Basaltic magmatism and the bulk composition of the moon

A. E. Ringwood

doi:10.1007/bf00577901

Basaltic magmatism and the bulk composition of the moon

Ringwood, A. E. 1977-07-01 00:00:00 I. Mafor and Heat-Producing Elements A. E. RINGWOOD Research School of Earth Sciences, Australian National University, Canberra, Australia (Received 21 January, 1977) Abstract. The lunar interior is comprised of two major petrological provinces: (1) an outer zone several hundred km thick which experienced partial melting and crystallization differentiation 4.4- 4.6 b.y. ago to form the lunar crust together with an underlying complementary zone of ultramafic cumulates and residua, and (2) the primordial deep interior which was the source region for mare basalts (3.2-3.8 b.y.) and had previously been contaminated to varying degrees with highly fractionated material derived from the 4.4-4.6 b.y. differentiation event. In both major petrologic provinces, basaltic magmas have been produced by partial melting. The chemical characteristics and high-pressure phase relationships of these magmas can be used to constrain the bulk compositions of their respective source regions. Primitive low-Ti mare basalts (e.g., 12009, 12002, 15555 and Green Glass) possessing high norma- tive olivine and high Mg and Cr contents, provide the most direct evidence upon the composition of the primordial deep lunar interior. This composition, as estimated on the basis of high pressure equi- libria displayed by the above basalts, combined with other geochemical criteria, is found to consist of orthopyroxene + clinopyroxene + olivine with total pyroxenes > olivine, 100 MgO/(MgO + FeO) = 75-80, about 4% of CaO and A120 a and 2X chondritic abundances of REE, U and Th. This compo- sition is similar to that of the earth's mantle except for a higher pyroxene/olivine ratio and lower 100 MgO/(MgO + FeO). The lunar crust is believed to have formed by plagioclase elutriation within a vast ocean of parental basaltic magma. The composition of the latter is found experimentally by removing liquidus plagioclase from the observed mean upper crust (gabbroic anorthosite) composition, until the resulting compo- sition becomes multiply saturated with plagioclase and a ferromagnesian phase (olivine). This parental basaltic composition is almost identical with terrestrial oceanic tholeiites, except for partial depletion in the two most volatile components, Na20 and SiO 2. Similarity between these two most abundant classes of lunar and terrestrial basaltic magmas strongly implies corresponding similarities between their source regions. The bulk composition of the outer 400 km of the Moon as constrained by the 4.6-4.4 b.y. parental basaltic magma is found to be peridotitic, with olivine > pyroxene, 100 MgO/ (MgO + FeO) ~ 86, and about 2X chondritic abundances of Ca, A1 and REE. The Moon thus appears to have a zoned structure, with the deep interior (below 400 km) possessing somewhat higher contents of FeO and SiO 2 than the outer 400 kin. This zoned model, derived exclusively on petrological grounds, provides a quantitative explanation of the Moon's mean density, moment of inertia and seismic velocity profile. The bulk composition of the entire Moon, thus obtained, is very similar to the pyrolite model com- position for the Earth's mantle, except that the Moon is depleted in Na (and other volatile elements) and somewhat enriched in iron. The similarity in major element composition extends also to the abundances of REE, U and Th. These compositional similarities, combined with the identity in oxygen isotope ratios between the Moon and the Earth's mantle, are strongly suggestive of a common genetic relationship. 1. Introduction One of the most significant results arising from the Apollo project was the demonstration that the lunar maria are composed of rocks resembling terrestrial basalts. Moreover it is The Moon 16 (1977) 389-423. All Rights Reserved. Copyright © 1977 by D. Reidel Publishing Company, Dordrecht-Holland. 390 A.E. RINGWOOD becoming increasingly clear that the lunar crust was ultimately derived by differentiation from a parental magma ocean of basaltic affinities (Section 5). In recent years, important progress has been made towards understanding the origin of terrestrial basaltic rocks. It is now widely accepted that they formed by partial melting of an ultramafic source rock in the mantle. The observed spectrum of basaltic compositions is interpreted primarily in terms of the degree of partial melting of the source material, the depth of partial melting and the subsequent crystallization history during ascent to the surface. High pressure, high temperature investigations of the crystallization behaviour of terrestrial basalts using the methods of experimental petrology have been successfully employed to constrain the nature of their source regions (Green and Ringwood, 1967; Green, 1970; Ringwood, 1975a). It is natural to attempt to understand the origin of lunar basalts in terms of processes analogous to those which have operated in the petrogenesis of terrestrial basalts and like- wise, to use experimental petrology to provide information on the nature of their source regions. The first comprehensive attempt in this direction was presented by Ringwood and Essene (1970a, b) at the Apollo 11 Lunar Science Conference in January, 1970. In view of controversies which later developed in this field, it is worth remarking that several of the key conclusions reached by these workers have subsequently proved to be well founded. A principal conclusion (see also, Ringwood, 1970)was that Apollo 11 high K and low K magmas had been produced mainly by partial melting processes in the lunar interior rather than by extensive fractional crystallization in huge lava lakes, as advocated, for example, by O'Hara et al. (1970a). The former interpretation was at first very much a minority viewpoint, as can be seen in the official summary of the Apollo 11 Conference (LSAPT, 1970, p. 450) but has since become widely accepted. Although varying degrees of partial melting were believed to be mainly responsible for the chemical diversity, it should be noted that Ringwood and Essene (1970b, pp. 784,791) did not exclude the possibility of moderate degrees (up to 30 percent) of fractional crystallization of olivine and pyroxene. Subsequent researches by several groups (e.g. James and Jackson, 1970; Compston et al., 1971; Kushiro and Haramura, 1971; Green et al., 1971a; James and Wright, 1972; Chappell and Green, 1973; Rhodes and Hubbard, 1973; Walker et al., 1975a; Longhi et al., 1974; Shih et al., 1975) have greatly clarified the relative import- ance of partial melting versus fractional crystallization in explaining the chemical diversity among mare basalts. The widespread occurrence of near-surface fractionation controlled by the separation of olivine-+ Cr spinel + Fe-Ti oxides has been clearly documented but in the great majority of cases, this has involved the crystallization of less than 30% of these minerals. The samples representing the most primitive, i.e., the least fractionated magmas can be recognised by their high Mg and Cr contents and, in the cases of low-Ti basalts, by high contents of normative olivine. In many such samples textural and mineralogical evidence indicates that they were erupted at the surface as liquids (Green et al., 1971a, 1975; Longhi et aI., 1972; Kesson, 1975; Walker et al., 1975a, 1976). These primitive liquids are believed to represent close BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 391 approximations to primary magmas formed by direct partial melting in the lunar interior. A second controversial area concerned the role of plagioclase in mare basalt petro- genesis. Several workers including Ringwood and Essene (1970b) concluded that Apollo 11 magmas were undersaturated with plagioclase, and therefore that plagioclase could not have been present as a residual phase in their source regions. This conclusion was extended to cover most Apollo 12 basalts by Green et al. (1971a) but was strongly disputed by O'Hara et al. (1970a, b), and Biggar et al. (1971). It was opposed also by trace element geochemists, (e.g. Gast et al., 1970; Haskin et al., 1970; Philpotts and Schnetzler, 1970; Gast, 1972) who required the presence of plagioclase, in the source regions in order to explain the negative europium anomalies of mare basalts. The conflict has since been resolved by several detailed experimental investigations (Green et al., ,1971b; Ringwood and Green, 1972; Green and Ringwood, 1973; Grove et al., 1973 ;Kushiro, 1972; Longhi et al., 1972, 1974; Walker et al., 1975a, 1976; Kesson, 1975). Whilst a few basalts (e.g. 12038, 14053, Luna 16) and several Apollo 11 ophitic basalts) were probably plagioclase-saturated upon eruption, the vast majority of Apollo 11, 12, 15 and 17 basalts were definitely undersaturated with plagloclase. Moreover, those magmas which could be demonstrated to have undergone the least near-surface crystallization differentiation and to have reached the surface essentially as liquids (e.g. 12009, 12002, 15555, 15016, Green Glass, 70215, 74275) were among those most under- saturated with plagioclase. These results firmly demonstrated that plagioclase was not a residual phase remaining in the source region after partial melting, so that the europium anomaly must be a characteristic inherited from the source region. This conclusion has had far-reaching implications for subsequent petrogenetic hypotheses. Moreover, the observed degrees of plagioclase undersaturation provided an additional limit to the permissible degrees of preeruptive crystal fractionation. A third conclusion reached by Ringwood and Essene (1970a, b) was that the source region of Apollo 11 basalts was composed mainly of orthopyroxene and subcalcic clino- pyroxene with olivine possibly present as a subsidiary phase. The source region contained 3 to 5% of CaO and A12 03 and had an Mg number (100 MgO/(MgO + FeO)) of 75 to 80. A corresponding study of Apollo 12 low-Ti basalts by Green etal. (1971a)and Green et al. (1971b) led to the conclusion that their source region likewise contained similar amounts of CaO and A12Oa and possessed a similar Mg number. It seemed likely, however, that olivine was also an important mineral in the Apollo 12 source region. The source was thus inferred to be an olivine pyroxenite. This initial interpretation of the source region of Apollo 11 and 12 basalts was extended to Apollo 15 and 17 basalts (Green and Ringwood, 1973; Chappell and Green, 1973; Green et al., 1975) and was regarded as representative of the bulk of the lunar interior (Ringwood, 1970, 1975b, 1976a). The conclusion that the bulk Moon contained only 3 to 5% of CaO and A1203, i.e., about twice the chondritic abundances, has been generally disregarded by most lunar scientists. Lunar bulk compositional models containing much higher abundances of CaO and Al:O3 have been advocated by many workers in the field. 392 A.E. RINGWOOD Models which have received a considerable degree of favourable attention include those of Taylor and Jake~, 1974 (8.1% Al203, 6.6% CaO), Ganapathy and Anders, 1974 (11.6% Al203, 9.3% CaO), Wgnke et al., 1974 (17.4% A1203, 13.6% CaO) and Anderson, 1973 (27.2% Al203, 22.1% CaO). We will demonstrate in the present paper that the early compositional models contain- ing relatively smaller amounts of CaO and Al203 are much closer to reality. 2. Mare Basalt Petrogenesis Early hypotheses of mare basalt petrogenesis based upon experimental petrology proposed that all classes of mare basalts had formed by varying degrees of partial melting of a common olivine-pyroxenite source at depths of 150-500 km (e.g., Ringwood and Essene, 1970b). Although this straightforward single-stage hypothesis provided an adequate explanation of many aspects of the major element chemistry of mare basalts and their source regions, it was unable to provide satisfactory explanations of other geochemical characteristics of high-Ti and low-Ti mare basalts, e.g., their REE patterns, Eu-anomalies, and TiO2 contents. In addition, the earlier differentiation event implied for mare basalt source regions by Sm-Nd (Lugmair et al., 1975) and U-Pb isotope systematics (Tera and Wasserburg, 1975) was not accounted for. To meet these difficulties, a second class of hypotheses was developed (e.g., Schnetzler and Philpotts, 1971) which maintained that mare basalts could have formed by the remelt- ing of chemically and mineralogically inhomogeneous olivine + pyroxene + ilmenite cumu- lates formed during the early differentiation of the Moon around 4.6-4.4 b.y., as the complement of the plagioclase-rich crust. Specifically, it was suggested (e.g. Walker et al., 1975a) that low-Ti basalts had formed by the partial melting of early olivine + pyroxene cumulates at considerable depths (200-400 kin) whereas the high-Ti basalts formed by partial melting of a zone of late-stage olivine + pyroxene + ilmenite cumulates at relatively shallow depths (aroun d 100 kin). The cumulate remelting hypothesis offers, in principle, an explanation for several important geochemical characteristics of mare basalts, e.g., their different TiO2 contents, the two-stage history recorded by Sm-Nd and Pb-U isotopes and the Eu anomalies and their varying magnitudes. However, this hypothesis still encounters a number of fatal difficulties. For example, it is unable to explain the obser- vation that the least fractionated high-Ti and low-Ti basalts have similar Cr contents and Mg numbers. This and other shortcomings have been discussed in detail by Ringwood (1975b) and Kesson and Ringwood (1976a). Although the cumulate-remelting hypothesis is not acceptable in its present form, its success in explaining several key features Of the trace-element and isotopic geochemistry of mare basalts strongly suggests that some form of reprocessing of cumulates is involved in mare basalt petrogenesis. This raises the possibility of combining the more attractive aspects of the single-stage, uniform, primordial source-region hypothesis with those of the cumulate-remelting hypothesis. Kesson and Pdngwood (1976a) and Ringwood and Kesson (1976a)have proposed a BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 393 ~ 10049 10017] Hi K ]0058 "'~-- ' 10020 4O Q; ~ 700]7 e~ ' ~" ...... 12051 -0 6 2c ~'~.'_-. =.-- ..... ]2002 ¢- .=_ uJ ill • "~ ,v LU ev, ~AVERAGE APOLLO 15 .~" ...... . .L.j.~.~.~.. Q. E ]( O. O. "~'" 15555 ....... ~ ...... GREEN GLASS I I I I I I I I I La Ce Nd Sm Eu Gd Dy Er Yb Fig. 1. Abundances of rare earths in a representative suite of mare basalt samples. After Haskin et al. (1970), Gast et al. (1970), Hubbard and Gast (1971), Helmke et al. (1972, 1973), Philpotts et al. (1972), Ridley et al. (1973), Shih et al. (1975). hypothesis in this category which appears to avoid most of the difficulties associated with previous petrogenetic hypotheses• This involves hybridization or assimilation of 4.6-4.4 b.y. differentiates by primordial material from the deep lunar interior• Early versions of hybridization hypotheses were proposed by Anderson (1971) and Hubbard and Minear (1975) and involve assimilation/hybridization at shallow levels• The Kesson- Ringwood model differs from these in that the hybridization is believed to have occurred at depth (about 400 km) and was caused by sinking of pods of 4.6-4•4 b.y. dense residual cumulates through the underlying early differentiates into the primordial interior. Com- plex assimilative interactions between the sinking cumulates and the primordial interior, followed by re-equilibration with residual phases, are believed to have produced hybrid source regions. Subsequent partial melting of these local hybrid source regions between 3.8 and 3.2 b.y. (approx.) produced mare basalts. A key impfication of the Kesson-Ringwood hypothesis is that high-Ti basalts were derived from more highly contaminated (or hybridized) local source regions than low-Ti basalts. Although the latter do contain a small but important component derived from 394 A.E. RINGWOOD I ! I I I lO Sm/Eu EU 5 anomaly XGree o • high-Ti CHONDRITES x Iow-Ti I i I I I 5 10 15 20 25 ppm Sm Fig. 2. Graph showing samarium/europium ratios versus samarium concentrations for Apollo 11, 12, 15 and 17 basalts. After Haskin et al. (1970, 1971), Helmke et al. (1973) and Shih et al. (1975). the early (4.6-4.4 b.y.) differentiation process, they are nevertheless mainly composed of material derived from the primordial deep lunar interior. It follows that low-Ti basalts are better able to provide more direct information on the composition of the primordial lunar interior than are high-Ti basalts. This general interpretation is supported by several specific lines of evidence. Rare earth element (REE) distributions for some representative high- and low-Ti basalts are shown in Figure 1. Two features which are attributable to participation of 4.6-4.4 b.y. differentiated cumulates are the magnitude of the europium anomalies and the depletion of light REE. Depletions of light REE are generally much smaller in low-Ti basalts than in high-Ti basalts. Likewise, the magnitudes of europium anomalies are generally much greater in high-Ti basalts than in low-Ti basalts, although there is a degree of overlap between some dif- ferentiated Apollo 12 low-Ti basalts and some of the more primitive of the Apollo 17 high-Ti basalts. Another feature implying a more complex history for high-Ti basalts is their low nickel contents (less than 10 ppm) as compared to low-Ti basalts, which contain up to 170 ppm Ni. For further discussion of this point, see Ringwood and Kesson (1976a) and Part II of this series. Figure 1 shows that there is a general trend for the size of the Eu-anomaly to decrease as the absolute abundances of the remaining trivalent REE decrease. The trend is shown more clearly in Figure 2 which is based upon similar diagrams presented by Haskin et al. (1970) and Helmke et al. (1973). The overall linear relationship between Sm/Eu ratios (a measure of the Eu-anomaly) and Sm contents is quite impressive. This may be caused by the combination of two factors: (1) varying degrees of mixing between a primordial component from the deep interior (no Eu-anomaly) and a fractionated component from the subcrustal cumulates possessing a deep Eu-anomaly, and, (2) varying degrees of partial melting of the source region, negatively correlated with the degrees of concentration of incompatible elements (trivalent REE in this case) in the resultant magmas. BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 395 Figures 1 and 2 clearly demonstrate the more primitive nature of most low-Ti basalts (particularly from Apollo 15) as compared to most high-Ti basalts. The fact that the Sm/ Eu ratio versus Sm extrapolates directly through Green Glass to the chondritic ratio is of considerably significance, strongly implying that the 'primitive' component of mare basalts from the deep primordial interior did not possess a Eu anomaly. Likewise, the 'flat' pattern of trivalent REE in Green Glass implies the existence of chondritic relative abundances in its source region. On the basis of the above discussion, we have chosen to employ compositional data and phase relationships displayed by low-Ti basalts only, in an attempt to constrain the composition of the primordial lunar interior. 3. Source Region of Low-Titanium Basalts A. IDENTIFICATION OF PRIMITIVE LOW-Ti BASALT COMPOSITIONS Low-Ti basalts which best qualify as primitive magmas are those possessing relatively high Cr and Mg contents, high contents of normative olivine, and which can furthermore be demonstrated to have reached the lunar surface essentially as liquids containing negligible or small amounts of previously crystallized phenocrysts. Of the suite of natural samples, these therefore represent magmas which experienced the least amount (possibly negligible) of olivine fractionation during ascent. Green et al. (1971b) carried out a detailed study of 12009 and demonstrated that this basalt satisfied these criteria. Likewise, Grove et al. (1973) and Walker et al. (1976) demonstrated that the picritic basalt 12002 belonged to this category. Detailed studies of the primitive Apollo 15 composition 15555 by Kesson (1975) and Longhi et al. (1972) and of 15016 by Kushiro (1972) also showed that this olivine-rich magma had arrived at the lunar surface in a totally liquid condition, although limited previous crystallization of olivine could not be precluded. Compositions of some of these primitive samples are given in Table I. Experimental studies on these compositions are of considerable value in determining the nature of their source regions. They contain 10 to 20 times the chondritic abundances of rare earths and other incompatible elements (Figure 1) and are similar in this respect to terrestrial oceanic tholeiites. The relative abundances of trivalent REE and other incompatible involatile elements are, moreover, quite close to chondritic relative abundances. Nevertheless, signi- ficant deviations including the Eu-anomaly (small in the case of 15555) testify to a degree of complexity in the chemical history of the source region prior to magma generation. Isotopic and chemical fractionations (e.g., depletion of light REE) also occurred in the source regions of terrestrial oceanic tholeiites prior to magma generation (Gast, 1968). This has not prevented the overall composition of the latter from being used effectively to constrain the compositions of their source regions (Green and Ringwood, t 967; Green, 1970; Ringwood, 1975a). Although the existence of relatively small prior chemical dis- turbances in the source regions of low-Ti primitive basalts is well established, the overall 396 A.E. RINGWOOD TABLE I Compositions of some primitive low-Ti basalts which are believed to have reached the surface in a completely molten condition. Rock 12040 may have possessed some cumulus olivine and is shown for comparison with Green Glass. Sample No. 12002 12009 15555 12040 Green Glass SiO2 43.6 45.0 44.6 43.9 45.2 TiO~ 2.6 2.9 2.1 2.5 0.4 AI~O 3 7.9 8.6 8.7 7.3 7.6 Cr203 1.0 0.6 0.6 0.6 0.4 FeO 21.7 21.0 22.5 21.1 19.7 MnO 0.2 0.8 0.3 0.3 n.a. MgO 14.9 11.6 11.4 16.5 17.9 CaO 8.3 9.4 9.4 8.0 8.1 Na20 0.2 0.2 0.3 0.7 0.1 Sum 100.3 99.6 99.9 100.3 99.6 100 MgO 55 49 47 58 62 MgO + FeO 2 REE 20 - 20 - 10 - 20 - 4.5 Chondrites Ab 2.0 2.0 2.3 1.4 1.1 An 20.4 22.6 22.6 19.1 20.3 Di 17.0 20.4 20.3 17.0 16.6 Hy 29.3 38.3 33.0 29.7 29.5 Ol 25.0 10.4 17.0 27.2 31.1 Chr 1.4 0.8 0.9 0.9 0.6 llm 4.9 5.6 4.0 4.7 0.8 References 1 2 2 2 3 1. J.M. Rhodes, personal communication. 2. Papike et al. (1976). 3. Ridley et al. (1973). bulk compositions of these rocks may nevertheless be used as effectively as in the case of terrestrial oceanic tholeiites, to limit the compositions of their source regions. The Apollo 15 Green Glass is believed to possess considerable petrogenetic significance in the latter respect (Green and Ringwood, 1973). Its bulk composition is seen to be very similar to the crystalline rock 12040 (Table I) so that there can be no doubt of its essen- tial basaltic nature and affinities. Rock 12040 has been interpreted (e.g., Newton et al., 1971) as consisting of a cumulate of 20% olivine and minor chrome spinel in a parental magma which may have possessed a major element composition similar to 12009 (Green et al., 1971b). On the other hand, an origin as a slowly cooled primary magma is not excluded (Green et al., 1971b). Green Glass could likewise consist of a cumulate of 20% olivine in a primitive basaltic magma which had crystallized, and subsequently been remelted by an impact event. Alternatively, it could be interpreted as a primary magma erupted as a liquid from the lunar interior and dispersed into fine droplets by a process of BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 397 fire-fountaining (Green and Ringwood, 1973). The fact that the age of Green Glass appears to be similar to that of other Apollo 15 mare basalts (Hunecke et al., 1974) favours the latter hypothesis, which is also supported by the absence of anomalous quan- tities of siderophile elements in Green Glass which might be expected to result from impact remelting. Green Glass possessing similar compositions to the Apollo 15 variety has also been found at the Apollo 11, 14 and Luna 16 sites (Bunch etal., 1972b)and is evidently derived from a widely distributed and important rock type. Regardless of which hypothesis may be correct, the composition of Green Glass testi- fies to the widespread existence of an extremely primitive class of basaltic or picritic magma at the lunar surface. The basaltic component of Green Glass has an extremely regular REE pattern, closely parallel to the chondritic pattern and possesses only a very small Eu-anomaly (Figure 1). The REE and other incompatible elements (including titanium) are present at levels of only 4 to 5 times the chondritic abundances. Whereas the least fractionated Apollo 12 and 15 magmas have liquidus olivines FO74_76 , the liquidus olivine of Green Glass is Fo84 (Green and Ringwood, 1973). The chemical relationships between Green Glass and other Apollo 12 and 15 magmas appear to be analogous to those between terrestrial peridotitic komatiites and basaltic komatiites on the one hand, and oceanic tholeiites on the other. Komatiites are believed to form from the same mantle source regions as tholeiites (Green, 1975; Nesbitt and Sun, 1976; Sun and Nesbitt, 1976) but represent a much higher degree of partial melting, (40-60%) of the terrestrial mantle as compared to oceanic tholeiites (10-20% partial melting). B. COMPOSITIONAL RELATIONSHIPS Rare earth abundances (relative to chondrites) in a suite of mare basalts are shown in Figure 1. They are seen to vary over a twenty-fold range from Green Glass to high-K Apollo 11 basalts. Despite the large range of absolute abundances, the relative abundances within this group rarely deviate from chondritic relative abundances by more than a factor of two. This behaviour is also shared by many other involatile incompatible elements, e.g., Li, Ba, Y, U, Th, Zr, hf, Ta, Nb (Schnetzler and Philpotts, 1971 ; Wanke et al., 1972; Church, 1972; Willis et al., 1972). Moreover, among those members of Apollo 11, 12, 15 and 17 suites which have been subjected to the least near-surface fractionation, Mg num- bers cluster around 74-78 (apart from Green Glass) and CaO, A1203 and Cr203 vary within narrow limits. If we ignore the modest internal fractionations of incompatible elements and the large variations in Eu and Ti, attributing these to prior fractionations as discussed previously, the general pattern suggests that these magmas have been formed by widely varying degrees of partial melting during which incompatible elements were strongly partitioned into the liquid phase. Among the low-Ti primitive basalts, REE and incompatible elements generally vary over a range of 10x to 30x chondritic abundances whilst Mg numbers, and Ca, A1 and Cr contents remain approximately constant. These characteristics suggest that these basalts have formed by rather limited degrees of melting (e.g., 5-20%) of the source material. With greater degrees of partial melting, one would expect to find a marked increase in Mg 398 A.E. RINGWOOD DEPTH KM 5oo I I I I I LIQUID ~.w-me ,-.",... -:.~... : : : : : : : .: :MELTING • ,i INTERVAL °°°°.,.°.° RANGE OF LUNAR TEMPERATU.E DISTRIBUTIONS oL) iii :) I,- n,,, iii 1000 o.. ul i m OLIVINE ECLOGITE +Ga + OI ~ 3"59g/cm 3 GREEN GLASS I ~/////J41 I I I I I I 0 10 20 30 40 PRESSURE KILOBARS Fig. 3. Stability fields and densities of mineral assemblages displayed by Green Glass in relation to the probable range of lunar internal temperature depth distributions. Melting interval and liquidus phases are also shown. After Green and Ringwood (1973). BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 399 number. This characteristic is displayed by Green Glass (Mg number of 62). The much lower abundances of incompatible elements in Green Glass also strongly suggest a much greater degree of partial melting of source material. An observation of profound import- ance is that as the apparent degree of partial melting increases, for example, from 12009 (and 12002) to 15555 to Green Glass, the mean abundances of incompatible elements decrease, the relative abundances of incompatible elements resemble more closely those of chondrites and the Eu anomalies also decrease, becoming very small in the case of Green Glass. These relationships strongly suggest that the 'primitive' component of low-Ti basalts possesses chondritic relative abundances and that the primordial source region contains appreciably smaller abundances (< 5-fold) of incompatible elements than Green Glass, assuming this material to represent a partial melt. That Green Glass is likely to represent a partial rather than a total melt of the lunar interior can readily be seen from its sub-solidus phase equilibria in relation to the probable range of temperature distributions within the Moon. The mean density of the lunar interior below the crust is 3.39 g cm -3 (Kaula et al., 1974). It is seen from Figure 3 that if the lunar interior possessed the composition of Green Glass, its mean density would be in the vicinity of 3.5-3.6 g cm -3. It is therefore clear that Green Glass must be a differentiate of a more primitive source material possessing substantially smaller abundances of incom- patible elements. If we choose to ignore the fine structure of incompatible element abundances and con- sider their mean concentrations relative to chondrites, a self-consistent model could be constructed if 12009 represented a 10% partial melt, 15555 a 20% partial melt and Green Glass a 40-60% partial melt of an essentially common source region containing approxi- mately twice the chondritic abundances of Ca, A1, Ti and incompatible elements (Green and Ringwood, 1973). These estimates are approximate but the degrees of freedom are limited. If one chooses a source region containing higher abundances (e.g., 3x or 4x chon- dritic), then 15555 and Green Glass must be assumed to represent much more extensive degrees of partial melting (> 75% for Green Glass) than appears reasonable. Moreover, this would imply a higher mean lunar density than is observed and the differences in Mg numbers between the two compositions would be unexplained. On the other hand, if the source region possessed appreciably smaller abundances of these elements (e.g., l x chondritic), then Green Glass would represent an unrealistically small degree of partial melting when considered in relation to its higher Mg number as compared to other mare basalts. Returning to the Earth's mantle, we note that the source regions of terrestrial basalts probably contain about 4% of CaO and Al203, i.e., about twice the (ordinary) chondritic abundances of these components. The Earth's mantle also appears to contain about twice the chondritic abundances of incompatible elements (Frey and Green, 1974; Ringwood, 1975a; Loubet et al., 1975). Although important differences will be seen to exist, there is nevertheless an interesting analogy between oceanic tholeiites and primitive low-Ti basalts, each representing 10-20% partial melts of source regions containing approximately twice chondritic abundances of Ca, Al, Ti and incompatible involatile elements, and which had 400 A.E. RINGWOOD been subjected to episodes of minor chemical fractionation prior to generation of the primary magmas. The analogy extends to peridotitic komatiites and lunar Green Glass, each believed to represent magmas formed by 40-60% of partial melting of the same respective source regions. Once again, we emphasize the significance of the near-chondritic relative abundances of incompatible elements in Green Glass, the low absolute abundances of incompatible elements in this material and its overall 'basaltic' chemistry, indicative of an origin by partial melting from a more primitive source. C. DEPTH OF ORIGIN The following lines of evidence indicate that the source region of mare basalt partial melts was situated deep within the lunar interior: (i) Mare basalts were generated over a period extending at least to 1.4 billion years after the formation of the Moon. Thermal history considerations employing a wide range of boundary conditions (e.g., Urey, 1962; Toks6z et al., 1976) show that deep-seated cooling would occur to a depth beyond 200 km over a period of 109 years. It is extremely difficult to understand how mare basalt magmas might be formed by partial melting within this outer cool shell some 3.2 billion years ago. Crater counts indeed imply that mare volcanism extended to even younger ages (Boyce et al., 1974). The thermal problem has been compounded by the recent downward revision of lunar heat flow (Langseth et al., 1976a, b), implying lower abundances of U, Th and K than were postulated in many earlier studies of lunar thermal history. It follows that the source regions lay at depths exceeding 200 kin. (ii) The mascons were presumably formed before or during the flooding of the mare basins. Their continued existence for up to 3.8 b.y. implies the existence of a strong, cool and thick (> 150kin) lithosphere at the time of mare volcanism (e.g., Kaula, 1969). An origin for mare basalts by partial melting in this region implies loss of strength and destruction of the lithosphere. Preservation of mascons would be inexplicable unless mare basalts had been derived from deeper regions. (iii) The composition and size of the lunar crust requires an extensive differentiation of the outer few hundred km of the moon (Section 5). The region beneath the crust and extending to these depths, most probably consists of barren cumulates, mainly of olivine and pyroxene, which could not have been parental to mare basalts (Ringwood, 1975b, 1976a). The compositions of mare basalts require the existence of a primordial or only slightly fractionated source region below 400 km which had not participated extensively in the early (4.4-4.6 b.y.) major differentiation which was responsible for the formation of the lunar crust (Kesson and Ringwood, 1976a; Ringwood and Kesson, 1976@ (iv) Primitive low-Ti mare basalt magmas crystallize olivine as a liquidus phase up to moderate pressures, where olivine is joined by subcalcic clinopyroxene or orthopyroxene. The olivine-pyroxene cotectics occur at the following pressures: 12009- 7 kb (Green et al., 1971b), 12002- 14kb (Grove et al., 1973), 15555 - 12kb (Kesson, 1975), Green Glass- 20kb (Walker et al., 1975b). It is possible that all of these magmas have crystal- lized some olivine during their ascent to the surface, so that these are minimum pressures. BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 401 Constraints on the degree of partial melting imposed by trace element abundances (Kesson, 1975) combined with other geochemical criteria, e.g., Ca/A1 ratios, (Ringwood and Essene, 1970b) require that pyroxenes remained in the residuum after partial melting and magma segregation. Accordingly, the pyroxene-olivine cotectic pressures are likely to represent the minimum depths at which the magmas were generated. The above pressures correspond to depths of 140-400 kin. It is possible that the ultimate sources were deeper, below 400 kin, owing to varying degrees of olivine crystallization during ascent. It is also quite likely that partial melting occurred on release of pressure, as diapirs ascended from deeper source regions (Green and Ringwood, 1967). The experimental cotectic pressures correspond to the depths of magma segregation rather than the ultimate depth of origin of the source diapirs. D. MINERALOGICAL NATURE OF SOURCE REGION By experimentally determining the nature and compositions of liquidus and near-liquidus phases of mare basalts over a wide range of pressures, powerful constraints may be placed upon the nature of their source regions. The principle involved is that at fixed P and T, chemical equilibrium between crystals and liquid is independent of the proportion of either phase. Thus, if a magma is found to crystallize phases A and B on its liquidus under specified P, T conditions, (say 2% crystals, 98% liquid) the equilibrium would be unaltered if the system consisted 98% of crystals A and B, and 2% of liquid. This would represent the case of partial melting of a source material to yield 2% magma of the observed com- position and 98% of residual crystals A and B. The compositions of experimental phases A and B can be accurately determined by electron probe microanalysis. In applying this method, we seek to choose basalts (applying criteria previously dis- cussed) which have undergone the minimum amount of fractionation en route to the sur- face. The method is capable in principle of characterizing the residual phases remaining in the source region after magma segregation. It is not, however, directly capable of yield- ing the proportions of mineral phases remaining in the residuum, which is essential if the bulk composition of the source region is to be estimated. There are, however, additional constraints which can be applied to facilitate a solution to the latter problem. One of these is to study comparative high pressure liquidus equilibria in a series of primary magma compositions which are believed to have formed from an approximately uniform source by widely va/ying degrees of partial melting e.g., 12009- 12002- 15555 -Green Glass (e.g., Green and Ringwood, 1973). These magmas have abundances of incom- patible elements varying by a factor of 5, suggesting corresponding variations in their degrees of partial melting. Whereas 12009, 12002, and 15555 are saturated or near- saturated olivine + clinopyroxene -+ orthopyroxene at pressures of 7-14kb and have liquidus phases with Mg numbers near 74-76, Green Glass is multiply saturated with olivine + orthopyroxene at 20kb (Walker et al., 1975b) and its liquidus phases have Mg numbers of 84-86. Thus, 12009, 12002 and 15555 may represent increasing degrees of partial melting during which olivine + clinopyroxene-+ orthopyroxene remained in the source region. Green Glass, on the other hand, represents a much greater degree of partial 402 A.E. RINGWOOD TABLE II Construction of model lunar olivine-pyroxenite source composition from 12009 plus near-liquidus phases at 15 kb. Compositional data from Green et al. (1971). Refractory Model lunar Residue source 10% 12009 40% cpx modified + Liquidus Liquidus Liquidus 30% opx 12009 90% 10% Fo~s Opx Cpx Olivine 30% O1 residua SiO 2 44.6 54.0 50.3 38.5 47.9 47.6 TiO 2 2.6 0.3 0.7 - 0.4 0.6 A1203 7.8 2.3 5.0 - 2.7 3.2 Cr203 0.5 0.8 0.9 0.4 0.5 0.5 FeO 21.2 13.0 15.9 22.2 16.9 17.3 MnO 0.3 0.3 0.3 0.2 0.3 0.3 MgO 14.1 27.7 21.2 38.0 28.2 26.8 CaO 8.6 2.0 6.1 0.2 3.0 3.6 Na20 0.2 0.06 0.1 - 0.06 0.07 100 MgO 54 79 71 75 75 73 MgO + FeO Orthopyroxene was crystallized from modified 12009 composition containing 10% additional olivine (Fo75). Clinopyroxene represents average of near-liquidus phases in 12009 and 12009 + 2% enstatite (Enso) at 1390 °C and 15 kb. Olivine analysis represents liquidus phase at atmospheric pressure but would be unchanged at higher pressure. melting (as shown by its higher Mg number and lower incompatible element abundances) and clinopyroxene was eliminated from its source region, leaving only olivine + ortho- pyroxene. The increasing tendency of the abundances of incompatible elements to approach chondritic relative abundances as the degree of partial melting increases (e.g., Green Glass) provides yet another constraint. It appears likely that the CaO/A1203 ratio of the source region was therefore close to the chondritic ratio (Ringwood and Essene, 1970b; Ringwood, 1970). In most mare basalts, olivine is joined at the cotectic by subcalcic clinopyroxene as pressure is increased. Because of the high CaO/A1203 ratio of the clino- pyroxene (typically > 2 as compared to the chondritic ratio of 0.8), there is no way of constructing a source region consisting of a mixture of magma and liquidus olivine plus clinopyroxene which has a CaO/A1203 ratio approaching the chondritic value. Ringwood and Essene (1970b) pointed out that orthopyroxene containing A1203> CaO would necessarily be a major phase in the source region if this constraint was to be met. They showed that Apollo 11 basalts, were, in fact, very nearly saturated with orthopyroxene at their ol-cpx cotectic pressures. Likewise Green et al. (1971b) demonstrated that whereas in 12009, olivine was joined on the liquidus by clinopyroxene at 7kb, a com- position consisting of 12009 + 10% olivine (which may well have crystallized out during ascent) instead had orthopyroxene on its liquidus at 15 kb. Moreover, Green Glass had olivine + orthopyroxene as liquidus phases over a wide pressure interval between 13 and 22 kb (Green and Ringwood, 1973). BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 403 TABLE III Construction of model lunar olivine-pyroxenite source composition from Green Glass plus near-liquidus phases at 15 kb. Compositional data from Green and Ringwood (1973). Model lunar Refractory source com- Liquidus Liquidus residue position 50% Green olivine orthopyroxene 60% ol GG 50% R. Glass 15 kb 15 kb 40% opx residue SiO 2 45.2 40.6 56.1 46.8 46.0 TiO 2 0.4 - 0.1 - 0.2 A1203 7.6 - 2.1 0.8 4.2 Cr203 0.4 0.4 0.7 0.5 0.5 FeO 19.7 14.8 8.8 12.4 16.1 MnO 0.2 0.2 0.1 0.2 0.2 MgO 17.9 43.5 30.2 38.2 28.1 CaO 8.1 0.4 1.8 1.0 4.5 Na20 0.1 - - - 0.05 100 MgO 62 84 86 85 76 MgO + FeO Green (1976)has recently clarified the roles of orthopyroxene and clinopyroxene in mare basalt petrogenesis. In many runs of relatively short duration, (<4hr) olivine is joined by subcalcic clinopyroxene with increasing pressure (e.g., 12009, 12002, 15555, 70215, 74275). In the case of 74275, Green demonstrated in long runs (> 300hr) that these subcalcic clinopyroxenes (4-7% CaO) were metastable relative to an equilibrium assemblage of orthopyroxene + medium Ca (14-17% CaO) clinopyroxene. He has since observed comparable behaviour displayed by high pressure liquidus subcalcic pyroxenes from other mare basalts including low-Ti varieties (pers. comm.). In view of these results, it seems likely that the most primitive mare basalts would be multiply-saturated with ol + opx + cpx at modest pressures and that these phases were residual after partial melt- ing at the appropriate depths within the lunar interior. These phase relationships permit the construction of model source regions possessing near-chondritic CaO/A12Oa ratios. Applying the principles and constraints discussed above, it is possible to construct reasonably self-consistent compositional models for possible mare basalt source regions. Exercises of this nature, using acceptable combinations of primitive basalt compositions together with analysed near-liquidus phase compositions, are shown in Tables II and III. It is emphasized that the source compositions so derived are not unique and that a great deal of further work upon the crystallization behaviour of primitive mare basalts, particu- larly the compositions of their equilibrium near-liquidus phases, will be required before the compositions of the source region can be regarded as tightly constrained. It will be particularly important to clarify the pyroxene equilibrium relationships, a difficult experimental challenge. At present, it has not been possible to produce a model source region matching the 404 A.E. RINGWOOD chondritic CaO/AI203 ratio of 0.8. The source regions derived in Tables II and III possess CaO/Al203 ratios of 1.1. This may be a consequence of the prior minor chemical distur- bances in the source region as discussed earlier (see also, Ringwood and Kesson, 1976a). Alternatively, it might be a primary property of the source region as discussed in Section 5. The conclusion that pyroxenes were more abundant than olivine in the mare basalt source region rests upon two observations. Primitive mare basalts contain only ~ 8.5% Al203, whereas primitive terrestrial oceanic tholeiites formed by similar degrees of partial melting contain ~ 16% Al203, yet the total abundances of A1, Ca and other involatfle elements in their respective source regions are believed to have been similar. (Certainly, no case has ever been made for believing that the moon contains less Ca, A1, U and Ti than the Earth's mantle.) The Al contents of terrestrial tholeiitic and lunar mare basaltic magmas are buffered and controlled by equilibrium with residual aluminous pyroxenes in the source. If mare basalt magmas have less Al203 than terrestrial magmas, then the residual pyroxenes in their source regions also have correspondingly less A1203. If the total Al203 content of the mare basalt source region is not less than that of the Earth's mantle, then there must be relatively more total pyroxene in the lunar source region than in the Earth's mantle. The second argument concerns the degrees of partial melting. In terrestrial basalt petrogenesis, as the degree of partial melting increases, as from primitive oceanic tholeiite (12x chondritic REE)to basaltic komatiite (4-8x chondritic REE)to perido- titic komatiite (3-6x chondritic REE- Sun and Nesbitt, 1976), the magmas become much richer in normative olivine (e.g., peridotitic komatiite containing 5x chondritic REE may contain over 50% of normative olivine). In the Earth, olivine would be the liquidus phase in these magmas up to very high pressures, probably > 60 kb. On the other hand, for the comparable sequence in the moon from 12009 - 12002 - 15555 to Green Glass (the analogue of komatiite), the normative olivine content does not increase nearly so drastically (from 11% to 32%) as the degree of partial melting increases. Moreover, olivine remains alone on the liquidus of Green Glass only to 20 kb where it is joined by orthopyroxene (Walker et al., 1975b). E. SUMMARY Despite the uncertaintities referred to above, it is believed that some firm and important conclusions can be drawn from the results of investigations to date. The source region of low-Ti mare basalts consisted of a mineral assemblage of orthopyroxene + clino- pyroxene + olivine. Plagioclase was absent. The abundance of pyroxenes exceeded that of olivine, contrary to the situation in the olivine-rich Earth's mantle. The Mg number in the lunar source region was 75-80 as compared to an Mg number close to 88 in the Earth's mantle (Ringwood, 1975a). The contents of CaO and A1203 in the lunar basalt source region were certainly each smaller than 5%, and probably in the vicinity of 3.5 to 4%, i.e., similar to their abundances in the Earth's mantle (Ringwood, 1975a) or about twice the (ordinary) chondritic abundances. The involatile, incompatible elements (e.g., REE, Ti, BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 405 Zr, U) in the lunar source region were probably present at levels about twice those of ordinary chondrites, and similar to those of the Earth's mantle (Ringwood, 1975a). 4. Further Limits on the Composition of the Lunar Interior Several authors have proposed models requiring that the bulk composition of the Moon contains much more A1203 (and CaO) than mare basalt source regions as deduced in the previous section. Models in the former category which have been widely discussed include those of Taylor and Jake~ (1974) (8.1% A1203); Ganapathy and Anders (1974) (11.6% A12Oa); W~inke et al. (1974) (17.4% A1203) and Anderson (1973) (27.2% A1203). Accord- ing to most models of this type, the outer regions of the Moon melted and differentiated about 4.6-4.4 b.y. ago to form a plagioclase-rich lunar crust underlain by a sequence of complementary olivine + pyroxene cumulates, some hundreds of kilometers thick. Mare basalts were interpreted as being formed by subsequent partial melting of the cumulates (e.g., Walker et al., 1975a; Taylor and Jake~, 1974) or by assimilative interactions between the cumulates and the primordial lunar interior (e.g., Hubbard and Minear, 1975). Ringwood (1976a, b) carried out a detailed experimental investigation of melting equi- libria displayed by the above compositions over a wide range of pressures and tempera- tures. This made it possible to determine the chemistry and mineralogy of the cumulate layers for each of these bulk compositions under conditions (a) of melting and differen- tiation of the entire Moon, and (b) melting and differentiation of an outer layer, a few hundred kilometers thick. The results of these investigations showed unequivocally that mare basalts could not have formed by partial melting of ferromagnesian cumulate layers appropriate to these compositional models. Nor could mare basalts have formed by direct partial melting of the primordial interior, or by assimilative interactions between underlying primordial material and the overlying cumulate zone (for the cases of melting of an outer layer only). Several distinct difficulties were encountered by each of these compositional models and the reader is referred to the detailed paper for a full discussion (Ringwood, 1976a). One of the most fundamental difficulties was that each of these models necessarily pro- duced mare basalt magmas containing much higher contents of alumina (12-18% A1203) than observed in the more primitive natural samples (~ 8.5% A1203). The only way in which this problem could be alleviated was by reducing the A1203 content of the bulk composition well below the range (minimum 8.1% A1203) investigated. Ringwood (1976a, b) concluded that an acceptable compositional model would contain only about 4% Al203. An analogous difficulty appears in explaining the composition of the pyroxene com- ponent of the lunar crust in terms of fractional crystallization of these high Ca, A1 compo- sitions. The bulk pyroxenes from highland anorthositic gabbros and gabbroic anorthosites tend to be relatively poor in CaO. The mean normative pyroxenes (diopside + hypersthene) in the W~nke et al. (1975) and Taylor (1973a) model compositions for the highlands con- tain only 6 and 3 mol.% of CaO respectively. 406 A.E. RINGWOOD For the Taylor-Jake] composition (6.6% CaO, 8.1% A1203) the pyroxene(s) crystal- lizing at the stage of plagioclase saturation contain a bulk average of at least 11% CaO and this increases with further fractionation. It does not seem that models of this kind which postulate extended fractional crystallization prior to, and during plagioclase precipitation, can account for the low mean CaO content of pyroxenes in the lunar highlands. This problem becomes increasingly acute for other bulk composition models even richer in calcium, such as those for Ganapathy and Anders (1974), W~inke ,et al., (1974) and Anderson (1973). In an attempt to avoid some of the above difficulties with high Ca, AI compositions, Taylor and Bence (1975) have recently proposed that the bulk moon contains only 6% A1203 and 4.9% CaO. This model is currently being tested by Kesson and Ringwood (1976b); Ringwood and Kesson (1976b) using the methods of experimental petrology. The preliminary results indicate rather strongly that even these levels of A1203 and CaO are too high to be acceptable. The basic conclusion arising from these investigations is that the bulk composition of the Moon contains abundances of CaO and A1203 which are generally similar to the Earth's mantle, i.e., probably in the vicinity of 3.5 to 4% of these compounds (Ringwood, and Kesson, 1976b; Kesson and Ringwood, 1976b). 5. Petrogenesis of the Lunar Crust The predominant rock types occurring in the upper regions of the lunar highland crust consist of a suite of breccias with compositions ranging between anorthositic gabbro, gabbroic anorthosite and anorthosite. Wgnke et al. (1974, 1975)have demonstrated that after minor corrections for the presence of KREEP component and meteoritic nickel- iron, the highland breccia compositions lie upon well-defined mixing lines, with pure anorthosite as an end component. Mixing relationships between different components of highland breccias have also been studied by Taylor (1973a), Taylor and Jake~ (1974) and Taylor and Bence (1975). The latter authors have used the observed breccia com- positions together with orbital XRF and gamma-ray data on abundances of Ca, A1 and Th to estimate the mean composition of the upper layer of the lunar highlands (Table IV). This composition corresponds to that of an anorthositic gabbro. The high MgO/(MgO + FeO) ratio (Mg number) and the substantial Cr content should be noted. These features have been interpreted to imply that the lunar crust has a substantial 'primitive' component and has not been subjected to extensive fractional crystallization which would result in a lower Mg number and Cr content (Wgnke et al., 1974, 1976; Taylor and Jakeg, 1974; Walker et al., 1975b). On the Earth, anorthositic rocks are often formed during the crystallization of large stratiform intrusions of basaltic magma, e.g., the Bushveld and Stillwater complexes (Jackson, 1967). Plagioclase is elutriated upwards to become concentrated in layers towards the top of the system whilst olivine and pyroxene sink to form basal layers. It is widely believed that the anorthositic suite of the lunar crust formed in an analogous BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 407 TABLE IV Derivation of composition of basaltic magma multiply saturated with plagioclase and olivine, which may have been parental to the lunar crust (PLC magma). I II Mean composition of Column I minus 57% upper crust of plagioclase Angs (Taylor and Jake~, 1974) PLC magma SiO 2 44.8 47.3 TiO 2 0.55 0.88 A1203 24.6 18.4 Cr203 0.1 0.13 FeO 6.6 9.4 MgO 8.6 12.3 CaO 14.2 11.3 Na:O 0.45 0.16 100 MgO 70 70 MgO + FeO manner, from a huge ocean of marie or ultramafic magma (e.g., Wood, 1970, 1972, 1975; Walker et al., 1975b). In the Moon, subsequent mixing of rock types by meteorite impact has presumably obscured the original regular stratiform structure. The average thickness of the lunar crust is about 60kin (Kaula etal., 1974). Seismic P velocity in the uppermost 10 km is quite low owing to intense fracturing. However Vp rises to 6.7 km s -1 at about 20 km depth and increases only slightly to about 6.8 km s -1 at 55 km depth (e.g., Toks6z et al., 1974; Dainty et al., 1974). Anorthositic gabbro breccias posses- sing compositions similar to the mean near-surface lunar crust composition (Table IV) display P velocities of 6.7-6.9 km s -x at confining pressures of 5-10 kb (e.g., Wang et al., 1973). The correspondence between experimentally-measured sample velocities and observed in situ velocities in the lower crust has been widely interpreted to imply that the entire lunar crust possesses a composition similar to that of the Observed mean near-surface composition (e.g., Taylor and Bence, 1975). This conclusion, however, is unwarranted. Liebermann and Ringwood (1976) measured the P velocity of pure anorthite, and, using existing velocity data for other relevant minerals, demonstrated that a pore-free gabbroic anorthosite containing 69% plagioclase (Angs), with orthopyroxene > olivine >> clinopyroxene >> ilmenite (mean Mg number = 72), as advocated by Taylor and Bence (1975), would have a Vp velocity of 7.4kms -x. Similar calculations showed that a lunar gabbro (40% plagioclase Angs, 50% ortho- pyroxene, 10% olivine) would have Vp = 7.5 km s -x. These results show firstly that P velocity is insensitive to large changes in the relative proportions of plagioclase to pyroxenes in lunar gabbroic anorthosite-anorthositic gabbro-gabbro compositions. Secondly, the intrinsic velocities of these rocks (7.4-7.5 km s -~) are substantially higher than the observed lower crust velocities of 6.7-6.8 km s -1. Liebermann and Ringwood (1976) pointed out that the discrepancy was most likely caused by shock damage and 408 A.E. RINGWOOD microfracturing in the rocks of the lower crust, caused by large meteoritic impacts and cratefing. The observed lunar velocities therefore do not justify the conclusion that the lower crust is of gabbroic anorthosite composition. The basic ambiguity in matching seismic velocities to rock compositions (also recognized by Wang et al., 1973 and by Toks6z et al., 1974) would equally permit the lower crust to consist of a mafic gabbro containing 40% or less of anorthite. In this case, the bulk composition of the whole crust could be equivalent to that of a normal gabbro (basalt) containing about 50% of plagio- clase and 18% of A1203. The latter interpretation would be consistent with the hypothesis that the crust represents a former gabbroic magma derived by extensive partial melting of the Moon's upper mantle. During crystallization of this parental magma, a limited amount of plagioclase elutriation occurred, resulting in a modest relative enrichment of plagioclase in the upper crust, and a corresponding modest depletion in the lower crust, as compared with the initial composition of the parent magma (Liebermann and Ringwood, 1976). Such behaviour is commonly observed in terrestrial mafic stratiform intrusions. It is plausible that it should also have occurred on the Moon. In view of the efficient differ- entiation which has occurred to form the bulk lunar crust, it would be surprising if this differentiation had not continued within the lunar crust, resulting in a significant degree of enrichment of plagioclase in the upper 20-30 kin. Comparisons with the differentiation behaviour displayed by the Bushveld and Stfllwater Complexes (Jackson, 1967) are of particular relevance in this connection. Taylor (1973b) has suggested that subsequent remixing by cratering processes would have homogenized any initially non-uniform lunar crust. This, however, appears very doubtful in view of the strong lateral compositional heterogeneties in the highland crust as revealed by orbital X-ray and gamma-ray spectroscopy (Adler et al., 1973; Metzger et al., 1974). Whilst it seems likely that the top 10kin or so of the crust have been mixed and overturned by saturation bombardment of 50-100 km diameter craters, the depths of the original craters represented by the much rarer large ring basins (> 200 km radius) and the degree of mixing caused by these events are poorly known. Head et al. (1975) believe that the maximum depth of excavation was only 20 kin. Chao et al. (1975)have argued on other grounds that the widespread anorthositic breccias prevalent at the lunar highlands surface represent a shallow layer formed by excavation and redistribution of aluminous material produced by the Orientale impact. A. NATURE OF THE LUNAR CRUST'S PARENTAL MAGMA The formation of the lunar crust is generally believed to have involved a large scale melting and differentiation process which affected an outer zone of the Moon, some hundreds of kilometers thick (e.g., Wood, 1970, 1972, 1975; Walker et al., 1975b). The energy source may have been supplied by partial conservation of the gravitational energy of accretion of the Moon (Ringwood, 1966, 1970). It is possible that an outer layer, perhaps 400kin thick, was totally melted, thereby forming an ultramafic parent magma (e.g., Walker et al., 1975b). However, consideration of available energy sources and heat balances and the large temperature interval between the liquidus and solidus (Ringwood, 1976a) indicate a BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 409 greater likelihood of producing extensive (e.g., 20-30%) partial melting of the outer regions of the Moon (see also Brett, 1977). This would produce a magma of basaltic bulk composition (parental to the lunar crust) overlying a thick zone of residual, refractory ferromagnesian minerals. Even in the event that total melting of the outer Moon occurred, thereby forming an ultramafic magma, extensive fractional crystallization of olivine + pyroxene would be necessary before the A1203 content of the residual magma was sufficiently high (17-19% A1203) to precipitate plagioclase (Ringwood, 1976a). Thus, in this case also, by the time that the plagioclase-rich lunar upper crust began to form, the parental magma must have been of a mafic or basaltic composition. Schonfeld (1975) has pointed out that the composition of the upper lunar crust can be interpreted in terms of a mush of cumulus plagioclase crystals plus trapped inter- cumulus parental gabbroic magma from which the plagioclase had crystallized. This is a simple and attractive concept. The high Mg number and Cr contents of the lunar crust suggest that the parental mafic liquid had not evolved extensively via fractional crystal- lization when it became trapped in the plagioclase cumulate. Melting relationships of relevant lunar gabbroic anorthosite compositions at low pres- sures (< 5 kb) have shown that plagioclase crystallizes over a wide temperature interval before being joined at a cotectic by olivine and/or pyroxene (Kesson and Ringwood, 1976c). The cotectic liquid is of an overall basaltic composition. In the case of the lunar crust, we obtain the composition of the parental basaltic composition by removing increasing amounts of tiquidus plagioclase (Angs) from the mean upper crust composition (Table IV), and determining the stage at which the residual liquid becomes multiply satu- rated at its liquidus by plagioclase and a ferromagnesian mineral (olivine and/or pyroxene). This composition has been experimentally determined (Table IV, column 2). The liquidus phases at the cotectic at atmospheric pressure are plagioclase (An > 95) and olivine (Fo88). This composition is believed to approximate that of the magma parental to the lunar crust, and, in terms of the previous discussion, to represent the bulk composition of the lunar crust. Thus, it may be of major volumetric and petrogenetic significance. It is of interest to compare the composition of this magma with that of primitive terrestrial oceanic tholeiites, which represent the most abundant rock type erupted at the earth's surface. This comparison is made in Table V, in which sample 3-18 recovered from Leg 3 of the Deep Sea Drilling Project (Frey et al., 1974) is chosen to represent a typical primitive oceanic tholeiite. It is well known that lunar basalts are depleted in volatile ele- ments relative to terrestrial basalts. Of the major elements shown in Table V, the most volatile are Na and Si (Grossman, 1972). In Table V, column 3, we have removed 7% of SiO2 and 1.8% Na20. We see that the composition of the residual modified terrestrial tholeiite is very similar to that of the lunar parental highland magma.* The resemblance between the magma parental to the lunar crust and terrestrial oceanic tholeiites extends also to key trace elements such as the rare earths. Hubbard et al. (1971) The higher content of Cr~O 3 in the PLC magma (0.13%) compared to the terrestrial tholeiite (0.05% Cr203) can be atttributed to the effect of differing oxygen fugacity conditions upon the partition of chromium between magma and residual olivines and pyroxenes as discussed in Part II. 410 A.E. RINGWOOD TABLE V Comparison of composition of basaltic magma which could have been parental to the lunar crust, with composition of a typical primitive terrestrial oceanic tholeiite modified by partial loss of volatile components. I II III Primitive terrestrial Parental lunar Column I minus oceanic tholeiite a crust magma (7% SiO 2"+ 1.8% Na 2 O) SiO 2 50.3 47.3 47.5 TiO 2 0.73 0.9 0.8 AI~O 3 16.6 18.4 18.2 FeO 7.99 9.4 8.8 MgO 10.2 12.3 11.2 CaO 13.2 11.3 14.5 Na20 2.00 0.2 0.2 100 MgO 69 70 69 MgO + FeO REE 9 -10 ~10 Chondrites a Sample DSDP 3-18 from Frey et al. (1974). CALC. LIQUID I.U 10- r~ "t- ° / ~.o - 15415 14.1 ~. ANOR. :E .< u') 0.1 Ba LaCe Nd Sm Eu Gd Dy Er Yb Lu Sr Fig. 4. Calculated rare earth element abundances in parental (basaltic) liquid from which lunar anorthosite 15415 is believed to have crystallized (after Hubbard et al., 1971). BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 411 calculated the REE abundances of the parental basaltic magma from which the anorthosite 15415 has crystallized, using experimental plagioclase-liquid partition coefficients. The parental liquid was found to possess chondritic relative abundances of the REE elements at about 10 times the absolute chondritic abundances (Figure 4). It is similar in this respect to many primitive terrestrial tholeiites including the example chosen in Table V. Hubbard et al. (1971) also showed that other low-K lunar anorthosites possessing positive Eu anomalies had likewise crystallized from parental liquids with generally similar REE abundances. Laul et al. (1974) demonstrated that lunar anorthosites possessing marked positive Eu anomalies were derived from parental magmas containing 8 to 15 times the chondritic abundances. McCaUum et al. (1975) found that the parent liquids in equilibrium with analogous lunar anorthosites contained 4 to 12 times the chondritic REE abundances. The similarity in major element and REE abundances between the most abundant kind of primitive lunar basaltic magma and the most abundant class of primitive terrestrial basaltic magma, modified only by the partial loss of two of the most volatile components (Na20 and SiO2) is believed to be of considerable genetic significance. We will return to this point subsequently. B. COMPOSITION OF SOURCE REGION FROM WHICH PLC MAGMA WAS DERIVED The composition of the (PLC) magma which is believed to be parental to the lunar crust is given in Table IV. Olivine Fo88 plus plagioclase crystallize simultaneously on its liquidus at 1250°C (in vacuum). The absence of a negative europium anomaly (Figure 4)implies that plagioclase was not a residual phase in the source region, and that the magma had not crystallized substantial amounts of plagioclase after segregating from its source region. Olivine, however, remains on the liquidus to about 7 kb (Kesson and Ringwood, t976c) and was probably present as a residual phase in the source region. It is possible, therefore, that the composition of the PLC magma has been modified by crystallization of olivine at relatively shallow depths and that the primary magma was richer in normative olivine than the composition given in Table IV. The primary PLC magma may have formed by extensive partial melting of the outer few hundred kilometers of the Moon. After segregating from the residual phases in its former source region, the vast ocean of primary PLC magma would have precipitated olivine until plagioclase saturation was reached. Most probably, therefore, a layer of olivine cumulates underlies the crust (Figure 5). If the PLC magma segregated from residual refractory ferromagnesian phases at average pressures less than 7 kb (150 km), the experimental phase equilibria show that the residual phase consisted of olivine (Foa8). At this pressure, reconnaissance runs showed that olivine is joined by orthopyroxene (Kesson and Ringwood, 1976c). For a model primary PLC magma containing ~> 10% more normative olivine than the composition given in Table IV, olivine remains on the liquidus to pressures exceeding 15kb and is joined by aluminous orthopyroxene which crystallizes over a wide pressure interval. These relations suggest that the residual, refractory phases remaining in the source region after segregation of the primary PLC magma consisted of olivine alone, or of olivine + orthopyroxene. 412 A. E. RINGWOOD HIGHLAND MARE CRUS" lOq "1- 4C 5C Fig. 5. PetIological structure of Moon at stage of generation of mare basalts. The PLC magma possesses a CaO/A1203 ratio of 0.61, substantially smaller than the chondritic ratio of 0.8. If the source region possessed a near-chondritic CaO/A1203 ratio, as suggested by its near-chondritic REE abundances (Figure 4), then the residual phases remaining behind after magma generation must have possessed a high net CaO/A1203 ratio. However, the liquidus orthopyroxenes observed in PLC and (PLC + olivine) compositions have CaO/AI203 ratios smaller than 0.4. The presence of orthopyroxene as a residual phase would exacerbate the discrepancy with the chondritic CaO/A12Os ratio. Ortho- pyroxene does not therefore appear likely to be a major residual phase in the source region. On the other hand, olivine containing 0.4 to 0.5% CaO but no/11203, is a required residual phase and would modify the CaO/A1203, in the desired direction, although not sufficiently far as to produce an overall chondritic ratio. If the latter were characteristic of the bulk system from which the lunar crust was BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 413 TABLE VI Derivation of composition of source region of lunar crust, equivalent to the mean composition of the outer few hundred kilometers of the Moon. I II III PLC Liquidus olivine Bulk comp. of outer Moon: magma FoBs 20% column I plus 80% column II SiO 2 47.3 41.1 42.3 TiO: 0.9 - 0.2 A1203 18.4 - 3.7 Cr:Oa 0.13 0.3 0.3 FeO 9.4 11.5 11.1 MgO 12.3 46.1 39.3 CaO 11.3 0.5 2.7 Na20 0.2 - 0.04 100 Mg 70 88 86 Mg + Fe REE Chondrites derived, then large quantities of clinopyroxene with CaO >> AlzOa must be assumed to be residual in the source region. Although this possibility cannot yet be finally excluded, it seems improbable in the light of existing phase equilibria data. It seems likely therefore, that the residual, refractory component remaining in the lunar mantle after extraction of the primary PLC magma consisted of pure dunite, similar to, or somewhat more magnesian than the observed liquidus PLC magma olivine (Fo88, 0.5% CaO). If we assume that the PLC magma represented a 20% partial melt of its source region, (comparable to primitive oceanic tholeiites in the Earth's mantle) then the CaO/ Al203 ratio of the bulk system would be 0.72, significantly, but not excessively below the chondritic ratio. The mean composition of the outer few hundred kilometers of the Moon (crust + mantle) on the above assumptions is given in Table VI. In some respects it is complementary to that inferred for the source regions of mare basalts, possessing a CaO/A12Oa ratio smaller than chondritic, whereas the corresponding ratio in the mare basalt source is higher. It is possible, therefore, that the bulk Moon possesses a chondritic mean CaO/A1203 ratio. In comparison to the Earth's mantle, the outer 400 km of the Moon appears to be richer in olivine, whilst the deep interior is richer in pyroxene. This may well reflect an intrinsic zonation in SiO2 content. The average olivine/pyroxene ratio (and SiO2 content) of the entire Moon could well be similar to that of Earth's mantle. The mean Mg value of 86 obtained for the outer 400 km of the Moon may be a slight underestimate, depending upon the degree of fractional crystallization of olivine which may have occurred follow- ing magma segregation. The mean Mg number of othe outer regions of the Moon is thus probably not significantly different from that of the Earth's mantle (89 - Table VII). 414 A.E. RINGWOOD TABLE VII Comparison of estimated bulk composition of entire Moon with model pyrolite composition of the terrestrial mantle. I II III IV Mare basalt Bulk comp. Bulk comp. Pyrolite d source region of outer of entire comp. a Moon b Moon e SiO 2 46.8 42.3 44.6 45.1 TiO~ 0.4 0.2 0.3 0.2 AI203 3.7 3.7 3.7 3.9 Cr203 0.5 0.3 0.4 0.3 FeO 16.7 11.1 13.9 7.9 MgO 27.5 39.3 33.4 38.1 CaO 4.1 2.7 3.4 3.1 Na20 0.06 0.04 0.05 0.4 100MgO 75 86 81 89 MgO + FeO REE 2 ~ 2 ~ 2 - 2 Chondrites Ab 0.5 0.3 0.4 3.4 An 9.9 10.0 9.9 8.9 Di 8.5 2.8 5.6 5.2 Hy 41.8 10.0 25.8 18.4 O1 37.9 76.1 57.2 63.3 Chr 0.7 0.4 0.6 0.5 Ilm 0.8 0.4 0.6 0.4 a Average from Tables II and III. b From Table VI. e Obtained by averaging columns I and II. d From Ringwood (1975a), Table 5-2, column 8, with A1203 adjusted to chondritic CaO/A1203 ratio as discussed in footnote 8 of the Table. However, the Mg number of the source regions of mare basalts (75-80) is smaller than that of the Earth's mantle (Section 3). Thus, the entire Moon appears to contain slightly more total FeO than the Earth's mantle. A sketch of the structure of the Moon as inferred from the preceding discussions is given in Figure 5. It should be mentioned that although the author would have preferred to interpret the available data so as to yield similar bulk compositions for the PLC and mare basalt source regions, it has not been possible to accomplish this objective. It seems that the source regions indeed differ in composition but in a complementary manner. Whilst neither individual region is identical to the inferred pyrolite composition of the Earth's mantle (Ringwood, 1975a), it is a remarkable fact that if these compositions are combined in similar proportions, the pyrolite composition (minus elements more volatile than sodium) is closely approached (Table VII). The principal difference between the mean bulk moon composition and pyrolite (apart from volatile elements) seems to be BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 415 that the Moon is significantly richer in FeO. (However the difference need not be as large as is indicated in Table VII). 6. Heat-producing Elements Although uranium, thorium and potassium are present as trace and minor elements in the Moon, a topic which is treated in Part II of this series, the fact that these elements control an important global property of the bulk Moon, namely the heat flow, justifies a discussion of their significance in Part I. The lunar heat flux values reported earlier (Langseth et al., 1972, 1973) implied, if correct, that the Moon possessed more than five times the chondritic abundances of uranium and thorium, and were widely held to justify lunar models in which related involatile elements, e.g., Ca, A1, REE were enriched to similar degrees. These models conflicted with the evidence from experimental petrology discussed in previous sections. A major reduction in estimates of lunar heat flow has since been made on the basis of a much more reliable estimate of the bulk thermal conductivity of the lunar regolith (Langseth etal., 1976a, b). The revised heat flow value at the Apollo 17 site is 1.4/lW cm-2 and at the Apollo 15 site, 2.1/~W cm-2. The large difference between these values empha- sized the need for considerable caution in using arguments based upon estimated 'mean' lunar heat fluxes. Any such estimates, based upon only two measurements, are likely to possess large uncertainties, as pointed out by Langseth et al. (1976b). The higher heat flow at the Apollo 15 site was also correlated with much higher than average concentrations of thorium (and by inference, U and K) than in the Apollo 17 region. It seems likely that magnetic fractionation processes have resulted in a very strong upward concentration of radioactive elements in the near-surface layer, as has happened in the terrestrial crust. Thus, the higher heat flow at the Apollo 15 site is most reasonably interpreted as being due to a layer of rocks abnormally rich in radioactive elements (Langseth et al., 1976b). This would justify an analysis similar to that which is commonly applied to measurements of terrestrial heat flow and surface radioactivity in which heat flow is plotted against surface radioactivity. This is done in Figure 6, in which the slope of the line defined by the two data points defines the thickness of the surface layer of variable radioactivity. On the basis of this model, the mean lunar heat flow could be esti- mated if the mean thorium content of the surface layer over the entire Moon were known. Langseth et al. (1976b) estimated on the basis of the orbital gamma ray Th measure- ments by Metzger et al. (1974) and the heat flow versus radioactivity plot shown in Figure 6, that the mean lunar heat flow is 1.76/JW cm-2. If all this heat was generated by radioactivity (taking Th/U = 3.7, K/U = 2000) a mean bulk lunar concentration of 45 ppb would be required. They also point out that a steady-state thermal regime for the Earth (with all radioactivity concentrated in the mantle) would require a mean uranium concentration of 42 ppb. The agreement between lunar and terrestrial mantle uranium concentrations is most striking. In the author's opinion, the mean terrestrial and lunar heat flows attributable to 416 A, E. RINGWOOD :E D 0 Z uJ LU "5 i 5 16 THORIUM CONCENTRATION ppm Fig. 6 Plot of heat flow observed at Apollo 15 and 17 sites versus mean thorium contents esti- mated for the soil and rocks of these regions based upon orbital gamma ray measurements and labora- tory measurements of samples. After Langseth et al. (1976b). Estimates of Th content in major lunar petrologic provinces (Metzger et al., 1974) are included. radioactivity may both require substantial reductions, but these do not affect the signi- ficance of the agreement. It now seems clear that the mean temperature of the earth was very high (~ 2500°C) soon after its formation, owing to the large amount of energy released by rapid formation of the core (e.g., Ringwood, 1975a). Under these conditions, 20 to 30% of the present terrestrial heat flow could result from the original heat content, of non-radiogenic origin (MacDonald, 1959, 1965) dependent on the efficiency of con- vective heat transport. Thus, the uranium content of the Earth's mantle may be closer to 30ppb. In the case of the Moon, it seems likely that the orbital gamma ray data of Metzger et al. (1974) sampled a region in which the highly radioactive mare regions in the western plains of the Moon's nearside were over-represented. It is believed that if the Th content of the highland provinces, particularly on the far side, had been weighted according to the actual areas which they occupy, rather than by the accidental sectors traversed by the orbiters, then a substantially lower mean surface abundance of Th might be derived. An analysis of this kind is currently in progress (Metzger, personal communi- cation). In the meantime, we note that if Taylor and Jake~ (1974) estimate of 1.5 ppm for the mean surface abundance of Th in the lunar crust is adopted, then, on the basis of Figure 6, the mean lunar heat flux would be 1.37/aW cm -3, corresponding to a mean lunar uranium abundance of 35 ppb. Corrections for the presence of a small amount of BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 417 original heat component in the lunar heat flow,* and heat refraction effects because of greater thermal conductivity in the sub-mare crust (Conel and Morton, 1975) could easily reduce this estimate to 30 ppb. We concluded earlier that both the Moon and the Earth's mantle contain about twice the (ordinary) chondritic abundances of Ca, AI and REE. The mean uranium content of ordinary chondrites is 13 ppb (Morgan, 1971) which would suggest a mean lunar uranium concentration of 26 ppb. The value of 30-35 ppb inferred from the lunar heat flow is somewhat higher but in view of uncertainties introduced by the 3-fold variation of uranium concentrations in ordinary chondrites on which the average of 0.013 ppb was based, and a twofold dispersion among primitive Type 1 carbonaceous chondrites, the differences can- not be considered to be significant. 7. Oxygen Isotopes About half of the Moon and of the Earth's mantle (by weight) is composed of oxygen, which possesses three isotopes 160, 170 and 180. Variations in oxygen isotope ratios between planets and various classes of meteorites are caused by two factors: (1) primary isotopic inhomogeneities in the oxygen in various regions of the solar nebula prior to accretion, presumably caused by the survival and inhomogeneous distribution of a particular class of interstellar grains rich in 160 (Clayton et al., 1973, 1976); (2) chemical isotopic fractionations, caused for example, by differing temperatures at which the solid matter which accreted to form planets and meteoritic parent bodies, equlibrated and became separated from the gases in the parental solar nebula. The oxygen isotope compositions of both lunar and terrestrial basalts are identical (e.g., Clayton et al., 1976). This implies that the material from which Earth and Moon was formed was homogeneous with respect to the ~60-rich interstellar grain component and also separated from the nebula at the same mean temperature. In contrast, the oxygen in most classes of meteorites possesses distinctly different porportions of the 160-rich grain component. Some differentiated meteorites - the eucrites and howardites - possess oxygen which differs only very slightly in this respect from terrestrial and lunar oxygen and may perhaps be identical (Clayton et al., 1976). However, the oxygen in these meteorites is distinctly different from terrestrial-lunar oxygen owing to a chemical fractionation effect (Taylor et al., 1965). The only classes of meteorites which possess identical oxygen isotopic compositions to the Earth and Moon are the enstatite chon- drites and achondrites. Meteorites provide clear evidence that the 160 component of dust grains was not uni- formly distributed within the solar nebula prior to accretion into meteorite parent bodies and planets. Moreover, the existence of chemically-derived oxygen isotope fractionations Tokst~z et al. (1976) assumes this to be negligible because of heat transfer by convection in the lunar interior which would lead to a steady state heat generation-dissipation situation. However, the author believes that the strong chemical heterogeneities in the lunar interior discussed in the paper would inhibit thermal convection. 418 A.E. RINGWOOD indicates that the parental material of different kinds of meteorites and planets separated from the solar nebula at differing temperatures. In the light of these observations, the identity in oxygen isotopic compositions between Moon and Earth acquires an added significance. When considered in combination with evidence assembled in previous sections of a similarity in the bulk compositions of the Moon and the Earth's mantle, a close genetic relationship between Moon and Earth is strongly suggested. This conclusion is reinforced by the consideration of trace element data in Part II of this series. 8. Physical Properties of the Moon The present model of the internal structure and composition of the Moon (Figure 5) has been developed exclusively from petrologic-geochemical evidence. Its consistency with lunar geophysical evidence is now examined. A. DENSITY The density of the lunar crust is estimated to be 2.95 g cm -3 (Kaula et al., 1974). This is in agreement with the density of the derived crustal composition (Table IV), providing that a small amount of porosity is assumed. The structure of lunar breccias support this latter assumption. The mean density of the lunar interior below the crust is 3.39 g cm -3 (Kaula et al., 1974). According to the presence model, the mean density (at atmospheric pressure) of the upper mantle (to 400km) is 3.34gcm -3 and that of the mantle below 400kin is 3.44 g cm -3. These are estimates based upon inferred mineral assemblages, but without corrections for the effects of thermal expansion and compressibility. These effects almost cancel out under lunar conditions and can be ignored in a first approximation. The mean density of the lunar upper and lower mantles of Figure 5 is thus 3.39 gcm -3, in agree- ment with observations. A metallic core is not required. B. MOMENT OF INERTIA Kaula et al. (1974) estimate the moment of inertia coefficient I/MR 2 of the Moon as 0.395 -+ 0.005. The moment of inertia coefficient of the model shown in Figure 5 is 0.394, in excellent agreement with the preferred observational value. A more detailed dis- cussion of density and moment of inertia constraints is provided by Kaula et al. (1974). C. SEISMIC VELOCITY DISTRIBUTION Nakamura et al. (1974) found that the seismic P velocity in the upper mantle from 60 to 300kin is very close to 8.0km s -1 and its velocity-density relationship is matched better by a layer dominantly composed of olivine Fo8o-84, rather than of pyroxene. Our petrological model (Figure 5) also has a layer of olivine between 60-400 km. Although its composition (FoBs) is slightly more magnesian than the preferred composition of Nakamura et al. (Foso_s4), the difference is within the limits of uncertainty. BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 419 Around a depth of 300km, Nakamura et al. (1974) and Latham et al. (1975) find evidence for a significant decrease in P and S wave velocities. The latter authors suggest that this marks the transition from the differentiated outer layer to the primordial mantle below. This interpretation agrees with our present model (Figure 5) except that the transition is placed closer to 400 km. The decrease in seismic velocities is readily explained by the increases in pyroxene relative to olivine and in iron content, which occur at this depth in the petrologic model. 9. Conclusion Although not identical, the major element composition of the bulk Moon is similar to that of the Earth's mantle. 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Publisher: Springer Journals
Copyright: Copyright © D. Reidel Publishing Company 1977
ISSN: 2948-2941
eISSN: 1573-0794
DOI: 10.1007/bf00577901
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Abstract

I. Mafor and Heat-Producing Elements A. E. RINGWOOD Research School of Earth Sciences, Australian National University, Canberra, Australia (Received 21 January, 1977) Abstract. The lunar interior is comprised of two major petrological provinces: (1) an outer zone several hundred km thick which experienced partial melting and crystallization differentiation 4.4- 4.6 b.y. ago to form the lunar crust together with an underlying complementary zone of ultramafic cumulates and residua, and (2) the primordial deep interior which was the source region for mare basalts (3.2-3.8 b.y.) and had previously been contaminated to varying degrees with highly fractionated material derived from the 4.4-4.6 b.y. differentiation event. In both major petrologic provinces, basaltic magmas have been produced by partial melting. The chemical characteristics and high-pressure phase relationships of these magmas can be used to constrain the bulk compositions of their respective source regions. Primitive low-Ti mare basalts (e.g., 12009, 12002, 15555 and Green Glass) possessing high norma- tive olivine and high Mg and Cr contents, provide the most direct evidence upon the composition of the primordial deep lunar interior. This composition, as estimated on the basis of high pressure equi- libria displayed by the above basalts, combined with other geochemical criteria, is found to consist of orthopyroxene + clinopyroxene + olivine with total pyroxenes > olivine, 100 MgO/(MgO + FeO) = 75-80, about 4% of CaO and A120 a and 2X chondritic abundances of REE, U and Th. This compo- sition is similar to that of the earth's mantle except for a higher pyroxene/olivine ratio and lower 100 MgO/(MgO + FeO). The lunar crust is believed to have formed by plagioclase elutriation within a vast ocean of parental basaltic magma. The composition of the latter is found experimentally by removing liquidus plagioclase from the observed mean upper crust (gabbroic anorthosite) composition, until the resulting compo- sition becomes multiply saturated with plagioclase and a ferromagnesian phase (olivine). This parental basaltic composition is almost identical with terrestrial oceanic tholeiites, except for partial depletion in the two most volatile components, Na20 and SiO 2. Similarity between these two most abundant classes of lunar and terrestrial basaltic magmas strongly implies corresponding similarities between their source regions. The bulk composition of the outer 400 km of the Moon as constrained by the 4.6-4.4 b.y. parental basaltic magma is found to be peridotitic, with olivine > pyroxene, 100 MgO/ (MgO + FeO) ~ 86, and about 2X chondritic abundances of Ca, A1 and REE. The Moon thus appears to have a zoned structure, with the deep interior (below 400 km) possessing somewhat higher contents of FeO and SiO 2 than the outer 400 kin. This zoned model, derived exclusively on petrological grounds, provides a quantitative explanation of the Moon's mean density, moment of inertia and seismic velocity profile. The bulk composition of the entire Moon, thus obtained, is very similar to the pyrolite model com- position for the Earth's mantle, except that the Moon is depleted in Na (and other volatile elements) and somewhat enriched in iron. The similarity in major element composition extends also to the abundances of REE, U and Th. These compositional similarities, combined with the identity in oxygen isotope ratios between the Moon and the Earth's mantle, are strongly suggestive of a common genetic relationship. 1. Introduction One of the most significant results arising from the Apollo project was the demonstration that the lunar maria are composed of rocks resembling terrestrial basalts. Moreover it is The Moon 16 (1977) 389-423. All Rights Reserved. Copyright © 1977 by D. Reidel Publishing Company, Dordrecht-Holland. 390 A.E. RINGWOOD becoming increasingly clear that the lunar crust was ultimately derived by differentiation from a parental magma ocean of basaltic affinities (Section 5). In recent years, important progress has been made towards understanding the origin of terrestrial basaltic rocks. It is now widely accepted that they formed by partial melting of an ultramafic source rock in the mantle. The observed spectrum of basaltic compositions is interpreted primarily in terms of the degree of partial melting of the source material, the depth of partial melting and the subsequent crystallization history during ascent to the surface. High pressure, high temperature investigations of the crystallization behaviour of terrestrial basalts using the methods of experimental petrology have been successfully employed to constrain the nature of their source regions (Green and Ringwood, 1967; Green, 1970; Ringwood, 1975a). It is natural to attempt to understand the origin of lunar basalts in terms of processes analogous to those which have operated in the petrogenesis of terrestrial basalts and like- wise, to use experimental petrology to provide information on the nature of their source regions. The first comprehensive attempt in this direction was presented by Ringwood and Essene (1970a, b) at the Apollo 11 Lunar Science Conference in January, 1970. In view of controversies which later developed in this field, it is worth remarking that several of the key conclusions reached by these workers have subsequently proved to be well founded. A principal conclusion (see also, Ringwood, 1970)was that Apollo 11 high K and low K magmas had been produced mainly by partial melting processes in the lunar interior rather than by extensive fractional crystallization in huge lava lakes, as advocated, for example, by O'Hara et al. (1970a). The former interpretation was at first very much a minority viewpoint, as can be seen in the official summary of the Apollo 11 Conference (LSAPT, 1970, p. 450) but has since become widely accepted. Although varying degrees of partial melting were believed to be mainly responsible for the chemical diversity, it should be noted that Ringwood and Essene (1970b, pp. 784,791) did not exclude the possibility of moderate degrees (up to 30 percent) of fractional crystallization of olivine and pyroxene. Subsequent researches by several groups (e.g. James and Jackson, 1970; Compston et al., 1971; Kushiro and Haramura, 1971; Green et al., 1971a; James and Wright, 1972; Chappell and Green, 1973; Rhodes and Hubbard, 1973; Walker et al., 1975a; Longhi et al., 1974; Shih et al., 1975) have greatly clarified the relative import- ance of partial melting versus fractional crystallization in explaining the chemical diversity among mare basalts. The widespread occurrence of near-surface fractionation controlled by the separation of olivine-+ Cr spinel + Fe-Ti oxides has been clearly documented but in the great majority of cases, this has involved the crystallization of less than 30% of these minerals. The samples representing the most primitive, i.e., the least fractionated magmas can be recognised by their high Mg and Cr contents and, in the cases of low-Ti basalts, by high contents of normative olivine. In many such samples textural and mineralogical evidence indicates that they were erupted at the surface as liquids (Green et al., 1971a, 1975; Longhi et aI., 1972; Kesson, 1975; Walker et al., 1975a, 1976). These primitive liquids are believed to represent close BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 391 approximations to primary magmas formed by direct partial melting in the lunar interior. A second controversial area concerned the role of plagioclase in mare basalt petro- genesis. Several workers including Ringwood and Essene (1970b) concluded that Apollo 11 magmas were undersaturated with plagioclase, and therefore that plagioclase could not have been present as a residual phase in their source regions. This conclusion was extended to cover most Apollo 12 basalts by Green et al. (1971a) but was strongly disputed by O'Hara et al. (1970a, b), and Biggar et al. (1971). It was opposed also by trace element geochemists, (e.g. Gast et al., 1970; Haskin et al., 1970; Philpotts and Schnetzler, 1970; Gast, 1972) who required the presence of plagioclase, in the source regions in order to explain the negative europium anomalies of mare basalts. The conflict has since been resolved by several detailed experimental investigations (Green et al., ,1971b; Ringwood and Green, 1972; Green and Ringwood, 1973; Grove et al., 1973 ;Kushiro, 1972; Longhi et al., 1972, 1974; Walker et al., 1975a, 1976; Kesson, 1975). Whilst a few basalts (e.g. 12038, 14053, Luna 16) and several Apollo 11 ophitic basalts) were probably plagioclase-saturated upon eruption, the vast majority of Apollo 11, 12, 15 and 17 basalts were definitely undersaturated with plagloclase. Moreover, those magmas which could be demonstrated to have undergone the least near-surface crystallization differentiation and to have reached the surface essentially as liquids (e.g. 12009, 12002, 15555, 15016, Green Glass, 70215, 74275) were among those most under- saturated with plagioclase. These results firmly demonstrated that plagioclase was not a residual phase remaining in the source region after partial melting, so that the europium anomaly must be a characteristic inherited from the source region. This conclusion has had far-reaching implications for subsequent petrogenetic hypotheses. Moreover, the observed degrees of plagioclase undersaturation provided an additional limit to the permissible degrees of preeruptive crystal fractionation. A third conclusion reached by Ringwood and Essene (1970a, b) was that the source region of Apollo 11 basalts was composed mainly of orthopyroxene and subcalcic clino- pyroxene with olivine possibly present as a subsidiary phase. The source region contained 3 to 5% of CaO and A12 03 and had an Mg number (100 MgO/(MgO + FeO)) of 75 to 80. A corresponding study of Apollo 12 low-Ti basalts by Green etal. (1971a)and Green et al. (1971b) led to the conclusion that their source region likewise contained similar amounts of CaO and A12Oa and possessed a similar Mg number. It seemed likely, however, that olivine was also an important mineral in the Apollo 12 source region. The source was thus inferred to be an olivine pyroxenite. This initial interpretation of the source region of Apollo 11 and 12 basalts was extended to Apollo 15 and 17 basalts (Green and Ringwood, 1973; Chappell and Green, 1973; Green et al., 1975) and was regarded as representative of the bulk of the lunar interior (Ringwood, 1970, 1975b, 1976a). The conclusion that the bulk Moon contained only 3 to 5% of CaO and A1203, i.e., about twice the chondritic abundances, has been generally disregarded by most lunar scientists. Lunar bulk compositional models containing much higher abundances of CaO and Al:O3 have been advocated by many workers in the field. 392 A.E. RINGWOOD Models which have received a considerable degree of favourable attention include those of Taylor and Jake~, 1974 (8.1% Al203, 6.6% CaO), Ganapathy and Anders, 1974 (11.6% Al203, 9.3% CaO), Wgnke et al., 1974 (17.4% A1203, 13.6% CaO) and Anderson, 1973 (27.2% Al203, 22.1% CaO). We will demonstrate in the present paper that the early compositional models contain- ing relatively smaller amounts of CaO and Al203 are much closer to reality. 2. Mare Basalt Petrogenesis Early hypotheses of mare basalt petrogenesis based upon experimental petrology proposed that all classes of mare basalts had formed by varying degrees of partial melting of a common olivine-pyroxenite source at depths of 150-500 km (e.g., Ringwood and Essene, 1970b). Although this straightforward single-stage hypothesis provided an adequate explanation of many aspects of the major element chemistry of mare basalts and their source regions, it was unable to provide satisfactory explanations of other geochemical characteristics of high-Ti and low-Ti mare basalts, e.g., their REE patterns, Eu-anomalies, and TiO2 contents. In addition, the earlier differentiation event implied for mare basalt source regions by Sm-Nd (Lugmair et al., 1975) and U-Pb isotope systematics (Tera and Wasserburg, 1975) was not accounted for. To meet these difficulties, a second class of hypotheses was developed (e.g., Schnetzler and Philpotts, 1971) which maintained that mare basalts could have formed by the remelt- ing of chemically and mineralogically inhomogeneous olivine + pyroxene + ilmenite cumu- lates formed during the early differentiation of the Moon around 4.6-4.4 b.y., as the complement of the plagioclase-rich crust. Specifically, it was suggested (e.g. Walker et al., 1975a) that low-Ti basalts had formed by the partial melting of early olivine + pyroxene cumulates at considerable depths (200-400 kin) whereas the high-Ti basalts formed by partial melting of a zone of late-stage olivine + pyroxene + ilmenite cumulates at relatively shallow depths (aroun d 100 kin). The cumulate remelting hypothesis offers, in principle, an explanation for several important geochemical characteristics of mare basalts, e.g., their different TiO2 contents, the two-stage history recorded by Sm-Nd and Pb-U isotopes and the Eu anomalies and their varying magnitudes. However, this hypothesis still encounters a number of fatal difficulties. For example, it is unable to explain the obser- vation that the least fractionated high-Ti and low-Ti basalts have similar Cr contents and Mg numbers. This and other shortcomings have been discussed in detail by Ringwood (1975b) and Kesson and Ringwood (1976a). Although the cumulate-remelting hypothesis is not acceptable in its present form, its success in explaining several key features Of the trace-element and isotopic geochemistry of mare basalts strongly suggests that some form of reprocessing of cumulates is involved in mare basalt petrogenesis. This raises the possibility of combining the more attractive aspects of the single-stage, uniform, primordial source-region hypothesis with those of the cumulate-remelting hypothesis. Kesson and Pdngwood (1976a) and Ringwood and Kesson (1976a)have proposed a BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 393 ~ 10049 10017] Hi K ]0058 "'~-- ' 10020 4O Q; ~ 700]7 e~ ' ~" ...... 12051 -0 6 2c ~'~.'_-. =.-- ..... ]2002 ¢- .=_ uJ ill • "~ ,v LU ev, ~AVERAGE APOLLO 15 .~" ...... . .L.j.~.~.~.. Q. E ]( O. O. "~'" 15555 ....... ~ ...... GREEN GLASS I I I I I I I I I La Ce Nd Sm Eu Gd Dy Er Yb Fig. 1. Abundances of rare earths in a representative suite of mare basalt samples. After Haskin et al. (1970), Gast et al. (1970), Hubbard and Gast (1971), Helmke et al. (1972, 1973), Philpotts et al. (1972), Ridley et al. (1973), Shih et al. (1975). hypothesis in this category which appears to avoid most of the difficulties associated with previous petrogenetic hypotheses• This involves hybridization or assimilation of 4.6-4.4 b.y. differentiates by primordial material from the deep lunar interior• Early versions of hybridization hypotheses were proposed by Anderson (1971) and Hubbard and Minear (1975) and involve assimilation/hybridization at shallow levels• The Kesson- Ringwood model differs from these in that the hybridization is believed to have occurred at depth (about 400 km) and was caused by sinking of pods of 4.6-4•4 b.y. dense residual cumulates through the underlying early differentiates into the primordial interior. Com- plex assimilative interactions between the sinking cumulates and the primordial interior, followed by re-equilibration with residual phases, are believed to have produced hybrid source regions. Subsequent partial melting of these local hybrid source regions between 3.8 and 3.2 b.y. (approx.) produced mare basalts. A key impfication of the Kesson-Ringwood hypothesis is that high-Ti basalts were derived from more highly contaminated (or hybridized) local source regions than low-Ti basalts. Although the latter do contain a small but important component derived from 394 A.E. RINGWOOD I ! I I I lO Sm/Eu EU 5 anomaly XGree o • high-Ti CHONDRITES x Iow-Ti I i I I I 5 10 15 20 25 ppm Sm Fig. 2. Graph showing samarium/europium ratios versus samarium concentrations for Apollo 11, 12, 15 and 17 basalts. After Haskin et al. (1970, 1971), Helmke et al. (1973) and Shih et al. (1975). the early (4.6-4.4 b.y.) differentiation process, they are nevertheless mainly composed of material derived from the primordial deep lunar interior. It follows that low-Ti basalts are better able to provide more direct information on the composition of the primordial lunar interior than are high-Ti basalts. This general interpretation is supported by several specific lines of evidence. Rare earth element (REE) distributions for some representative high- and low-Ti basalts are shown in Figure 1. Two features which are attributable to participation of 4.6-4.4 b.y. differentiated cumulates are the magnitude of the europium anomalies and the depletion of light REE. Depletions of light REE are generally much smaller in low-Ti basalts than in high-Ti basalts. Likewise, the magnitudes of europium anomalies are generally much greater in high-Ti basalts than in low-Ti basalts, although there is a degree of overlap between some dif- ferentiated Apollo 12 low-Ti basalts and some of the more primitive of the Apollo 17 high-Ti basalts. Another feature implying a more complex history for high-Ti basalts is their low nickel contents (less than 10 ppm) as compared to low-Ti basalts, which contain up to 170 ppm Ni. For further discussion of this point, see Ringwood and Kesson (1976a) and Part II of this series. Figure 1 shows that there is a general trend for the size of the Eu-anomaly to decrease as the absolute abundances of the remaining trivalent REE decrease. The trend is shown more clearly in Figure 2 which is based upon similar diagrams presented by Haskin et al. (1970) and Helmke et al. (1973). The overall linear relationship between Sm/Eu ratios (a measure of the Eu-anomaly) and Sm contents is quite impressive. This may be caused by the combination of two factors: (1) varying degrees of mixing between a primordial component from the deep interior (no Eu-anomaly) and a fractionated component from the subcrustal cumulates possessing a deep Eu-anomaly, and, (2) varying degrees of partial melting of the source region, negatively correlated with the degrees of concentration of incompatible elements (trivalent REE in this case) in the resultant magmas. BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 395 Figures 1 and 2 clearly demonstrate the more primitive nature of most low-Ti basalts (particularly from Apollo 15) as compared to most high-Ti basalts. The fact that the Sm/ Eu ratio versus Sm extrapolates directly through Green Glass to the chondritic ratio is of considerably significance, strongly implying that the 'primitive' component of mare basalts from the deep primordial interior did not possess a Eu anomaly. Likewise, the 'flat' pattern of trivalent REE in Green Glass implies the existence of chondritic relative abundances in its source region. On the basis of the above discussion, we have chosen to employ compositional data and phase relationships displayed by low-Ti basalts only, in an attempt to constrain the composition of the primordial lunar interior. 3. Source Region of Low-Titanium Basalts A. IDENTIFICATION OF PRIMITIVE LOW-Ti BASALT COMPOSITIONS Low-Ti basalts which best qualify as primitive magmas are those possessing relatively high Cr and Mg contents, high contents of normative olivine, and which can furthermore be demonstrated to have reached the lunar surface essentially as liquids containing negligible or small amounts of previously crystallized phenocrysts. Of the suite of natural samples, these therefore represent magmas which experienced the least amount (possibly negligible) of olivine fractionation during ascent. Green et al. (1971b) carried out a detailed study of 12009 and demonstrated that this basalt satisfied these criteria. Likewise, Grove et al. (1973) and Walker et al. (1976) demonstrated that the picritic basalt 12002 belonged to this category. Detailed studies of the primitive Apollo 15 composition 15555 by Kesson (1975) and Longhi et al. (1972) and of 15016 by Kushiro (1972) also showed that this olivine-rich magma had arrived at the lunar surface in a totally liquid condition, although limited previous crystallization of olivine could not be precluded. Compositions of some of these primitive samples are given in Table I. Experimental studies on these compositions are of considerable value in determining the nature of their source regions. They contain 10 to 20 times the chondritic abundances of rare earths and other incompatible elements (Figure 1) and are similar in this respect to terrestrial oceanic tholeiites. The relative abundances of trivalent REE and other incompatible involatile elements are, moreover, quite close to chondritic relative abundances. Nevertheless, signi- ficant deviations including the Eu-anomaly (small in the case of 15555) testify to a degree of complexity in the chemical history of the source region prior to magma generation. Isotopic and chemical fractionations (e.g., depletion of light REE) also occurred in the source regions of terrestrial oceanic tholeiites prior to magma generation (Gast, 1968). This has not prevented the overall composition of the latter from being used effectively to constrain the compositions of their source regions (Green and Ringwood, t 967; Green, 1970; Ringwood, 1975a). Although the existence of relatively small prior chemical dis- turbances in the source regions of low-Ti primitive basalts is well established, the overall 396 A.E. RINGWOOD TABLE I Compositions of some primitive low-Ti basalts which are believed to have reached the surface in a completely molten condition. Rock 12040 may have possessed some cumulus olivine and is shown for comparison with Green Glass. Sample No. 12002 12009 15555 12040 Green Glass SiO2 43.6 45.0 44.6 43.9 45.2 TiO~ 2.6 2.9 2.1 2.5 0.4 AI~O 3 7.9 8.6 8.7 7.3 7.6 Cr203 1.0 0.6 0.6 0.6 0.4 FeO 21.7 21.0 22.5 21.1 19.7 MnO 0.2 0.8 0.3 0.3 n.a. MgO 14.9 11.6 11.4 16.5 17.9 CaO 8.3 9.4 9.4 8.0 8.1 Na20 0.2 0.2 0.3 0.7 0.1 Sum 100.3 99.6 99.9 100.3 99.6 100 MgO 55 49 47 58 62 MgO + FeO 2 REE 20 - 20 - 10 - 20 - 4.5 Chondrites Ab 2.0 2.0 2.3 1.4 1.1 An 20.4 22.6 22.6 19.1 20.3 Di 17.0 20.4 20.3 17.0 16.6 Hy 29.3 38.3 33.0 29.7 29.5 Ol 25.0 10.4 17.0 27.2 31.1 Chr 1.4 0.8 0.9 0.9 0.6 llm 4.9 5.6 4.0 4.7 0.8 References 1 2 2 2 3 1. J.M. Rhodes, personal communication. 2. Papike et al. (1976). 3. Ridley et al. (1973). bulk compositions of these rocks may nevertheless be used as effectively as in the case of terrestrial oceanic tholeiites, to limit the compositions of their source regions. The Apollo 15 Green Glass is believed to possess considerable petrogenetic significance in the latter respect (Green and Ringwood, 1973). Its bulk composition is seen to be very similar to the crystalline rock 12040 (Table I) so that there can be no doubt of its essen- tial basaltic nature and affinities. Rock 12040 has been interpreted (e.g., Newton et al., 1971) as consisting of a cumulate of 20% olivine and minor chrome spinel in a parental magma which may have possessed a major element composition similar to 12009 (Green et al., 1971b). On the other hand, an origin as a slowly cooled primary magma is not excluded (Green et al., 1971b). Green Glass could likewise consist of a cumulate of 20% olivine in a primitive basaltic magma which had crystallized, and subsequently been remelted by an impact event. Alternatively, it could be interpreted as a primary magma erupted as a liquid from the lunar interior and dispersed into fine droplets by a process of BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 397 fire-fountaining (Green and Ringwood, 1973). The fact that the age of Green Glass appears to be similar to that of other Apollo 15 mare basalts (Hunecke et al., 1974) favours the latter hypothesis, which is also supported by the absence of anomalous quan- tities of siderophile elements in Green Glass which might be expected to result from impact remelting. Green Glass possessing similar compositions to the Apollo 15 variety has also been found at the Apollo 11, 14 and Luna 16 sites (Bunch etal., 1972b)and is evidently derived from a widely distributed and important rock type. Regardless of which hypothesis may be correct, the composition of Green Glass testi- fies to the widespread existence of an extremely primitive class of basaltic or picritic magma at the lunar surface. The basaltic component of Green Glass has an extremely regular REE pattern, closely parallel to the chondritic pattern and possesses only a very small Eu-anomaly (Figure 1). The REE and other incompatible elements (including titanium) are present at levels of only 4 to 5 times the chondritic abundances. Whereas the least fractionated Apollo 12 and 15 magmas have liquidus olivines FO74_76 , the liquidus olivine of Green Glass is Fo84 (Green and Ringwood, 1973). The chemical relationships between Green Glass and other Apollo 12 and 15 magmas appear to be analogous to those between terrestrial peridotitic komatiites and basaltic komatiites on the one hand, and oceanic tholeiites on the other. Komatiites are believed to form from the same mantle source regions as tholeiites (Green, 1975; Nesbitt and Sun, 1976; Sun and Nesbitt, 1976) but represent a much higher degree of partial melting, (40-60%) of the terrestrial mantle as compared to oceanic tholeiites (10-20% partial melting). B. COMPOSITIONAL RELATIONSHIPS Rare earth abundances (relative to chondrites) in a suite of mare basalts are shown in Figure 1. They are seen to vary over a twenty-fold range from Green Glass to high-K Apollo 11 basalts. Despite the large range of absolute abundances, the relative abundances within this group rarely deviate from chondritic relative abundances by more than a factor of two. This behaviour is also shared by many other involatile incompatible elements, e.g., Li, Ba, Y, U, Th, Zr, hf, Ta, Nb (Schnetzler and Philpotts, 1971 ; Wanke et al., 1972; Church, 1972; Willis et al., 1972). Moreover, among those members of Apollo 11, 12, 15 and 17 suites which have been subjected to the least near-surface fractionation, Mg num- bers cluster around 74-78 (apart from Green Glass) and CaO, A1203 and Cr203 vary within narrow limits. If we ignore the modest internal fractionations of incompatible elements and the large variations in Eu and Ti, attributing these to prior fractionations as discussed previously, the general pattern suggests that these magmas have been formed by widely varying degrees of partial melting during which incompatible elements were strongly partitioned into the liquid phase. Among the low-Ti primitive basalts, REE and incompatible elements generally vary over a range of 10x to 30x chondritic abundances whilst Mg numbers, and Ca, A1 and Cr contents remain approximately constant. These characteristics suggest that these basalts have formed by rather limited degrees of melting (e.g., 5-20%) of the source material. With greater degrees of partial melting, one would expect to find a marked increase in Mg 398 A.E. RINGWOOD DEPTH KM 5oo I I I I I LIQUID ~.w-me ,-.",... -:.~... : : : : : : : .: :MELTING • ,i INTERVAL °°°°.,.°.° RANGE OF LUNAR TEMPERATU.E DISTRIBUTIONS oL) iii :) I,- n,,, iii 1000 o.. ul i m OLIVINE ECLOGITE +Ga + OI ~ 3"59g/cm 3 GREEN GLASS I ~/////J41 I I I I I I 0 10 20 30 40 PRESSURE KILOBARS Fig. 3. Stability fields and densities of mineral assemblages displayed by Green Glass in relation to the probable range of lunar internal temperature depth distributions. Melting interval and liquidus phases are also shown. After Green and Ringwood (1973). BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 399 number. This characteristic is displayed by Green Glass (Mg number of 62). The much lower abundances of incompatible elements in Green Glass also strongly suggest a much greater degree of partial melting of source material. An observation of profound import- ance is that as the apparent degree of partial melting increases, for example, from 12009 (and 12002) to 15555 to Green Glass, the mean abundances of incompatible elements decrease, the relative abundances of incompatible elements resemble more closely those of chondrites and the Eu anomalies also decrease, becoming very small in the case of Green Glass. These relationships strongly suggest that the 'primitive' component of low-Ti basalts possesses chondritic relative abundances and that the primordial source region contains appreciably smaller abundances (< 5-fold) of incompatible elements than Green Glass, assuming this material to represent a partial melt. That Green Glass is likely to represent a partial rather than a total melt of the lunar interior can readily be seen from its sub-solidus phase equilibria in relation to the probable range of temperature distributions within the Moon. The mean density of the lunar interior below the crust is 3.39 g cm -3 (Kaula et al., 1974). It is seen from Figure 3 that if the lunar interior possessed the composition of Green Glass, its mean density would be in the vicinity of 3.5-3.6 g cm -3. It is therefore clear that Green Glass must be a differentiate of a more primitive source material possessing substantially smaller abundances of incom- patible elements. If we choose to ignore the fine structure of incompatible element abundances and con- sider their mean concentrations relative to chondrites, a self-consistent model could be constructed if 12009 represented a 10% partial melt, 15555 a 20% partial melt and Green Glass a 40-60% partial melt of an essentially common source region containing approxi- mately twice the chondritic abundances of Ca, A1, Ti and incompatible elements (Green and Ringwood, 1973). These estimates are approximate but the degrees of freedom are limited. If one chooses a source region containing higher abundances (e.g., 3x or 4x chon- dritic), then 15555 and Green Glass must be assumed to represent much more extensive degrees of partial melting (> 75% for Green Glass) than appears reasonable. Moreover, this would imply a higher mean lunar density than is observed and the differences in Mg numbers between the two compositions would be unexplained. On the other hand, if the source region possessed appreciably smaller abundances of these elements (e.g., l x chondritic), then Green Glass would represent an unrealistically small degree of partial melting when considered in relation to its higher Mg number as compared to other mare basalts. Returning to the Earth's mantle, we note that the source regions of terrestrial basalts probably contain about 4% of CaO and Al203, i.e., about twice the (ordinary) chondritic abundances of these components. The Earth's mantle also appears to contain about twice the chondritic abundances of incompatible elements (Frey and Green, 1974; Ringwood, 1975a; Loubet et al., 1975). Although important differences will be seen to exist, there is nevertheless an interesting analogy between oceanic tholeiites and primitive low-Ti basalts, each representing 10-20% partial melts of source regions containing approximately twice chondritic abundances of Ca, Al, Ti and incompatible involatile elements, and which had 400 A.E. RINGWOOD been subjected to episodes of minor chemical fractionation prior to generation of the primary magmas. The analogy extends to peridotitic komatiites and lunar Green Glass, each believed to represent magmas formed by 40-60% of partial melting of the same respective source regions. Once again, we emphasize the significance of the near-chondritic relative abundances of incompatible elements in Green Glass, the low absolute abundances of incompatible elements in this material and its overall 'basaltic' chemistry, indicative of an origin by partial melting from a more primitive source. C. DEPTH OF ORIGIN The following lines of evidence indicate that the source region of mare basalt partial melts was situated deep within the lunar interior: (i) Mare basalts were generated over a period extending at least to 1.4 billion years after the formation of the Moon. Thermal history considerations employing a wide range of boundary conditions (e.g., Urey, 1962; Toks6z et al., 1976) show that deep-seated cooling would occur to a depth beyond 200 km over a period of 109 years. It is extremely difficult to understand how mare basalt magmas might be formed by partial melting within this outer cool shell some 3.2 billion years ago. Crater counts indeed imply that mare volcanism extended to even younger ages (Boyce et al., 1974). The thermal problem has been compounded by the recent downward revision of lunar heat flow (Langseth et al., 1976a, b), implying lower abundances of U, Th and K than were postulated in many earlier studies of lunar thermal history. It follows that the source regions lay at depths exceeding 200 kin. (ii) The mascons were presumably formed before or during the flooding of the mare basins. Their continued existence for up to 3.8 b.y. implies the existence of a strong, cool and thick (> 150kin) lithosphere at the time of mare volcanism (e.g., Kaula, 1969). An origin for mare basalts by partial melting in this region implies loss of strength and destruction of the lithosphere. Preservation of mascons would be inexplicable unless mare basalts had been derived from deeper regions. (iii) The composition and size of the lunar crust requires an extensive differentiation of the outer few hundred km of the moon (Section 5). The region beneath the crust and extending to these depths, most probably consists of barren cumulates, mainly of olivine and pyroxene, which could not have been parental to mare basalts (Ringwood, 1975b, 1976a). The compositions of mare basalts require the existence of a primordial or only slightly fractionated source region below 400 km which had not participated extensively in the early (4.4-4.6 b.y.) major differentiation which was responsible for the formation of the lunar crust (Kesson and Ringwood, 1976a; Ringwood and Kesson, 1976@ (iv) Primitive low-Ti mare basalt magmas crystallize olivine as a liquidus phase up to moderate pressures, where olivine is joined by subcalcic clinopyroxene or orthopyroxene. The olivine-pyroxene cotectics occur at the following pressures: 12009- 7 kb (Green et al., 1971b), 12002- 14kb (Grove et al., 1973), 15555 - 12kb (Kesson, 1975), Green Glass- 20kb (Walker et al., 1975b). It is possible that all of these magmas have crystal- lized some olivine during their ascent to the surface, so that these are minimum pressures. BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 401 Constraints on the degree of partial melting imposed by trace element abundances (Kesson, 1975) combined with other geochemical criteria, e.g., Ca/A1 ratios, (Ringwood and Essene, 1970b) require that pyroxenes remained in the residuum after partial melting and magma segregation. Accordingly, the pyroxene-olivine cotectic pressures are likely to represent the minimum depths at which the magmas were generated. The above pressures correspond to depths of 140-400 kin. It is possible that the ultimate sources were deeper, below 400 kin, owing to varying degrees of olivine crystallization during ascent. It is also quite likely that partial melting occurred on release of pressure, as diapirs ascended from deeper source regions (Green and Ringwood, 1967). The experimental cotectic pressures correspond to the depths of magma segregation rather than the ultimate depth of origin of the source diapirs. D. MINERALOGICAL NATURE OF SOURCE REGION By experimentally determining the nature and compositions of liquidus and near-liquidus phases of mare basalts over a wide range of pressures, powerful constraints may be placed upon the nature of their source regions. The principle involved is that at fixed P and T, chemical equilibrium between crystals and liquid is independent of the proportion of either phase. Thus, if a magma is found to crystallize phases A and B on its liquidus under specified P, T conditions, (say 2% crystals, 98% liquid) the equilibrium would be unaltered if the system consisted 98% of crystals A and B, and 2% of liquid. This would represent the case of partial melting of a source material to yield 2% magma of the observed com- position and 98% of residual crystals A and B. The compositions of experimental phases A and B can be accurately determined by electron probe microanalysis. In applying this method, we seek to choose basalts (applying criteria previously dis- cussed) which have undergone the minimum amount of fractionation en route to the sur- face. The method is capable in principle of characterizing the residual phases remaining in the source region after magma segregation. It is not, however, directly capable of yield- ing the proportions of mineral phases remaining in the residuum, which is essential if the bulk composition of the source region is to be estimated. There are, however, additional constraints which can be applied to facilitate a solution to the latter problem. One of these is to study comparative high pressure liquidus equilibria in a series of primary magma compositions which are believed to have formed from an approximately uniform source by widely va/ying degrees of partial melting e.g., 12009- 12002- 15555 -Green Glass (e.g., Green and Ringwood, 1973). These magmas have abundances of incom- patible elements varying by a factor of 5, suggesting corresponding variations in their degrees of partial melting. Whereas 12009, 12002, and 15555 are saturated or near- saturated olivine + clinopyroxene -+ orthopyroxene at pressures of 7-14kb and have liquidus phases with Mg numbers near 74-76, Green Glass is multiply saturated with olivine + orthopyroxene at 20kb (Walker et al., 1975b) and its liquidus phases have Mg numbers of 84-86. Thus, 12009, 12002 and 15555 may represent increasing degrees of partial melting during which olivine + clinopyroxene-+ orthopyroxene remained in the source region. Green Glass, on the other hand, represents a much greater degree of partial 402 A.E. RINGWOOD TABLE II Construction of model lunar olivine-pyroxenite source composition from 12009 plus near-liquidus phases at 15 kb. Compositional data from Green et al. (1971). Refractory Model lunar Residue source 10% 12009 40% cpx modified + Liquidus Liquidus Liquidus 30% opx 12009 90% 10% Fo~s Opx Cpx Olivine 30% O1 residua SiO 2 44.6 54.0 50.3 38.5 47.9 47.6 TiO 2 2.6 0.3 0.7 - 0.4 0.6 A1203 7.8 2.3 5.0 - 2.7 3.2 Cr203 0.5 0.8 0.9 0.4 0.5 0.5 FeO 21.2 13.0 15.9 22.2 16.9 17.3 MnO 0.3 0.3 0.3 0.2 0.3 0.3 MgO 14.1 27.7 21.2 38.0 28.2 26.8 CaO 8.6 2.0 6.1 0.2 3.0 3.6 Na20 0.2 0.06 0.1 - 0.06 0.07 100 MgO 54 79 71 75 75 73 MgO + FeO Orthopyroxene was crystallized from modified 12009 composition containing 10% additional olivine (Fo75). Clinopyroxene represents average of near-liquidus phases in 12009 and 12009 + 2% enstatite (Enso) at 1390 °C and 15 kb. Olivine analysis represents liquidus phase at atmospheric pressure but would be unchanged at higher pressure. melting (as shown by its higher Mg number and lower incompatible element abundances) and clinopyroxene was eliminated from its source region, leaving only olivine + ortho- pyroxene. The increasing tendency of the abundances of incompatible elements to approach chondritic relative abundances as the degree of partial melting increases (e.g., Green Glass) provides yet another constraint. It appears likely that the CaO/A1203 ratio of the source region was therefore close to the chondritic ratio (Ringwood and Essene, 1970b; Ringwood, 1970). In most mare basalts, olivine is joined at the cotectic by subcalcic clinopyroxene as pressure is increased. Because of the high CaO/A1203 ratio of the clino- pyroxene (typically > 2 as compared to the chondritic ratio of 0.8), there is no way of constructing a source region consisting of a mixture of magma and liquidus olivine plus clinopyroxene which has a CaO/A1203 ratio approaching the chondritic value. Ringwood and Essene (1970b) pointed out that orthopyroxene containing A1203> CaO would necessarily be a major phase in the source region if this constraint was to be met. They showed that Apollo 11 basalts, were, in fact, very nearly saturated with orthopyroxene at their ol-cpx cotectic pressures. Likewise Green et al. (1971b) demonstrated that whereas in 12009, olivine was joined on the liquidus by clinopyroxene at 7kb, a com- position consisting of 12009 + 10% olivine (which may well have crystallized out during ascent) instead had orthopyroxene on its liquidus at 15 kb. Moreover, Green Glass had olivine + orthopyroxene as liquidus phases over a wide pressure interval between 13 and 22 kb (Green and Ringwood, 1973). BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 403 TABLE III Construction of model lunar olivine-pyroxenite source composition from Green Glass plus near-liquidus phases at 15 kb. Compositional data from Green and Ringwood (1973). Model lunar Refractory source com- Liquidus Liquidus residue position 50% Green olivine orthopyroxene 60% ol GG 50% R. Glass 15 kb 15 kb 40% opx residue SiO 2 45.2 40.6 56.1 46.8 46.0 TiO 2 0.4 - 0.1 - 0.2 A1203 7.6 - 2.1 0.8 4.2 Cr203 0.4 0.4 0.7 0.5 0.5 FeO 19.7 14.8 8.8 12.4 16.1 MnO 0.2 0.2 0.1 0.2 0.2 MgO 17.9 43.5 30.2 38.2 28.1 CaO 8.1 0.4 1.8 1.0 4.5 Na20 0.1 - - - 0.05 100 MgO 62 84 86 85 76 MgO + FeO Green (1976)has recently clarified the roles of orthopyroxene and clinopyroxene in mare basalt petrogenesis. In many runs of relatively short duration, (<4hr) olivine is joined by subcalcic clinopyroxene with increasing pressure (e.g., 12009, 12002, 15555, 70215, 74275). In the case of 74275, Green demonstrated in long runs (> 300hr) that these subcalcic clinopyroxenes (4-7% CaO) were metastable relative to an equilibrium assemblage of orthopyroxene + medium Ca (14-17% CaO) clinopyroxene. He has since observed comparable behaviour displayed by high pressure liquidus subcalcic pyroxenes from other mare basalts including low-Ti varieties (pers. comm.). In view of these results, it seems likely that the most primitive mare basalts would be multiply-saturated with ol + opx + cpx at modest pressures and that these phases were residual after partial melt- ing at the appropriate depths within the lunar interior. These phase relationships permit the construction of model source regions possessing near-chondritic CaO/A12Oa ratios. Applying the principles and constraints discussed above, it is possible to construct reasonably self-consistent compositional models for possible mare basalt source regions. Exercises of this nature, using acceptable combinations of primitive basalt compositions together with analysed near-liquidus phase compositions, are shown in Tables II and III. It is emphasized that the source compositions so derived are not unique and that a great deal of further work upon the crystallization behaviour of primitive mare basalts, particu- larly the compositions of their equilibrium near-liquidus phases, will be required before the compositions of the source region can be regarded as tightly constrained. It will be particularly important to clarify the pyroxene equilibrium relationships, a difficult experimental challenge. At present, it has not been possible to produce a model source region matching the 404 A.E. RINGWOOD chondritic CaO/AI203 ratio of 0.8. The source regions derived in Tables II and III possess CaO/Al203 ratios of 1.1. This may be a consequence of the prior minor chemical distur- bances in the source region as discussed earlier (see also, Ringwood and Kesson, 1976a). Alternatively, it might be a primary property of the source region as discussed in Section 5. The conclusion that pyroxenes were more abundant than olivine in the mare basalt source region rests upon two observations. Primitive mare basalts contain only ~ 8.5% Al203, whereas primitive terrestrial oceanic tholeiites formed by similar degrees of partial melting contain ~ 16% Al203, yet the total abundances of A1, Ca and other involatfle elements in their respective source regions are believed to have been similar. (Certainly, no case has ever been made for believing that the moon contains less Ca, A1, U and Ti than the Earth's mantle.) The Al contents of terrestrial tholeiitic and lunar mare basaltic magmas are buffered and controlled by equilibrium with residual aluminous pyroxenes in the source. If mare basalt magmas have less Al203 than terrestrial magmas, then the residual pyroxenes in their source regions also have correspondingly less A1203. If the total Al203 content of the mare basalt source region is not less than that of the Earth's mantle, then there must be relatively more total pyroxene in the lunar source region than in the Earth's mantle. The second argument concerns the degrees of partial melting. In terrestrial basalt petrogenesis, as the degree of partial melting increases, as from primitive oceanic tholeiite (12x chondritic REE)to basaltic komatiite (4-8x chondritic REE)to perido- titic komatiite (3-6x chondritic REE- Sun and Nesbitt, 1976), the magmas become much richer in normative olivine (e.g., peridotitic komatiite containing 5x chondritic REE may contain over 50% of normative olivine). In the Earth, olivine would be the liquidus phase in these magmas up to very high pressures, probably > 60 kb. On the other hand, for the comparable sequence in the moon from 12009 - 12002 - 15555 to Green Glass (the analogue of komatiite), the normative olivine content does not increase nearly so drastically (from 11% to 32%) as the degree of partial melting increases. Moreover, olivine remains alone on the liquidus of Green Glass only to 20 kb where it is joined by orthopyroxene (Walker et al., 1975b). E. SUMMARY Despite the uncertaintities referred to above, it is believed that some firm and important conclusions can be drawn from the results of investigations to date. The source region of low-Ti mare basalts consisted of a mineral assemblage of orthopyroxene + clino- pyroxene + olivine. Plagioclase was absent. The abundance of pyroxenes exceeded that of olivine, contrary to the situation in the olivine-rich Earth's mantle. The Mg number in the lunar source region was 75-80 as compared to an Mg number close to 88 in the Earth's mantle (Ringwood, 1975a). The contents of CaO and A1203 in the lunar basalt source region were certainly each smaller than 5%, and probably in the vicinity of 3.5 to 4%, i.e., similar to their abundances in the Earth's mantle (Ringwood, 1975a) or about twice the (ordinary) chondritic abundances. The involatile, incompatible elements (e.g., REE, Ti, BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 405 Zr, U) in the lunar source region were probably present at levels about twice those of ordinary chondrites, and similar to those of the Earth's mantle (Ringwood, 1975a). 4. Further Limits on the Composition of the Lunar Interior Several authors have proposed models requiring that the bulk composition of the Moon contains much more A1203 (and CaO) than mare basalt source regions as deduced in the previous section. Models in the former category which have been widely discussed include those of Taylor and Jake~ (1974) (8.1% A1203); Ganapathy and Anders (1974) (11.6% A12Oa); W~inke et al. (1974) (17.4% A1203) and Anderson (1973) (27.2% A1203). Accord- ing to most models of this type, the outer regions of the Moon melted and differentiated about 4.6-4.4 b.y. ago to form a plagioclase-rich lunar crust underlain by a sequence of complementary olivine + pyroxene cumulates, some hundreds of kilometers thick. Mare basalts were interpreted as being formed by subsequent partial melting of the cumulates (e.g., Walker et al., 1975a; Taylor and Jake~, 1974) or by assimilative interactions between the cumulates and the primordial lunar interior (e.g., Hubbard and Minear, 1975). Ringwood (1976a, b) carried out a detailed experimental investigation of melting equi- libria displayed by the above compositions over a wide range of pressures and tempera- tures. This made it possible to determine the chemistry and mineralogy of the cumulate layers for each of these bulk compositions under conditions (a) of melting and differen- tiation of the entire Moon, and (b) melting and differentiation of an outer layer, a few hundred kilometers thick. The results of these investigations showed unequivocally that mare basalts could not have formed by partial melting of ferromagnesian cumulate layers appropriate to these compositional models. Nor could mare basalts have formed by direct partial melting of the primordial interior, or by assimilative interactions between underlying primordial material and the overlying cumulate zone (for the cases of melting of an outer layer only). Several distinct difficulties were encountered by each of these compositional models and the reader is referred to the detailed paper for a full discussion (Ringwood, 1976a). One of the most fundamental difficulties was that each of these models necessarily pro- duced mare basalt magmas containing much higher contents of alumina (12-18% A1203) than observed in the more primitive natural samples (~ 8.5% A1203). The only way in which this problem could be alleviated was by reducing the A1203 content of the bulk composition well below the range (minimum 8.1% A1203) investigated. Ringwood (1976a, b) concluded that an acceptable compositional model would contain only about 4% Al203. An analogous difficulty appears in explaining the composition of the pyroxene com- ponent of the lunar crust in terms of fractional crystallization of these high Ca, A1 compo- sitions. The bulk pyroxenes from highland anorthositic gabbros and gabbroic anorthosites tend to be relatively poor in CaO. The mean normative pyroxenes (diopside + hypersthene) in the W~nke et al. (1975) and Taylor (1973a) model compositions for the highlands con- tain only 6 and 3 mol.% of CaO respectively. 406 A.E. RINGWOOD For the Taylor-Jake] composition (6.6% CaO, 8.1% A1203) the pyroxene(s) crystal- lizing at the stage of plagioclase saturation contain a bulk average of at least 11% CaO and this increases with further fractionation. It does not seem that models of this kind which postulate extended fractional crystallization prior to, and during plagioclase precipitation, can account for the low mean CaO content of pyroxenes in the lunar highlands. This problem becomes increasingly acute for other bulk composition models even richer in calcium, such as those for Ganapathy and Anders (1974), W~inke ,et al., (1974) and Anderson (1973). In an attempt to avoid some of the above difficulties with high Ca, AI compositions, Taylor and Bence (1975) have recently proposed that the bulk moon contains only 6% A1203 and 4.9% CaO. This model is currently being tested by Kesson and Ringwood (1976b); Ringwood and Kesson (1976b) using the methods of experimental petrology. The preliminary results indicate rather strongly that even these levels of A1203 and CaO are too high to be acceptable. The basic conclusion arising from these investigations is that the bulk composition of the Moon contains abundances of CaO and A1203 which are generally similar to the Earth's mantle, i.e., probably in the vicinity of 3.5 to 4% of these compounds (Ringwood, and Kesson, 1976b; Kesson and Ringwood, 1976b). 5. Petrogenesis of the Lunar Crust The predominant rock types occurring in the upper regions of the lunar highland crust consist of a suite of breccias with compositions ranging between anorthositic gabbro, gabbroic anorthosite and anorthosite. Wgnke et al. (1974, 1975)have demonstrated that after minor corrections for the presence of KREEP component and meteoritic nickel- iron, the highland breccia compositions lie upon well-defined mixing lines, with pure anorthosite as an end component. Mixing relationships between different components of highland breccias have also been studied by Taylor (1973a), Taylor and Jake~ (1974) and Taylor and Bence (1975). The latter authors have used the observed breccia com- positions together with orbital XRF and gamma-ray data on abundances of Ca, A1 and Th to estimate the mean composition of the upper layer of the lunar highlands (Table IV). This composition corresponds to that of an anorthositic gabbro. The high MgO/(MgO + FeO) ratio (Mg number) and the substantial Cr content should be noted. These features have been interpreted to imply that the lunar crust has a substantial 'primitive' component and has not been subjected to extensive fractional crystallization which would result in a lower Mg number and Cr content (Wgnke et al., 1974, 1976; Taylor and Jakeg, 1974; Walker et al., 1975b). On the Earth, anorthositic rocks are often formed during the crystallization of large stratiform intrusions of basaltic magma, e.g., the Bushveld and Stillwater complexes (Jackson, 1967). Plagioclase is elutriated upwards to become concentrated in layers towards the top of the system whilst olivine and pyroxene sink to form basal layers. It is widely believed that the anorthositic suite of the lunar crust formed in an analogous BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 407 TABLE IV Derivation of composition of basaltic magma multiply saturated with plagioclase and olivine, which may have been parental to the lunar crust (PLC magma). I II Mean composition of Column I minus 57% upper crust of plagioclase Angs (Taylor and Jake~, 1974) PLC magma SiO 2 44.8 47.3 TiO 2 0.55 0.88 A1203 24.6 18.4 Cr203 0.1 0.13 FeO 6.6 9.4 MgO 8.6 12.3 CaO 14.2 11.3 Na:O 0.45 0.16 100 MgO 70 70 MgO + FeO manner, from a huge ocean of marie or ultramafic magma (e.g., Wood, 1970, 1972, 1975; Walker et al., 1975b). In the Moon, subsequent mixing of rock types by meteorite impact has presumably obscured the original regular stratiform structure. The average thickness of the lunar crust is about 60kin (Kaula etal., 1974). Seismic P velocity in the uppermost 10 km is quite low owing to intense fracturing. However Vp rises to 6.7 km s -1 at about 20 km depth and increases only slightly to about 6.8 km s -1 at 55 km depth (e.g., Toks6z et al., 1974; Dainty et al., 1974). Anorthositic gabbro breccias posses- sing compositions similar to the mean near-surface lunar crust composition (Table IV) display P velocities of 6.7-6.9 km s -x at confining pressures of 5-10 kb (e.g., Wang et al., 1973). The correspondence between experimentally-measured sample velocities and observed in situ velocities in the lower crust has been widely interpreted to imply that the entire lunar crust possesses a composition similar to that of the Observed mean near-surface composition (e.g., Taylor and Bence, 1975). This conclusion, however, is unwarranted. Liebermann and Ringwood (1976) measured the P velocity of pure anorthite, and, using existing velocity data for other relevant minerals, demonstrated that a pore-free gabbroic anorthosite containing 69% plagioclase (Angs), with orthopyroxene > olivine >> clinopyroxene >> ilmenite (mean Mg number = 72), as advocated by Taylor and Bence (1975), would have a Vp velocity of 7.4kms -x. Similar calculations showed that a lunar gabbro (40% plagioclase Angs, 50% ortho- pyroxene, 10% olivine) would have Vp = 7.5 km s -x. These results show firstly that P velocity is insensitive to large changes in the relative proportions of plagioclase to pyroxenes in lunar gabbroic anorthosite-anorthositic gabbro-gabbro compositions. Secondly, the intrinsic velocities of these rocks (7.4-7.5 km s -~) are substantially higher than the observed lower crust velocities of 6.7-6.8 km s -1. Liebermann and Ringwood (1976) pointed out that the discrepancy was most likely caused by shock damage and 408 A.E. RINGWOOD microfracturing in the rocks of the lower crust, caused by large meteoritic impacts and cratefing. The observed lunar velocities therefore do not justify the conclusion that the lower crust is of gabbroic anorthosite composition. The basic ambiguity in matching seismic velocities to rock compositions (also recognized by Wang et al., 1973 and by Toks6z et al., 1974) would equally permit the lower crust to consist of a mafic gabbro containing 40% or less of anorthite. In this case, the bulk composition of the whole crust could be equivalent to that of a normal gabbro (basalt) containing about 50% of plagio- clase and 18% of A1203. The latter interpretation would be consistent with the hypothesis that the crust represents a former gabbroic magma derived by extensive partial melting of the Moon's upper mantle. During crystallization of this parental magma, a limited amount of plagioclase elutriation occurred, resulting in a modest relative enrichment of plagioclase in the upper crust, and a corresponding modest depletion in the lower crust, as compared with the initial composition of the parent magma (Liebermann and Ringwood, 1976). Such behaviour is commonly observed in terrestrial mafic stratiform intrusions. It is plausible that it should also have occurred on the Moon. In view of the efficient differ- entiation which has occurred to form the bulk lunar crust, it would be surprising if this differentiation had not continued within the lunar crust, resulting in a significant degree of enrichment of plagioclase in the upper 20-30 kin. Comparisons with the differentiation behaviour displayed by the Bushveld and Stfllwater Complexes (Jackson, 1967) are of particular relevance in this connection. Taylor (1973b) has suggested that subsequent remixing by cratering processes would have homogenized any initially non-uniform lunar crust. This, however, appears very doubtful in view of the strong lateral compositional heterogeneties in the highland crust as revealed by orbital X-ray and gamma-ray spectroscopy (Adler et al., 1973; Metzger et al., 1974). Whilst it seems likely that the top 10kin or so of the crust have been mixed and overturned by saturation bombardment of 50-100 km diameter craters, the depths of the original craters represented by the much rarer large ring basins (> 200 km radius) and the degree of mixing caused by these events are poorly known. Head et al. (1975) believe that the maximum depth of excavation was only 20 kin. Chao et al. (1975)have argued on other grounds that the widespread anorthositic breccias prevalent at the lunar highlands surface represent a shallow layer formed by excavation and redistribution of aluminous material produced by the Orientale impact. A. NATURE OF THE LUNAR CRUST'S PARENTAL MAGMA The formation of the lunar crust is generally believed to have involved a large scale melting and differentiation process which affected an outer zone of the Moon, some hundreds of kilometers thick (e.g., Wood, 1970, 1972, 1975; Walker et al., 1975b). The energy source may have been supplied by partial conservation of the gravitational energy of accretion of the Moon (Ringwood, 1966, 1970). It is possible that an outer layer, perhaps 400kin thick, was totally melted, thereby forming an ultramafic parent magma (e.g., Walker et al., 1975b). However, consideration of available energy sources and heat balances and the large temperature interval between the liquidus and solidus (Ringwood, 1976a) indicate a BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 409 greater likelihood of producing extensive (e.g., 20-30%) partial melting of the outer regions of the Moon (see also Brett, 1977). This would produce a magma of basaltic bulk composition (parental to the lunar crust) overlying a thick zone of residual, refractory ferromagnesian minerals. Even in the event that total melting of the outer Moon occurred, thereby forming an ultramafic magma, extensive fractional crystallization of olivine + pyroxene would be necessary before the A1203 content of the residual magma was sufficiently high (17-19% A1203) to precipitate plagioclase (Ringwood, 1976a). Thus, in this case also, by the time that the plagioclase-rich lunar upper crust began to form, the parental magma must have been of a mafic or basaltic composition. Schonfeld (1975) has pointed out that the composition of the upper lunar crust can be interpreted in terms of a mush of cumulus plagioclase crystals plus trapped inter- cumulus parental gabbroic magma from which the plagioclase had crystallized. This is a simple and attractive concept. The high Mg number and Cr contents of the lunar crust suggest that the parental mafic liquid had not evolved extensively via fractional crystal- lization when it became trapped in the plagioclase cumulate. Melting relationships of relevant lunar gabbroic anorthosite compositions at low pres- sures (< 5 kb) have shown that plagioclase crystallizes over a wide temperature interval before being joined at a cotectic by olivine and/or pyroxene (Kesson and Ringwood, 1976c). The cotectic liquid is of an overall basaltic composition. In the case of the lunar crust, we obtain the composition of the parental basaltic composition by removing increasing amounts of tiquidus plagioclase (Angs) from the mean upper crust composition (Table IV), and determining the stage at which the residual liquid becomes multiply satu- rated at its liquidus by plagioclase and a ferromagnesian mineral (olivine and/or pyroxene). This composition has been experimentally determined (Table IV, column 2). The liquidus phases at the cotectic at atmospheric pressure are plagioclase (An > 95) and olivine (Fo88). This composition is believed to approximate that of the magma parental to the lunar crust, and, in terms of the previous discussion, to represent the bulk composition of the lunar crust. Thus, it may be of major volumetric and petrogenetic significance. It is of interest to compare the composition of this magma with that of primitive terrestrial oceanic tholeiites, which represent the most abundant rock type erupted at the earth's surface. This comparison is made in Table V, in which sample 3-18 recovered from Leg 3 of the Deep Sea Drilling Project (Frey et al., 1974) is chosen to represent a typical primitive oceanic tholeiite. It is well known that lunar basalts are depleted in volatile ele- ments relative to terrestrial basalts. Of the major elements shown in Table V, the most volatile are Na and Si (Grossman, 1972). In Table V, column 3, we have removed 7% of SiO2 and 1.8% Na20. We see that the composition of the residual modified terrestrial tholeiite is very similar to that of the lunar parental highland magma.* The resemblance between the magma parental to the lunar crust and terrestrial oceanic tholeiites extends also to key trace elements such as the rare earths. Hubbard et al. (1971) The higher content of Cr~O 3 in the PLC magma (0.13%) compared to the terrestrial tholeiite (0.05% Cr203) can be atttributed to the effect of differing oxygen fugacity conditions upon the partition of chromium between magma and residual olivines and pyroxenes as discussed in Part II. 410 A.E. RINGWOOD TABLE V Comparison of composition of basaltic magma which could have been parental to the lunar crust, with composition of a typical primitive terrestrial oceanic tholeiite modified by partial loss of volatile components. I II III Primitive terrestrial Parental lunar Column I minus oceanic tholeiite a crust magma (7% SiO 2"+ 1.8% Na 2 O) SiO 2 50.3 47.3 47.5 TiO 2 0.73 0.9 0.8 AI~O 3 16.6 18.4 18.2 FeO 7.99 9.4 8.8 MgO 10.2 12.3 11.2 CaO 13.2 11.3 14.5 Na20 2.00 0.2 0.2 100 MgO 69 70 69 MgO + FeO REE 9 -10 ~10 Chondrites a Sample DSDP 3-18 from Frey et al. (1974). CALC. LIQUID I.U 10- r~ "t- ° / ~.o - 15415 14.1 ~. ANOR. :E .< u') 0.1 Ba LaCe Nd Sm Eu Gd Dy Er Yb Lu Sr Fig. 4. Calculated rare earth element abundances in parental (basaltic) liquid from which lunar anorthosite 15415 is believed to have crystallized (after Hubbard et al., 1971). BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 411 calculated the REE abundances of the parental basaltic magma from which the anorthosite 15415 has crystallized, using experimental plagioclase-liquid partition coefficients. The parental liquid was found to possess chondritic relative abundances of the REE elements at about 10 times the absolute chondritic abundances (Figure 4). It is similar in this respect to many primitive terrestrial tholeiites including the example chosen in Table V. Hubbard et al. (1971) also showed that other low-K lunar anorthosites possessing positive Eu anomalies had likewise crystallized from parental liquids with generally similar REE abundances. Laul et al. (1974) demonstrated that lunar anorthosites possessing marked positive Eu anomalies were derived from parental magmas containing 8 to 15 times the chondritic abundances. McCaUum et al. (1975) found that the parent liquids in equilibrium with analogous lunar anorthosites contained 4 to 12 times the chondritic REE abundances. The similarity in major element and REE abundances between the most abundant kind of primitive lunar basaltic magma and the most abundant class of primitive terrestrial basaltic magma, modified only by the partial loss of two of the most volatile components (Na20 and SiO2) is believed to be of considerable genetic significance. We will return to this point subsequently. B. COMPOSITION OF SOURCE REGION FROM WHICH PLC MAGMA WAS DERIVED The composition of the (PLC) magma which is believed to be parental to the lunar crust is given in Table IV. Olivine Fo88 plus plagioclase crystallize simultaneously on its liquidus at 1250°C (in vacuum). The absence of a negative europium anomaly (Figure 4)implies that plagioclase was not a residual phase in the source region, and that the magma had not crystallized substantial amounts of plagioclase after segregating from its source region. Olivine, however, remains on the liquidus to about 7 kb (Kesson and Ringwood, t976c) and was probably present as a residual phase in the source region. It is possible, therefore, that the composition of the PLC magma has been modified by crystallization of olivine at relatively shallow depths and that the primary magma was richer in normative olivine than the composition given in Table IV. The primary PLC magma may have formed by extensive partial melting of the outer few hundred kilometers of the Moon. After segregating from the residual phases in its former source region, the vast ocean of primary PLC magma would have precipitated olivine until plagioclase saturation was reached. Most probably, therefore, a layer of olivine cumulates underlies the crust (Figure 5). If the PLC magma segregated from residual refractory ferromagnesian phases at average pressures less than 7 kb (150 km), the experimental phase equilibria show that the residual phase consisted of olivine (Foa8). At this pressure, reconnaissance runs showed that olivine is joined by orthopyroxene (Kesson and Ringwood, 1976c). For a model primary PLC magma containing ~> 10% more normative olivine than the composition given in Table IV, olivine remains on the liquidus to pressures exceeding 15kb and is joined by aluminous orthopyroxene which crystallizes over a wide pressure interval. These relations suggest that the residual, refractory phases remaining in the source region after segregation of the primary PLC magma consisted of olivine alone, or of olivine + orthopyroxene. 412 A. E. RINGWOOD HIGHLAND MARE CRUS" lOq "1- 4C 5C Fig. 5. PetIological structure of Moon at stage of generation of mare basalts. The PLC magma possesses a CaO/A1203 ratio of 0.61, substantially smaller than the chondritic ratio of 0.8. If the source region possessed a near-chondritic CaO/A1203 ratio, as suggested by its near-chondritic REE abundances (Figure 4), then the residual phases remaining behind after magma generation must have possessed a high net CaO/A1203 ratio. However, the liquidus orthopyroxenes observed in PLC and (PLC + olivine) compositions have CaO/AI203 ratios smaller than 0.4. The presence of orthopyroxene as a residual phase would exacerbate the discrepancy with the chondritic CaO/A12Os ratio. Ortho- pyroxene does not therefore appear likely to be a major residual phase in the source region. On the other hand, olivine containing 0.4 to 0.5% CaO but no/11203, is a required residual phase and would modify the CaO/A1203, in the desired direction, although not sufficiently far as to produce an overall chondritic ratio. If the latter were characteristic of the bulk system from which the lunar crust was BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 413 TABLE VI Derivation of composition of source region of lunar crust, equivalent to the mean composition of the outer few hundred kilometers of the Moon. I II III PLC Liquidus olivine Bulk comp. of outer Moon: magma FoBs 20% column I plus 80% column II SiO 2 47.3 41.1 42.3 TiO: 0.9 - 0.2 A1203 18.4 - 3.7 Cr:Oa 0.13 0.3 0.3 FeO 9.4 11.5 11.1 MgO 12.3 46.1 39.3 CaO 11.3 0.5 2.7 Na20 0.2 - 0.04 100 Mg 70 88 86 Mg + Fe REE Chondrites derived, then large quantities of clinopyroxene with CaO >> AlzOa must be assumed to be residual in the source region. Although this possibility cannot yet be finally excluded, it seems improbable in the light of existing phase equilibria data. It seems likely therefore, that the residual, refractory component remaining in the lunar mantle after extraction of the primary PLC magma consisted of pure dunite, similar to, or somewhat more magnesian than the observed liquidus PLC magma olivine (Fo88, 0.5% CaO). If we assume that the PLC magma represented a 20% partial melt of its source region, (comparable to primitive oceanic tholeiites in the Earth's mantle) then the CaO/ Al203 ratio of the bulk system would be 0.72, significantly, but not excessively below the chondritic ratio. The mean composition of the outer few hundred kilometers of the Moon (crust + mantle) on the above assumptions is given in Table VI. In some respects it is complementary to that inferred for the source regions of mare basalts, possessing a CaO/A12Oa ratio smaller than chondritic, whereas the corresponding ratio in the mare basalt source is higher. It is possible, therefore, that the bulk Moon possesses a chondritic mean CaO/A1203 ratio. In comparison to the Earth's mantle, the outer 400 km of the Moon appears to be richer in olivine, whilst the deep interior is richer in pyroxene. This may well reflect an intrinsic zonation in SiO2 content. The average olivine/pyroxene ratio (and SiO2 content) of the entire Moon could well be similar to that of Earth's mantle. The mean Mg value of 86 obtained for the outer 400 km of the Moon may be a slight underestimate, depending upon the degree of fractional crystallization of olivine which may have occurred follow- ing magma segregation. The mean Mg number of othe outer regions of the Moon is thus probably not significantly different from that of the Earth's mantle (89 - Table VII). 414 A.E. RINGWOOD TABLE VII Comparison of estimated bulk composition of entire Moon with model pyrolite composition of the terrestrial mantle. I II III IV Mare basalt Bulk comp. Bulk comp. Pyrolite d source region of outer of entire comp. a Moon b Moon e SiO 2 46.8 42.3 44.6 45.1 TiO~ 0.4 0.2 0.3 0.2 AI203 3.7 3.7 3.7 3.9 Cr203 0.5 0.3 0.4 0.3 FeO 16.7 11.1 13.9 7.9 MgO 27.5 39.3 33.4 38.1 CaO 4.1 2.7 3.4 3.1 Na20 0.06 0.04 0.05 0.4 100MgO 75 86 81 89 MgO + FeO REE 2 ~ 2 ~ 2 - 2 Chondrites Ab 0.5 0.3 0.4 3.4 An 9.9 10.0 9.9 8.9 Di 8.5 2.8 5.6 5.2 Hy 41.8 10.0 25.8 18.4 O1 37.9 76.1 57.2 63.3 Chr 0.7 0.4 0.6 0.5 Ilm 0.8 0.4 0.6 0.4 a Average from Tables II and III. b From Table VI. e Obtained by averaging columns I and II. d From Ringwood (1975a), Table 5-2, column 8, with A1203 adjusted to chondritic CaO/A1203 ratio as discussed in footnote 8 of the Table. However, the Mg number of the source regions of mare basalts (75-80) is smaller than that of the Earth's mantle (Section 3). Thus, the entire Moon appears to contain slightly more total FeO than the Earth's mantle. A sketch of the structure of the Moon as inferred from the preceding discussions is given in Figure 5. It should be mentioned that although the author would have preferred to interpret the available data so as to yield similar bulk compositions for the PLC and mare basalt source regions, it has not been possible to accomplish this objective. It seems that the source regions indeed differ in composition but in a complementary manner. Whilst neither individual region is identical to the inferred pyrolite composition of the Earth's mantle (Ringwood, 1975a), it is a remarkable fact that if these compositions are combined in similar proportions, the pyrolite composition (minus elements more volatile than sodium) is closely approached (Table VII). The principal difference between the mean bulk moon composition and pyrolite (apart from volatile elements) seems to be BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 415 that the Moon is significantly richer in FeO. (However the difference need not be as large as is indicated in Table VII). 6. Heat-producing Elements Although uranium, thorium and potassium are present as trace and minor elements in the Moon, a topic which is treated in Part II of this series, the fact that these elements control an important global property of the bulk Moon, namely the heat flow, justifies a discussion of their significance in Part I. The lunar heat flux values reported earlier (Langseth et al., 1972, 1973) implied, if correct, that the Moon possessed more than five times the chondritic abundances of uranium and thorium, and were widely held to justify lunar models in which related involatile elements, e.g., Ca, A1, REE were enriched to similar degrees. These models conflicted with the evidence from experimental petrology discussed in previous sections. A major reduction in estimates of lunar heat flow has since been made on the basis of a much more reliable estimate of the bulk thermal conductivity of the lunar regolith (Langseth etal., 1976a, b). The revised heat flow value at the Apollo 17 site is 1.4/lW cm-2 and at the Apollo 15 site, 2.1/~W cm-2. The large difference between these values empha- sized the need for considerable caution in using arguments based upon estimated 'mean' lunar heat fluxes. Any such estimates, based upon only two measurements, are likely to possess large uncertainties, as pointed out by Langseth et al. (1976b). The higher heat flow at the Apollo 15 site was also correlated with much higher than average concentrations of thorium (and by inference, U and K) than in the Apollo 17 region. It seems likely that magnetic fractionation processes have resulted in a very strong upward concentration of radioactive elements in the near-surface layer, as has happened in the terrestrial crust. Thus, the higher heat flow at the Apollo 15 site is most reasonably interpreted as being due to a layer of rocks abnormally rich in radioactive elements (Langseth et al., 1976b). This would justify an analysis similar to that which is commonly applied to measurements of terrestrial heat flow and surface radioactivity in which heat flow is plotted against surface radioactivity. This is done in Figure 6, in which the slope of the line defined by the two data points defines the thickness of the surface layer of variable radioactivity. On the basis of this model, the mean lunar heat flow could be esti- mated if the mean thorium content of the surface layer over the entire Moon were known. Langseth et al. (1976b) estimated on the basis of the orbital gamma ray Th measure- ments by Metzger et al. (1974) and the heat flow versus radioactivity plot shown in Figure 6, that the mean lunar heat flow is 1.76/JW cm-2. If all this heat was generated by radioactivity (taking Th/U = 3.7, K/U = 2000) a mean bulk lunar concentration of 45 ppb would be required. They also point out that a steady-state thermal regime for the Earth (with all radioactivity concentrated in the mantle) would require a mean uranium concentration of 42 ppb. The agreement between lunar and terrestrial mantle uranium concentrations is most striking. In the author's opinion, the mean terrestrial and lunar heat flows attributable to 416 A, E. RINGWOOD :E D 0 Z uJ LU "5 i 5 16 THORIUM CONCENTRATION ppm Fig. 6 Plot of heat flow observed at Apollo 15 and 17 sites versus mean thorium contents esti- mated for the soil and rocks of these regions based upon orbital gamma ray measurements and labora- tory measurements of samples. After Langseth et al. (1976b). Estimates of Th content in major lunar petrologic provinces (Metzger et al., 1974) are included. radioactivity may both require substantial reductions, but these do not affect the signi- ficance of the agreement. It now seems clear that the mean temperature of the earth was very high (~ 2500°C) soon after its formation, owing to the large amount of energy released by rapid formation of the core (e.g., Ringwood, 1975a). Under these conditions, 20 to 30% of the present terrestrial heat flow could result from the original heat content, of non-radiogenic origin (MacDonald, 1959, 1965) dependent on the efficiency of con- vective heat transport. Thus, the uranium content of the Earth's mantle may be closer to 30ppb. In the case of the Moon, it seems likely that the orbital gamma ray data of Metzger et al. (1974) sampled a region in which the highly radioactive mare regions in the western plains of the Moon's nearside were over-represented. It is believed that if the Th content of the highland provinces, particularly on the far side, had been weighted according to the actual areas which they occupy, rather than by the accidental sectors traversed by the orbiters, then a substantially lower mean surface abundance of Th might be derived. An analysis of this kind is currently in progress (Metzger, personal communi- cation). In the meantime, we note that if Taylor and Jake~ (1974) estimate of 1.5 ppm for the mean surface abundance of Th in the lunar crust is adopted, then, on the basis of Figure 6, the mean lunar heat flux would be 1.37/aW cm -3, corresponding to a mean lunar uranium abundance of 35 ppb. Corrections for the presence of a small amount of BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 417 original heat component in the lunar heat flow,* and heat refraction effects because of greater thermal conductivity in the sub-mare crust (Conel and Morton, 1975) could easily reduce this estimate to 30 ppb. We concluded earlier that both the Moon and the Earth's mantle contain about twice the (ordinary) chondritic abundances of Ca, AI and REE. The mean uranium content of ordinary chondrites is 13 ppb (Morgan, 1971) which would suggest a mean lunar uranium concentration of 26 ppb. The value of 30-35 ppb inferred from the lunar heat flow is somewhat higher but in view of uncertainties introduced by the 3-fold variation of uranium concentrations in ordinary chondrites on which the average of 0.013 ppb was based, and a twofold dispersion among primitive Type 1 carbonaceous chondrites, the differences can- not be considered to be significant. 7. Oxygen Isotopes About half of the Moon and of the Earth's mantle (by weight) is composed of oxygen, which possesses three isotopes 160, 170 and 180. Variations in oxygen isotope ratios between planets and various classes of meteorites are caused by two factors: (1) primary isotopic inhomogeneities in the oxygen in various regions of the solar nebula prior to accretion, presumably caused by the survival and inhomogeneous distribution of a particular class of interstellar grains rich in 160 (Clayton et al., 1973, 1976); (2) chemical isotopic fractionations, caused for example, by differing temperatures at which the solid matter which accreted to form planets and meteoritic parent bodies, equlibrated and became separated from the gases in the parental solar nebula. The oxygen isotope compositions of both lunar and terrestrial basalts are identical (e.g., Clayton et al., 1976). This implies that the material from which Earth and Moon was formed was homogeneous with respect to the ~60-rich interstellar grain component and also separated from the nebula at the same mean temperature. In contrast, the oxygen in most classes of meteorites possesses distinctly different porportions of the 160-rich grain component. Some differentiated meteorites - the eucrites and howardites - possess oxygen which differs only very slightly in this respect from terrestrial and lunar oxygen and may perhaps be identical (Clayton et al., 1976). However, the oxygen in these meteorites is distinctly different from terrestrial-lunar oxygen owing to a chemical fractionation effect (Taylor et al., 1965). The only classes of meteorites which possess identical oxygen isotopic compositions to the Earth and Moon are the enstatite chon- drites and achondrites. Meteorites provide clear evidence that the 160 component of dust grains was not uni- formly distributed within the solar nebula prior to accretion into meteorite parent bodies and planets. Moreover, the existence of chemically-derived oxygen isotope fractionations Tokst~z et al. (1976) assumes this to be negligible because of heat transfer by convection in the lunar interior which would lead to a steady state heat generation-dissipation situation. However, the author believes that the strong chemical heterogeneities in the lunar interior discussed in the paper would inhibit thermal convection. 418 A.E. RINGWOOD indicates that the parental material of different kinds of meteorites and planets separated from the solar nebula at differing temperatures. In the light of these observations, the identity in oxygen isotopic compositions between Moon and Earth acquires an added significance. When considered in combination with evidence assembled in previous sections of a similarity in the bulk compositions of the Moon and the Earth's mantle, a close genetic relationship between Moon and Earth is strongly suggested. This conclusion is reinforced by the consideration of trace element data in Part II of this series. 8. Physical Properties of the Moon The present model of the internal structure and composition of the Moon (Figure 5) has been developed exclusively from petrologic-geochemical evidence. Its consistency with lunar geophysical evidence is now examined. A. DENSITY The density of the lunar crust is estimated to be 2.95 g cm -3 (Kaula et al., 1974). This is in agreement with the density of the derived crustal composition (Table IV), providing that a small amount of porosity is assumed. The structure of lunar breccias support this latter assumption. The mean density of the lunar interior below the crust is 3.39 g cm -3 (Kaula et al., 1974). According to the presence model, the mean density (at atmospheric pressure) of the upper mantle (to 400km) is 3.34gcm -3 and that of the mantle below 400kin is 3.44 g cm -3. These are estimates based upon inferred mineral assemblages, but without corrections for the effects of thermal expansion and compressibility. These effects almost cancel out under lunar conditions and can be ignored in a first approximation. The mean density of the lunar upper and lower mantles of Figure 5 is thus 3.39 gcm -3, in agree- ment with observations. A metallic core is not required. B. MOMENT OF INERTIA Kaula et al. (1974) estimate the moment of inertia coefficient I/MR 2 of the Moon as 0.395 -+ 0.005. The moment of inertia coefficient of the model shown in Figure 5 is 0.394, in excellent agreement with the preferred observational value. A more detailed dis- cussion of density and moment of inertia constraints is provided by Kaula et al. (1974). C. SEISMIC VELOCITY DISTRIBUTION Nakamura et al. (1974) found that the seismic P velocity in the upper mantle from 60 to 300kin is very close to 8.0km s -1 and its velocity-density relationship is matched better by a layer dominantly composed of olivine Fo8o-84, rather than of pyroxene. Our petrological model (Figure 5) also has a layer of olivine between 60-400 km. Although its composition (FoBs) is slightly more magnesian than the preferred composition of Nakamura et al. (Foso_s4), the difference is within the limits of uncertainty. BASALTIC MAGMATISM AND THE BULK COMPOSITION OF THE MOON I 419 Around a depth of 300km, Nakamura et al. (1974) and Latham et al. (1975) find evidence for a significant decrease in P and S wave velocities. The latter authors suggest that this marks the transition from the differentiated outer layer to the primordial mantle below. This interpretation agrees with our present model (Figure 5) except that the transition is placed closer to 400 km. The decrease in seismic velocities is readily explained by the increases in pyroxene relative to olivine and in iron content, which occur at this depth in the petrologic model. 9. Conclusion Although not identical, the major element composition of the bulk Moon is similar to that of the Earth's mantle. Moreover, this similarity extends to the abundances of the rare earths, uranium, thorium and probably also to many other involatile, lithophile elements. A model of the internal structure of the Moon based exclusively upon petrological- geochemical considerations provides an excellent explanation of the depth-distributions of density and seismic velocities within the Moon. Acknowledgements The author is greatly indebted to Dr S. Kesson for assistance in preparing the manu- script and for constructive criticisms. Helpful comments by Dr R. Brett are also greatly appreciated. References Adler, I., Trombka, J., Schmadebeck, R., Lowman, P., Blodget, N., Yin, L., and Eller, E.: 1973, 'Results of the Apollo 15 and 16 X-ray Experiment', Proc. Lunar Sci. Conf. 4th, 2783-2791. Anderson, A. T., Jr.: 1971, 'Exotic Armalcolite and the Origin of Apollo 11 Ilmenite Basalts', Geo- chim. Cosmochim. Acta 35,969-973. Anderson, D. L.: 1973, 'The Composition and Origin of the Moon', Earth Planet. Sci. 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Discover Space – Springer Journals

Published: Jul 1, 1977

Keywords: Olivine; Bulk Composition; Basaltic Magma; Mare Basalt; Lunar Crust

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Basaltic magmatism and the bulk composition of the moon

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Basaltic magmatism and the bulk composition of the moon

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