Experimental petrology and petrogenesis of mare volcanics

0016-7037/92/$5.00

Geochimico et Cosmochimica Acta Vol. 56, pp. 2235-2251 Copyright Q 1992 Pergamon Press Ltd. Printed in U.S.A.

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Experimental petrology and petrogenesis of mare volcanics* JOHN L~NGHI Lament-Doherty Geological Observatory of Columbia University, Palisades, NY 10964, USA (Received February 15, 1991; accepted in revisedform March 20, 1992)

Abstract-Mare volcanics consist of basalts and picritic pyroclastic glasses spanning a wide range of TiOz concentration. The more primitive low-Ti basalts and picritic glasses have olivine alone on their lowpressure liquidi. Most of the chemical variation among the low-Ti basalts is the result of olivine fractionation in a series of parental MgO-rich liquids differing in TiO2 concentration. With one possible exception (Apollo 17 VLT) none of the picritic compositions is a suitable parent for any of the observed low-Ti basalts. Most of the chemical variation among the high-Ti basalts is the result of a series of magmas fractionated along the low-pressure olivine + armalcolite/ilmenite cotectic. All of the picritic high-Ti glasses have olivine alone on the liquidus, but none is a suitable parent for any of the basalts. Volcanics with intermediate Ti02 concentrations (5 to 10 wt% ) are widespread in the maria, even though they are not well represented in the sample collections; however, there is no evidence either among the samples or from remote sensing studies of basalts with > 13 wt% TiO2 that would be expected as differentiates of the picritic glasses with the highest Ti02 concentrations. Controlled-cooling-rate crystallization studies on a variety of mare compositions have provided the basis for reconstructing the size and, in some cases, stratigraphy of mare flows. Groundmass textures, crystal size, crystal morphology, nucleation density, and zoning patterns have all been employed to quantify cooling histories of mare basal& A single-stage linear cooling rate may produce a porphrytic texture. Rapid cooling may also cause plagioclase to crystallize after a mineral that it precedes during slow cooling. Mare basalts are highly reduced. Mineral assemblages and intrinsic oxygen fugacity measurements indicate fs below the wtistite-iron buffer and at or near iron metal saturation. Accordingly, experiments run in high-purity iron capsules gained or lost little iron. Most basalts are undersaturated with respect to sulphur, so reduction through sulphur volatilization cannot be invoked to explain the presence of iron in olivine phenocrysts. The low oxidation state is most likely the result of melting a reduced interior under fluid-absent conditions. The progressive reduction of Cr3+ --, Cr ‘+ and Ti4+ + Ti 3+in lunar melts permits the elimination of the Fe3+ that is present at iron-saturation in simple systems. Crystal-liquid partition coefficients determined from melting experiments have been used in a wide range of calculations of major and trace element evolution. The coefficient for Fe-Mg exchange between olivine and liquid apparently varies with TiOl concentration of the liquid and is particularly useful in assessing whether fine-g&red rocks have excess olivine. Nb and Ti have excess concentrations in mare basalts relative to adjacent REE in incompatibility diagrams. These excesses or positive anomalies are consistent with ilmenite accumulation in light of measured partition coefficients and imply continuous variation of accumulated ilmenite even in the low-Ti mare source regions. Pressures of multiple saturation (olivine + pyroxene + Cr-rich spine1 + ilmenite) are in the range of 5 to 12.5 kbar for primitive mare basalts and in the range of 18 to 25 kbar for the picritic glasses. LowCa pyroxene is the only pyroxene along the liquidus of the low-Ti basalts and glasses; however, augite is the pyroxene most commonly observed along the high-pressure liquidi of the high-Ti basalts: High-Ti picrites have augite in the subliquidus region at intermediate pressures where olivine is the liquidus phase, but orthopyroxene is the liquidus phase at multiple saturation. Because of the steep depth/pressure gradient in the outer portion of the Moon (20 km/ kbar), these pressures imply: (a) great depths of melting within the Moon and (b) some means of transporting magmas hundreds of kilometers to the surface without significant chemical modification, if both olivine and pyroxene were left in the residuum. Modeling of major elements during polybaric partial melting suggests that it is possible to reproduce the composition and high-pressure signature of the low-Ti (green) picritic glasses by accumulating small degrees of melt extracted from an upwelling source region. In such a case, melting of the differentiated source must begin at > 1000 km depth (40 kbar ) and cease at - 100 km ( 5 kbar ) . If this model is correct, then experimental determination of the pressure of multiple saturation gives an average pressure of melting: the onset of melting is at higher pressure and actual segregation of the melt from the mantle is at lower pressure.

INTRODUCI’ION * Presented at a workshop on Mare Volcanism and Basalt Petrogenesis held on October 27 and 28,1990, during the Annual Meeting of the Geological Society of America in Dallas, Texas, organized by Lawrence A. Taylor and John Longhi.

MARE VOLCANISMproduced a diverse suite of basalts and volcanic glasses. Perhaps the most dramatic feature of these volcanics 2235

is the wide range of Ti02 concentration

(0.3 to

2236

J. Longhi

16.4 wt% ) that dwarfs the range in common terrestrial basalts by more than a factor of three. Other prominent features are the absence of water, the low abundances of alkalies and other volatile elements, the absence of ferric iron and the presence of iron metal, and the nearly ubiquitous depletion of Eu relative to the middle rare earth elements (BVSP, 198 1; TAYLOR, 1982). In the aftermath of the first Apollo sample return, there developed the concept of a primordial, global melting event, a “magma ocean,” to account for a plagioclase( Eu)-enriched crust overlying a plagioclase( Eu)-depleted mare basalt source region (SMITH et al., 1970; WOOD et al., 1970). Since then a major goal of lunar science has been determining the scale of this melting event and experimental studies of mare volcanics have played an important role in constraining minimum depths of the primordial melting These studies have also helped to constrain the composition and oxidation state of the interior as well as the cause of the large variation in the TiOz concentration. However, the nature of sample collecting by the astronauts produced many uncertainties that terrestrial geologists typically do not face because all of the basalts were more or less random samples of Boat excavated from bedrock by meteorite impact and strewn about the surface. The stratigraphy of the samples had to be reconstructed by a combination of textural, chemical, and chronological data, the thickness of the source lava flows had to be calculated from estimates of cooling rates, and most importantly, some judgment had to be made as to which crystalline basahs had liquid compositions. Even with the eventual recognition of some mare glasses as volcanic (not impact generated), deciding which basalts had liquid compositions remained important because most basalts and glasses are apparently not directly related (see later). During the initial phase of sample return and analysis, most of the experimental work on mare basalts focused on determining melting relations over a range of pressure (e.g., O’HARA et al., 1970; RINGWOOD and ESSENE, 1970; GREEN et al., 197la, 1971b, 1975; KUSHIRO et al., 1971; HODGES and KUSHIRO, 1972; LONGHI et al., 1974; KESSON and LINDSLEY,1975; WALKER et al., 19’72, 1975, 1976a, 1977: GROVE and VANIMAN, 1978) with the aim of determining low-pressure crystallization sequences, parental liquid compositions, and pressures at which the liquidi were multiply saturated. Experimental studies of picritic volcanic glass compositions (DELANO, 1980; CHEN et al., 1982; CHEN and LINDSLEY,1983) represent the last of this type of work. Other studies were aimed at inducing silicate liquid immiscibility (RUTHER~RD et al., 1974), replicating textures (LOFGREN et al., 1974; WALKER et al., 1976a), and inducing specific zoning patterns in minerals (GROVE and WALKER, 1977) in controlled-cooling-rate experiments. The most recent experimental work has dealt with constraining the oxidation state of the lunar interior by measuring the spinel/liquid partitioning of Cr as a function of oxygen fugacity (DELANO, 199Oa). Data produced in experimental studies of mare basalts and related compositions have also provided a basis for various empirical models of liquidus equilibria (LONGHI, 1987, 199la) and crystal/liquid partitioning of various major and trace elements (MCCALLUM and CHARETTE, 1978; LONGHI et al., 1978; DELANO, 1980; MCKAY, 1982, 1986; MCKAY et al., 1990; COLSONet al., 1988). The various an-

alytical expressions developed in these studies have been widely used in calculations of major and trace element evolution during partial melting and fractional c~stallization processes. This review builds on an earlier review of experimental studies by KESSON and LINDSLEY,( 1976). Readers interested in a more thorough discussion of the problems encountered with sample containers, a more complete bibliography of the early experimental studies, or a review of devolatilization studies and hypotheses should consult this earlier paper. LOW PRESSURE Phase Relations Figure 1 illustrates the variation of TiOz vs. MgO in mare volcanics. Where possible, only the compositions of finegrained, hand-specimen-sized basalts are shown in order to avoid trends produced by crystal accumulation. High- and low-Ti basahs show different overall variations. The high-Ti basalts show an overall positive correlation of Ti9 and MgO, whereas the low-Ti basalts show no overall correlation. Upon closer examination, however, the array of low-Ti basalt compositions appears to be composed of a series of sub-parallel, sub-horizontal trends of basalts from a given landing site. Thus there appears to be two distinct trends for Ap12, two for Ap15, one for Apl7 with “orphans” from Ap14, Ap17, LUNA 16, and LUNA 24 sprinkled about. Chemical studies have shown that the well-defined sub-horizontal trends are olivine-control lines relating comagmatic basalts (RHODES and HUBBARD,1973; RHODESet al., 1977) and that apparent coincidences between basahs from different landing sites (e.g., the Ap 14 sample lying along one of the Ap15 trends) break down for other elements. As summarized in Fig. 2, a number of experimental inv~ti~tions have shown that divine (+ minor chromite) is the liquidus phase for the more MgOrich samples in these trends (Ap12: GREEN et al., 197 la& BIGGAR et al., 1971; KUSH~ROet al., 1971; GROVE et al., 1973; RHODES et al., 1979; Apl4: WALKERet al., 1972; Apl5: KESSON and LINDSLEY, 1975; WALKER et al., 1977). The presence of MgO-rich vitrophyres 12008 ( DUNGAN and BROWN, 1977) and 12009 (BRETT et al., 1971) plus applications of Fe-Mg olivine/liquid exchange coefficients (discussed later) to other MgO-rich compositions (e.g., WALKER et al., 1977) demonstrate the existence of olivine-saturated low-Ti liquids with 11 to 13 wt% MgO that were parental to less magnesian basal& This combination of ~~o~aphic, chemical, and experimental data obviates the hypothesis of BIGGAR et al. ( 197 1) that the parental magmas were low in MgO and that the olivine-control trends in Fig. 1 were the result of crystal accumulation. The only group for which olivine cannot be readily identified as the major liquidus phase is a series of very low-Ti (VLT) ferrobasaltic glasses from Luna 24. GROVE and VANIMAN( 1978) have shown plagioclase to be the sole liquidus phase for these compositions which are similar to the LUNA 24 crystalline basalt fragments (Fig. 1), albeit slightly higher in A1203. These LUNA 24 glasses have both low MgO and Mg’ (MgO/(MgO + FeO) - 0.3 1). If these glasses are undiffe~ntiat~ mantle melts (quenched primary magmas), then the phase relations re-

2231

Petrogenesis of mare basalt

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FIG. I. Variation of TiOz and MgO in mare bar&s and picritic glasses ( LONGHI, 1987 ). (a) Low-Ti compositions; (b) high-Ti compositions. Dashed lines are calculated fractional crystallization paths. Olivine is the first phase to crystallize; abbreviations mark appearance of additional phases: Sp = chromite; Pi = pigeonite; PI = plagioclase; Aug = augite; Ilm = ilmenite; Ar = annalcolite. Sample abbreviations: numbers refer to Apollo and LUNA landing sites; II = ilmenite basalts; OB = olivine basalts; PB = pigeonite basalts; QNB = quartz-normative basalts; L = LUNA; VLT = very low-Ti. Glass abbreviations: G = green; Y = yellow; 0 = orange; R = red, B = black, “A” and “V” refer to groupings of DELANO ( 1986); subscripts refer to column numbers in Table 3 of DELANO ( 1986).

ported by GROVE and VANIMAN ( 1978) show that source region was devoid of olivine as well as being unusually ferroan. Possible alternatives are either that the glasses are impact melts of a mixture of evolved basalt plus (highlands?) plagioclase or that delayed nucleation of plagioclase during rapid subsurface crystallization of a more primitive, olivine-saturated magma caused the residual liquid to overstep the olivine + plagioclase co&tic prior to eruption (GROVE and BENCE, 1979). More and larger samples would help shed some light on this problem. Whether the high-MgO, low-Ti parental basalts have primary compositions is problematical. As Fig. 1 illustrates, there are picritic volcanic glasses (DELANO, 1986) with higher MgO contents than the basalts (these glasses also have higher Mg ‘) . Figure 2 shows that the picritic glasses have much higher liquidus temperatures than the mare basal& These glasses also have olivine on the liquidus (STOLPER, 1974; GROVE and VANIMAN,1978) and their calculated fractionation paths (dashed lines are discussed later) are sub-parallel to the natural trends. However, when these fractionation paths are examined in terms of all the major elements, none of the paths links a known variety of mare basalt with a known picritic composition with the possible exception of the Ap17 VLT basalts and the Ap17 green glass ( LONGHI, 1987). Thus it is not clear whether hmited sampling has obscured the relation between the low-Ti basalts and their picritic parents or whether there are. distinct families of basaltic and picritic

volcanics. This uncertainty extends to the high-Ti compositions, although the phase relations are different. There is no obvious correlation of TiOz and MgO among the intermediate- to high-Ti picritic glasses; however, there is a distinct positive correlation among the basal& Chemical studies have shown five distinct compositional groups among the Apl 1 high-Ti basal@ 4 low-K, 1 high-K ( BEATY and ALBEE, 1978), and 3 distinct groups among the Ap17 highTi basalts (RHODES et al., 1976); however, unlike the lowTi basal& there is no evidence of simple olivine control of the intragroup variations. With the exception of studies of one sample, 74275 (GREEN et al., 1975), experiments from several laboratories have shown that the high-Ti bar&s are multi-saturated with silicate and Ti-oxide within 25°C oftheir liquidi and several are multi-saturated within 10” ( O’HARA et al., 1970, LONGHIet al., 1974;WALKERet al., 1975; GREEN et al., 1975 ) . Generally, olivine and armalcolite are the highest

temperature phases (*minor Cr-spinel) with armalcolite reacting to form ilmenite at intermediate temperatures; plagioclase, augite, and pigeonite appear only at temperatures below the stability range of armalcolite: Olivine certainly reacts out when pigeonite appears, but may also begin to react in the presence of augite ( O’HARA et al., 1970). Thus, more magnesian and titaniferous samples will have olivine and armalcolite as near-liquidus phases, whereas less magneSian and titaniferous samples may have ilmenite, pyroxene, and plagioclase as near-liquidus phases. Co-crystallization of

J. Longhi

2238

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ferromagnesian silicates ( fplagioclase) plus Ti-oxides produced a trend of decreasing TiOz and MgO very similar to the overall trend of the basalts themselves. Given that local heterogeneities and accumulations of microphenocrysts may alter liquidus relations even in fine-gmined samples (WALKER et al., 1976b), it is reasonable to conclude that as a group the high-Ti basalts are multi-saturated (silicate + Ti-oxide) and that variable fractionation produces the overall positive slope in Fig. 1. Basalt 74275, which is the one clear exception (it has nearly 12 wt% TiOz and plots to the right of the main group of high-Ti basalts in Fig. 1 ), has olivine alone below the liquidus for more than 100’ before armalcolite appears and it may well be transitional between basalts and the picritic glasses (GREEN et al., 1975). However, 74275 also contains xenoliths of disaggregated micro-dunite and xenocrysts of olivine ( MEYERand WILSHIRE, 1974; WALKERet al., 1976b; DELANOand LINDSLEY,1982), so its suitability as a liquid composition is suspect. Armalcolite or ferropseudobrookite, ( Fe,Mg)Ti20s, is commonly rimmed by ilmenite in high-Ti basalts (e.g., HAGGERTY et al., 1970). Experiments in melt-free simple oxide systems(LI~~~~~~etal., ~~~~;KESSON~~~LINDSLEY,1975)

showed armalcolite reacting to ilmenite + rutile with decreasing temperature and Mg ‘. Melting experiments on basalts showed a limited temperature interval ( 1170” to 1140°C) and compositional range of the melt ( 12.5 to 10.5 wt% TiO*) over which armalcolite and ilmenite coexist ( LONGHI et al., 1974; KJSSON, 1975; WALKER et al., 1975; DELANO, 1980). Even though no rutile was reported in the melting experiments (possibly it was overlooked), all three sets of observations are consistent with a reaction relation between armalcolite and basaltic liquid to produce ilmenite. As with low-Ti basal& there has been some disagreement as to the composition of the parental high-Ti liquids. Experiments on most of the high-Ti basalts show plagioclase to be unstable near the liquidus. Yet O’HARA et al. ( 1975 ) maintained that the compositions of the parent liquids should have been similar to the compositions of the impact-generated mare soils, which have plagioclase on their liquidi because they are more aluminous and have lower Mg’ than most of the fine-grained basalts. O’HAFU et al. (1975) ascribed the compositional differences between rocks and soil to accumulation of skeletal microphenocrysts of olivine and armalcolite or ilmenite into the evolved parental liquid. However,

Petrogenesis of mare basalt results of the melting experiments on natural compositions are more consistent with the conventional hypothesis that the fine-grained basalts with olivine and armalcolite/ ilmenite microphenocrysts have bulk compositions similar to their parental liquids. In these experiments liquidus olivines have Mg’ values similar to those in the cores of the zoned microphenocrysts, thus indicating that no net gain of settled crystals has occurred. Also, experiments on fine-grained rocks tend to reproduce the general order of crystallization inferred from petrographic studies. Finally, there is the evidence, cited above, that some primitive, fine-grained low-Ti basalts have compositions that closely approximate those of low-Ti vitrophyres. So there are good reasons to expect that the compositions of the fine-grained high-Ti basalts are very similar to those of the liquids from which they crystallized. The two high-Ti picritic glasses that have been investigated, 74220 (170range: WALKER~~al., 1975)and 15318 (15Red: KESSON, 1975; DELANO, 1980)) have olivine on their liquidi ( T zz1300°C). Minor chrome spine1 follows next. Armalcolite is the second major phase to crystallize and develops a reaction relation with the liquid when ilmenite appears. Augite appears before plagioclase in the red glass experiments; augite and plagioclase appear together in the orange glass experiments. Both equilibrium experiments and model calculations of fractional crystallization (Fig. 1, dashed lines) show that the residual liquids will follow paths of increasing TiOZ and decreasing MgO when olivine (Kr-spinel) crystallizes, but that TiOZ and MgO will decrease together when olivine and a Ti-oxide co-crystallize. This overall trend continues as other silicates appear. Thus, except for the extensive near-liquidus crystallization interval of olivine, the phase relations of the high-Ti picritic glasses are very similar to those of the high-Ti basalts. Pigeonite is the only low-Ca pyroxene observed in mare basalts. Experiments on analogs of mare basalt liquids in which compositional parameters were varied show that the liquidus boundary between orthopyroxene and pigeonite is a complex function of Mg ’ and the proportions of wollastonitc (or diopside) and plagioclase components in the liquid such that decreasing plagioclase content and Mg ’ favor pigeonite and decreasing wollastonite favors orthopyroxene ( HUEBNER and TURNOCK, 1980; LONGHI and PAN, 1988 ) . For the plagioclase and wollastonite contents of typical mare basalt liquids saturated with olivine and low-Ca pyroxene, the orthopyroxenefpigeonite transition occurs at Mg’ of 0.6 in the liquid, yet experiments on the natural compositions typically show Mg’ - 0.4 in liquids with coexisting olivine and lowCa pyroxene. Consequently, the absence of orthopyroxene in mare basalts is largely an accident of their relatively low Mg’. One of the uses of the phase equilibria data on mare basalts has been the development of empirical expressions for the compositional dependence of liquidus boundaries depicted in pseudo-ternary projections ( LONGHI, 199 1a). Integration of the liquidus boundary expressions with crystal-liquid partition coefficients and geothermometers (see later) makes it possible to calculate equilibrium and fractional crystallization paths of mare compositions. Comparison of calculated melting sequences with those produced in melting experiments (Fig. 2) provides a calibration of the computational model.

2239

LONGHI ( 1987) has applied a version of the model to the question of the relation between picritic glasses and the mare basal& He calculated crystallization paths for a representative set of the picritic glasses, and some of these results are shown in Fig. 1 in terms of TiOZ and MgO as dashed lines. As mentioned above, these calculations show that when all the major elements are considered the differentiates of the picritic glass magmas look like typical mare compositions, yet in detail none of the DELANO ( 1986) set of 24 picritic glasses is a possible parent magma for any of the known varieties of mare basalt with the exception of the Apollo 17 green glass. The simplest explanation is that there is nearly a continuum of mare compositions from very low to very high TiOz and our random sampling is incomplete. However, there is some evidence that suggests otherwise: DELANO( 199Ob)hasshown that all the picritic glasses have a FeO/Sc ratio of approximately 4500, whereas the Apollo 11, 14, and 17 mare basalts have much lower ratios: ratios that are too low to be explained by fractionation of the olivine available in the picritic compositions. Thus, the possibility of separate modes of basaltic and picritic volcanism remains. If picritic magmas are the parents of mare basal& then the fractionation paths illustrated in Fig. 1 also suggest that our sampling of mare basalts is not only incomplete, but non-representative. Our practical experience has focused on low-TiOz and high-TiOz compositions, but the fractionation paths lead as well to potential intermediate-TiOZ ( 5- 10 w-t%) basalts (of which we have only a few chips) and to very highTiOz (> 15 wt%) basalts (of which we have none; LONGHI, 1990a). Remote sensing of absorption features in the visible and near infrared spectra of mare surfaces shows large areas of intermediate-TiOz compositions in the western mare, but nowhere is there the suggestion of very high-TiOz compositions ( PIETERSet al., 1980). So more sampling is needed to answer the simplest of questions: what is the actual range of mare basalt compositions?

Kinetics Lacking the stratigraphic information available to terrestrial petrologists collecting samples even in a cursory manner, lunar petrologists set about to recreate the cooling histories of mare basalts with the hope of constraining the thickness of lava flows, locating samples within a given flow, evaluating different zoning patterns in minerals, and describing in situ crystallization in volcanic rocks. Among the first results were the observations that porphyritic textures, typically large skeletal crystals in groundmasses with radiating intergrowths of pyroxene, plagioclase, and ilmenite, could develop during crystallization induced by a linear cooling rate ( L~FGREN et al., 1974). WALKER et al. (1976a) reached an even more surprising conclusion in their study of the crystallization behavior of porphrytic olivine basalt 12002: they concluded that the cooling rate actually slowed down between the time of crystallization of olivine phenocrysts and later growth of groundmass pyroxene. They found that the log of olivine nucleation density and the log of spacing of growth instabilities in pigeonite crystals were continuous functions of cooling rate and they input petrographic data from the rock into these functions to estimate that 12002 cooled initially at

J. Longhi

2240

1975). Liquid immiscibility develops only at relatively slow cooling rates (s2”C/hr) and at temperatures below 1000°C in mare compositions. These experiments did not produce as much compositional separation between high-Fe and highSi liquids as was measured in the natural basalts; also, silica phases, nearly ubiquitous in the residua of medium- to coarsegrained mare basal& were not reported in the run products. In a series of isothermal crystallization experiments LONGHI ( 1990b) was able to show that development of liquid immiscibility in these compositions was a stable equilibrium feature, but that the absence of a silica phase in the coolingrate experiments was the result of its failure to nucleate. Here too, though, the synthetic liquids did not separate chemically as far as their natural counterparts. In addition to exploring kinetic effects on phase equilibria and texture, several studies examined kinetic effects on crystal/liquid partitioning and zoning patterns. WALKER et al. ( 1975) showed that the exchange of Fe and Mg between olivine rims and liquid was unaffected by cooling rate. GROVE and BENCE( 1977) showed that Fe-Mg-Ca relations for pyroxene and liquid were also mostly unaffected by cooling rate, but that partitioning of Cr, Ti, and Al increased markedly with cooling rate; there were stoichiometric effects as well with higher Al contents requiring octahedrally coordinated Al. These studies were consistent with many earlier petrographic observations that mare basalts with similar bulk compositions, but different textures, had pyroxenes with different stoichiometries and concentrations of nonquadrilateral components. Thus, although Al contents of pyroxenes might have important barometric significance in plutonic rocks, Al contents of volcanic pyroxenes are useful only as cooling rate indicators. Several authors also recognized that the final zon-

- 1 “C/hr as olivine crystallized and later slowed to 0.1 to O.Z’C/hr as pyroxene began to crystallize. Some uncertainty about inferences based on nucleation density exists, however, because different sample containers (Fe-capsules, Fe-wire loops, Pt-wire loops) give different nucleation densities ( LOFGREN et al., 1979) and because even in a volcanic cooling regime some of the crystals present in a given volume of a flow may have settled into that volume from above (WALKER et al., 1976~). GROVE and WALKER( 1977) showed that the log of the width of groundmass plagioclase was a linear function of the log of cooling rate and used this relationship to estimate the cooling rate of several Apollo 15basalts.WALKER et al. ( 1976a) also demonstrated that cooling rate could alfect order of crystallization of some phases: Fig. 3 shows that plagioclase appears before ilmenite during equilibrium crystallization, but after ilmenite during crystallization at cooling rates of 1“C/ hr or more. Figure 3 also shows that rapid cooling lowers the crystallization temperatures of all the phases, but has the least effect on olivine. US~ELMANet al. ( 1975) compared textures of Apollo 17 basalts generated in controlled-cooling-rate experiments with those of the natural rocks and concluded that two-stage cooling histories were required and that most of the rocks crystallized from lavas that erupted with olivine, armalcolite, ilmenite, and chromian tilvospinel phenocrysts. Their conclusions are consistent with those discussed above that plagioclase was a relatively latecrystallizing phase in the high-Ti basalts and that the intersample chemical variation was controlled largely by fractionation along the olivine + armalcolite / ilmenite cotectic. Perhaps the most dramatic replication of textures was the production of immiscible silicate liquids in controlled-cooling-rate experiments (RUTHERFORDet al., 1974; HESSet al.,

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2241

Petrogenesis of mare basalt ing patterns in minerals are the result of growth in a fractionating liquid followed by post-growth re-equilibration. In the case of plagioclase and pyroxene, post-growth re-equilibration appears minimal in volcanic regimes because of slow diffusive exchange rates. But faster exchange rates for olivine (e.g., BUENINGand BUSECK, 1973) make smoothing of original zoning patterns a distinct possibility. To this end BIANCO and TAYLOR ( 1977) reported Fe-Mg zoning profiles in olivines grown at specific cooling rates and TAYLORet al. ( 1977) used these data as basis for calculations of the post-growth cooling rates needed to simulate zoning profiles in rocks 12002 (0.4’C/hr) and 15555 (0.2”C/hr). These results agree well with the results of WALKERet al. ( 1975) cited above on 12002. WALKER et al. ( 1977) used a similar approach to calculate the cooling rate needed to homogenize olivine in rock 15065, the coarsest-grained mare sample. Their result, 0.01 “C/hr, is consistent with a lava flow with minimum thickness of 10 meters. Crystal/Liquid

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.P 7 ZP Y

0.32 -

Kd"q,KO&i,j

(1980)

w”

6.30 -

a

. .

0.26 0.26 -

,,f/

1.”,.,, 1

.,.,.,.,.,, 0

2

4

6

Fe

The nearly constant value of KgJiq = 0.30 made it possible to estimate Mg’ in the liquid from which a given olivine crystallized or the Mg ’ of the liquidus olivine given a basalt composition; the temperature dependences of the single element partition coefficients could be used as geothennometers when both olivine and liquid compositions were known. LONGHI et al. ( 1978) showed that Fe-Mg distribution between olivine and mare basalt liquids was similar in general, but distinct in detail. They found that the logarithms ofthe simple partition coefficients were also linear functions of inverse temperature, but offset systematically from the terrestrial values (at given temperature the lunar coefficients are lower). They also found that KgJiq was independent of temperature, but that there was a distinct correlation of K$liq with Ti02 concentration in the liquid at low pressure (Fig. 4) such that K$‘“q in low-Ti liquids was higher than the terrestrial average, whereas K$-‘iq in high-Ti liquids was lower. The data of DELANO (1980) on the Apollo 15 red glass extend the correlation to even higher TiOl contents. His data also show KzAiq at constant TiOz increasing by 0.0 1/ kbar at least up to 20 kbar. One common application of the K$liq data has been to check the compositions of crystalline basalts with zoned olivine phenocrysts to see if the rocks have liquid compositions: if the predicted liquidus olivine composition is the same as that of the most magnesian olivine cores in the rock, then the rock has a liquid composition; if the predicted olivine is more magnesian than that found in the rock, then the rock composition is a mixture of olivine plus liquid (e.g., GREEN et al., 1971a; GROVE et al., 1973). LONGHI et al. ( 1978) also reported on the olivine/liquid partitioning of several minor elements in low-Ti liquids: they found that Mn partitioning was temperature dependent and

Experimental studies of the partitioning of various elements between crystals and basaltic liquids have provided the basis for a large number of calculations of the compositions of basaltic liquids: some as simple as determining the concentration of a single element, others more elaborate involving the evolution of a whole set of elements during fractional crystallization. Quantitative treatment of major element partitioning may be traced to the landmark investigation of FeMg distribution between olivine and terrestrial basaltic liquids by ROEDER and EMSLIE( 1970). They showed that the logarithms of the simple molar partition coefficients for Fe0 and MgO, K$‘liqand K&F, were linear functions of inverse temperature and that the Fe-Mg exchange coefficient, Kgliq, was nearly independent of temperature and composition, where KO$‘J= (“go)o’ - 0””

(Feo)“‘(Mgo)‘iq = (FeQ)‘iq(MgG)”

= (1 - Mg’)“‘(Mg’)liq (1 - M,‘)eq(Mg’)o’ *

Partitioning

_-

=

6

10

12

14

16

16

20

wt% TiO2 (liq) FIG,4. Variation of Kidiq with TiO? concentration in the liquid at low pressure.

J.

2242

Longhi

very similar to Fe, whereas Ti, Cr, and Ca were independent of temperature with average values of their respective molar partition coefficients of0.042 f .008,0.98 -I-0.28, and 0.033 + 004. AKELLA et al. (1976) and HUEBNERet al. (1976) found similar values of K$:q with considerable scatter and no obvious correlation with experimental conditions. GROVE and BENCE ( 1977) observed that the Fe-Mg exchange coefficient for pigeonite, ,KEge"q, in low-Ti liquids is lower than K$liq and that Kpd""' correlates positively with temperature (0.28 at 12OO’C vs. 0.17 at 975°C). Also corre2zg (negatively) with temperature are Kfrliq, KyFm"q, and Kcr . As expected, typical values of these coefficients at 1150°C are higher (0.10, 0.16, and 2.0, respectively) than those of olivine-liquid pairs. HUEBNERet al. ( 1976) also observed that Cr is strongly ,compatible in pigeonite, but with appreciable scatter in KcFm"'. These coefficients for A1203, TiOa, and Cr,O, are generally consistent with the data reported by LONGHIand PAN ( 1988)on a wide range of mareanalog compositions. There is relatively little data for augiteliquid partitioning in isothermal experiments whose significance is not obscured by compositional heterogeneity. The data of HUEBNER et al. (1976) indicate that Kggeliq is comparable to KgFm"q. There have not been any direct experiments on mare compositions comparable to those of DRAKE ( 1976) to measure the exchange of Na, K, and Ca between plagioclase and liquid. However, LONGHI et al. ( 1976) have reported on Fe and Mg partitioning between plagioclase and mare analog liquids. Because of the small size of the plagioclase crystals, they found it necessary to correct their Fe analyses,for secondary fluorescence. Average coefficients. are Kt? = 0.033 + .OlO, K gz = 0.065 f .016, and K$'lq= 0.51 + .13; however, they noted that the single-element coefficients were composition dependent and the relative uncertainty could be reduced considerably by the use of coefficients with stoichiometrybased composition terms analogous to the method of DRAKE and WEILL ( 1975). The measured value of Kg-"'implies that Mg’ in plagioclase will be intermediate between those of olivine or pyroxene and coexisting liquid.

Much of the work of measuring partition coefficients of rare earth (BEE) and other trace elements between coexisting phases in experimental run products with the electron microprobe is based upon the pioneering study of DRAKE and WEILL ( 1975 ). These authors showed that a variety of trace elements had near constant crystal-liquid partition coefficients even when doped to the percent level. Several similar studies on low-Ti lunar analog compositions followed (e.g., WEILL andMcKA~, 1975; MCKAY and WEILL, 1976,1977). Some of the measurements were confirmed by beta-track methods (IRVING et al., 1978). However, some of these data, particularly those of the light BFE, were plagued by secondary fluorescence errors. These errors have been avoided in more recent studies of trace element partition coefficients through the growth and analysis of larger crystals (MCKAY, 1982). Automation and improved stability of electron microprobes have also made it possible to measure elemental concentrations down to the 10 ppm level (MCKAY, 1986), so that severe doping of trace elements is no longer necessary. Where overlap exists, determination of partition coefficients by mechanical separation of phases in experimental charges followed by isotope dilution analyses has tended to confirm the microprobe data (NAKAMURA et al., 1986). Despite much work, the compositional and thermal dependences of most of the elemental partition coefficients have yet to be worked out. Some exceptions are the work of COL~ONet al. ( 1988) on the effect of liquid composition on olivine-liquid partition coefficients and the work of MCKAY et al. (1986a) on the effect of the wollastonite content of pyroxene on pyroxeneliquid coefficients. Table 1 summarizes some of the current data for partition coefficients appropriate for mare (lunar where necessary) compositions. There have been many applications of these and other trace element partition coefficients to the problem of mare petrogenesis; the works of HUGHES et al. ( 1988, 1989), BROPHY and BASU( 1989), and SHEARERand PAPIKE( 1989) are some recent examples. Most of these studies involve complex modeling of the proportions of phases in the mare basalt source region. Examination of the trace element contents of

Table 1. Seleced mace ekment partitionccefficienu (WIratio) determined in melting experimentson mme and lunar analog cmnposirions* 01

SC

Rb Sr zr Nb Ba La Ce Nd Sill EU Cd

pl

0.018’

(0.0037

P%

aug

ilm

(0.017)3 (1.6s <0.015 <0.015 (0.15)3

0.0000019~c o.OOc079 0.000589

(0.048)‘0

0.001029

(0.021)‘2

@‘2$”

Lillll

2.22

0.65’ 0.0024

0.155 0.025

0.336 0.85

1.25 1.45

(0.01 I)3 O.OO1724 0.00584 0.0114 0.00504

0.0214

(0.030)s (0.049)8 (0.12)s (0.19)s (0.18)8 (0.22)8

0.0039” o.O019t* 0.0010” 0.0011” 0.00616 0.0106” 0.062 0.0753 0.073” 0.4196

0.0082 0.0066

Dy Tm 0.062 Yb 0.01949 (0.0095)‘2 0.0874 (0.30)” LU 0.02506 (0.30)s 0.0466 Hf 0.0486 1.476 *: only Fe&wing analog compositions ate included 0: non-mare-lunar e: extrapolated 1: COLSON et al. (1988) - zwera e ofruns 117aand 117b; 2: IRVING et al. (1978); 3: MCKAY andWBILL(1976)- 12&l’ C,. 4: MCKAY et al. (1990) - Wo5; 5: McCALLlJM and CHAREl-fX (1978); 6: MCKAY et al. (1986b); 7: MCKAY and WEILL (1977); 8: MCKAY et al. (1986a) - Wo35; 9: MCKAY (1986); 10: WEILL UUIMCKAY (1975); 11: NAKAMURA et al. (1986); 12: MCKAY (1982).

Petrogenesis of mare basalt

for such reduced ions as Ti 3+ and Cr ‘+ in order to maintain charge balance (e.g., HAGGERTY et al., 1970). Small inclusions of Fe-Ni metal in olivine phenocrysts in Apollo 12 basalts indicated crystallization conditions at or near metal saturation ( BRETT et al., 197 1) . Recognition of an assemblage of Fe-rich olivine, cristobalite, and metal in the mesostases of Apollo 14 basalts (EL GORESY et al., 1972) constrained the oxygen fugacity to be near the quartz-fayalite-iron buffer, which is more reducing than the wiistite-iron buffer. A practical application of this knowledge was the use of iron capsules sealed in evacuated silica tubes to contain samples of basalt during melting experiments. Unfortunately, some of the iron employed in the initial series of experiments (e.g., GREEN et al., 197 1a,b, GROVE et al., 1973 ) contained impurities which reduced the lunar samples further, causing iron-loss, enhancing pyroxene stability, and raising liquidus temperatures. Through trial and error some petrologists settled on capsules of ultra-high purity iron, which produced minimal net gain or loss of iron in the charges (LONGHI et al., 1974; GREEN et al., 1975; KESSON, 1975; WALKER et al., 1977). These experiments yielded crystallization sequences at iron saturation. Other experiments, conducted in gas-mixing furnaces with platinum wire loops to suspend the charges, showed that the liquidus phases and crystallization orders of mare basalts were best matched by oxygen fugacities no more than one log unit below the wiistite-iron buffer; however, even in this narrow interval the crystallization order of ilmenite and augite could be reversed by varying oxygen fugacity such that ilmenite crystallizes first near the wiistite-iron buffer and augite appears first at lower oxygen fugacity ( USSELMANet al., 1975; USSELMANand LOFGREN, 1976). The reason for this change is presumably the progressive destabilization of ilmenite as Ti4+ + Ti3+. Direct measurement of intrinsic oxygen fugacities of mare basalts by the double-cell solid electrolyte method have also showed fo, in the range of0.2 to 1.Olog unit below the wiistiteiron buffer ( SATO et al., 1973; SATO, 1976). Similar measurements by SATO ( 1979) on the Apollo 17 orange glass,

some typical mare basalts in light of these coefficients can provide some important insights into the nature of the mare source region that are not always highlighted. Figure 5 presents a version of the concentration diagram proposed by WOOD et al. ( 1979) in which elements are plotted in order of apparent incompatibility. The present version has been modified to eliminate volatile elements and to emphasize the REE. Figure 5a shows that eucrites, depleted MORB, and typical ocean island basalts have relatively featureless patterns. Figure 5b shows that mare basalts and KREEP have dramatic anomalies not only in Eu, indicating plagioclase fractionation, but also in Ti and Ta-Nb (Ta and Nb are assumed to partition similarly). The strong negative anomalies of Ta-Nb and Ti in the KREEP pattern result from the fact that Nb is nearly neutral and Ti strongly compatible during the crystallization of ilmenite (MCCALLUM and CHARETTE, 1978). As might be expected, the pattern of the high-Ti Apollo 17 basalt shows positive Ta-Nb and Ti anomalies indicating accumulation of ilmenite in the source region. Surprisingly, the Apollo 15 basalt with only 2.3 wt% TiOz also shows definite relative excesses of Ta-Nb and Ti, indicating ilmenite accumulation. Thus the large range of Ti02 in mare volcanics is apparently the result of a nearly continuous variation in the amount of ilmenite in the source region. Simple remelting of static cumulates (e.g., TAYLOR and JAKBS, 1974) does not account for this variation. A dynamic process in which high-level latestage ilmenite (&at&e) cumulates sink into a less dense, less differentiated interior (KESSON and RINGWOOD, 1976) is required. Oxidation State

The presence of metallic iron and the virtual absence of ferric iron ( AGRELLet al., 1970) in Apollo 11 basalts quickly led petrologists to the realization that the oxidation state of lunar basalts was much more reduced than terrestrial basal& much like that of the basaltic achondrites (e.g., KEIL et al., 1970). Microprobe analyses of minerals also showed the need

Juvlnas

2243

x 5

i

,.----

I

Sa Th Ta

Nb La Ce Nd

Hf Zr Sm Eu Gd Ti

Yb Lu

Ba Th Ta Nb La Ce Nd Hf Zr Sm Eu Gd Ti Yb Lu

PIG. 5. Representative chondrite-normalized incompatible element concentrations in terrestrial, meteoritic, and lunar compositions. General order of elements after WOOD et al., ( 1979). (a) HAW 10, Mauna Loa tholeiite ( BVSP, 1981);Juvinas(BVSP, 1981);TypeIMORB(SuNetal., 1979).(b)RREEP(WARREN, 1989); 15016(BVSP, 1981); 702 15( BVSP, 198 I ) . Dashed lines represent extrapolated REE or inferred Ta, Nb, Hf, or Zr where it is assumed that normalized Ta = Nb and Hf = Zr.

2244

J. Longhi

however, have shown a different behavior: during the first heating intrinsic oxygen fugacity remains approximately 0.25 log unit above the wiistite-iron buffer, but during cooling fo2 drops below the wtistite-iron buffer into the range of typ ical mare basalts and remains there during subsequent heating and cooling cycles. SATO ( 1979) ascribed this behavior to the combination of oxidizing volatile species deposited on the surface of the glass spherules and bits of graphite (not directly observed) contained within the glass itself. He proposed that C-O gases, exsolved from the ascending picritic magma as it passed through the crust, reacted with graphite suspended in the magma to buffer fo, at levels above the w~sti~-iron buffer; he also suggested that the exsolved gases became the propellant for the pyroclastic eruption that produced the glass spherules. SATO (1979) further suggested that graphite in the deep lunar interior maintained a relatively high ambient fo, near the wtistite-magnetite buffer. His reasons for this latter proposal are not clear inasmuch as reactions involving graphite can only reduce silicate minerals in the absence of a C-gas phase, and a free-gas phase in the Moon’s mantle seems unlikely. However, ifcorrect, SATO'Shypothesis would require that all the manifestations of very low ,fo, in mare basalts are the result of gas-melt reduction at low pressures. More recently, DELANO ( 1990a) has shown with experiments on the Apollo 15 yellow/brown volcanic glass composition that the Cr content of liquids coexisting with spine1 varies considerably with foZ, such that Cr is low ( < 1000 ppm) at terrestrial fugacities and high (>2500 ppm) at lunar fugacities: the reasons being that Cr3+ --)rCr2+ increases with decreasing fo, and that Cr2+ is less soluble in spine1 than Cr3+. He pointed out that lunar volcanic glasses have much higher Cr contents (2700 to 6300 ppm: DELANO, 1986) than primitive terrestrial basalts (c 1.500 ppm: BVSP, 198 1) and he interpreted these data to suggest that a relatively reduced oxidation state must apply in the lunar mantle, otherwise spine1 would hold back Cr as it apparently does during melting in the Earth’s mantle. The other ex~~rnent~ data bearing on the lunar redox state are the measurements of sulfur solubility by DANKWERTHet al. ( 1979). They showed that 3400 ppm of sulfur was soluble in a melt of 74275 in contrast to 1650 ppm actually present in the rock. They argued that other high-Ti basalts were similarly undersaturated in sulfur at the time of eruption and, therefore, sulfur vola~li~tion could not be invoked as a reduction mechanism that produced iron metal in the early stages of crystallization. Thus it appears that the low oxidation state of lunar basalts is the result of melting of a reduced interior and, although graphite may have been present in the orange glass magma, it seems possible that the relatively high intrinsic fo, obtained by SATO (1979) during the first heating cycle of the orange glass may have been due to the release of a volatile component that had fractionally lost some of its more reduced species as it condensed in the lunar vacuum. It is also possible that much of the iron metal present in the mesostases of mare basalts is not the result of gas loss-induced reduction, but may simply be the result of the c~st~li~tion of oxygen-bung minerals from an oxygen-poor liquid. In simple systems Fe-rich olivine melts incongruently to iron metal plus an oxygen-rich ( Fe3+-bearing) liquid at fo, near the wiistite-iron buffer ( BOWEN and

SCHAIRER,1935 ) . The presence of Ti3+ and Cr2+ in various minerals and the experimental verification of substantial proportions of these ions in synthetic silicate liquids at lunar redox conditions ( SCHREIBER,1977) allows for the possibility that there is a range of conditions under which extracting oxygen from the liquid (by lowering foZ externally or by crystallizing oxygen-~a~ng phases in a closed system) first eliminates Fe3’ and then increases Ti3+/Ti4+ and Cr2+/Cr3+ without precipitating iron metal: iron metal ultimately precipitates, but by mechanisms that produce Ti4+ and Cr3+ instead of Fe3+ (SCHREIBERet al., 1982). HIGH PRESSURE

Phase Relations Figures 6 and 7 illustrate the high-pressure phase relations of some typical olivine-normative mare basalts with nearliquid compositions. The most important aspect of the various T-P diagrams is the pressure at which olivine and pyroxene F spine1 appear together on or near the liquidus Spine1 in the near-liquidus runs is typically Cr-rich and is apparently stabilized by graphite and molybdenum containers (LESSON and LINDSLEY, 1976), so its significance is uncertain. For reasons discussed below petrologists have traditionally interpreted the pressure of multiple saturation as the pressure at which the magma segregated from a crystalline residuum. This interpretation rests on the assumption that the composition being investigated is that of a primary liquid; however, if the magma is fractionated (removal of olivine is the most likely cause of fractionation) then the pressure of multiple saturation is likely to be a minimum pressure of segregation. In either case the pressure of multiple saturation provides an important constraint on the depth of melting. Another important aspect of the diagrams pertains to the composition of mantle minerals: if the magma is truly pri-

DEPTH 100 200 I I I EQUILIBRIUM

KM 300 I I

PRESSURE

400 I 0

P

KB

FIG.6.High-pressure liquidus relations of low-Ti basalt 12002 (WALKERet al., 197ha).

2245

Petrogenesis of mare hasalt KM

DEPTH 100

300

200

EXPERIMENTS IN THE MELTING INTERVAL -ALL

RUNS WITH

LIQUID1300

-

MOLYBDENUM

-

LIQUID

T “C

T “C

PRESSURE FIG. 7.

15

IO

5

20

KB

High-pressure liquidus relations of high-Ti basalt 702 15( LQNGHI et al., 1974).

mar-y, then the compositions of the minerals at the point of multiple saturation are the compositions of the residual minerals in the source region. For the case of the Earth, where we have some independent knowledge of upper mantle mineral compositions from the study of nodules, petrologists are able to select probable primary magmas on the basis of the composition of the liquidus olivine (Mg’ r 0.89).BVSP ( 198 1) has developed a series of compositional models for the lunar mantle to satisfy seismic, mean density, and moment of inertia constraints that have Mg’ 2 0.8. The compositions (Mg’) of liquidus olivines, measured in experimental run products or computed from K$‘iq, are >0.8 for the picritic glasses and ~0.8for the mare basal& suggesting that the picritic glasses may be primary but not the basalts. However, the large range of TiOz in the suite of mare volcanics (Fig. 1) indicates substantial chemical heterogeneity in the mare source region and hence there remains the possibility of heterogeneity in Mg’ as well. Consequently, Mg ’ is not a firm constraint and the appropriateness of the basalts as primary magmatic compositions remains an open question. Regardless of which compositions are truly primary, there are two features of the T-P diagrams with important petrogenetic consequences. The first is that plagioclase is not stable along or near the high-pressure liquidus of any olivine-normative basalt or volcanic glass. Plagioclase is stable to ~5 kbar along the liquidus of LUNA 24 ferrobasalt glass (GROVE and VANIMAN, 1978), but with only 6 w-t%MgO and Mg’ = 0.3this is a very unlikely primary magmatic composition. The absence of plagioclase as a residual phase in the source

region requires that an episode of plagioclase depletion occurred prior to the melting in order to explain the characteristic negative Eu-anomaly of mare basalts mentioned in the Introduction. A popular explanation is that mare volcanics were generated by the remelting of plagioclasedepleted cumulates of olivine, pyroxene, and ilmenite; these cumulates were themselves formed during global differentiation as complements to the plagioclase-enriched crust (TAYLORand JAK&, 1974). An important variation of this hypothesis is that KREEP, which developed very late in the global differentiation epoch and which has a very strong negative Euanomaly (WARREN and WASSON, 1979)) contributes to the incompatible element signature of the mare volcanics (HUGHES et al., 1988). Some of the original cumulates may not have crystal&d from plagioclase-saturated magmas, but KREEP has either been mechanically mixed into the cumulates or has been assimilated into ascending magmas. In either event, the experimental results require a large-scale episode of plagioclase separation prior to mare basalt genesis. The other significant feature of the T-P diagrams is the magnitude of the pressure of multiple saturation along the liquidus. Although the pressures evident in Figs. 6 and 7 (5 to 12.5 kbar) are hardly unusual, these pressures translate into depths of 100 to 250 km in the outer third of the Moon where the depth/pressure gradient is 20 km/ kbar. Table 2 summarizes the temperature, pressure, and mineralogy at multiple liquidus saturation for the picritic and more primitive mare basaltic compositions. Experiments performed in molybdenum capsules undoubtedly lost some iron to the

J. Longhi

2246

Table 2. Temuexatwe and pressureof mtdtiule saturationin experiments on mate basalts and Sample

T(oC)

phases

P&bat)

___ low_Ti

coatainet

tcfuwlce

___

01, opx Fe(h), MO cl, pig Fez) 01, pig I::15 cl, pig Fe(l) 20-25 01, opx Fe(l) MO 10 01, opx, sp 01, opx Fe(h) 01, pig 11 (F$Pt) 15016(b) 01, pig, sp 12 cl. pig Fe(h) ;$ll(s) Fe(h), MO <5 01, pig, sp 8.5 15555(b) cl, pig Fe(h) 17-18 01, opx MO 15 Gteen(sp) 01, opx(?) 17 VLT(sp) 18 Fe(h) 5 LUNA 24(s) aug, pl Fe(h) ___high_Tt __. 10017 aug, ilm 1220 10 Fe(l) avg. Apll 01, aug, arm(?) <1200 <5 Fe(l) 01, opx 1460 25 15 Rexi Fe(h) 70017(b) MO 1230 5-7 01, aug, sp G 1240 7.5 01, opx, sp 70215(b) 1210 01, attg, ilm, sp MO 5-6 16 1230 6 G 16 01, aug, sp 1250 7-9 01, aug 17 Fe(h) 9 70215(s) 1240 8 01, aug, ilm Fe(h) 17 1490 01, opx 7422O(pg) F:(h) ._ 1410 01, opx 12 ;7” 74275(b) 1320 01, aug F$) * = olivine-enriched composition; b = naturalbasalt; s = synthetic basalt; pg = naturalpictitic glass; sp = synthetic pictite; Fe(l) = low purity iron or mild steel; Fe(h) = high purity iron; G = graphite; MO = molybdenum; (Fe,Pt) = iron-platinum alloy 12002(b) 12009(b) “B” (s) 12022(b)* 12040(b)* 14072(b) 14 VLT(sp)

1380 1270 1300 >1360 >1400 1310 1490 >1500 1310 1350 21270 1300 1450 1500 1200

12.5 5-7

2

1: WALKER et al. (1976a); 2: GREEN et al. (1971b): 3: KUSHIRO (1972); 4: GREEN et al. (1971a); 5: WALKER et al. (1972); 6: CHEN et al. (1982); CHEN and LINDSLEY (1983); 8: HODGES and KUSHIRO(1972);9:KESSON(1975); 1O:WALKERetal. (1977); 11: STOLPER (1974); 12: GROVE and VANIMAN (1978); 13: O’HARA et al. (1970); 14: RINGWOOD attd ESSENE (1970); 15: DELANO (1980); 16: LONGHI et al. (1974); 17: GREENet al. (1975); 18: WALKERet al. (1975).

capsules ( KESSON and LINDSLEY, 1976 ). Likewise, comparison of experiments performed in capsules of pure iron and Fe-Pt alloys indicates that some experiments in pure iron capsules may have lost iron by reduction ( CHEN and LINDSLEY, 1983 ) . Consequently, many of the pressures of multiple saturation are too low by 1 to 3 (?) kbar because of the decrease in olivine component induced by iron loss. Taken as they are, the data show that the pressures of multiple saturation for basalts extend to more than 20 kbar (400 km) and to 25 kbar (>500 km) for the picritic glasses. Comparison of experimental and natural olivine compositions indicates that the basalts with the highest pressures of multiple saturation, 12022 and 12040, are probably olivine-enriched compositions and so their multiple saturation pressures probably have no direct petrogenetic significance (GREEN et al., 197 lb). Nonetheless, the remaining picritic glass and basalt data imply not only that the Moon melted to great depths but that some transport mechanism existed that was capable of maintaining the chemical signature of high pressure during extended transit to the surface. In addition, if the basalts and picritic glasses represent distinct magmatic styles, then their source regions were also distinct: 100 to 250 km for the basaltic magmas (with possible residual ilmenite in the case of the high-TiOz basalts) and 360 to >500 for the picritic magmas ( no residual ilmenite). If only the glasses are primary, then all of the melting and segregation took place below 360 km. The type of pyroxene at multiple saturation along the highpressure liquidus is another important concern. Augite and orthopyroxene are the liquidus pyroxenes for the high-Ti02 compositions, whereas pigeonite and orthopyroxene are the

liquidus pyroxenes for the low-Ti02 compositions. To a first approximation, the distinction between pigeonite and orthopyroxene is the result of differences in Mg’ of the liquid with higher Mg’ favoring orthopyroxene (HUEBNER and TURNOCK, 1980; LONGHI and PAN, 1988). Some compositions may crystallize orthopyroxene on the liquidus at high pressure but only pigeonite at low pressure because at low pressure pigeonite is usually preceded by a considerable interval of olivine crystallization which drives Mg’ down in the liquid. The presence of augite along the high-pressure liquidus of several of the high-TiOz compositions is significant not only because it suggests a mineralogical difference between the low- and high-Ti02 source regions but also because trace element models of the high-TiOz source regions typically require residual augite (e.g., NYQUIST et al., 1977; HUGHES et al., 1989). The presence of orthopyroxene along the high-pressure liquidus of some compositions (e.g., 15 Red, 74220), even though augite is the first pyroxene to crystallize at low pressure, is a result of the expansion of orthopyroxene liquidus field with pressure (e.g., STOLPER, 1980).Interestingly, the preferred model of HUGHES et al., 1989) shows no orthopyroxene (or pigeonite) in the 74220 source region. If their model is accurate, then the presence of orthopyroxene at the multiple saturation point on the liquidus may be an indication that the 74220 magma was not multiply saturated when it segregated. If the same holds true for the Apollo 15 Red glass, then the depth of melting for the picritic high-TiOz glasses may be considerably less than the 400 to 500 km implied by their pressures of multiple saturation. However, similar trace element modeling for low-Ti02 picritic glasses shows residual

2241

Petrogenesis of mare basalt olivine and orthopyroxene in the source (NYQUIST et al., 1977; HUGHES et al., 1988), so there is no incentive to relax the multiple saturation assumption for low-Ti02 compositions. Modeling It should be clear that the high-pressure phase relations have important implications for the composition and thermal structure of the lunar interior during mare volcanism. However, these implications depend on the correctness of the multiple saturation hypothesis. The major alternative to the multiple saturation hypothesis has been to invoke relatively large degrees of partial melting at some low pressure that leaves behind a residuum consisting of only olivine (e.g., BINDER, 1982, 1985). For a source with 2.0 wt% A1203 and 80% olivine, such a process would require approximately 30% melting to produce a green glass composition (Table 3); sources with more A1203 or less olivine would require larger degrees of melting. Since geochemical models have consistently maintained the need for significant amounts (>20%) of orthopyroxene in the mare source regions (e.g., NYQUIST et al., 1977; HUGHES et al., 1988), the prospect of large degrees of melting of source regions with small amounts of pyroxene has remained unpopular even though petrologists have not explained how magmas could retain their primary compositions through transits of several hundred kilometers to the surface. Very rapid modes of magma transport, such as gas-charged (kimberlitic) eruptions from depth or transport of melt through a crack, seem unlikely on the Moon. For example, given the volatile-depleted nature of the Moon (TAYLOR, 1982), saturating a magma at depth with a volatile species ( COZ?) is difficult to justify. Also, SPERA( 1992) has calculated that melt must travel approximately 10 m/set through a crack 40 m in diameter in order to avoid fractionating during a 400 km transit to the surface, but as yet there is no explanation of how a crack of such dimensions would form or stay open. Recently, some workers have concluded that the major element compositions of terrestrial basalts develop by the pooling of melt generated over a range of pressures (KLEIN and LANGMUIR,1987; MCKENZIE and BICIUE, 1988). These models are based in part upon investigations of porous flow of silicate melt through a deforming and compacting crystalline matrix (MCKENZIE, 1984; RICHTERand MCKENZIE,

1984). Polybaric melting models may have some applicability to mare basalt petrogenesis. Mare basalts were produced by partial melting of low-Al mafic to ultramafic sources. During the time of mare petrogenesis the Moon had a thin but growing elastic lithosphere that eventually supported the loading of mare basalts in the great basins to form mascons: stress modeling suggests elastic lithosphere thicknesses of -50 km at 3.6-3.8 b.y., and increasing to - 100 km afterwards (SOL OMON and HEAD, 1979 ). A thermal lithosphere, probably of comparable thickness, extended below the elastic lithosphere. Beneath the base of the lithosphere was the mare basalt source region, partially molten in places and probably convecting. In such a physical regime mare basalt magmas generated over a range of pressures might pool at the base of a lithosphere thinned by thermal erosion and thereafter rise to the surface along fractures caused by impacts. A major test of the applicability of the polybaric melting model is its ability to generate composite melts ponded at - 100 km ( 5 kbar) that appear to be multi-saturated at higher pressure and hence to be derived from greater depths. Figure 8 ( LONGHI, 199 1b) illustrates this situation: the field of the picritic green volcanic glasses (DELANO, 1986) sits astride the calculated 20 kbar olivine + orthopyroxene liquidus boundary, well below the position of the boundary at 5 kbar. The low-pressure liquidus boundaries, which are appropriate for green glass compositions, are drawn from the empirical expressions of LONGHI ( 199 1a) ; the high-pressure boundaries are drawn from modifications of the low-pressure expressions to fit the 8 to 35 kbar data of FALLOONand GREEN ( 1988). The curved paths in Fig. 8 represent the composition of pooled melts extracted from an ascending partially molten source region. In each case melting begins at some initial pressure (indicated at the lower end of the curve) and proceeds in a series of -2 kbar steps down to 5 or 6 kbar with the extracted melt pooling in a reservoir. The source also becomes progressively depleted in response to melt extraction, but this change is not shown. To simulate the porous flow process, some percentage of the melt generated in each step was retained in the source. Retaining some melt in the source is necessary to avoid rapidly depleting the source in basaltic components and to keep the temperature of the solidus from increasing. More details of the effects of varying the source composition, the size of the pressure step, the extent of melting in each step, and the fraction of melt retained in the source are given by LONGHI (1992).

Table 3. Comparison of gnxn glasses and mode1 compositions in weight J oxides

KzO Na20 Mg’

1

2

3

ii2 2.88

44.2 0.26 1.87 0.7 15.6 35.1 0.17 2.05 0.003 0.03 0.80

43.9 1.44 7.73 0.59 20.1 15.1 .25 10.7 0.02 0.18 0.57

0.7 I 11.7 37.3 0.17 2.62 0.002 0.05 0.85

4 45.1 0.94 6.06 0.61 21.8 17.8 0.26 7.22 0.02 0.11 0.59

5

6

zifd: 24; 6.71 6.89 20.2 19.5 0.23 7.40 0.10 0.63

0.56 23.1 16.6 0.28 7.94 0.05 0.19 0.56

1: Lunar mantle (TAYLOR, 1982) ; 2 : Model soume.region (LONGHI, 1991b); 3: pooled melt composition - 40 kbar initial melting, 1.2% melt/kbar, 0.4 melt extraction efficiency, final melting at 6 kbaq 4: pooled melt composition - 40 kbar Initialmelting, 1.2% meIVkbar, 0.7 melt extraction efficiency, tinal melting at 6 kbq 5: Apollo 17 green glass (DELANO, 1986): 6 : Apollo 14 green glass B (DELANO, 1986).

2248

J. Longhi

Qtz

A

IWO3 oxygen units rig'= NAB = NOR =

6 .07 ,005

without chemical mo~fication hundreds of kilometers from the melting region to the base. of the lithosphere where they pond or that a separate melt pool ascends in close proximity to the melting region. This latter scenario requires that the compaction zone is larger than the melt zone, otherwise the melt will escape quickly as a porosity wave (SCOTT and STEVENSON,1986). Qu~itatively, the need for nearly contiguous regions of melt generation and accumulation suggests diapiric upwelling, which may be better suited to producing the large range in TiOz of the mare volcanics from a heterogeneous source than pervasive two-phase porous flow. FOURS

b

40

kb

FIG. 8. Comparison of green glass compositions (DELANO, I986 ) and calculated polybaric pooled melt compositions. Low-pressure liquidus boundaries are appropriate for green glass compositional parameters shown in the upper right: NAB and NOR are the albite and orthoclase factions of the normative feldspar; QWo is the wollastonite coordinate of the model system 01-PI-Wo-Qtz (LONGHI, 1991a). High-pressure Iiquidus boundaries are drawn from modifications to the empirical expressions of LONGHI( 1991a) based on the 8 to 35 kbar data of FALLOONand GREEN( 1988). Curvedlines describe the trace of pooled melt compositions; open squares show liquid composition at the onset of melting; numbers next to curves are the percentages of melt present in each cycle that are extracted from the source and added to the pool (melt extraction efficiency). The final pooled melt compositions (Table 3) from the 40 kbar calculation have the requisite high-pressure signature and plot well into the olivine liquidus field at 5 kbar even though there was never any more than 2.4% melt in the melting region. The pooled melt compositions are functions not only of initial pressure and melt extraction elhciency but also of initial source composition. Source regions with low Mg’ and A&O3 (plagioclase component) facilitate the match of pooled melt with green glass compositions. Table 3 lists two of the final pooled melt compositions, some green glasses from the DELANO ( 1986) compilation for comparison, the TAYLOR f 1982) estimated composition of the lunar mantle, and the model source composition employed in these calculations. The model source composition is consistent with an origin by accumulation of predominantly olivine and lowCa pyroxene, but is significantly lower in Mg’ and Al203 than the estimate of TAYLOR. The possibility of melting highly d~ffe~ntiated material at the depths that 40 kbar imply (i.e., > 1000 km) has important

implications for the scale of the Moon’s early differentiation as well as its subsequent thermal regime. But before these implications are considered there is much work to do. One project is to determine the positions of the olivine + pyroxene liquidus boundaries for lunar compositions in the range of 25 to 50 kar. Second, there is a need to integrate these chemical models with dynamical physical models to ensure plausibility. For example, the manner in which the melt is pooled implies either that small degrees of melt are able to travel

STUDIES

The general phase relations of mare volcanics are fairly well known, having been measured directly or calculated. Primitive mare basalts have olivine on the liquidus and plagioclase as a late-appearing phase. What is not known is the proper range of compositions for which phase relations need to be determined. Are primitive mare basalts primary or are they differentiates of liquids similar in olivine content to the picritic glasses? If mare basalts are primary, then their highpressure liquidus relations are petrogenetically significant and there may have been two distinctive modes of volcanism, one that yielded lavas and one that yielded pyroclastic deposits. Furthermore, there may be no basaltic counterparts to the very high-Ti red and black picritic glasses. If all mare basalts are differentiates, then only the high-pressure liquidus relations of the picritic glasses have any direct significance and very high-Ti basalts should be present on the Moon. The most obvious way to resolve this uncertainty is through new sampling missions. Recent telescope observations by COOMBS and HAWKE ( i 992 ) suggest that transitions from pyroclastic deposits to lava flows occur on the western limb ofthe Moon. If so, then this would be a likely place to sample. Even in the absence of new samples, there remains much experimental work to do in the areas of volatiles, partition coeflicients, and general phase equilibria. New work is needed to identify the volatile species that propelled the pyroclastic eruptions and to determine what effect, if any, these volatiles had on petrogenesis. The effect of volatiles on oxidation state is particularly vexing because if, as SATO( 1979) has suggested, graphite was present in the mare source regions and remained only partially dissolved in lunar magmas as they rose through the crust, then exsolution of CO, would have reduced the mare magmas and removed the obvious signatures of a relatively oxidized interior. However, if graphite was not suspended in the mare magmas, then any loss of CO* would probably have oxidized the magmas ( MATHEZ, 1984). Preliminary experiments by DELANO ( 199Oa) on Cr partitioning tend to favor a reduced mare source region. As petrogenetic modeling becomes more

complicated,

there is a continuing need for accurate measurements of partition coefficients for a wide range of elements, especially those elements with dual behavior (incompatible for some phases and compatible for others). Comparison of the concentrations of REE and hip-meld-st~ngth elements (Ta-NbHf-Zr-Ti) shows that ilmenite may have been added even to the source of the low-I? basal&, a phenomena consistent with the hypothesis of sinking ilmenite-clinopyroxene blobs pro-

Petrogenesisof mare basalt posed by KESSON and RINGWOOD (1976). Further measurements of partition coefficients are needed, however, to resolve uncertainties in the role of trapped instanteous (SNYDER et al., 1991) or admixed late-stage (SHERVAIS and ERHLICH, 1991) liquids in the formation of the mare basalt source region from a lunar magma ocean. Finally, there is a need to integrate the physics of the lunar interior to the chemistry of mare basalt genesis. Basalt petrogenesis is undoubtedly a polybaric process that involves the buoyancy and wetting properties of silicate melts, adiabatic gradients, and heats of fusion as well as the array of liquidus relations and partition coefficients. Some progress has been made in this area (DELANO, 199Oc; HESS et al., 1991; LONGHI, 1991b), but much remains to be done. Acknowledgments-1 thank J. Delano, D. H. Green, and D. Walker for helpful reviews. This research was supported by NASA grant NAG 9-329. Lamont-Doherty Geological Observatory Contribution #4926. Editorial handling: P. C. Hess

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