge, J.F., 1980. Secular variations in composition of the solar wind: Evidence and causes. In: Pepin, R.O., Eddy, J.A., M errill, R.B. (Eds.), The Ancient Sun: Fossil Record in the Earth, Moon, and Meteorites. Pergamon, New York, pp. 475–489. Kerridge, J.F., 1989. What has caused the secular increase in solar nitrogen-15? Science 245, 480–486. Kerridge, J.F., Signer, P., Wieler, R., Becker, R.H., Pepin, R.O., 1991. Long-term changes in composition of solar particles implanted in extraterrestrial materials. In: Sonett, C.P., Giampapa, M.S., Matthews, M.S. (Eds.), The Sun in Time. University of Arizona Press, Tucson, pp. 98–117. Korotev, R.L., 1997. Some things we can infer from the Moon from the composition of the Apollo 16 regolith. Meteor. Planet. Sci. 32, 447–478. Lange, R.A., Cashman, K.V., Navrotsky, A., 1994. Direct measurements of latent heat during crystallization and melting of a ugandite and an olive basalt. Contrib. Mineral. Petrol. 118, 169–181. Lange, R.A., Navrotsky, A., 1992. Heat capacities of Fe2O3-bearing silicate liquids. Contrib. Mineral. Petrol. 110, 311–320. Langevin, Y., Arnold, J.R., 1977. The evolution of the lunar regolith. Annu. Rev. Earth Planet. Sci. 5, 449–489. Langseth, M.G., Keihm, S.J., Peters, K., 1976. Revised lunar heat-flow values. Proc. Lunar Sci. Conf. 7, 3143–3171. Lorenzetti, S., Busemann, H., Eugster, O., 2005. Regolith history of lunar meteorites. Meteor. Planet. Sci. 40, 315–327. Lucey, P.G., and 17 colleagues, 2006. Understanding the lunar surface and space– Moon interaction. In: New Views of the Moon. Rev. Min. Geochem. 60, 82–219. McKay, D.S., 2009. Do lunar pyroclastic deposits contain the secrets of the Solar System? Lunar Reconnaissance Orbiter Science Targeting Meeting, June 9–11, 2009. Tempe, Arizona, pp. 77–78 (LPI Contribution No. 1483). McKay, D.S., Fruland, R.M., Heiken, G.H., 1974. Grain size and evolution of lunar soils. Proc. Lunar Sci. Conf. 5, 887–906. McKay, D.S., Bogard, D.D., Morris, R.V., Korotev, R.L., Wentworth, S.J., Johnson, P., 1989. Apollo 15 regolith breccias – Window to a KREEP regolith. Proc. Lunar Planet. Sci. Conf. 19, 19–41. McKay, D.S., Heiken, G.H., Basu, A., Blanford, G., Simon, S., Reedy, R., French, B.M., Papike, J., 1991. The lunar regolith. In: Heiken, G.H., Vaniman, D.T., French, B.M. (Eds.), The Lunar Sourcebook: A User’s Guide to the Moon. Cambridge University Press, Cambridge, pp. 285–356. Murase, R., McBirney, A.T., 1970. Viscosity of lunar lavas. Science 167, 1491–1492. Murase, R., McBirney, A.T., 1973. Properties of some common igneous rocks and melts at high temperatures. Bull. Geol. Soc. Am. 84, 3562–3592. Navrotsky, A., 1995. Energetics of silicate melts. In: Stebbins, J.F., McMillan, P.F., Dingwell, D.B. (Eds.), Structure, Dynamics and Properties of Silicate Melts. Mineral. Soc. Am., Washington, DC, pp. 121–142.

604

S.A. Fagents et al. / Icarus 207 (2010) 595–604

NRC, 2007. The Scientific Context for Exploration of the Moon. National Research Council, Washington, DC. . Nishiizumi, K., Arnold, J.R., Kohl, C.P., Caffee, M.W., Masarik, J., Reedy, R.C., 2009. Solar cosmic ray records in lunar rock 64455. Geochim. Cosmochim. Acta 73, 2163–2176. Norris, S.J., Swart, P.K., Wright, I.P., Grady, M.M., Pillinger, C.T., 1983. A search for collectable, isotopically light carbon and nitrogen components in lunar soils and breccias. Proc. Lunar Sci. Conf. 14, J. Geophys. Res. 88, B200–210. Oberbeck, V.R., Quaide, W.L., 1968. Genetic implications of lunar regolith thickness variations. Icarus 9, 446–465. Ono, T., Kumamoto, A., Nakagawa, H., Yamaguchi, Y., Oshigami, S., Yamaji, A., Kobayashi, T., Kasahara, Y., Oya, H., 2009. Lunar Radar Sounder observations of subsurface layers under the nearside maria of the Moon. Science 323, 90–92. Ozima, M., Seki, K., Terada, N., Miura, Y.N., Podosek, F.A., Shinagawa, M., 2005. Terrestrial nitrogen and noble gases in lunar soils. Nature 436, 655–659. Ozima, M., Yin, Q.-Z., Podosek, F.A., Miura, Y.N., 2008. Toward prescription for approach from terrestrial noble gas and light element records in lunar soils understanding early Earth evolution. Proc. Natl. Acad. Sci. 105, 17654–17658. Pepin, R.O., 1980. Rare gases in the past and present solar wind. In: Pepin, R.O., Eddy, J.A., Merrill, R.B. (Eds.), The Ancient Sun: Fossil Record in the Earth, Moon, and Meteorites. Pergamon, New York, pp. 1135–1146. Pepin, R.O., Nyquist, L.E., Phinney, D., Black, D.C., 1970. Rare gases in Apollo 11 lunar material. Proc. Apollo 11 Lunar Sci. Conf., 1435–1454. Pieri, D.C., Glaze, L.S., Abrams, M.J., 1990. Thermal radiance observations of an active lava flow during the June 1984 eruption of Mount Etna. Geology 18, 1018–1022. Pieters, C.M., and 28 colleagues, 2009. Character and spatial distribution of OH/H2O on the surface of the Moon seen by M3 on Chandrayaan-1. Science 326, 568– 572. doi:10.1126/science.1178658. Pinkerton, H., James, M., Jones, A., 2002. Surface temperature measurements of active flows on Kilauea volcano, Hawaii. J. Volcanol. Geotherm. Res. 113, 159– 176. Reedy, R.C., Arnold, J.R., Lal, D., 1983. Cosmic-ray record in Solar System matter. Annu. Rev. Nucl. Particle Sci. 33, 505–537. Retallack, G.J., 2005. Palaeosols. In: Selley, R.C., Cocks, L.R.M., Plimer, I.R. (Eds.), Encyclopedia of Geology. Academic Press, London, pp. 203–208. Robie, R.A., Hemingway, B.S., Wilson, W.H., 1970. Specific heats of lunar surface materials from 90 to 350 K. Proc. Lunar Sci. Conf. 2, Geochim. Cosmochim. Acta 3 (Suppl. 1), 2361–2367. Saal, A.E., Hauri, E.H., Lo Cascio, M., Van Orman, J.A., Rutherford, M.C., Cooper, R.F., 2008. Volatile content of lunar volcanic glasses and the presence of water in the Moon’s interior. Nature 454, 192–195. doi:10.1038/nature07047. Sackmann, I.-J., Boothroyd, A.I., 2003. Our Sun V: A bright young Sun consistent with helioseismology and warm temperatures on ancient Earth and Mars. Astrophys. J. 583, 1024–1039. Sagan, C., Mullen, G., 1972. Earth and Mars: Evolution of atmospheres and surface temperatures. Science 177, 52–56. Sharpton, V.L., Head, J.W., 1982. Stratigraphy and structural evolution of southern Mare Serenitatis: A reinterpretation based on Apollo Lunar Sounder Experiment data. J. Geophys. Res. 87, 10963–19998. Shoemaker, E.M., Morris, E.C., Batson, R.M., Holt, H.E., Larsen, K.B., Montgomery, D.R., Rennilson, J.J., Whitaker, E.A., 1968. Television Observations from Surveyor,

Surveyor Project Final Report, Part II. Jet Propulsion Laboratory, Pasadena, pp. 22–136. Simoneit, B.R., Christiansen, P.C., Burlingame, A.L., 1973. Volatile element chemistry of selected lunar, meteoritic, and terrestrial samples. Proc. Lunar Sci. Conf. 4, Geochim. Cosmochim. Acta 2 (Suppl. 4), 1635–1650. Spudis, P.D., 1996. The Once and Future Moon. Smithsonian Institution Press, Washington, DC. 308 pp. Spudis, P.D., 2001. The case for renewed human exploration of the Moon. Earth, Moon, Planets 87, 159–171. Spudis, P.D., Taylor, G.J., 2009. A major KREEP-basalt–mare basalt unconformity on the Moon. Lunar Planet. Sci. Conf. 40. Abstract #1039. Srinivasan, B., Hennecke, E.W., Sinclair, D.E., Manuel, O.K., 1972. A comparison of noble gases released from lunar fines (#15601.64) with noble gases in meteorites and in the Earth. Proc. Lunar Sci. Conf. 3, 1927–1945. Sunshine, J.M., Farnham, T.L., Feaga, L.M., Groussin, O., Merlin, F., Milliken, R.E., A’Hearn, M.F., 2009. Temporal and spatial variability of lunar hydration as observed by the Deep Impact spacecraft. Science 326, 565–568. doi:10.1126/ science.1179788. Taylor, G.J., 1985. The need for a lunar base: Answering basic questions about planetary science. In: Mendell, W.W. (Ed.), Lunar Bases and Space Activities of the 21st Century. Lunar and Planetary Institute, Houston, pp. 189–197. Taylor, G.J., Warren, P.H., Ryder, G.C., Delano, J.W., Pieters, C.M., Lofgren, G.W., 1991. Lunar rocks. In: Heiken, G.H., Vaniman, D.T., French, B.M. (Eds.), The Lunar Sourcebook: A User’s Guide to the Moon. Cambridge University Press, Cambridge, pp. 183–284. Taylor, S.R., 1982. Planetary Science: A Lunar Perspective. Lunar and Planetary Institute, Houston. 481 pp. Thiemans, M.H., Clayton, R.N., 1980. Solar and cosmogenic nitrogen in the Apollo 17 deep drill core. Proc. Lunar Sci. Conf. 11, 1435–1451. Touloukian, Y.S., Judd, W.R., Roy, R.F., 1981. Physical Properties of Rocks and Minerals. Data Series on Material Properties, vols. 1–2. McGraw-Hill, New York. 548 pp. Turcotte, D.L., Schubert, G., 2002. Geodynamics. Cambridge University Press. 456pp. USGS, 1999. Groundwater Atlas of the United States: Alaska, Hawaii, Puerto Rico and the US Virgin Islands. USGS HA730-N. . Vaniman, D.T., Reedy, R., Heiken, G.H., Olhoeft, G., Mendell, W.W., 1991. The lunar environment. In: Heiken, G.H., Vaniman, D.T., French, B.M. (Eds.), The Lunar Sourcebook: A User’s Guide to the Moon. Cambridge University Press, Cambridge, pp. 27–60. Wieler, R., Kehm, K., Meshik, A.P., Hohenberg, C.M., 1996. Secular changes in the xenon and krypton abundances in the solar wind recorded in single lunar grains. Nature 384, 46–49. Whitmire, D.P., Doyle, L.R., Reynolds, R.T., Matese, J.J., 1995. A slightly more massive young Sun as an explanation for warm temperatures on early Mars. J. Geophys. Res. 100 (E3), 5457–5464. Wilhelms, D., 1987. Geologic History of the Moon. US Geol. Surv. Prof. Paper 1348. Wood, B.E., Muller, H.-R., Zank, G., Linsky, J.L., 2002. Measured mass loss rates of solar-like stars as a function of age and activity. Astrophys. J. 574, 412–425. Williams, D.A., Fagents, S.A., Greeley, R., 2000. A reassessment of the emplacement and erosional potential of turbulent, low-viscosity lavas on the Moon. J. Geophys. Res. 105, 20189–20205.