Geology and petrology of enormous volumes of …Geology and petrology of enormous volumes of impact...

Icarus 223 (2013) 749–765

Contents lists available at SciVerse ScienceDirect

Icarus

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

Geology and petrology of enormous volumes of impact melt on the Moon:A case study of the Orientale basin impact melt sea

William M. Vaughan a,⇑, James W. Head a, Lionel Wilson b, Paul C. Hess a

a Department of Geological Sciences, Brown University, Providence, RI 02912, USAb Lancaster Environment Centre, Lancaster University, Lancaster LA1 4YQ, UK

a r t i c l e i n f o a b s t r a c t

Article history:Received 6 August 2012Revised 11 January 2013Accepted 16 January 2013Available online 30 January 2013

Keywords:MoonCrateringImpact processes

0019-1035/$ - see front matter � 2013 Elsevier Inc. Ahttp://dx.doi.org/10.1016/j.icarus.2013.01.017

⇑ Corresponding author. Fax: +1 401 863 3978.E-mail address: [email protected] (W

Lunar basin-forming impacts produce enormous volumes (>105 km3) of impact melt. All known basin-forming impacts combined may produce �108 km3 of impact melt, �1/20th the volume of the lunar crust.Despite their volumetric importance, the geology and petrology of massive deposits of impact melt on theMoon have been little studied, in part because most basin impact melt deposits are old and have beenobscured or buried by subsequent impact cratering and mare infill. We investigate the geology and modelthe petrology of fresh massive impact melt deposits in the relatively young 930 km diameter Orientalebasin. Models of impact melt production combined with geologic analyses based on new LOLA topo-graphic data suggest that most of the impact melt (�2/3) produced by the Orientale-forming impactoccurs in a�15 km thick impact melt sheet (better described as an impact melt sea) �350 km in diameterwith a volume of �106 km3. We anticipate that the Orientale melt sea has undergone large-scale igneousdifferentiation, since terrestrial impact melt sheets (such as Manicouagan, Sudbury, and Morokweng) lessthan a tenth of the thickness and a hundredth of the volume of the Orientale melt sea have differentiated.We develop a model for the cumulate stratigraphy of the solidified Orientale impact melt sea. A modeledcumulate stratigraphy (occurring below a quench crust and anorthositic fallback breccia) with an �8 kmthick layer of norite overlying a �4 km layer of pyroxenite and a basal �2 km thick layer of dunite pro-duced by equilibrium crystallization of a homogenized melt sea, consistent with vigorous convection inthat melt sea, is supported by remotely-sensed norite excavated by the central peak of Maunder craterfrom �4 km depth. Generally, we predict that very large basin-forming impacts, including the SouthPole-Aitken (SPA) basin-forming impact, produce melt seas with a cumulate stratigraphy similar to thatof the Orientale melt sea. Impact melt differentiation may explain apparently anomalous lithologies exca-vated in the SPA basin interior. We note that impact melt differentiates are slow-cooled (the Orientalemelt sea took on the order of 105 years to solidify) and, if meteoritic siderophiles fractionate into metalor sulfide layers, may not be siderophile-enriched; therefore, impact melt differentiates may pass forpristine highland plutonic rocks in the lunar sample suite. These predictions can be tested with currentand future mission data.

� 2013 Elsevier Inc. All rights reserved.

1. Introduction

Simple and complex lunar craters <150 km in diameter containrelatively small volumes (�104 km3) of impact melt (Cintala andGrieve, 1998). The geology of the small impact melt deposits in thesecraters is relatively well understood: impact melt floods crater floorsand drapes crater central peaks, walls, and rims (Bray et al., 2010;Howard et al., 1974; Howard and Wilshire, 1973, 1975; Plesciaand Cintala, 2012; Spray and Thompson, 2008); melt distributionis controlled by impact angle, post-impact ejection related to slump-ing, and pre-existing topography (Hawke and Head, 1977). Thinsheets of impact melt occurring in these craters freeze quickly (War-

ll rights reserved.

.M. Vaughan).

ren et al., 1996) to produce glasses or cryptocrystalline lithologieschemically similar to target rocks.

Field and experimental evidence shows that the volume of shockmelt produced by an impact scales approximately as the fourthpower of impact crater transient cavity diameter (Abramov et al.,2012; Cintala and Grieve, 1998; Grieve and Cintala, 1992, 1997).So although craters <150 km in diameter contain relatively smallvolumes (�104 km3) of impact melt, basin-forming impacts creat-ing basins >300 km in diameter are predicted by these analytic scal-ing relationships (Abramov et al., 2012; Cintala and Grieve, 1998;Grieve and Cintala, 1992) as well as hydrocode models (Collinset al., 2002; Ivanov, 2005) to produce enormous volumes(>105 km3) of impact melt. The volumes of impact melt producedby basin-forming impacts are well in excess of the volumes of manyterrestrial bodies of water and igneous intrusions (Fig. 1). In fact, all

http://dx.doi.org/10.1016/j.icarus.2013.01.017

mailto:[email protected]


http://www.sciencedirect.com/science/journal/00191035

http://www.elsevier.com/locate/icarus

Fig. 1. The volume of impact melt in a �30 km lunar crater, estimated using thescaling laws of Cintala and Grieve (1998), and the volume of the Orientale impactmelt sea, calculated in Section 2, compared to selected terrestrial igneous intrusions(Bonini, 1982; Nielsen, 2004) and bodies of water.

750 W.M. Vaughan et al. / Icarus 223 (2013) 749–765

basin-forming impacts combined are modeled to produce on the or-der of 108 km3 of impact melt (see Kring et al. (2012) and our Sec-tion 4.4), so that as much as �1/20th of the lunar crust may beimpact melt.

Despite their volumetric importance, the geology and petrologyof massive impact melt deposits produced by basin-forming impactsare poorly understood. This is in part because basin interiors wheremassive impact melt deposits occur (Head, 1974; Melosh, 1989;Spudis, 1993) have been obscured and modified by impact crateringand mare basalt infill (Head and Wilson, 1992). For example, no im-pact melt deposits are exposed in the mare basalt-flooded Imbrium,Serenitatis, and Crisium basin interiors (Head et al., 1993). More-over, although the thick (>1 km) sheets of impact melt predictedto occur in lunar basins (Cintala and Grieve, 1998) cool slowly andmay undergo igneous differentiation (Cintala and Grieve, 1998;Grieve et al., 1991), the cumulate stratigraphy of these thick impactmelt sheets has never been modeled, although several authors havecarefully considered some particular petrologic consequences of im-pact melt differentiation (Morrison, 1998; Warren et al., 1996).

Where on the Moon should we look to improve our understand-ing of massive impact melt deposits? The relatively young (�3.7–3.8 Ga) (Baldwin, 1974; Whitten et al., 2011) Orientale basin (Head,1974; Howard et al., 1974; McCauley, 1977), which has relativelyfew impact craters and little mare basalt infill (Greeley et al.,1993; Whitten et al., 2010, 2011), is an ideal location to investigatethe distribution, thickness, and volume of massive impact meltdeposits. Moreover, the relatively well-constrained structure ofthe farside crust (Wieczorek et al., 2006) and upper mantle (Khanet al., 2006) probably melted by the Orientale impact provides nec-essary compositional data (Shearer et al., 2006) to model the cumu-late stratigraphy of an Orientale impact melt sheet and to considerthe petrologic implications of impact melt differentiation.

In this paper, we (1) investigate the distribution, thickness, and vol-ume of impact melt deposits produced by the Orientale-forming im-pact using models of impact melt production combined with recenttopographic data and (2) model the cumulate stratigraphy of a solidi-fied, differentiated Orientale melt sheet, considering the general impli-cations for the lunar highlands crust and the lunar sample suite.

2. Distribution, thickness, and volume of massive impact meltdeposits in Orientale

2.1. General geology of the Orientale basin

The deposits associated with the �930 km diameter Orientaleimpact basin (Fig. 2) have been mapped and subdivided into three

units (Fassett et al., 2011; Head, 1974; Howard et al., 1974; McCau-ley, 1977): the radially-textured Hevelius Formation, lying outsidethe �930 km diameter Cordillera Ring; the Montes Rook Forma-tion, comprising knobby terrain between the Cordillera Ring andthe �620 km diameter Outer Rook Ring; and the Maunder Forma-tion, which occurs inside the Outer Rook Ring, around the �480 kmdiameter Inner Rook Ring, and on the floor of the �350 km diam-eter central depression. These three major rings (the Inner RookRing, the Outer Rook Ring, and the Cordillera Ring) have each beeninterpreted as the Orientale basin crater rim crest (see Spudis(1993) for a review). We interpret the Orientale basin as a modifiedpeak-ring basin (Head, 2010), taking the Inner Rook Ring to be thebasin peak ring (Baker et al., 2011a, 2011b, 2012) and the OuterRook Ring to be the basin rim crest (Head, 1974, 2010; Headet al., 1993). In the following, we take the current diameter ofthe Outer Rook Ring (620 km) as the best estimate of the Orientaletransient cavity diameter, since there is no clear evidence for ter-racing within this rim that would have substantially increased itsdiameter (Head, 2010; Head, 2012).

2.2. Geology of impact melt-related facies in Orientale

The facies of the Maunder Formation (Fig. 3), named for theprominent 55 km diameter complex crater Maunder superposedon the north edge of the central depression of the Orientale basin,have historically been interpreted as impact melt-related (Head,1974; Head et al., 1993; Howard et al., 1974; Spudis, 1993). Thecentral smooth facies (Fig. 3, left), occurring within the innerdepression of the Orientale basin exposed through thin mare fill(Head, 2010; Whitten et al., 2010, 2011), has been interpreted aspure impact melt (Head, 1974; Howard et al., 1974; McCauley,1977). Wrinkle ridges (Fig. 3), fractures, and polygonal cracks areapparent in the smooth facies and under the overlying mare. Atan average distance of 175 km from the Orientale basin center(Fig. 2), the topography abruptly rises �1.75 km along a series ofmarginal normal faults to a corrugated and fissured rough facies(Fig. 3, right). This outer corrugated, fissured facies has been inter-preted as clast-rich impact melt formed from mixing of pure im-pact melt with brecciated debris from the Orientale impact(Head, 1974; Head et al., 1993; Howard et al., 1974; McCauley,1977).

2.3. How much impact melt was produced by the Orientale-formingimpact?

Impact melt volume scaling relationships adjusted for lunargravity derived by Cintala and Grieve (1998) predict that a chon-dritic projectile with an asteroidal velocity of 10 km/s and a trajec-tory normal to the lunar surface giving rise to an Orientale-sizedtransient cavity �620 km in diameter produces �6 � 106 km3 ofimpact melt. This volume is probably an overestimate: manyauthors have suggested that the Orientale-forming impact was ob-lique (Schultz and Papamarcos, 2010; Scott et al., 1977; Wilhelmset al., 1987) and oblique impacts are known to produce less meltthan vertical impacts of the same size (Pierazzo and Melosh,2000). Transverse widening of the Cordillera Ring of Orientale(Schultz et al., 2012) may indicate that the Orientale impactorhad an impact angle of 15–30�. The volume of melt produced byan impact with an impact angle of 15–30� is 10–50% of the volumeof melt produced by a vertical impact (Pierazzo and Melosh, 2000);taking 22.5�, the center of this impact angle range, to be the mostlikely impact angle of the Orientale-forming projectile in theabsence of additional constraints, we estimate that theOrientale-forming impact produced �1.5 � 106 km3 (25 vol.% of�6 � 106 km3) of impact melt. We emphasize that the compositionand velocity of the Orientale-forming projectile are not well

Fig. 2. Lunar Orbiter Laser Altimeter (LOLA)-derived topographic map of the Orientale basin indicating major basin rings and geologic units.

Fig. 3. The smooth (left) and rough (right) facies of Orientale’s Maunder Formation from an LROC WAC monochrome mosaic of the Orientale basin.

W.M. Vaughan et al. / Icarus 223 (2013) 749–765 751

constrained: variations in these parameters can change the impactmelt volume by a factor of two in the Cintala and Grieve (1998)scaling model.

An independent estimate of the volume of impact melt pro-duced by the Orientale-forming impact can be made using the im-pact melt volume scaling relationship for oblique impacts derivedby Abramov et al. (2012). This relationship, which explicitly in-cludes impact angle as a variable, predicts that a chondritic projec-tile with an impact angle of 22.5� giving rise to an Orientale-sizedtransient cavity 620 km in diameter produces about�2.8 � 106 km3 of melt, about twice the volume estimated aboveby coupling the models of Cintala and Grieve (1998) and Pierazzoand Melosh (2000). Discrepancies between impact melt volumespredicted by the scaling relationships of Cintala and Grieve(1998) and Abramov et al. (2012) have previously been noted byAbramov et al. (2012). In order to assess the Cintala and Grieve

(1998) and Abramov et al. (2012) estimates of the volume of im-pact melt produced by the Orientale-forming impact, we latercompare melt deposit thicknesses determined from these calcu-lated melt volumes (Section 2.4) to an independent determinationof melt deposit thickness (Sections 2.5–2.7). For simplicity, in thefollowing section we consider only the melt volume of�1.5 � 106 km3 calculated from the scaling model of Cintala andGrieve (1998).

2.4. Accounting for the �1.5 � 106 km3 of impact melt produced by theOrientale-forming impact

Where does the estimated �1.5 � 106 km3 of impact melt pro-duced by the Orientale-forming impact end up? A priori, becausethe surface area of the Moon is only �3.8 � 107 km2, and sincethere is no evidence that the Moon is covered in a �40 m (i.e.,


1.5 � 106 km3/3.8 � 107 km2) thick layer of Orientale impact melt,most of the impact melt produced by the Orientale-forming impactmust occur in or near the Orientale basin.

This is in accordance with what is known about the process ofimpact melt generation (Fig. 4). Impact melt forms in a melt cavityinterior to (although, in multi-ring basins, likely extending below)what will become the crater transient cavity (Fig. 4a) (Cintala andGrieve, 1998). Some of the generated impact melt is excavated informing the transient cavity; additional melt is ejected from thecrater interior during modification (Fig. 4b) (Hawke and Head,1977; Osinski et al., 2011). Studies of large terrestrial craters suchas Sudbury demonstrate that a relatively small proportion of melt(<10% by volume) is ejected from the crater interior during modi-fication (Grieve and Cintala, 1992). Excavation flow modeling sug-gests that up to 70–80 vol.% of impact melt remains inside thetransient cavity in large lunar craters (Cintala and Grieve, 1998).Based on these models, we estimate that �25 vol.% of impact meltwas ejected from the Orientale basin (perhaps incorporated inHevelius Formation ejecta). This implies that �1.1 � 106 km3

(75 vol.% of 1.5 � 106 km3) of impact melt produced in the Orien-tale-forming impact remains inside the Orientale transient cavityrim, here taken to be the Outer Rook Ring (Head, 1974, 2010,2012).

Fig. 4. The material melted by the Orientale-forming impact (a), the Orientale melt cavitythe final crater geometry of Orientale (c) (Andrews-Hanna and Stewart, 2011; Hikida andet al., 2006; Wieczorek et al., 2006; Zuber et al., 1994). Melt fills a truncated sphere withtotal volume of 1.5 � 106 km3 (Cintala and Grieve, 1998), following our discussion in Se(Wieczorek and Phillips, 1999) and the transient cavity diameter has been decreased by amodel taking Z = 3 (Melosh, 1989). Note that (1) the transient cavity is gravitationally unfinal basin diameter (Fig. 2) is much larger than the transient cavity diameter in (b). (For ithe web version of this article.)

Occurring inside the Outer Rook Ring are the impact melt-re-lated rough and smooth facies (Fig. 3) of the Maunder Formationdescribed in Section 2.2 above. About 1.1 � 106 km3 of impact meltmust be apportioned between these two facies. How much impactmelt is contained in each facies?

The rough facies mantles topography similar to the peaks of theInner Rook Ring (Head, 1974), which stand up to 2–3 km abovesurrounding terrain (Figs. 2 and 3). This constrains the volume ofimpact melt contained in this facies: the relief of the mantledtopography is <1 km (Fig. 5), but the original relief was likely nogreater than the relief of the nearby peaks of the Inner Rook Ring,2–3 km; therefore, the mantling deposit is �2 km thick at most.The area of the rough facies—approximately the area of the annu-lus between the edge of the central depression (175 km from thebasin center) and the Outer Rook Ring (310 km from the basin cen-ter)—is �2 � 105 km2. If the clast-rich rough facies is 2 km thick(which for 50 vol.% clasts is the equivalent of a 1 km thick layerof pure impact melt) it contains �1 � 105 km3 (50 vol.% puremelt � 2 km � �2 � 105 km2) of impact melt. The remaining�1.0 � 106 km3 (�1.1 � 106 km3 � �1 � 105 km3) of impact meltmust be contained in the central smooth facies, thought to be apure impact melt sheet (Section 2.2). This smooth facies is175 km in radius (Fig. 5) or �105 km2 in area, so the impact melt

and transient cavity geometry immediately following hypervelocity impact (b), andWieczorek, 2007) superposed on the lunar highlands crust and upper mantle (Khanradius/truncation height taken as 0.9/1.1, following Barr and Citron, (2011), and a

ctions 2.3–2.7. The basin transient cavity depth-diameter ratio is taken to be 1/10factor of ½ for clarity. The excavation flow curve (in blue) is given by the Maxwell Z-stable, so the geometry shown in (b) is never fully realized at any time and (2) the

nterpretation of the references to color in this figure legend, the reader is referred to

Fig. 5. Average radial topographic profile of Maunder Formation facies. Radial profiles crossing superposed Maunder and Kopff craters were excluded. See Baker et al. (2011a,2011b, 2012) for methodology.


sheet must be on the order of 10 km thick to account for theremaining �1.0 � 106 km3 of impact melt.

A �10 km thick impact melt sheet is not necessarily unrealistic:the Sudbury Igneous Complex, a terrestrial impact melt sheetoccurring in a �250 km diameter crater (rim diameter 140 km),is �2.5–3 km thick (Thierrault et al., 2002). More compellingly,there is independent geologic evidence for a �10 km thick impactmelt sheet occurring in the Orientale basin’s central depression.

2.5. Geologic evidence for a �10 km thick melt sheet

As noted in Section 2.2, the surface of the smooth facies is onaverage offset �1.75 km (Fig. 5) below the rough facies along mar-ginal normal faults. This offset has been previously explained asvertical subsidence caused by thermal stresses resulting from im-pact-generated heat and uplift of crustal isotherms (Bratt et al.,1985). This model predicts a �2 km drop over 100 km of radial dis-tance (Fig. 14 from Bratt et al. (1985)). However, new, high-resolu-tion LOLA altimetry (Smith et al., 2010) shows that the verticalsubsidence of the central depression is abrupt: the topographydrops �2 km over a radial distance of �20 km along the west edgeof the depression (Fig. 6), five times more steeply than predicted bythe model of Bratt et al. (1985).

Wilson and Head (2011) suggest that the vertical subsidence ofthe smooth facies, an impact melt sheet, is largely the result of ver-tical shrinkage upon solidification and cooling of impact melt.There is evidence for lateral shrinkage due to the same process:fractures on the smooth facies resemble deformations on the sur-

Fig. 6. Four average radial topographic profiles of Maunder Formation facies over differewere excluded. See Baker et al. (2011a, 2011b, 2012) for methodology.

faces of terrestrial lava lakes (Barberi and Varet, 1970) known toform due to lateral contraction upon solidification and cooling.We now calculate the amount of vertical subsidence that resultsfrom solidification and cooling of impact melt sheets on the Moonin order to constrain the thickness and volume of the Orientalemelt sheet.

2.6. Estimating melt sheet thickness from vertical subsidence

The vertical shrinkage of a melt sheet upon solidification andcooling is related to its total thickness. We now determine the nat-ure of this relationship in order to calculate the thickness of theOrientale melt sheet from its observed vertical subsidence. Thederivation follows Wilson and Head (2011).

A melt sheet emplaced at a temperature Te that solidifies andcools to the ambient temperature Tc will decrease in volume bothduring solidification and while cooling to Tc. What is the magni-tude of the total proportional decrease in volume? Volume is inver-sely proportional to density, so dV/V = �dq/q. A simple rule ofthumb is that the density of silicate melt increases (and volumeof melt decreases) by �10% upon solidification. This rule may bemisleading for melts of the lunar crust, since the density of anor-thite melt increases only �5.5% upon solidification (based on cal-culations of silicate liquid density following Nelson andCarmichael (1979)); however, as will be shown in Section 3.3,the Orientale melt sheet contains abundant mafic material, so wecontinue to use the �10% estimate. The proportional change in vol-ume of a material at constant pressure, dV/V, due to subsequent

nt azimuthal ranges. Radial profiles crossing superposed Maunder and Kopff craters


cooling is equal to adT, where a is the volumetric coefficient ofthermal expansion, �3 � 10�5 K�1, for silicate rock (Turcotte andSchubert, 2002). The dT is 1328 K, the difference between Te (takento be 1568 K, the average of the lava liquidus, 1713 K, and solidus,1423 K, temperatures suggested by Williams et al. (2000)) and Tc

(taken to be the average of the lunar day and night temperatures,�240 K); therefore, the volume of the melt sheet must decreaseby �4% upon cooling. A �10% volume decrease compounded witha �4% volume decrease results in approximately a �14% volumedecrease, since these percentages are both small.

What portion of this 14% volume decrease results in verticalsubsidence as opposed to lateral shrinkage? If the melt sheet con-tracts equally in all directions, the percentage thickness decreasewill be one-third of the percentage volume decrease, about�4.7%. However, this is physically and geologically unrealistic:the melt sheet does not cool at an equal rate at each of its bound-aries, but cools primarily through radiation at its top, and willtherefore not contract equally in all directions; moreover, the con-traction features (cracks and depressed polygons) present in the Sand SW portion of the smooth facies occupy only 1–2% of the basinfloor and are not consistent with a 9.3% decrease in surface area. Amore detailed treatment of lava lake solidification and cooling(Worster et al., 1990, 1993) gives a better estimate of vertical sub-sidence. The upper surface of a lava lake cools by radiation andloses heat much faster than the other boundaries of the lava lake.Initially, lava exposed to the surface solidifies to form plates; theseplates must undergo mainly vertical contraction, since lateral con-traction would cause lava to well up and the solidified plates tofounder into the liquid beneath, settling toward the base of thepond and probably being partially remelted. Eventually a stablecrust will form; evidence from drilling of the Kilauea Iki lava lake(Hardee, 1980) confirms that the stable crust forms well beforethe mean temperature of the pond has approached the solidus.The partitioning between vertical shrinkage and lateral shrinkagefrom the 4% volume decrease due to the cooling of this solid crustcan be determined from the area of depressions and fractures onthe pond surface; for simplicity, we assume that between one-third and all of the volume decrease upon cooling results in verticalsubsidence. The melt beneath the stable crust freezes progres-sively; therefore, almost all of its 10% volume decrease upon solid-ification will result in vertical contraction.

Fig. 7. Vertical subsidence in meters as a function of impact melt sheet thicknessfor the 11.3% and 14% limits derived in this paper and in Wilson and Head (2011).

In short, we anticipate that nearly all the �10% volume decreaseupon solidification results in vertical subsidence and that betweenone-third and all of the �4% volume decrease upon cooling resultsin vertical subsidence. Therefore, we predict that the verticalshrinkage of a lunar melt sheet is between �11.3% and �14% ofits depth. Fig. 7 shows the subsidence predicted as a function ofmelt sheet depth for these 11.3% and 14% limits.

2.7. Thickness and volume of the Orientale melt sheet

The Orientale melt sheet has an observed average vertical reliefof �1.75 km (Fig. 5). As stated above, the observed vertical subsi-dence is a factor of five greater than predicted by Bratt et al.(1985), so subsidence due to solidification and cooling of meltmust dwarf subsidence due to other thermal stresses. Therefore,we treat all the observed vertical subsidence as the result of meltsolidification and cooling. The 11.3% and 14% limits due to coolingand contraction derived above imply that the melt sheet was ini-tially between �12.5 and �15.5 km thick (1.75 km/14% or 11.3%).Lateral contraction features are observed in the smooth facies, so,anticipating that much of the �4% volume decrease upon solidifi-cation resulted in lateral shrinkage, we choose the �11.3% boundas more realistic and use the corresponding thickness of 15.5 kmas the thickness of the Orientale melt sheet in all followingcalculations.

The diameter (350 km) and thickness (15.5 km) of the meltsheet can be used to calculate the volume of the melt sheet. (Notethat the thickness of the solidified, cooled melt sheet is not15.5 km, but �14 km – i.e., 15.5 km � (100–11.3)%; its volume is84% of the original melt sheet volume.) Unfortunately, the geome-try of the melt sheet is poorly constrained due to the difficulty inpredicting the events following the formation of the transient cav-ity (Fig. 4) leading up to cavity collapse (Head, 2010). The approx-imation that the melt sheet is shaped like a spherical cap gives amelt sheet volume of �8 � 105 km3. Inward movement and up-ward translation during transient cavity collapse could potentiallyproduce a cylindrical melt sheet (Head et al., 1993) with a volumeof �1.5 � 106 km3. The true melt sheet volume is probably some-where between the extremes of �8 � 105 km3 and�1.5 � 106 km3. These bounds are in striking agreement with themelt sheet volume �106 km3 estimated from the melt productionmodel of Cintala and Grieve (1998) but are inconsistent with amelt sheet volume of �1.9 � 106 km3 estimated from the melt pro-duction model of Abramov et al. (2012) and our calculations in Sec-tion 2.4. Therefore, we take �1.5 � 106 km3 as the volume ofimpact melt produced by the Orientale-forming impact, since thisvolume determined from the melt scaling model of Cintala andGrieve (1998) is most consistent with our estimates of melt sheetvolume. We hereafter describe this 15.5 km thick melt sheet with avolume of �106 km3 as a melt sea to emphasize its enormousdimensions (Fig. 1).

In summary, most impact melt (�2/3) produced by the Orien-tale-forming impact occurs in a �15.5 km thick (now �14 kmthick) impact melt sea �350 km in diameter with a volume of�106 km3. We now model the petrology of this impact melt sea.

3. Cumulate stratigraphy of the Orientale melt sea

3.1. Did the Orientale melt sea differentiate?

Mounting petrologic evidence suggests that impact melt sheetsassociated with the large terrestrial craters Manicouagan(O’Connell-Cooper and Spray, 2011), Sudbury (Grieve et al., 1991;Thierrault et al., 2002; Zieg and Marsh, 2005), and Morokweng (Hartet al., 2002) have undergone large-scale igneous differentiation

Table 1Bulk composition model parameters.

Layer Thickness (km) Density (g/cm3) An (wt.%) En (wt.%) Fo (wt.%)

Upper crust 35 2.8 86 10.5 3.5Lower crust 35 3.0 60 30 10Upper mantle >70 3.3 0 75 25


resulting in cumulate stratigraphies similar to those of well-knowndifferentiated plutons such as the Skaergaard and the Stillwater. TheSudbury Igneous Complex, the largest of these impact melt sheets,has a thickness of�2.5–3 km and a volume of�104 km3 (Thierraultet al., 2002). On the basis of our analysis in Section 2, the solidifiedOrientale melt sea has a thickness of �14 km and a volume of�106 km3. Terrestrial impact melt sheets less than a tenth of the thick-ness of the Orientale melt sea and a hundredth of its volume have under-gone large-scale igneous differentiation. We therefore anticipate thatthe Orientale melt sea too has undergone large-scale igneousdifferentiation.

The idea that lunar impact melt sheets have undergone igneousdifferentiation is by no means new: see, e.g., Basilevsky and Neu-kum (2010), Cintala and Grieve (1998), Grieve et al. (1991), andMorrison (1998). However, this idea deserves to be revisited nowmore than ever, as recent and ongoing missions (the Moon Miner-alogy Mapper (M3) instrument on India’s Chandrayaan-1 missionand the Lunar Reconnaissance Orbiter (LRO) and Gravity Recoveryand Interior Laboratory (GRAIL) missions) collect high-resolutioncompositional and geophysical data that make it possible to inves-tigate the lunar highlands crust (and massive impact melt sheets)in unprecedented detail.

3.2. Modeling the cumulate stratigraphy of the Orientale melt sea

In the following sections, we develop a model for the cumulatestratigraphy of the Orientale melt sea (indeed, for any impact meltsea formed from melting of the lunar highlands crust and/or uppermantle: see Section 4.2). This model makes predictions aboutobservables: namely, the lithologies present at depth in a solidifiedmelt sea, which may be sampled by central peaks and walls ofsuperposed craters, and the density structure of this solidified meltsea, which may be determined by seismic and/or gravitymeasurements.

A general model of cumulate stratigraphy can be subdividedinto three component models: (1) a mixing model to recover thecomposition of the differentiating liquid, if unknown; (2) a thermo-dynamic model to determine the crystallization sequence of this li-quid composition; and (3) a fluid mechanical model to convert thiscrystallization sequence into a stratigraphic sequence. Our modelof the cumulate stratigraphy of the Orientale melt sea (1) deter-mines the composition of the differentiating liquid by mixing lunarcrust and upper mantle compositions weighted by their mass frac-tions in a modeled melt cavity, (2) finds the crystallization se-quence of this liquid on the Fo–An–Qz pseudoternary phasediagram, and (3) sinks or floats crystals according to crystal–liquiddensity contrasts in order to convert the crystallization sequenceinto a stratigraphic sequence.

3.3. Composition of liquid(s) in the Orientale melt sea (bulkcomposition)

We calculate the bulk composition of the Orientale melt sea bymixing lunar crust and upper mantle compositions weighted bytheir mass fractions in a modeled melt cavity (Fig. 4a).

The geometry of the melt cavity (Fig. 4a) has been approxi-mated as a truncated sphere (Barr and Citron, 2011; Pierazzoet al., 1997). For simplicity, we treat the melt cavity as a spheretangent to the lunar surface. The volume of this spherical melt cav-ity must be the volume of impact melt produced, which was calcu-lated in Section 2.3 to be to be �1.5 � 106 km3. Therefore, theradius of the melt cavity is �70 km (i.e., the cube root of1.5 � 106 km3/(4p/3)), rounded to an even 70 km for the followingcalculations. The diameter of the melt cavity is approximatelytwice the average crustal thickness of the lunar farside (Wieczoreket al., 2006; Zuber et al., 1994), implying that the Orientale impact

melted the lunar upper mantle. The Orientale impact apparentlydid not excavate the lunar upper mantle—no mafic mantle materialhas been detected in Orientale basin deposits (Head, 2010; Headet al., 2010a; Pieters et al., 2009; Yamamoto et al., 2010)—but thisis as expected, since in Orientale-size lunar basins with a transientcavity 620 km in diameter the depth of melting is modeled to ex-ceed the depth of excavation by a factor of two (see Fig. 22 of Cint-ala and Grieve (1998)).

We model the lunar highlands crust and upper mantle as com-posed of entirely anorthite, enstatite, and forsterite in varying pro-portions (Shearer et al., 2006). This is a good approximation for thecrust—lunar highland rocks generally have modal plagio-clase + pyroxene + olivine > 99 (Shearer et al., 2006; Wieczoreket al., 2006)—and a reasonable approximation for the upper mantlefrom the base of the crust to 200 km, which has been modeled asorthopyroxenite by thermodynamic models matching measuredand modeled seismic velocities (Khan et al., 2006; Kuskov, 1997).Ti-rich phases are neglected, but these are probably more impor-tant in the nearside crust and upper mantle (Wieczorek et al.,2006). We take the farside crustal thickness to be 70 km (Zuberet al., 1994). This is somewhat thicker than suggested by recentmodeling (Wieczorek et al., 2006), but using the number of70 km affords two conveniences: first, in this case the modeleddepth of melting is twice the thickness of the lunar crust, so thatexactly half the volume of the melt cavity can be taken to be crustalmaterial; second, approximating the melt cavity geometry as asphere rather than as a truncated sphere (Barr and Citron, 2011)somewhat decreases the volume of crustal material in the meltsea, which is counteracted by increasing the modeled crustal thick-ness. The lunar crust is compositionally stratified: the upper crustis anorthosite and the lower crust, exposed in basin ejecta, is prob-ably anorthositic norite (Pieters et al., 2009; Wieczorek et al.,2006). Therefore, we separate the crust into two planar layers ofequal 35 km thickness; the thicknesses of these crustal layers areequal, as is approximately the case in the dual-layered crustalthickness model of Wieczorek et al. (2006). The densities and com-positions of these layers are based on the dual-layered crustalthickness model of Wieczorek et al. (2006) with the additional con-straint that enstatite and forsterite are present in the proportion3:1 in both these layers, justified by the modal mineralogy of anor-thosite and anorthositic norite samples (some lunar sample modesare tabulated in Wieczorek et al. (2006)). The upper mantle is takento have a density of 3.3 g/cm3 and a composition of 75 wt.% ensta-tite and 25 wt.% forsterite (Khan et al., 2006). All bulk compositionmodel parameters are presented in Table 1.

These parameters (Table 1) capture in the An–Fo–Qz system thesalient features of the lunar crust and upper mantle—an anortho-site upper crust, an anorthositic norite lower crust, and a pyroxe-nite upper mantle. Minor changes in these parameters do notchange the crystallization sequences in Section 3.5.

The mass fraction of each layer melted by a 140 km diameterspherical melt cavity tangent to the lunar surface is calculated fromthe volume of the spherical melt cavity intersecting the layer and thedensity of the layer (Table 2). The bulk composition of the melt sea isthen given by the compositions of each layer weighted by their massfractions in the melt (Table 2). The resulting melt sea is fairly mafic,with an olivine noritic bulk composition (Stöffler et al., 1980). Thismodeled mafic bulk composition is not necessarily inconsistent with

Table 2Bulk composition of the Orientale melt sea.

Layer (vol.%) (wt.%)

Proportion of melt seaUpper crust 15.6 14.0Lower crust 34.4 33.1Upper mantle 50.0 52.9

An (wt.%) En (wt.%) Fo (wt.%)

Bulk composition of melt sea31.9 51.1 17.0


remote sensing observations suggesting that surface or near-surfaceportions of the melt sea are anorthositic or noritic (see Section 4.1),since melt sea differentiates need not have the bulk composition ofthe melt sea (see the discussion of South Pole-Aitken basin inSection 4.2.2).

3.4. Initial composition of liquid(s) in the Orientale melt sea (meltmixing)

We now investigate how the bulk composition of the Orientalemelt sea is related to the initial composition of differentiating li-quid(s) in the sea.

Vigorous convection, driven by the thermal gradient betweenhot melt and a cold free surface, is anticipated in a large melt sea(Morrison, 1998). Convection may mix and homogenize the meltsea so that the composition of any small volume of melt is identicalto the bulk composition of the melt. In this case, the initial compo-sition of the differentiating liquid in the Orientale melt sea is sim-ply the bulk composition of the Orientale melt sea given in Table 2above.

However, even a well-mixed melt sea may not be homogenous:Zieg and Marsh (2005) suggest that the relatively homogenousgranophyre and norite layers of the Sudbury Igneous Complex, aterrestrial impact melt sheet, formed due to physical liquid immis-cibility. In their scenario, mafic and felsic liquids (melts of differenttarget lithologies) were prevented from rapidly mixing by sheardue to their order-of-magnitude-different viscosities. Before theseliquids could mix by diffusion, they separated according to theirdensities, a process taking only �2 weeks in Sudbury (Zieg andMarsh, 2005), giving rise to a density-stratified melt sea with alight granophyre layer overlying a dense norite layer.

Similarly, anorthosite and mafic liquids could have separated inthe Orientale melt sea, since plagioclase-rich crustal melts and ma-fic upper mantle melts have very different viscosities and densities.The density contrast between anorthosite (�2.6 g/cm3) and pyrox-enite liquids (�2.9 g/cm3) (Nelson and Carmichael, 1979) is nearlythree times greater than the density contrast between the grano-phyre (�2.7 g/cm3) and norite (�2.8 g/cm3) layers of the SudburyIgneous Complex (Zieg and Marsh, 2005). The threefold greaterdensity contrast of lunar felsic and mafic liquids compared withSudbury liquids could partially offset 1/6th lunar gravity so thatdensity separation of immiscible liquids could be almost as effec-tive on the Moon as in the Sudbury Igneous Complex. Therefore,we also model the relative volume of anorthositic and mafic liquidsin a density-stratified melt sea (Table 3).

Table 3Density-stratified melt sea parameters. Layer thicknesses are for a cyli

Layer Density (g/cm3) P

Anorthositic liquid (100 wt.% An) 2.61 3Mafic liquid (75 wt.% En, 25 wt.% Fo) 2.92 6

One-third of the melt produced by the Orientale impact mayhave been ejected from the Orientale melt sea (Section 2.4) to beincorporated in the rough facies of the Maunder Formation andHevelius Formation ejecta. Separation of physically immiscible liq-uids takes weeks to months (Zieg and Marsh, 2005) after impact,much longer than the processes that drive melt ejection, transientcavity excavation and subsequent crater modification (Melosh,1989; Hawke and Head, 1977). On this basis, we take the melt tohave been homogeneous during melt cavity collapse: in this case,even though ejection decreases the volume of the melt sea byone-third, it does not affect its bulk composition.

3.5. Crystallization sequences in the Orientale melt sea

We now model the crystallization sequences of homogenousand density-stratified melt seas. Both equilibrium crystallization(crystals always in equilibrium with liquid) and perfect fractionalcrystallization (crystals only instantaneously in equilibrium withliquid) are modeled in order to bound melt sea lithologies,although vigorous convection is likely to have kept crystals in equi-librium with liquid in the Orientale melt sea.

Crystallization sequences are determined using the An–Fo–Qzpseudoternary phase diagram (Fig. 8a). The 1 atm phase diagramis used since pressures are <1 kbar even at the bottom of the meltsea, 15.5 km deep in the crust (Mizutani and Osako, 1974). Themelt has an initial temperature >2000 �C (Melosh, 1996), wellabove the liquidus of the homogenous melt sea at �1550 �C. (Theliquidus of the mafic liquid in the density-stratified melt sea is sim-ilar, �1650 �C.)

3.5.1. Equilibrium crystallization of a homogenous melt seaOlivine-saturated liquid crystallizes olivine until the liquid be-

comes sufficiently silicic to react with previously-crystallized olivineto form orthopyroxene. The liquid evolves along the olivine–orthopy-roxene boundary curve to the peritectic, where plagioclase comes onthe liquidus; the liquid is consumed before all olivine has been re-sorbed. The final crystallization sequence is olivine ? orthopyrox-ene � olivine ? plagioclase + orthopyroxene � olivine (Fig. 8b).

3.5.2. Fractional crystallization of a homogenous melt seaOlivine-saturated liquid crystallizes olivine until the liquid be-

comes sufficiently silicic to crystallize orthopyroxene. No olivinesolids are in equilibrium with the liquid, which leaves the oliv-ine–orthopyroxene boundary curve, crystallizing orthopyroxeneuntil the liquid becomes plagioclase-saturated. The liquid evolvesalong the plagioclase-orthopyroxene boundary curve to the eutec-tic, where plagioclase, orthopyroxene, and quartz co-crystallizeuntil the liquid is consumed. The final crystallization sequence isolivine ? orthopyroxene ? plagioclase + orthopyroxene ? plagio-clase + orthopyroxene + quartz (Fig. 8c).

3.5.3. Equilibrium crystallization of a density-stratified melt seaThe pure plagioclase-composition liquid does not undergo dif-

ferentiation. The crystallization sequence of the mafic liquid (aFo + En mixture) follows from the well-known Fo–Qz binary phasediagram: olivine-saturated liquid crystallizes olivine until the li-quid evolves to the peritectic, where olivine reacts with liquid toform orthopyroxene; the liquid is consumed before all olivine

ndrical melt sea 15.5 km thick.

roportion of melt sea (wt.%)|(vol.%) Thickness (km)

1.9 34.4 5.38.1 65.6 10.2

Fig. 8. One-atmosphere An–Fo–Qz ternary (a) phase diagram after Winter (2010) with (b) equilibrium crystallization and (c) fractional crystallization sequences forcomposition X, the Orientale melt sea.


has been resorbed. The final crystallization sequence of the maficliquid is olivine ? orthopyroxene � olivine.

3.5.4. Fractional crystallization of a density-stratified melt seaThe pure plagioclase-composition liquid does not undergo dif-

ferentiation. The crystallization sequence of the mafic liquid (aFo + En mixture) follows from the well-known (see, e.g., Winter(2010)) Fo–Qz binary phase diagram: olivine-saturated liquid crys-tallizes olivine until the liquid evolves to the peritectic. All olivinehas been removed from the liquid, so the liquid compositionleaves the peritectic, crystallizing orthopyroxene and finally ortho-pyroxene and quartz at the eutectic. The final crystallizationsequence of the mafic liquid is olivine ? orthopyroxene ?orthopyroxene + quartz.

The crystallization sequences above are probably a good firstapproximation to the true crystallization sequence of the Orien-tale melt sea. However, certain phases are not modeled, such asclinopyroxene and Ti-bearing phases. Also, the composition ofsolids is assumed to be constant, whereas in reality plagioclase,olivine, and orthopyroxene are all solid solutions. The Mg or Annumber of the first olivine/orthopyroxene or plagioclase solids isalmost certainly greater than the Mg or An number of the lastolivine/orthopyroxene or plagioclase solids. Since the appearanceof clinopyroxene or changes in Mg or An number may provideimportant constraints on melt sheet differentiation, we plan touse more sophisticated thermodynamic models to investigatethe crystallization sequences of lunar impact melt sheets ingreater detail.

Fig. 9. Anorthite crystal sink-float diagram for forsterite Mg numbe

3.6. Crystal–liquid density contrasts in the Orientale melt sea

We compare crystal and liquid densities to determine whethercrystallizing solids will sink or float in order to convert the crystal-lization sequences in Section 3.5 to stratigraphic sequences.

An–Fo–Qz liquids are substantially less dense (always >0.1 g/cm3, calculated from Nelson and Carmichael (1979)) than anorthitecrystals (q = 2.75 g/cm3). Even in an iron-rich liquid with a con-stant Mg number of 40 (Fig. 9a), more ferroan than even the inter-stitial mafic minerals in ferroan anorthosites (Shearer et al., 2006),anorthite never floats in a plagioclase-saturated liquid (Fig. 9b).Our conclusion is that anorthite crystals do not float in (Ti-poor)impact melts of the crust and upper mantle. The critical implica-tions of this idea for the cumulate stratigraphies and density struc-tures in impact melts of the crust will be considered inSection 4.2.1. Since plagioclase crystals are substantially less densethan orthopyroxene and olivine crystals (q � 3.2 g/cm3), we con-clude that all crystallizing solids sink in their coexisting liquids.Therefore, the crystallization sequences in Section 3.5 above areeffectively stratigraphic sequences: for example, co-crystallizingplagioclase and orthopyroxene both sink to form norite.

3.7. Normative calculations and stratigraphic columns for theOrientale melt sea

Finally, we calculate the weight proportions and thicknesses ofeach co-crystallizing assemblage to create stratigraphic columnsfor solidified melt seas. (Although crystal settling is likely to be

r = 40 (a) superposed on An–Fo–Qz ternary phase diagram (b).

Fig. 10. Model Orientale cumulate stratigraphies and density profiles (to scale)produced by equilibrium and fractional crystallization of homogenous and density-


kinetically incompatible with perfect equilibrium crystallization,we nonetheless convert equilibrium crystallization sequences tostratigraphic sequences in order to bound melt sea stratigraphies.Real stratigraphic sequences may approach sequences based onequilibrium crystallization.) The following sections briefly describeour normative calculations, which are based on the An–Fo–Qz andFo–Qz phase diagrams.

3.7.1. Equilibrium crystallization of a homogenous melt seaThe liquid crystallizes to give monomineralic olivine (dunite)

which then reacts with liquid to form an overlying layer of pyrox-enite and finally norite as plagioclase comes on the liquidus at theperitectic. The weight proportion of the dunite layer is given sim-ply by the initial wt.% Fo of the liquid; the norite layer has an An/Enratio of �3/2 as measured from the phase diagram and, accord-ingly, its weight proportion is given by the sum of the initialwt.% An and 2/3 that wt.% of En; the wt.% of the pyroxenite layeris equal to the remaining wt.% En (Table 4).

3.7.2. Fractional crystallization of a homogenous melt seaThe weight proportions of the four crystallizing lithologies—

monomineralic olivine (dunite), monomineralic orthopyroxene(pyroxenite), orthopyroxene + anorthite (norite), and orthopyrox-ene + anorthite + quartz (quartz diorite)—are constrained by fourequations, the first three relating the compositions of each lithol-ogy (measured from the An–Fo–Qz ternary) to the bulk composi-tion and the fourth relating the relative proportions of duniteand pyroxenite as determined from olivine and orthopyroxenecrystallization path length on the An–Fo–Qz ternary (Fig. 8c). Theresulting system of linear equations can be written as the followingmatrix equation:

1 0:7 0:32 0:160 0 0:55 0:510 0:3 0:13 0:33�1 1:5 0 0

0BBB@

1CCCA

XDu

XPx

XNr

XQz

0BBB@

1CCCA ¼

0:530:320:15

0

0BBB@

1CCCA ð1Þ

We solve this matrix equation for the weight proportions of theselithologies (Table 4).

3.7.3. Equilibrium crystallization of a density-stratified melt seaThe overlying anorthositic liquid crystallizes to anorthosite. The

mafic liquid crystallizes to form monomineralic olivine (dunite)and monomineralic orthopyroxene (pyroxenite) layers. The weight

Table 4Layer compositions and thicknesses in solidified cylindrical melt seas 14 km thick.

Layer C

Equilibrium crystallization, homogenous melt sea Norite 6Pyroxenite 1Dunite 1

Fractional crystallization, homogenous melt sea Quartz diorite 5Q

Anorthositic norite 5Pyroxenite 1Dunite 1

Equilibrium crystallization, density-stratified melt sea Anorthosite 1Pyroxenite 1Dunite 1

Fractional crystallization, density-stratified melt sea Anorthosite 1Quartz pyroxenite 2Pyroxenite 1Dunite 1

proportions (Table 4) of all three layers are given simply by the ini-tial wt.% An, Fo, and En of the liquid.

3.7.4. Fractional crystallization of a density-stratified melt seaThe weight proportions of the three crystallizing lithologies—

monomineralic olivine (dunite), monomineralic orthopyroxene(pyroxenite), and orthopyroxene + quartz (quartz pyroxenite)—are constrained by three equations, the first two relating the com-position of each lithology (measured from the Fo–Qz binary) to thebulk composition and the third relating the relative proportions ofdunite and pyroxenite as determined from olivine and orthopyrox-ene crystallization path length on the Fo–Qz binary. The resultingsystem of linear equations can be written as the following matrixequation:

1 0:7 0:550 0:3 0:45�1 3:3 0

0B@

1CA

XDu

XPx

XQz

0B@

1CA ¼

0:780:22

0

0B@

1CA ð2Þ

We solve this matrix equation for the weight proportions of theselithologies (Table 4).

omposition Density(g/cm3)

Proportion of meltsea (wt.%) (vol.%)

Thickness(km)

1.6 wt.% An, 38.4 wt.% En 2.95 51.8 54.0 7.600 wt.% En 3.20 31.2 30.0 4.200 wt.% Fo 3.25 17.0 16.0 2.2

1.0 wt.% An, 26.1 wt.%z, 22.9 wt.% En

2.85 13.7 14.6 2.0

5.0 wt.% An, 45.0 wt.% En 2.95 45.3 46.7 6.500 wt.% En 3.20 16.4 15.6 2.200 wt.% Fo 3.25 24.6 23.0 3.2

00 wt.% An 2.75 31.9 35.4 5.000 wt.% En 3.20 51.1 48.7 6.800 wt.% Fo 3.25 17.0 15.9 2.2

00 wt.% An, 78.6 wt.% En 2.75 31.9 35.1 4.91.4 wt.% Qz 3.10 27.8 27.2 3.800 wt.% En 3.20 9.4 8.9 1.200 wt.% Fo 3.25 30.9 28.8 4.0

stratified melt seas.


Finally, the weight proportions are converted to volume propor-tions based on layer density (a weighted average of componentmineral density); these volume proportions are directly propor-tional to the thicknesses in a cylindrical melt sea (Table 4). Wechoose to model melt seas as cylindrical since this geometry hasbeen predicted as the consequence of inward movement and up-ward translation during transient cavity collapse (Head, 2010,2012; Head et al., 1993). (The effect of a melt sea with a curvedbase, which has more of its volume near the surface, is to increasethe thickness of deep layers and decrease the thickness of shallowlayers.) The cumulate stratigraphies and density structures of thesemodeled melt seas are shown in Fig. 10.

4. Discussion

We now assess our modeled cumulate stratigraphies (Fig. 10)and consider more generally the implications of impact melt differ-entiation for the lunar highlands crust.

4.1. Constraints on Orientale melt sea cumulate stratigraphy:implications for processes once operating in the Orientale melt sea

The modeled cumulate stratigraphies (Fig. 10) produced bymelt homogenization/density stratification and equilibrium crys-tallization/fractional crystallization have observable differences.Constraints on the lithologies present at depth in the solidified Ori-entale melt sea make it possible to rule out some of these modeledstratigraphies and determine what processes operated in the mix-ing and crystallization of the Orientale melt sea.

There are currently two important constraints on the cumulatestratigraphy of the Orientale melt sea: (1) The top �1 km of thesolidified impact melt sea is anorthosite. M3 spectral observationsof the Maunder Formation (Head, 2010; Head et al., 2010a; Pieterset al., 2009) indicate that the upper surface of the smooth facies isanorthosite. Moreover, small (�10 km diameter) craters superpos-ing the impact melt sheet, which depth/diameter ratios predictexcavate material from a depth of up to �1 km, uncover moreanorthosite. (2) Norite occurs at �4 km depth. The central peak ofMaunder, a 55 km diameter crater superposed on the Orientalemelt sheet, is predicted by depth/diameter ratios to have sampledthe melt sea at a depth of �4 km (Whitten et al., 2011). This peakhas a noritic composition (Whitten et al., 2011). These two con-straints can themselves be interpreted as evidence for the differen-tiation of the Orientale melt sea: near-surface lithologies are notthe same as those present at depth.

Which modeled cumulate stratigraphies (Fig. 10) are consistentwith these constraints? Fractional crystallization of a homogenousmelt sea produces a �2 km thick layer of quartz diorite at the topof the melt sea, clearly inconsistent with constraint (1). Densitystratification produces a �5 km thick layer of pure anorthosite atthe top of the melt sea, inconsistent with constraint (2). Equilib-rium crystallization of a homogenous melt sea produces a�7.5 km thick layer of norite at top of the melt sea, consistent withconstraint (2). This is not necessarily inconsistent with constraint(1): the anorthositic roof of the Orientale melt sheet could be anearly-formed quench crust; a fallback breccia, as in Sudbury (Thi-errault et al., 2002); or the result of modal separation of plagioclasefrom pyroxene in late-crystallizing norite.

Therefore, we conclude that the Orientale melt sea was homog-enized by melt mixing and underwent equilibrium crystallization.Indeed, both homogenization and equilibrium crystallization sug-gest vigorous convection, consistent with a cooling lava lake (Wor-ster et al., 1990) and particularly with a cooling impact melt sea(Morrison, 1998).

It is possible that high-resolution gravity measurements fromthe ongoing GRAIL mission could provide additional constraintson Orientale melt sea cumulate stratigraphy. Certainly, gravitymeasurements from the GRAIL mission can provide constraintson melt sea thickness. We note in passing that, in light of our evi-dence for a thick sea of impact melt at the center of the Orientalebasin, crustal thickness models for the Orientale basin need to bereassessed (Fig. 11). We predict that the ‘‘crust’’ occurring at thecenter of the Orientale basin is actually a solidified melt sea witha density of �3.05 g/cm3, not anorthosite or brecciated anorthositewith a density of �2.75 g/cm3, as has been traditionally assumed(Andrews-Hanna and Stewart, 2011; Hikida and Wieczorek,2007). The �10 km thick anorthositic crust modeled to occur atthe center of the Orientale basin (Andrews-Hanna and Stewart,2011; Hikida and Wieczorek, 2007) becomes a �20 km thick im-pact melt sheet (Fig. 11) using this revised density, roughly consis-tent with (although �1.3 times greater than) our melt sheetthickness estimates in Section 2.

4.1.1. Is crystal settling an operative differentiation mechanism inimpact melt seas?

In this paper, we interpret the differentiation of the Orientaleimpact melt sea (and other lunar impact melt seas; see Section 4.2)in terms of crystal settling: we transform crystallization sequencesinto stratigraphic sequences under the assumption that gravityefficiently segregates crystals from melt to form igneous strata.We are by no means the first authors to make this assumption:Morrison (1998) and Nakamura et al. (2009) used crystal settlingto interpret the differentiation of the South Pole-Aitken Basin im-pact melt sheet. However, field studies and theoretical analysesof terrestrial impact melt sheets and magma chambers (Marsh,1989; Marsh, 1996; Warren et al., 1996; Zieg and Marsh, 2005)may indicate that crystal settling does not operate in the differen-tiation of impact melt. Field studies of the Sudbury Igneous Com-plex, a terrestrial impact melt sheet, show that massivegranophyre and norite layers in the Sudbury Igneous Complexare homogenous; crystal settling did not take place in their parentmagmas (Zieg and Marsh, 2005). This is perhaps unsurprising:crystal settling operates on isolated phenocrysts in a magma body,not the dense mat of crystals along the cooling margin of a magmabody (Marsh, 1989; Marsh, 1996); impact melts, however, aresuperheated (>2000 �C) (Melosh, 1996), and contain no pheno-crysts to settle. Marsh (1996) writes, ‘‘. . .initially crystal-free mag-mas [such as impact melts] cannot undergo strong chemicaldifferentiation by fractional crystallization of singly settlingcrystals.’’

What are the implications for solidified melt sea stratigraphy ifcrystal settling indeed does not operate in the differentiation of im-pact melt seas? Impact melt bodies are then either (1) homogenousor (2) heterogeneous but differentiated by another mechanism,such as viscous emulsion differentiation (Zieg and Marsh, 2005).However, the constraints on Orientale melt sea cumulate stratigra-phy described in Section 4.1 are inconsistent with large-scalehomogeneity and not entirely consistent with viscous emulsiondifferentiation (Fig. 10, lower left). Perhaps our model of viscousemulsion differentiation is colored by our assumption that the cra-ter excavation flow does not change the bulk composition of theimpact melt sea: if the excavation flow removed only impact meltof upper crustal composition, an uppermost anorthosite layer(Fig. 10, lower left) would be thinned, bringing noritic lithologiestowards the surface. In this case, the noritic central peak of Maun-der represents recrystallized lower crustal material.

If viscous emulsion differentiation is indeed the operative dif-ferentiation mechanism in impact melt seas, solidified impact meltseas recapitulate the structure of the lunar crust and upper mantle,just as the Sudbury impact melt sheet recapitulates the preexisting

Fig. 11. Traditional (a) crustal thickness models for the center of the Orientale basin, following Andrews-Hanna and Stewart (2011) and Hikida and Wieczorek (2007), and therevised model (b) described here.

Fig. 12. Bulk composition of an impact melt sea as a function of basin size (km).


crustal structure of the Canadian shield (Zieg and Marsh, 2005).Moreover, since impact melt seas have a very different aspect ratiofrom the initial melt cavity (Fig. 4), solidified impact melt bodiesmay bring lower crust and upper mantle materials to the near sur-face where they can be excavated by complex craters of moderatesize (such as Maunder in Orientale). Further investigation of lunarimpact melt sea stratigraphy will clarify the differentiation of lunarimpact melts and the roles of crystal settling and viscous emulsiondifferentiation.

4.2. Cumulate stratigraphy of other lunar impact melt seas

There are at least 30 lunar basins (craters >300 km diameter)besides Orientale, as determined from the database of Head et al.(2010b). Impact melt scaling relations (Cintala and Grieve, 1998)predict that the impacts that formed these basins each produced>105 km3 of melt. Very probably, these basins, like Orientale, con-tain solidified impact melt seas. What is the cumulate stratigraphyof these melt seas and how might it differ from the cumulate stra-tigraphy of Orientale? Since target crust/upper mantle structure isnot thought to change substantially across the farside highlands(Wieczorek et al., 2006), the important variable controlling meltsea bulk composition is depth of melting, which is a function of ba-sin size. The compositional endmembers are (1) relatively smallbasin-forming impacts that melt only the lunar crust and (2) giantbasin-forming impacts that melt mainly the lunar mantle (Fig. 12).The compositional watershed is the plagioclase-olivine cotectic(Fig. 12): small basin-forming impacts form melt seas that becomeplagioclase-saturated early, whereas large basin-forming impactsform melt seas that become plagioclase-saturated late in crystalli-zation. Having considered the cumulate stratigraphy of a melt seain a moderately-sized basin, Orientale, we now proceed to modelthe cumulate stratigraphy of melt seas occurring in a small lunarbasin and the very large South-Pole Aitken basin.

4.2.1. Cumulate stratigraphy of a melt sea in a small lunar basinWe consider a hypothetical basin-forming impact that melts the

entire thickness of the lunar crust and none of the upper mantle,

corresponding to a spherical melt cavity 70 km in diameter, whichwe take to be the thickness of the farside crust (Wieczorek et al.,2006; Zuber et al., 1994). A spherical melt cavity 70 km in diametercontains �1.7 � 105 km3 of melt, which, according to the scalingrelations of Cintala and Grieve (1998), corresponds to a basin tran-sient cavity �250 km in diameter and a final basin diameter of�415 km (Croft, 1985), almost exactly the diameter of the basinKorolev (Baker et al., 2011a). We reuse our modeled bulk composi-tion parameters (Table 1) to calculate the bulk composition of thiscrustal melt sea: 72.5 wt.% An, 20.5 wt.% En, and 7.0 wt.% Fo. Wemodel equilibrium crystallization of a homogenized crustal meltsea since these processes best reproduced the observed cumulatestratigraphy of the Orientale melt sea (Section 4.1).

This crustal melt is initially plagioclase-saturated (Fig. 13a). Pla-gioclase crystallizes until the liquid becomes olivine-saturated; theliquid then evolves along the plagioclase-olivine cotectic to the

Fig. 13. One-atmosphere An–Fo–Qz ternary (a) phase diagram after Winter (2010) with (b) the equilibrium crystallization sequence for composition X0 , a crustal melt.


peritectic, where orthopyroxene and plagioclase co-crystallize; theliquid is consumed before all olivine has been resorbed. The crys-tallization sequence of this liquid (Fig. 13b) is plagioclase ? plagio-clase + olivine ? plagioclase + orthopyroxene � olivine. Plagio-clase sinks in impact melts of the crust (Fig. 9), so thiscrystallization sequence is effectively the cumulate stratigraphy.We calculate the weight proportions of these co-crystallizingassemblages using methods similar to those in presented in Sec-tion 3.7. The cumulate stratigraphy of a crustal melt sea occurringin a Korolev-sized basin is compared with the likely cumulate stra-tigraphy of the solidified Orientale melt sea in Fig. 14.

The lithologies exposed at or near the surface or in the centralpeaks of complex craters are modeled to be norite in both smalland large melt seas. The critical observable difference betweenthese two melt seas is their density structure. The crustal meltsea has an inverted density profile—the result of early-crystallizingplagioclase solids sinking to form an anorthosite layer beneathdenser troctolite and norite layers—whereas the Orientale andother mantle melt seas have a regular density profile with a densedunite layer underlying less dense pyroxenite and norite. Futureseismic measurements on solidified melt sea surfaces could poten-tially resolve these different density structures.

4.2.2. South Pole-Aitken basin melt sea cumulate stratigraphyThe enormous 2500 km diameter South Pole-Aitken (SPA) basin

(Wilhelms et al., 1979) almost certainly melts huge volumes ofultramafic mantle material (Morrison, 1998). However, the SPAfloor, possibly the surface of a melt sea, is noritic, not ultramafic,and craters superposing this floor excavate pyroxene, not mantleolivine (Borst et al., 2012; Moriarty et al., 2011, 2012; Nakamuraet al., 2009). Morrison (1998) suggested that differentiation ofthe SPA melt sheet could reconcile these apparently conflictingobservations: a basal ultramafic layer could sequester mantle oliv-ine and an upper noritic layer could concentrate crustal plagio-clase. We estimate the thickness and model the cumulatestratigraphy of the South Pole-Aitken basin melt sea to investigatethis idea in greater detail.

How thick is the SPA melt sea? We take the SPA transient cavitydiameter to be 2099 km (Wieczorek and Phillips, 1999). Using theimpact melt volume scaling relationship derived by Abramov et al.(2012) and the model parameters described in Kring et al. (2012),we calculate that the South Pole-Aitken basin-forming impact pro-duced �6.8 � 107 km3 of impact melt. About 25 vol.% of this meltwas excavated from the basin transient cavity, leaving5.1 � 107 km3 of impact melt inside the inner basin ring to form

a melt sea. The thickness of this melt sea is constrained by its arealextent. Morrison (1998) calculated the areal extent of the SPA meltsea by taking the inner ring of SPA to be a circle �1800 km in diam-eter (Wilhelms et al., 1979). We calculate the areal extent of theSPA melt sea using the recent measurements by Garrick-Bethelland Zuber (2009) suggesting that the inner ring of SPA is an ellipsewith a semimajor axis of 970 km and a semiminor axis of 720 km.A prismatic melt sea with this areal extent containing�5.1 � 107 km3 of impact melt is �25 km thick. However, the Ori-entale melt sea probably occurs in a central depression with adiameter �70% of that of the Outer Rook Ring (350 km vs.480 km). If this ratio holds for SPA, so that the semimajor andsemiminor axes of the surface of the SPA melt sea are �680 km(70% of 970 km) and �500 km (70% of 970 km), then the SPA meltsea may be �50 km thick.

What is the cumulate stratigraphy of a hypothetical solidifiedSPA melt sea? A spherical melt cavity containing the estimated�6.8 � 107 km3 SPA melt volume has a radius of �250 km, corre-sponding to a depth of melting of �500 km. The bulk compositionof the SPA melt sea, reusing the bulk composition model parame-ters in Table 1, is nearly that of the upper mantle, with 4 wt.%An, 72 wt.% En, and 24 wt.% Fo. The crystallization sequence of thismelt and the modal calculations are the same as for the Orientalemelt lake described in Section 3. The solidified South Pole-Aitkenbasin melt sea, if�25 km thick, is thus predicted to consist of a thin�2.6 km layer of norite overlaying a massive �16.5 km layer ofpyroxenite and a basal 5.9 km thick dunite cumulate (Fig. 14).

What are the implications of this cumulate stratigraphy? First,the noritic floor of SPA could very well be the roof of a massive dif-ferentiated impact melt sea. Second, observations of crater centralpeaks that determine the depth at which norite transitions topyroxenite can be used to constrain the thickness of the SPA meltsea. Third, Nakamura et al. (2009) argue that the SPA impact meltsea has not undergone differentiation on the basis that all centralpeaks observed in their study excavate pyroxenite. We suggest in-stead that the SPA impact melt sea has undergone differentiationand that all central peaks observed in their study (which excavatematerial from depths of �6–10 km) excavate a 16.5 km thickhomogenous pyroxenite cumulate (Fig. 14).

4.3. Impact melt differentiates in the lunar sample suite?

Huge volumes of impact melt are produced by basin-formingimpacts. Much of this impact melt may have differentiated: theOrientale-forming impact alone may have given rise to �106 km3

Fig. 14. Model cumulate stratigraphies and density profiles for melt seas in basins of three different sizes: a ‘‘crustal’’ melt sea occurring in a small basin approximately thesize of Korolev, the Orientale melt sea, and the SPA melt sea. Layers in each melt sea are to scale, although melt seas are not to the same scale.


of impact melt differentiates. For comparison, the total volume oflunar mare basalts is estimated to be �107 km3 (Head and Wilson,1992). Could impact melt differentiates excavated from massiveimpact melt seas be represented in the lunar sample suite?

4.3.1. Can impact melt differentiates pass as pristine?Pristine lunar rocks (Warren, 1993), which have coarse-grained

textures and low siderophile element concentrations, are thoughtto represent endogenous lunar igneous rocks largely unaltered byimpact processes. We suggest that impact melt differentiatesformed in massive impact melt seas also have coarse-grained tex-tures and low siderophile element concentrations and so may beable to pass for pristine lunar rocks (Basilevsky and Neukum,2010).

Massive impact melt seas cool slowly. The diameter of the Ori-entale melt sea (�175 km) exceeds its probable depth (�15.5 km)by an order of magnitude, so we approximate a (non-convecting)melt sea as a cooling slab similar to an intruded sill. The solidifica-tion time t of a sill cooling by conduction can be approximated bythe formula

t ¼ 0:694L2

� �2

Kð3Þ

where L is the half-thickness of the sill and K is its thermal diffusiv-ity, �10�2 cm2/s for rock (Jaeger, 1968). From (3), we estimate thatthe Orientale melt sea took on the order of 105 years to solidify, longenough for impact melt differentiates formed in that sea to developcoarse-grained textures (Zieg and Marsh, 2002). These calculationsare supported by the observation that impact melt differentiatesformed at depth in the �3 km thick terrestrial Sudbury impact meltsheet are also coarse-grained (Zieg and Marsh, 2002) and were longmistaken for endogenous plutonic rocks (Thierrault et al., 2002).

Also, although impact melt seas may have a relatively high bulkconcentration of meteoritic siderophiles, these siderophile ele-ments may fractionate into basal deposits of metal or sulfides un-mixed from metal- or sulfur-saturated impact melt, so thatsilicates from these impact melt seas may have nearly lunar, notmeteoritic, siderophile element concentrations.

4.3.2. Excavating impact melt differentiatesMost impact melt differentiates form at depth. Certain impact

melt differentiates, such as dunites and troctolites, may form atdepths below 10 km. For these lithologies to be exposed at the sur-face and introduced into the lunar sample suite, they must be exca-vated by the impact cratering process. Very deep lithologies canonly be excavated by very large craters; we can see that deep dif-ferentiates cannot possibly have been excavated at Orientale (nocraters >65 km superpose the Orientale melt sheet, so dunite pos-

sibly occurring below 12 km in the Orientale melt sea has neverbeen excavated and cannot occur in the lunar sample suite). Atother basin melt sheet locations, shallow impact melt differenti-ates, probably noritic in composition, are the most likely to beexcavated by impacts and introduced into the lunar sample suite.Old basins (Fig. 15) with many superposed large craters and basins,such as the South Pole-Aitken basin (Head, 2010; Head et al., 1993,2010a; Pieters et al., 1997) are the most likely source regions fordeep impact melt differentiates, and certain norites or pyroxenitescould come from SPA.

Sections 4.3.1 and 4.3.2 suggest that impact melt differentiatesmay exist in the lunar sample suite and may be able to pass forhighland plutonic rocks. We now consider three highland plutonicrock suites—Mg-suite lithologies, young ferroan anorthosites, andMg-spinel lithologies—to determine whether these puzzling lithol-ogies could potentially be impact melt differentiates.

4.3.3. Mg-suite lithologiesMg-suite rocks—dunites, norites, and troctolites characterized

by high Mg numbers (70–95) and high rare earth element (REE)content (up to 400 times chondritic) (Hess, 1994; Shearer et al.,2006)—are traditionally thought to have formed from intrusionsof mafic magmas into ferroan anorthosite crust, since dunites, nor-ites, and troctolites can be formed from fractional crystallization ofa Fo-rich liquid assimilating plagioclase (see Fig. 8a for reference).The Mg-suite chemical signature is harder to explain: simple mix-ing of a high Mg number liquid with a low Mg number, REE-en-riched (KREEP) liquid produces a liquid with insufficiently highMg number to be the Mg-suite parent magma; Hess (1994) arguedthat the Mg-suite parent magma is not a mixture of melts, but amelt of a mixture: a mixture of KREEP, olivine-rich dunite, andanorthosite produces a partial melt with high Mg number andREE content. The process giving rise to this mixture is unclear; pro-cesses such as cumulate overturn have been invoked (Sheareret al., 2006).

Could differentiation of impact melt seas produce Mg-suitelithologies? Certainly, dunites, troctolites, and norites can form inimpact melt seas: the Orientale melt sea contains norite and prob-ably dunite, and smaller impacts that melt mainly the lunar crust(Section 4.2) will produce troctolite. But these lithologies are mix-tures of melts and for the reasons described above cannot have thefull Mg-suite chemical signature. Also, thermobarometry of Mg-suite rocks suggests that these rocks formed 40–50 km deep inthe crust (McCallum and Schwartz, 2001). It is possible that remo-tely-sensed Mg-suite lithologies are really unrelated Mg-suite-likedunites, troctolites, and norites lacking the Mg-suite chemical sig-nature. These Mg-suite-like lithologies could be impact melt differ-entiates; however, until Mg-suite-like lithologies are identified in

Fig. 15. Lunar basins greater than 300 km in diameter from Head et al. (2010b).


the lunar sample suite, the existence of these counterfeits must beregarded as speculation.

4.3.4. Young ferroan anorthositesCertain ferroan anorthosites with ages <4.4 Ga (Borg et al.,

2011) are much too young to have crystallized in the lunar magmaocean. These ferroan anorthosites are thought to be cumulatesformed during early plutonism (Shearer et al., 2006). Since impactmelting generally resets geochronometers, young ferroan anortho-sites could have formed in massive impact melt sheets (formationin thin, flash-frozen impact melt sheets is probably incompatiblewith the large grain sizes of these anorthosites). One difficulty withthis idea is that anorthite is negatively buoyant in impact melts ofthe crust and mantle (Fig. 9), so anorthosite generally lines the bot-toms of melt seas and may be difficult to excavate in quantity. Ifthe anorthositic roof of the Orientale melt sheet indeed formeddue to igneous differentiation, young anorthosites from impactmelt sheets may be more common, and easier to excavate, thanhas been imagined.

4.3.5. Mg-spinel lithologiesRecently discovered Mg-spinel lithologies (Pieters et al., 2011),

rocks with �30 wt.% spinel containing <5 wt.% Fe, have been ex-plained as the product of melt-wallrock reactions of Mg-suite mag-mas with anorthosite crust (Prissel et al., 2012). Potentially, spinelcould form from reaction of anorthite liquids formed by impactmelting with mafic minerals or melts. In density-stratified meltseas, anorthosite liquids overlie mafic liquids; large volumes of spi-nel could potentially be produced at this interface. One difficultywith this idea is that impact melts of the ferroan crust and mantleare relatively ferroan and could produce ferroan spinels.

4.4. Volumetric importance of impact melt in the lunar highlands crust

How much impact melt is incorporated in the lunar highlandscrust? The database of lunar craters P20 km in diameter compiledby Head et al. (2010b) includes 30 basins >300 km in diameter.Using the impact melt volume scaling laws of Cintala and Grieve(1998), we estimate that collectively these 30 basins contain onthe order of 108 km3 of impact melt, 10 times the total volume ofmare basalt (Head and Wilson, 1992) and about 1/20th the volumeof the lunar crust. These basins, and the massive impact meltsheets that they contain, are distributed widely across the Moon

(Fig. 15). Understanding the geology and the petrology of the108 km3 of massive impact melt deposits occurring in these ba-sins—work we have just begun to undertake here—is criticallyimportant for understanding the lunar highlands crust.

5. Conclusions

Basin-forming impacts forming craters >300 km in diameterproduce enormous volumes of impact melt (>105 km3). We haveinvestigated the geology and petrology of these huge volumes ofimpact melt, focusing especially on the relatively young lunar ba-sin Orientale, and conclude that:

1. Models of impact melt production combined with recent geo-logic and topographic data suggest that most impact melt(�2/3) produced by the Orientale-forming impact occurs in a�15.5 km thick (solidified to �14 km thick) impact melt sea�350 km in diameter with a volume of �106 km3.

2. The Orientale basin impact melt sea has probably undergonelarge-scale igneous differentiation, since terrestrial impact meltsheets (such as Manicouagan, Sudbury, and Morokweng) atenth of its thickness and a hundredth of its volume are knownto have differentiated.

3. A modeled cumulate stratigraphy (occurring below a quenchcrust and anorthositic fallback breccia, with an �8 km thicklayer of norite overlying a �4 km thick pyroxenite with a basal�2 km thick layer of dunite) produced by equilibrium crystalli-zation of a homogenized melt sea is consistent with remotely-sensed norite excavated from �4 km depth in the central peakof Maunder crater.

4. Small basins melting mainly the lunar crust produce initiallyplagioclase-saturated melt seas that first crystallize negativelybuoyant anorthosite and therefore develop inverted densitystructures. The noritic floor of the South-Pole Aitken basinmay be the roof of a differentiated impact melt lake; crater cen-tral peaks superposing SPA may excavate a massive layer ofpyroxenite.

5. Thirty lunar basins >300 km in diameter together contain on theorder of 108 km3 of impact melt, one-twentieth the total volumeof the lunar crust. Impact melt differentiates excavated fromthese solidified impact melt seas may well be represented inthe lunar sample suite, mimicking pristine highland plutoniclithologies.


6. High-resolution compositional and gravity data collected by theM3 instrument on India’s Chandrayaan-1 mission and the ongo-ing LRO and GRAIL missions make it possible to probe the struc-ture of the lunar highlands crust in unprecedented detail;analyses based on these datasets, such as constraining thecumulate stratigraphy of melt seas from careful compositionalanalyses of crater central peaks, will further our understandingof the geology and petrology of massive impact melts.

Acknowledgments

Thoughtful reviews by John Spray and Bruce Marsh improvedthis manuscript. This work was supported under NASA Lunar Sci-ence Institute (Brown-MIT) grant NNA09DB34A, which is grate-fully acknowledged.

References

Abramov, O., Wong, S.M., Kring, D.A., 2012. Differential melt scaling for obliqueimpacts on terrestrial planets. Icarus 218, 905–916. http://dx.doi.org/10.1016/j.icarus.2011.12.022.

Andrews-Hanna, J.C., Stewart, S.T., 2011. The crustal structure of Orientale andimplications for basin formation. Lunar Planet. Sci. XLII, 2194 (abstract).

Baker, D.M.H. et al., 2011a. The transition from complex crater to peak-ring basin onthe Moon: New observations from the Lunar Orbiter Laser Altimeter (LOLA)instrument. Icarus 214, 377–393. http://dx.doi.org/10.1016/j.icarus.2011.05.030.

Baker, D.M.H. et al., 2011b. The transition from complex crater to peak-ring basin onMercury: New observations from MESSENGER flyby data and constraints onbasin-formation models. Planet. Space Sci. 59, 1932–1948. http://dx.doi.org/10.1016/j.pss.2011.05.010.

Baker, D.M.H., Head, J.W., Neumann, G.A., Smith, D.E., Zuber, M.T., 2012. Thetransition from complex craters to multi-ring basins on the Moon: Quantitativegeometric properties from Lunar Reconnaissance Orbiter Laser Altimeter (LOLA)data. J. Geophys. Res. 117, E00H16. http://dx.doi.org/10.1029/2011JE004021.

Baldwin, R.B., 1974. Was there a ‘‘terminal lunar cataclysm’’ 3.9–4.0 � 109 yearsago? Icarus 23, 157–166. http://dx.doi.org/10.1016/0019-1035(74)90003-7.

Barberi, F., Varet, J., 1970. The Erta Ale volcanic range (Danakil depression, northernAfar, Ethiopia). Bull. Volcan. 34, 848–917.

Barr, A.C., Citron, R.I., 2011. Scaling of melt production in hypervelocity impactsfrom high-resolution numerical simulations. Icarus 211, 913–916. http://dx.doi.org/10.1016/j.icarus.2010.10.022.

Basilevsky, A.T., Neukum, G., 2010. Are All Lunar Highland Pristine Rocks ReallyPristine? Nördlingen 2010: The Ries Crater, the Moon, and the Future of HumanSpace Exploration, 7015 (abstract).

Bonini, W.E., 1982. The size of the Stillwater complex: An estimate from gravitydata. In: Walker, D., McCallum, I.S. (Eds.), Workshop on Magmatic Processes ofEarly Planetary Crusts: Magma Oceans and Stratiform Layered Intrusions. Lunarand Planetary Institute, Houston, TX, pp. 53–55.

Borg, L.E., Connelly, J.N., Boyet, M., Carlson, R.W., 2011. Chronological evidence thatthe Moon is either young or did not have a global magma ocean. Nature 477,70–72. http://dx.doi.org/10.1038/nature10328.

Borst, A.M., Foing, B.H., Davies, G.R., van Westrenen, W., 2012. Surface mineralogyand stratigraphy of the lunar South Pole-Aitken basin determined fromClementine UV/VIS and NIR data. Planet. Space. Sci. 68, 76–85. http://dx.doi.org/10.1016/j.pss.2011.07.020.

Bratt, S.R., Solomon, S.C., Head, J.W., 1985. The evolution of impact basins: Cooling,subsidence, and thermal stress. J. Geophys. Res. 90, 12415–12433. http://dx.doi.org/10.1029/JB090iB14p12415.

Bray, V.J. et al., 2010. New insight into lunar impact melt mobility fromthe LRO camera. Geophys. Res. Lett. 37, L21202. http://dx.doi.org/10.1029/2010GL044666.

Cintala, M.J., Grieve, R.A.F., 1998. Scaling impact melting and crater dimensions:Implications for the lunar cratering record. Meteorit. Planet. Sci. 33, 889–912.http://dx.doi.org/10.1111/j.1945-5100.1998.tb01695.x.

Collins, G.S., Melosh, H.J., Morgan, J.V., Warner, M.R., 2002. Hydrocode simulationsof Chicxulub crater collapse and peak-ring formation. Icarus 157, 24–33. http://dx.doi.org/10.1006/icar.2002.6822.

Croft, S.K., 1985. The scaling of complex craters. J. Geophys. Res., Suppl. 90, C828–C842.

Fassett, C.I., Head, J.W., Smith, D.E., Zuber, M.T., Neumann, G.A., 2011. Thickness ofproximal ejecta from the Orientale basin from Lunar Orbiter Laser Altimeter(LOLA) data: Implications for multi-ring basin formation. Geophys. Res. Lett. 38,L17201. http://dx.doi.org/10.1029/2011GL048502.

Garrick-Bethell, I., Zuber, M.T., 2009. Elliptical structure of the lunar SouthPole-Aitken basin. Icarus 204, 399–408. http://dx.doi.org/10.1016/j.icarus.2009.05.032.

Greeley, R. et al., 1993. Galileo imaging observations of lunar maria and relateddeposits. J. Geophys. Res. 98, 17183–17205. http://dx.doi.org/10.1029/93JE01000.

Grieve, R.A.F., Cintala, M.J., 1992. An analysis of differential impact melt-craterscaling and implications for the terrestrial cratering record. Meteoritics 27,526–538.

Grieve, R.A.F., Cintala, M.J., 1997. Planetary differences in impact melting. Adv.Space Res. 20, 1551–1560. http://dx.doi.org/10.1016/S0273-1177(97)00877-6.

Grieve, R.A.F., Stöffler, D., Deutsch, A., 1991. The Sudbury structure: Controversial ormisunderstood? J. Geophys. Res. 96, 22753–22764. http://dx.doi.org/10.1029/91JE02513.

Hardee, H.C., 1980. Solidification in Kilauea Iki lava lake. J. Volcanol. Geoth. Res. 7,211–223. http://dx.doi.org/10.1016/0377-0273(80)90030-X.

Hart, R.J., Cloete, M., McDonald, I., Carlson, R.W., Andreoli, M.A.G., 2002. Siderophile-rich inclusions from the Morokweng impact melt sheet, South Africa: Possiblefragments of a chondritic meteorite. Earth Planet. Sci. Lett. 198, 49–62. http://dx.doi.org/10.1016/S0012-821X(02)00497-1.

Hawke, B.R., Head, J.W., 1977. Impact melt on lunar crater rims. In: Roddy, D.J.,Pepin, R.O., Merrill, R.B. (Eds.), Impact and Explosion Cratering: Planetary andTerrestrial Implications. Pergamon Press, New York, NY, pp. 815–841.

Head, J.W., 1974. Orientale multi-ringed basin interior and implications for thepetrogenesis of lunar highland samples. Moon 11, 327–356. http://dx.doi.org/10.1007/BF00589168.

Head, J.W., 2010. Transition from complex craters to multi-ringed basins onterrestrial planetary bodies: Scale-dependent role of the expanding melt cavityand progressive interaction with the displaced zone. Geophys. Res. Lett. 37,L02203. http://dx.doi.org/10.1029/2009GL041790.

Head, J.W., 2012. Lunar Orientale Basin: Characterization and Insights into Multi-Ringed Basin Formation. Moscow Sol. Syst. P. 3.

Head, J.W., Wilson, L., 1992. Lunar mare volcanism: Stratigraphy, eruptionconditions, and the evolution of secondary crusts. Geochim. Cosmochim. Acta56, 2155–2175. http://dx.doi.org/10.1016/0016-7037(92)90183-J.

Head, J.W. et al., 1993. Lunar impact basins: New data for the western limb and farside (Orientale and South Pole-Aitken basins) from the first Galileo flyby. J.Geophys. Res. 98 (17149–17181), 1993. http://dx.doi.org/10.1029/93JE01278.

Head, J.W. et al., 2010a. The lunar Orientale basin: Structure and crustal mineralogyfrom Chandrayaan-1 Moon Mineralogy Mapper (M3) data. Lunar Planet. Sci. XLI,1030 (abstract).

Head, J.W. et al., 2010b. Global distribution of large lunar craters: Implications forresurfacing and impactor populations. Science 329, 1504–1507. http://dx.doi.org/10.1126/science.1195050.

Hess, P.C., 1994. Petrogenesis of lunar troctolites. J. Geophys. Res. 99, 19083–19093.http://dx.doi.org/10.1029/94JE01869.

Hikida, H., Wieczorek, M.A., 2007. Crustal thickness of the Moon: New constraintsfrom gravity inversions using polyhedral shape models. Icarus 192, 150–166.http://dx.doi.org/10.1016/j.icarus.2007.06.015.

Howard, K.A., Wilshire, H.G., 1973. Flows of impact melt at lunar craters. LunarPlanet. Sci. VI, 2805–2829 (abstract).

Howard, K.A., Wilshire, H.G., 1975. Flows of impact melt at lunar craters. J. Res. U.S.Geol. Surv. 3, 237–251.

Howard, K.A., Wilhelms, D.E., Scott, D.H., 1974. Lunar basin formation and highlandstratigraphy. Rev. Geophys. 12, 309–327. http://dx.doi.org/10.1029/RG012i003p00309.

Ivanov, B.A., 2005. Numerical modeling of the largest terrestrial meteorite craters.Solar Syst. Res. 39, 381–409. http://dx.doi.org/10.1007/s11208-005-0051-0.

Jaeger, J.C., 1968. Cooling and solidification of igneous rocks. In: Hess, H.H.,Poldervaart, A. (Eds.), Basalts. Interscience, New York, NY, pp. 503–536.

Khan, A., Connolly, J.A.D., Maclennan, J., Mosegaard, K., 2006. Joint inversion ofseismic and gravity data for lunar composition and thermal state. Geophys. J.Int. 168, 243–258. http://dx.doi.org/10.1111/j.1365-246X.2006.03200.x.

Kring, D.A., Abramov, O., Marchi, S., 2012. Impact melt production during the basin-forming epoch. Lunar Planet. Sci. XLIII, 1615 (abstract).

Kuskov, O.L., 1997. Constitution of the Moon: 4. Composition of the mantle fromseismic data. Phys. Earth Planet Inter. 102, 239–257. http://dx.doi.org/10.1016/S0031-9201(96)03259-1.

Marsh, B.D., 1989. Magma chambers. Annu. Rev. Earth Planet. Sci. 17, 439–474.http://dx.doi.org/10.1146/annurev.ea.17.050189.002255.

Marsh, B.D., 1996. Solidification fronts and magmatic evolution. Mineral. Mag. 60,5–40. http://dx.doi.org/10.1180/minmag.1996.060.398.03.

McCallum, I.S., Schwartz, J.M., 2001. Lunar Mg suite: Thermobarometry andpetrogenesis of parental magmas. J. Geophys. Res. 106, 27969–27983. http://dx.doi.org/10.1029/2000JE001397.

McCauley, J.F., 1977. Orientale and Caloris. Phys. Earth Planet. Inter. 15, 220–250.http://dx.doi.org/10.1016/0031-9201(77)90033-4.

Melosh, H.J., 1989. Impact Cratering: A Geologic Process. Oxford University Press,London.

Mizutani, H., Osako, M., 1974. Elastic-wave velocities and thermal diffusivities ofApollo 17 rocks and their geophysical implications. Lunar Planet. Sci. V, 2891–2901 (abstract).

Moriarty, D.P. et al., 2011. Finsen and alder: A compositional study of lunar centralpeak craters in the South Pole-Aitken basin. Lunar Planet. Sci. XLII, 2564(abstract).

Moriarty, D.P., Pieters, C.M., Isaacson, P.J., 2012. Compositional heterogeneity withinlunar central peaks. Lunar Planet. Sci. XLIII, 2399 (abstract).

Morrison, D.A., 1998. Did a thick South Pole-Aitken basin melt sheet differentiate toform cumulates? Lunar Planet Sci. XXIX, 1657 (abstract).





http://dx.doi.org/10.1016/j.pss.2011.05.010

http://dx.doi.org/10.1029/2011JE004021

http://dx.doi.org/10.1016/0019-1035(74)90003-7


http://dx.doi.org/10.1038/nature10328

http://dx.doi.org/10.1016/j.pss.2011.07.020

http://dx.doi.org/10.1029/JB090iB14p12415

http://dx.doi.org/10.1029/2010GL044666

http://dx.doi.org/10.1029/2010GL044666

http://dx.doi.org/10.1111/j.1945-5100.1998.tb01695.x

http://dx.doi.org/10.1006/icar.2002.6822

http://dx.doi.org/10.1029/2011GL048502



http://dx.doi.org/10.1029/93JE01000


http://dx.doi.org/10.1016/S0273-1177(97)00877-6



http://dx.doi.org/10.1016/0377-0273(80)90030-X

http://dx.doi.org/10.1016/S0012-821X(02)00497-1

http://dx.doi.org/10.1007/BF00589168

http://dx.doi.org/10.1029/2009GL041790

http://dx.doi.org/10.1016/0016-7037(92)90183-J


http://dx.doi.org/10.1126/science.1195050



http://dx.doi.org/10.1029/RG012i003p00309

http://dx.doi.org/10.1029/RG012i003p00309

http://dx.doi.org/10.1007/s11208-005-0051-0

http://dx.doi.org/10.1111/j.1365-246X.2006.03200.x

http://dx.doi.org/10.1016/S0031-9201(96)03259-1

http://dx.doi.org/10.1016/S0031-9201(96)03259-1

http://dx.doi.org/10.1146/annurev.ea.17.050189.002255

http://dx.doi.org/10.1180/minmag.1996.060.398.03

http://dx.doi.org/10.1029/2000JE001397

http://dx.doi.org/10.1016/0031-9201(77)90033-4


Nakamura, R. et al., 2009. Ultramafic impact melt sheet beneath the South Pole-Aitken basin on the Moon. Geophys. Res. Lett. 36, L22202. http://dx.doi.org/10.1029/2009GL040765.

Nelson, S.A., Carmichael, I.S.E., 1979. Partial molar volumes of oxide components insilicate liquids. Contrib. Mineral. Petrol. 71, 117–124. http://dx.doi.org/10.1007/BF00375427.

Nielsen, T.F.D., 2004. The shape and volume of the Skaergaard intrusion, Greenland:Implications for mass balance and bulk composition. J. Petrol. 45, 507–530.http://dx.doi.org/10.1093/petrology/egg092.

O’Connell-Cooper, C.D., Spray, J.G., 2011. Geochemistry of the impact-generatedmelt sheet at Manicouagan: Evidence for fractional crystallization. J. Geophys.Res. 116, B06204. http://dx.doi.org/10.1029/2010JB008084.

Osinski, G.R., Tornabene, L.L., Grieve, R.A.F., 2011. Impact ejecta emplacement onterrestrial planets. Earth Planet. Sci. Lett. 310, 167–181. http://dx.doi.org/10.1016/j.epsl.2011.08.012.

Pierazzo, E., Melosh, H.J., 2000. Understanding oblique impacts from experiments,observations, and modeling. Annu. Rev. Earth Planet. Sci. 28, 141–167. http://dx.doi.org/10.1146/annurev.earth.28.1.141.

Pierazzo, E., Vickery, A.M., Melosh, H.J., 1997. A reevaluation of impact meltproduction. Icarus 127, 408–423. http://dx.doi.org/10.1006/icar.1997.5713.

Pieters, C.M. et al., 1997. Mineralogy of the mafic anomaly in the South Pole-Aitkenbasin: Implications for excavation of the lunar mantle. Geophys. Res. Lett. 24,1903–1906. http://dx.doi.org/10.1029/97GL01718.

Pieters, C.M. et al., 2009. Mineralogy of the lunar crust in spatial context: Firstresults from the Moon Mineralogy Mapper (M3). Lunar Planet. Sci. XL, 2052(abstract).

Pieters, C.M. et al., 2011. Mg-spinel lithology: A new rock type on the lunarfarside. J. Geophys. Res. 116, E00G08. http://dx.doi.org/10.1029/2010JE003727.

Plescia, J.B., Cintala, M.J., 2012. Impact melt in small lunar highland craters. J.Geophys. Res. 117, E00H12. http://dx.doi.org/10.1029/2011JE003941.

Prissel, T.C. et al., 2012. Melt–wallrock reactions on the Moon: Experimentalconstraints on the formation of newly discovered Mg-spinel anorthosites. LunarPlanet. Sci. XLIII, 2743 (abstract).

Schultz, P.H., Papamarcos, S., 2010. Evolving flowfields from the Imbrium andOrientale impacts. Lunar Planet. Sci. XLI, 2480 (abstract).

Schultz, P.H., Stickle, A.M., Crawford, D.A., 2012. Effect of asteroid decapitation oncraters and basins. Lunar Planet. Sci. XLIII, 2428 (abstract).

Scott, D.H., McCauley, J.F., West, M.N., 1977. Geologic Map of the West Side of theMoon. U.S. Geol. Survey Misc. Invest. Map I-1034.

Shearer, C.K. et al., 2006. Thermal and magmatic evolution of the Moon. Rev.Mineral. Geochem. 60, 365–518. http://dx.doi.org/10.2138/rmg.2006.60.4.

Smith, D.E. et al., 2010. Initial observations from the Lunar Orbiter Laser Altimeter(LOLA). Geophys. Res. Lett. 37, L18204. http://dx.doi.org/10.1029/2010GL043751.

Spray, J.G., Thompson, L.M., 2008. Constraints on central uplift structure from theManicouagan impact crater. Meteorit. Planet. Sci. 43, 2049–2057. http://dx.doi.org/10.1111/j.1945-5100.2008.tb00660.x.

Spudis, P.D., 1993. The Geology of Multi-Ring Impact Basins: The Moon and OtherPlanets. Cambridge University Press, Cambridge.

Stöffler, D., Knoll, H.-D., Marvin, U.B., Simonds, C.H., Warren, P.H., 1980.Recommended classification and nomenclature of lunar highland rocks – Acommittee report. In: Papike, J.J., Merrill, R.B. (Eds.), Proceedings of the

Conference on the Lunar Highlands Crust. Pergamon Press, New York, NY, pp.51–70.

Thierrault, A.M., Fowler, A.D., Grieve, R.A.F., 2002. The Sudbury Igneous Complex: Adifferentiated impact melt sheet. Econ. Geol. 97, 1521–1540. http://dx.doi.org/10.2113/gsecongeo.97.7.1521.

Turcotte, D.L., Schubert, G., 2002. Geodynamics. Cambridge University Press,Cambridge.

Warren, P.H., 1993. A concise compilation of petrologic information on possiblypristine nonmare Moon rocks. Am. Mineral. 78, 360–376.

Warren, P.H., Claeys, P., Esteban, C.-P., 1996. Mega-impact melt petrology(Chicxulub, Sudbury, and the Moon): Effects of scale and other factors onpotential for fractional crystallization and development of cumulates. Geol. Soc.Am. Spec. Pap. 307, 105–124.

Whitten, J.S. et al., 2010. Characteristics, affinities and ages of volcanic depositsassociated with the Orientale basin from Chandrayaan-1 Moon MineralogyMapper (M3) data: Mare stratigraphy. Lunar Planet. Sci. XLI, 1841 (abstract).

Whitten, J.S. et al., 2011. Lunar mare deposits associated with the Orientale impactbasin: New insights into mineralogy, history, mode of emplacement, andrelation to Orientale basin evolution from Moon Mineralogy Mapper (M3) datafrom Chandrayaan-1. J. Geophys. Res. 116, E00G09. http://dx.doi.org/10.1029/2010JE003736.

Wieczorek, M.A., Phillips, R.J., 1999. Lunar multiring basins and the crateringprocess. Icarus 139, 246–259. http://dx.doi.org/10.1006/icar.1999.6102.

Wieczorek, M.A. et al., 2006. The constitution and structure of the lunar interior.Rev. Mineral. Geochem. 60, 221–364. http://dx.doi.org/10.2138/rmg.2006.60.3.

Wilhelms, D.E., Howard, K.A., Wilshire, H.G., 1979. Geologic Map of the South Poleof the Moon. U.S. Geol. Survey. Map I-1162.

Wilhelms, D.E., McCauley, J.F., Trask, N.J., 1987. The Geologic History of the Moon.U.S. Geol. Survey. Professional Paper 1348.

Williams, D.A., Fagents, S.A., Greeley, R., 2000. A reassessment of the emplacementand erosional potential of turbulent, low-viscosity lavas on the Moon. J.Geophys. Res. 105, 20189–20205. http://dx.doi.org/10.1029/1999JE001220.

Wilson, L., Head, J.W., 2011. Impact melt sheets in lunar basins: Estimatingthickness from cooling behavior. Lunar Planet. Sci. XLII, 1345 (abstract).

Winter, J.D., 2010. An Introduction to Igneous and Metamorphic Petrology. PrenticeHall, New Jersey.

Worster, M.G., Huppert, H.E., Sparks, R.S.J., 1990. Convection and crystallization inmagma cooled from above. Earth Planet. Sci. Lett. 101, 7–89. http://dx.doi.org/10.1016/0012-821X(90)90126-I.

Worster, M.G., Huppert, H.E., Sparks, R.S.J., 1993. The crystallization of lava lakes. J.Geophys. Res. 98, 15891–15901. http://dx.doi.org/10.1029/93JB01428.

Yamamoto, S. et al., 2010. Possible mantle origin of olivine around lunar impactbasins detected by SELENE. Nat. Geosci. 3, 533–536. http://dx.doi.org/10.1038/ngeo897.

Zieg, M.J., Marsh, B.D., 2002. Crystal size distributions and scaling laws in thequantification of igneous textures. J. Petrol. 43, 85–101. http://dx.doi.org/10.1093/petrology/43.1.85.

Zieg, M.J., Marsh, B.D., 2005. The Sudbury Igneous Complex: Viscous emulsiondifferentiation of a superheated impact melt sheet. Geol. Soc. Am. Bull. 117,1427–1450. http://dx.doi.org/10.1130/B25579.1.

Zuber, M.T., Smith, D.E., Lemoine, F.G., Neumann, G.A., 1994. The shape and internalstructure of the Moon from the Clementine mission. Science 266, 1839–1843.http://dx.doi.org/10.1126/science.266.5192.1839.

http://dx.doi.org/10.1029/2009GL040765

http://dx.doi.org/10.1007/BF00375427

http://dx.doi.org/10.1007/BF00375427

http://dx.doi.org/10.1093/petrology/egg092

http://dx.doi.org/10.1029/2010JB008084

http://dx.doi.org/10.1016/j.epsl.2011.08.012

http://dx.doi.org/10.1146/annurev.earth.28.1.141


http://dx.doi.org/10.1029/97GL01718

http://dx.doi.org/10.1029/2010JE003727

http://dx.doi.org/10.1029/2010JE003727

http://dx.doi.org/10.1029/2011JE003941

http://dx.doi.org/10.2138/rmg.2006.60.4

http://dx.doi.org/10.1029/2010GL043751

http://dx.doi.org/10.1029/2010GL043751

http://dx.doi.org/10.1111/j.1945-5100.2008.tb00660.x

http://dx.doi.org/10.2113/gsecongeo.97.7.1521

http://dx.doi.org/10.1029/2010JE003736

http://dx.doi.org/10.1029/2010JE003736


http://dx.doi.org/10.2138/rmg.2006.60.3

http://dx.doi.org/10.1029/1999JE001220

http://dx.doi.org/10.1016/0012-821X(90)90126-I

http://dx.doi.org/10.1029/93JB01428

http://dx.doi.org/10.1038/ngeo897

http://dx.doi.org/10.1038/ngeo897

http://dx.doi.org/10.1093/petrology/43.1.85

http://dx.doi.org/10.1130/B25579.1

http://dx.doi.org/10.1126/science.266.5192.1839

Geology and petrology of enormous volumes of …Geology and petrology of enormous volumes of impact...

Documents

Transcript of Geology and petrology of enormous volumes of …Geology and petrology of enormous volumes of impact...