Journal of Geophysical Research: Planets

Numerical modeling of the formation and structure of the Orientale impact basin

Authors

  • Ross W. K. Potter,

    1. Center for Lunar Science and Exploration, Lunar and Planetary Institute, Houston, Texas, USA
    2. NASA Lunar Science Institute, Moffett Field, California, USA
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  • David A. Kring,

    1. Center for Lunar Science and Exploration, Lunar and Planetary Institute, Houston, Texas, USA
    2. NASA Lunar Science Institute, Moffett Field, California, USA
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  • Gareth S. Collins,

    1. Department of Earth Science and Engineering, Imperial College London, London, UK
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  • Walter S. Kiefer,

    1. Center for Lunar Science and Exploration, Lunar and Planetary Institute, Houston, Texas, USA
    2. NASA Lunar Science Institute, Moffett Field, California, USA
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  • Patrick J. McGovern

    1. Center for Lunar Science and Exploration, Lunar and Planetary Institute, Houston, Texas, USA
    2. NASA Lunar Science Institute, Moffett Field, California, USA
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Corresponding author: R. W. K. Potter, Center for Lunar Science and Exploration, Lunar and Planetary Institute, 3600 Bay Area Blvd., Houston, TX 77058, USA. (potter@lpi.usra.edu)

Abstract

[1] The Orientale impact basin is the youngest and best-preserved lunar multi-ring basin and has, thus, been the focus of studies investigating basin-forming processes and final structures. A consensus about how multi-ring basins form, however, remains elusive. Here we numerically model the Orientale basin-forming impact with the aim of resolving some of the uncertainties associated with this basin. By using two thermal profiles estimating lunar conditions at the time of Orientale's formation and constraining the numerical models with crustal structures inferred from gravity data, we provide estimates for Orientale's impact energy (2–9  × 1025 J), impactor size (50–80 km diameter), transient crater size (∼320–480 km), excavation depth (40–55 km), and impact melt volume (∼106 km3). We also analyze the distribution and deformation of target material and compare our model results and Orientale observations with the Chicxulub crater to investigate similarities between these two impact structures.

1 Introduction

[2] Orientale, located on the western limb of the lunar nearside, was the last of the multi-ring basin-forming impacts, occurring 3.85 [Wilhelms, 1987] to 3.68 [Whitten et al., 2011] billion years ago. Due in part to its relatively young age, Orientale is the best preserved lunar multi-ring basin and has, therefore, been used as the archetype for investigating multi-ring basin formation and structure. Many formation theories have been suggested [e.g., Hartmann and Yale, 1968; Baldwin, 1974; Head, 1974; Hodges and Wilhelms, 1978; Melosh and McKinnon, 1978; Head, 2010]; however, a consensus has yet to be reached. Consequently, fundamental attributes of basins such as impact-generated melt volume, and basin rim and transient crater dimensions remain uncertain. In this paper, we aim to resolve uncertainties associated with Orientale's formation and structure by numerically modeling this basin-forming event. Using Orientale's gravity-inferred crustal structure as a constraint for our modeled basins, we provide estimates for a number of Orientale's attributes including impact energy and transient crater dimensions. We also analyze the distribution and deformation of target material in our models, compare these with observations of the Orientale basin, and draw parallels with the Chicxulub impact structure.

2 Geology of Orientale

[3] Orientale's basin rim is defined by the 930 km diameter Cordillera Ring (CR), the basin's most topographically prominent ring which sits ∼5 km above the basin center [Smith et al., 2010]. Three ring structures are found within the Cordillera Ring (Figure 1): the Outer Rook Ring (ORR: 620 km diameter), the Inner Rook Ring (IRR: 480 km diameter), and a 320 km diameter ring referred to here as the Inner Ring (IR). Two additional, though less certain, ring structures beyond the Cordillera Ring (diameters of 1300 km and 1900 km) have also been proposed [Pike and Spudis, 1987].

Figure 1.

Lunar Orbiter Laser Altimeter (LOLA) plan view of and cross sections through Orientale basin. Ring structures: CR (Cordillera Ring), ORR (Outer Rook Ring), IRR (Inner Rook Ring), and IR (Inner Ring). Distances on the cross sections are measured from the basin center.

[4] Gravity-inferred crustal structure beneath Orientale, as with the majority of other lunar basins [Wieczorek and Phillips, 1999; Hikida and Wieczorek, 2007], suggests a relatively thinned crust beneath the basin center which gradually thickens outward forming an annular bulge of crust. Outside this bulge, crustal thickness is relatively uniform and a fair representation of the pre-impact crustal thickness. Crustal thickness estimates beneath the center of Orientale range from 15 km [Wieczorek and Phillips, 1999] to less than 1 km [Hikida and Wieczorek, 2007]; Wieczorek et al. [2006] suggested all upper crustal material was removed from the basin center during formation. Orientale straddles the crustal dichotomy between the lunar nearside and farside with thicker crust on its western side (∼60 km) compared to its eastern side (∼40 km) [Wieczorek et al., 2013]. These crustal thicknesses, suggested from Gravity Recovery and Interior Laboratory (GRAIL) data, are 10–20 km thinner than previously thought.

[5] Of its surface materials, Orientale's Maunder Formation, which stretches out to the Outer Rook Ring, is interpreted as the impact melt sheet [Bussey and Spudis, 2000]. Recent studies estimate a melt volume on the order of 106 km3[Budney et al., 1998; Cintala and Grieve, 1998; Vaughan et al., 2013]. The volume of melt may have been great enough to undergo differentiation, though Budney et al. [1998] tentatively suggested that differentiation did not occur as their examination of crater-ejected material within the melt sheet showed no compositional variation. Beyond the Cordillera Ring lies the Hevelius Formation, which is compositionally similar to the Maunder Formation and interpreted as highly feldspathic impact ejecta [Spudis et al., 1984]. Between these two formations (primarily between the Outer Rook Ring and Cordillera Ring) lies the Montes Rook Formation which has a weak mafic signature [Head et al., 1993] suggesting a deeper crustal composition.

[6] All highlands units within Orientale appear to have an anorthositic-norite or noritic-anorthosite composition, apart from the Inner Rook Ring which appears to be composed of pure anorthosite [Hawke et al., 2003; Ohtake et al., 2009]. As well as the compositional differences [see also Cheek et al., 2012], the Inner Rook Ring is also structurally different [Nahm et al., 2013] compared to the Outer Rook Ring (and the Cordillera Ring), suggesting a different formation mechanism; the Inner Rook Ring could be analogous to a peak ring [Nahm et al., 2013]. Spectroscopic data suggest that peak rings and central peaks consist of material uplifted from deeper layers [Bussey and Spudis, 1997]. Hydrocode simulations of peak ring crater formation [e.g., Collins et al., 2002; Ivanov, 2005; Collins et al., 2008] suggest peak ring materials originate from a depth of approximately 0.1–0.15 of the transient crater diameter, slightly deeper than the maximum depth of excavation. It has, therefore, been suggested that the lunar crust is not simply composed of anorthosite overlaying norite but rather composed of an anorthositic-norite/noritic-anorthosite layer overlaying a pure anorthosite layer [Hawke et al., 2003]. No mantle signatures have been identified within Orientale basin [Yamamoto et al., 2010], suggesting mantle material was not excavated during impact.

[7] Despite the number of studies that have focused on the formation of Orientale [see Melosh, 1989, Chapter 9, and references therein], several aspects of its structure remain uncertain. For example, each ring structure up to, and including, the Cordillera Ring has been estimated as marking the extent of Orientale's transient crater [Nahm et al., 2013]. The transient crater is a fundamental feature of impact cratering; numerical models and laboratory experiments have shown it can be used to infer impact energy and momentum [Holsapple, 1982; Schmidt and Housen, 1987], impact melt volume [Cintala and Grieve, 1998], and excavation volume, with reasonable accuracy. Different transient crater sizes will, therefore, provide several estimates for those basin attributes and also affect estimates for the location of the basin rim. Uncertainty in the formation of Orientale also means uncertainty in its subsurface structure, including the nature of faulting throughout the basin [e.g., Head et al., 1993; Head, 2010; Nahm et al., 2013]. Orientale's structure can possibly be interpreted by comparison with the best-preserved large-scale terrestrial impact structure, Chicxulub; Orientale's Outer Rook Ring and Cordillera Ring have previously been compared with Chicxulub's main crater rim and outer ring, respectively [Morgan and Warner, 1999].

[8] The aim of this work is to resolve some of the uncertainties associated with Orientale's formation and structure by numerically modeling this basin-forming event. Through our modeling, we provide estimates for Orientale's impact energy, transient crater size, excavation depth, and impact melt volume. Analysis of the distribution and deformation of target material allows us to make further comparisons between our models and observations of the Orientale basin, and also allows us to compare Orientale with Chicxulub, for which there exists both hydrocode models and geophysical interpretations of the subsurface.

3 Methods

[9] We use the iSALE hydrocode [Collins et al., 2004; Wünnemann et al., 2006], one of several multimaterial, multirheology extensions of the SALE hydrocode [Amsden et al., 1980] to numerically model the Orientale basin-forming event. iSALE is well tested against other hydrocodes [Pierazzo et al., 2008] and has previously been used to model lunar basin-forming impacts [Potter et al., 2012a, 2012b].

[10] Over two dozen simulations were carried out (Table 1) using an infinite half-space target. The diameter of Orientale (930 km), relative to the radius of the Moon (∼1740 km), suggests it is likely to be near the size threshold at which modeling using a spherical target may be more appropriate; Ivanov et al. [2010] and Potter et al. [2012a] both used spherical targets when producing lunar basins >1000 km in diameter. Compared to a half-space setup, a spherical geometry setup may affect the ultimate distance that ejecta lands from the basin center (it would be further away due to the curving nature of the target surface), but this is not particularly significant in this study. We, therefore, believe a half-space target setup is suitable for Orientale but that spherical geometry should be considered for larger lunar basin modeling.

Table 1. Table of Model Run Resultsa
 ImpactImpactorCrustalCrustalTransientExcavationTransientMantle
ThermalVelocityDiameterAnnulus RadiusAnnulus DepthCrater RadiusDepthCrater VolumeMelt Volume
Profile(km/s)(km)(km)(km)(km)(km)(km3)(km3)
  1. a

    Representative best-fit impacts for each thermal profile are highlighted in bold.

1104016263122282.3 × 1067.1 × 104
1106024868158416.8 × 1067.3 × 105
11080383692014914 × 1063.0 × 106
110100479732295626 × 1065.4 × 106
110120550772576541 × 1061.2 × 107
1154019865140344.1 × 1063.6 × 105
1155025268162427.5 ×1061.1 ×106
11560293681814712 × 1062.8 × 106
11580487722275925 × 1068.5 × 106
115100564802596941 × 1061.5 × 107
1204022968154396.0 × 1069.3 × 105
12050286701764611 × 1063.2 × 106
12060328712085217 × 1066.1 × 106
12080503762476534 × 1061.6 × 107
2105013167139363.2 × 1063.6 × 102
2106015870158425.2 × 1069.5 × 103
21080211761934912 × 1062.6 × 105
210100264832195425 × 1061.1 × 106
210120308852636235 × 1063.0 × 106
2154013567140343.2 × 1067.9 × 103
2156020074182469.7 × 1064.4 × 105
21580263822335521 ×1062.4 ×106
215100326872666652 × 1066.8 × 106
215120386922937784 × 1061.6 × 107
22060233771975014 × 1061.5 × 106
22080307872536130 × 1066.4 × 106
220100369892847952 × 1061.7 × 107

[11] The target was divided into crustal and mantle layers with properties appropriate for the Moon. The thermodynamic behavior and compressibility of each material in the model is described by an equation of state (EoS). Tables generated using an Analytical EoS (ANEOS) [Thompson and Lauson, 1972] for dunite [Benz et al., 1989] were used to represent the mantle. As a suitable proxy for asteroidal material [Pierazzo et al., 1997], the ANEOS for dunite was also used to represent the impactor. A Tillotson EoS for low pressure phase gabbroic anorthosite [Ahrens and O'Keefe, 1997] was used to represent the crust. The low pressure phase was chosen over the high pressure phase, as its parameters are more suited to describing the lunar crust. The low pressure phase EoS will, however, underestimate gabbroic anorthosite temperatures in regions experiencing shock pressures between 15 GPa (the transition between low and high pressure phases) and 100 GPa (the critical shock pressure for incipient melting of the low pressure phase) [Ahrens and O'Keefe, 1997]. This will mostly affect hot crustal material around the basin center, however, most of this material is removed by the impact. Crustal material away from the basin center, outside the transient crater, will experience shock pressures ≲10 GPa and, therefore, be appropriately modeled by the low pressure gabbroic anorthosite phase.

[12] Material strength was accounted for using the strength model described by Collins et al. [2004] and the damage model described by Ivanov et al. [2010]. Strength and thermal parameters for the crust were calculated from experimental gabbro strength [Stesky et al., 1974; Shimada et al., 1983] and melting [Azmon, 1967] data. Mantle strength parameters were calculated from experimental rock strength data for dunite [Shimada et al., 1983; Ismail and Murrell, 1990] and peridotite [Stesky et al., 1974; McKenzie and Bickle, 1988] with curves for melt temperature as a function of pressure taken from Davison et al. [2010]. Based on gravity-derived crustal structures for Orientale [Wieczorek and Phillips, 1999; Hikida and Wieczorek, 2007], the modeled pre-impact crustal thickness was 60 km. As stated previously, recent GRAIL gravity observations indicate that the thickness of the lunar crust grades from ∼40 km on Orientale's eastern side to ∼60 km on its western side [Wieczorek et al., 2013]. Changing the pre-impact crustal thickness from 60 km to 40 km has been shown to have little effect on the radius of the maximum crustal thickness (the crustal annular bulge, discussed later on) and, therefore, basin size [Potter et al., 2012b]. It would, however, affect other features of the crustal structure, such as the maximum depth of crustal material in the annulus and the possible excavation of mantle material. Analysis of the GRAIL-derived crustal thickness north and south of Orientale, however, suggests the impact took place into crust closer to 60 km, rather than 40 km, thick. Our 60 km thick crust should, therefore, remain a suitable Orientale pre-impact crustal thickness estimate.

[13] Temperature gradients within the crust and mantle layers, approximating lunar conditions at the time of Orientale's formation ∼3.8 Ga, were based on temperature-depth profiles from Potter et al. [2012a] and Spohn et al. [2001] (Figure 2). Thermal profile 1 (TP1), from Potter et al. [2012a], had a crust and upper mantle temperature gradient of 10 K/km, temperatures at the mantle solidus between depths of 150–350 km, and a deep (below 800 km) mantle temperature of 1670 K. Thermal profile 2 (TP2), from Spohn et al. [2001], had a crustal gradient of 10 K/km, mantle temperatures which approached but did not reach, the solidus between depths of 300–500 km, and a deep mantle temperature of 1770 K.

Figure 2.

Pre-impact target temperatures for thermal profiles 1 and 2.

[14] Impactor diameter was varied between 40 and 120 km; however, the number of cells across the impactor radius was kept constant. A value of 20 cells was chosen (Table 2) as this provided a reasonable trade-off between resolution errors and computation time (as impact diameter decreased, cell size decreased thereby increasing computation time). Vapor production was not relevant to this study; therefore, material with a density <300 kg/m3was removed from the calculations to further expedite computation time. This threshold resulted in ∼1% of the mass of material excavated/displaced by the transient crater being removed from the calculations. Impact velocity was varied between 10 and 20 km/s. A spatially constant gravitational acceleration of 1.62 m/s2was used in all simulations.

Table 2. Resolution Test Data for a 60 km Diameter Impactor, Impacting at 15 km/s Using Thermal Profile 1
Cells PerCell SizeCrustal AnnulusTransient CraterTransient CraterMantle Melt
Projectile Radius(km)Radiusa(km)Radius (km)Volume (106 km3)Volume (106 km3)
  1. a

    The greater crustal annulus radius with 10 cells per projectile radius compared to 20 and 40 cells is probably due to the larger cell size.

562731650.821.74
1032951791.052.15
201.52931811.182.55
400.752931841.252.68

[15] Due to the axisymmetric nature of the 2-D hydrocode, all simulations were of vertical impacts. Though vertical impacts are unlikely in nature, they should provide a reasonable proxy for moderately oblique (approximately 45°) impacts. Previous numerical modeling and laboratory experiments have shown craters remain circular in shape down to impact angles of 15–20° [Gault and Wedekind, 1978; Davison et al., 2011]; numerical models have also shown that the ratio of melt volume to transient crater volume remains fairly constant between impact angles of 90–30° [Pierazzo and Melosh, 2000]. Highly oblique (<15°) impacts can alter the symmetry of ring features, reduce uplift, and concentrate ejecta and impactor material (through impactor decapitation) downrange of the impact point [Schultz et al., 2012]; some regions of Orientale suggest an impact of this nature [Schultz et al., 2012]. Nevertheless, the vertical impact models used here should provide a reasonable starting model for the Orientale basin-forming event, which can be refined by full 3-D models of oblique impacts in the future.

[16] The volume of impact-generated mantle melt was calculated using a method similar to that of Ivanov et al. [2010] and Potter et al. [2012a]. Tracer particles, initially placed in every grid cell, tracked the path followed by material that began in each cell. Throughout the simulation, the temperature and pressure at the location of the tracer was recorded. In post-processing, the pressure of the tracer at any given time was used to calculate the instantaneous melt temperature of the material tracked by the tracer using the Simon approximation [Poirier, 1991]. A mass of material was defined as completely molten if the temperature recorded by the tracer exceeded the instantaneous liquidus temperature. Material with temperatures between the solidus and liquidus was defined as partially molten, with the fraction of melt varying linearly from zero at the solidus to one at the liquidus. Total melt volumes were calculated once the basin-forming process was complete.

[17] Unfortunately, the ANEOS tables do not account for the latent heat of melting if the solid-solid phase transition is taken into account. Consequently, super-solidus mantle temperatures are overestimated resulting in anomalously large mantle melt volumes. Like Potter et al. [2012a], we include a post-processing correction for latent heat in our melt volume calculations. Latent heat is absorbed uniformly with temperature between the solidus and liquidus, as we assume a linear increase in melt production between these two reference temperatures. The effect of latent heat can, therefore, be accounted for by using an enhanced heat capacity inline imagewhen the temperature is between the solidus and liquidus, inline image [Onorato et al., 1978], where Cp is the heat capacity, L is the latent heat, TLis the liquidus, and TSis the solidus. Cpand L values of 1300 J/kg K and 700,000 J/kg, respectively, are used and assume an olivine and pyroxene-dominated mantle [Navrotsky, 1995]. Langmuir et al. [1992] and Kiefer [2003] used a similar correction approach when calculating adiabatic decompression in convecting systems. Without this correction, our melt volume predictions would be a factor of 4 larger. As stated above, this correction for latent heat is done only in post-processing when calculating melt volumes; the uncorrected ANEOS-derived temperatures are, therefore, used in the numerical simulations.

[18] As the vast majority of crustal material was removed from the basin centers in our Orientale-sized impacts, the volume of crustal melt was not included in our melt volume calculations. Although some crustal material was entrained in the mantle-dominated central melt pool, its volume was small relative to the mantle melt volume (<5% of the mantle melt volume) and, therefore, did not significantly affect our melt volume calculations. Crustal melt volume will, however, be important for smaller lunar basin-scale impacts which remove less crustal material from the basin center and, therefore, have a significant crustal layer beneath their basin centers. Additionally, in our models, crustal material outside of the transient crater did not experience peak shock pressures in excess of 100 GPa, the critical shock pressure for incipient melting of the analogous low pressure gabbroic anorthosite phase, and so would not affect the total impact-generated melt volume.

[19] Impact basins and complex craters contain interior topographic structures such as peak rings and/or central peaks that protrude from an otherwise flat and shallow crater floor. These structures are thought to form via the gravitational collapse of a much deeper, bowl-shaped transient crater. Numerical models of crater formation have shown that some material weakening mechanism is required to facilitate the necessary collapse of the crater. One such proposed mechanism is acoustic fluidization [Melosh, 1979; Melosh and Ivanov, 1999]. Here acoustic fluidization is implemented via the block model [Melosh and Ivanov, 1999; Wünnemann and Ivanov, 2003]. Scaling relations from Wünnemann and Ivanov [2003] were used to choose block model parameters for lunar craters and constrained by the subsurface structure of the Chicxulub impact structure [Collins et al., 2008].

[20] At low strain rates, geological materials lose all resistance to shear upon melting [e.g., Jaeger and Cook, 1969]. In iSALE, shear strength is a function of the ANEOS-derived temperature and decreases to zero at the solidus. Shear strength is, therefore, not affected by our latent heat correction, which only affects the temperature in regions that are above the solidus. The rheology and behavior of partially molten material under the high strain rate conditions experienced during impacts, however, is likely to be far more complex [Stewart, 2011]. Therefore, as a first approximation, we applied a partial melt viscosity of 1010 Pa s to super-solidus material [see Potter, 2012], providing this material with some resistance to shear as super-solidus material will be a mixture of melt and solid clasts and not an inviscid fluid.

[21] The basin-forming process, particularly basin collapse, is, however, sensitive to the chosen partial melt viscosity. Test simulations [Potter, 2012] showed that low partial melt viscosities (no viscosity to 105 Pa s) resulted in secondary central uplifts reaching heights in excess of 600 km; higher partial melt viscosities (1012 Pa s) failed to produce central uplifts—the transient crater floor slowly rose back to a flat surface. A partial melt viscosity of 1010 Pa s, however, produced a basin-forming event similar to that modeled for the Chicxulub impact [Collins et al., 2008]. Additionally, our chosen value (1010 Pa s) produced basins with features similar to those observed and inferred at Orientale, suggesting it is a reasonable first approximation. A list of model input parameters can be found in Table 3.

Table 3. Crust and Mantle Input Parameters
ParameterMantleCrust
Equation of StateANEOSTillotson
  1. a

    Ahrens and O'Keefe [1997].

  2. b

    see Collins et al. [2004], Ivanov et al. [2010], and Poirier [1991].

  3. c

    see equations (A2) and (A3) in Collins et al. [2004].

  4. d

    see Table A2 in Ivanov et al. [2010].

  5. e

    see equations (11) and (12) in Wünnemann and Ivanov [2003].

Tillotson (low pressure phase)a  
  a-0.5
  b-0.145
  A (GPa)-71
  B (GPa)-75
  E0 (MJ/kg)-489
Thermalb  
  Solidus (at zero pressure) (K)13731513
  Specific heat capacity (J/kg K)1300590
  Thermal softening parameter1.11.2
  Constant in Simon1.52 × 1091.84 × 109
   approximation (Pa)  
  Exponent in Simon4.057.27
   approximation  
Strengthc  
  Poisson ratio0.250.25
  Cohesion of material1 × 1051 × 105
   (damaged rock) (Pa)  
  Coefficient of internal0.630.71
   friction (damaged)  
  Limiting strength at high3.26 × 1092.47 × 109
   pressure (damaged) (Pa)  
  Cohesion of material5.07 × 1063.19 × 107
   (intact rock) (Pa)  
  Coefficient of internal friction1.581.1
   (intact)  
  Limiting strength at high3.26 × 1092.47 × 109
   pressure (intact) (Pa)  
Damaged  
  Failure strain at zero pressure1 × 10−41 × 10−4
  Increase in failure strain with1 × 10−111 × 10−11
   pressure (1/Pa)  
  Pressure above which failure is3 × 1083 × 108
   compressional (Pa)
Acoustic fluidizatione(TP1 best-fit  
   impactor—50 km diameter)  
  Acoustic vibration decay time (s)10001000
  Kinematic viscosity (acoustically1.25 × 1051.25 × 105
   fluidized region) (m2/s)  
Acoustic fluidizatione(TP2 best-fit  
   impactor—80 km diameter)  
  Acoustic vibration decay time (s)16001600
  Kinematic viscosity (acoustically2 × 1052 × 105
   fluidized region) (m2/s)  

4 Results and Discussion

4.1 Basin Formation

4.1.1 The Formation Process

[22] Figure 3 illustrates the basin-forming process, as predicted by the iSALE hydrocode, for two Orientale-sized impact events into targets with different initial thermal profiles. In the first few minutes after impact, the transient crater grows (Figure 3a) as material is displaced from the impact site. Some material stays within the transient crater and some material is ejected and draped over the surrounding landscape. The transient crater floor is lined by a thin layer of (melted) crust, and the ejecta curtain contains only crustal material. The transient crater floor begins to rise, just prior to the crater walls collapsing under the influence of gravity, eventually forming a central uplift (Figures 3b–3c). This uplift also collapses (Figures 3d–3e); depending on impact energy and target strength, additional uplift and collapse phases involving central basin material may occur. Within a couple of hours of initial impact, the impact energy has sufficiently attenuated, and the process is complete (Figure 3f).

Figure 3.

The basin-forming process for two Orientale-sized impacts into (left) TP1 and (right) TP2. Each impactor has a different energy (TP1 impactor: 50 km diameter at 15 km/s; TP2 impactor: 80 km diameter at 15 km/s). rtc is the transient crater radius; rca is the crustal annular bulge radius. See text for discussion.

[23] The two models in Figure 3 produce basins with similar crustal annulus radius sizes (discussed presently), despite their impact energies differing by a factor of ∼4. The thermal state of the target pre-impact, therefore, has a significant effect on the basin-forming process, especially crater collapse, as shown in Figure 3 and previously in Potter et al. [2012b]. Pre-impact, TP1 is hotter than TP2 at depths of 60–300 km and, therefore, weaker; the weaker target requires a lower impact energy to produce a basin of a given size.

[24] Additionally, these preheated targets allow ductile material flow in the upper mantle during basin formation by virtue of thermal softening (i.e., hotter materials are weaker and more susceptible to permanent deformation). At depths equivalent to the transient crater floor (100–200 km), the target material is already at, or near, the solidus and has little-to-no shear strength. Hence, a dynamic weakening mechanism, such as acoustic fluidization, is not necessary to enable uplift of the crater floor in Orientale-scale impact simulations. Acoustic fluidization is, however, important for facilitating collapse around the transient crater walls (see Videos S1 and S2 in the supporting information). This is because crust in this location is relatively cold and strong (Figure 3a) and not heated sufficiently by the impact. Without acoustic fluidization (see Videos S3 and S4), collapse of the transient crater wall still occurs, in response to uplift of the thermally softened crater floor, but the intricacies of crustal deformation differ because it is stronger. As a result, the final form of the basin (i.e., crustal structure, topography, and crustal annulus radius) in a simulation with acoustic fluidization is different to the final form in a simulation without. As our simulations with acoustic fluidization better reproduce the inferred structure of Orientale (see later sections), weakening of the crust (by acoustic fluidization or a weakening mechanism other than thermal softening) appears to be required in Orientale-scale basin formation.

[25] The models illustrated in Figure 3 represent two of the best-fit numerical models to Orientale's gravity-derived crustal structure. As discussed in greater detail below, these models produce a central zone of thinned crust surrounded by an annulus of thickened crust which resemble the actual crustal structure at Orientale [Wieczorek and Phillips, 1999; Hikida and Wieczorek, 2007]. As in Potter et al. [2012b], the dimensions of the crustal annulus, defined as the radial distance from the basin center to the local maximum crustal thickness, are used here as strong model constraints because the crustal annulus is a characteristic feature of the majority of lunar basins and is easily resolved in both gravity and present hydrocode models.

[26] The thickening of the crust forming the annular bulge in our models is caused by a combination of two factors: (a) the deposition of proximal, overturned ejecta on top of the transient crater rim and (b) the collapse of the transient crater. The latter is a complex process involving both lateral and vertical movement of the crust between the transient and final crater rim. Initially, during collapse, crustal motion is inward and downward, which thins the crust but also drops the base of the crust down. Inner crustal material is then lifted vertically, as the crater floor rises, forming a collar around the central uplift. With the collapse of the uplift, this crustal collar is thrust outward into and over cooler, stronger crustal material.

[27] Figure 4 plots the crustal annulus radius against impactor radius for all modeled impact scenarios into the two thermal profiles. The two representative best-fit impacts are highlighted by solid inverted triangles. The crustal profiles of these two best-fit models are compared to the gravity-inferred Orientale crustal annular bulge radius, rca, and depth, zca, in Figure 5. The best-fit impactor using TP1 has an energy of 2.4  × 1025 J (50 km diameter impactor; 15 km/s impact velocity); for TP2, the best-fit impactor has an energy of 9.9  × 1025 J (80 km diameter impactor; 15 km/s impact velocity). This latter estimate agrees well with preliminary models of Orientale impacts conducted by Stewart [2011] who used the same target thermal profile [Spohn et al., 2001] but a different hydrocode (CTH). Our different energy estimates for the Orientale impact, therefore, highlight the importance of the target's thermal state in the basin-forming process.

Figure 4.

Crustal annulus radius against impactor radius for all modeled TP1 (red icons) and TP2 (blue icons) impacts. Models are also subdivided by impact velocity (10 km/s: triangles, 15 km/s: inverted triangles, and 20 km/s: diamonds). The best-fit impact for each thermal profile is highlighted by a solid color icon. The model results show crustal annulus radius increases with increasing impactor radius and impact velocity. Orientale rca is Orientale's observed crustal annulus radius.

Figure 5.

Representative best-fit numerical models to the gravity-inferred radius, rca, and depth, dca, of Orientale's crustal annular bulge [Wieczorek and Phillips, 1999]. Thermal profile 1: 50 km diameter, 15 km/s impact velocity; thermal profile 2: 80 km diameter, 15 km/s impact velocity. Locations for Orientale's observed rings structures are also shown.

[28] The crustal thickness model from Wieczorek and Phillips [1999] that is used in Figure 5 assumes that the entire gravity anomaly in the basin is due to relief on the crust-mantle interface. An alternative gravity inversion includes an additional 8 km of mantle uplift at the center of the basin [Hikida and Wieczorek, 2007]. Both of these models are strongly super-isostatic, with the crust-mantle interface in the basin center occurring at much shallower depth than if the system were isostatic. Our numerical models show, however, that the central part of the basin is at least partially molten to a depth of several hundred kilometers (Figure 10) and has no long-term strength. The crust-mantle interface should, therefore, relax to its isostatic position in a geologically short period of time. A variety of post-impact processes may make important contributions to the gravity field, such as volcanic intrusions, uplift driven by the subisostatic crustal annulus, or collapse of the original basin structure [Kiefer et al., 2012; Andrews-Hanna, 2013; Dombard et al., 2013]. If this is the case, then the Wieczorek and Phillips [1999] crustal thickness profile in the region out to a radial distance of ∼200 km, where, in our models, the post-impact temperature is very hot and the viscosity is very low, cannot be used to constrain a preferred numerical model. Orientale, like many lunar impact basins, has a subisostatic gravity anomaly at larger distances from the basin center [Neumann et al., 1996; Hikida and Wieczorek, 2007; Andrews-Hanna, 2013; Andrews-Hanna et al., 2013]; our numerical models commonly produce an annulus of thickened crust surrounding the impact basin [Potter et al., 2012b].

[29] At a radial distance of ∼200–300 km from the basin center, there is no impact-induced heating of the lower crust and mantle; so if the impact created a crustal annulus, it might not experience subsequent viscous relaxation. Figure 5 suggests that impact parameters intermediate between our best TP1 and TP2 models may explain the location and amplitude of the Orientale crustal annulus.

[30] The predicted basin topography of the two best-fit impact models at the conclusion of the impact is also shown in Figure 5. In the center of the basin, the near-surface material is mostly molten (Figure 10), and the topography in this region will likely be modified by post-impact viscoelastic flow [Kiefer et al., 2012; Andrews-Hanna, 2013; Freed et al., 2013]. At greater distances from the basin center, the post-impact crustal temperature is lower, and the topography may suffer less post-impact deformation. The radial location of the topographic high for the TP1 best-fit impact, 445 km (with a high point of 3.1 km), is similar to that of the observed topographic high, the CR, at a radial distance of 465 km. For the TP2 best-fit impact, the topographic high appears well beyond the location of the CR at a radial distance of 631 km (with a high point of 7.5 km). The post-impact surface crosses the preimpact surface in our best-fit models at radial distances of 160 km for the TP1 impact, which is similar to the IR radius, and 536 km for the TP2 impact, which is beyond the location of the CR. Observed Orientale topography suggests the basin topography crosses an elevation of 0 km between the IRR and the IR at a distance of ∼240 km. Topography is, however, not a particularly robust measurement in our models, partly due to cell size and partly due to the hot, molten nature of material around the basin center (100–200 km radius); so we do not use it as a primary constraint for comparing our numerical models to Orientale observations.

4.1.2 Material Deformation

[31] A natural consequence of the basin-forming process is the deformation of target material. Due to the resolution of these iSALE models (inline image1 km), deformation features, such as faults, on a decimeter-scale to hectometer-scale, cannot be accurately modeled. However, large-scale bulk deformation can be determined. Figure 6 uses color to illustrate how accumulated plastic strain, an invariant measure of permanent deformation (a simple conceptual interpretation of a strain of one is a square block whose topside is sheared by a distance equal to its length; see Collins et al. [2004] for more details), evolves during the basin-forming process for the best-fit impacts. Overlaid on the strain plot is a grid of Lagrangian tracer particles; rows and columns of the grid are spaced 10 km apart. The thicker solid black line represents the crust/mantle boundary. The tracers track the progress of material through the basin-forming process and highlight: the initiation of stratigraphic overturn about the fold hinge as the transient cavity expands (Figure 6a); deformation within the upper hundred kilometers of the target as the crater floor rises (Figure 6b); a secondary overturn event as more ductile rocks in the over-heightened central uplift collapse outward interacting with more brittle rocks (Figures 6c–6d); and subsidence and inward displacement of the thickened crust (Figures 6e–6f). Videos of the impacts highlighted in Figure 6 can be found in the supporting information (Videos S5 and S6). Material out to distances equivalent to Orientale's ORR is affected by the two overturn events.

Figure 6.

Accumulated plastic strain distribution within the crust and upper mantle for each thermal profile's best-fit impact with an overlay of the Lagrangian tracer grid during the basin-forming process. Arrows highlight motion direction. The observed locations of Orientale's ring structures are also plotted: CR (Cordillera Ring), ORR (Outer Rook Ring), IRR (Inner Rook Ring), and IR (Inner Ring).

[32] The initial thermal states of the targets also affect the distribution of strain. Thermal profile 1 intersects the mantle solidus at ∼120 km (Figure 2), resulting in an abrupt rheological change at this depth with a cool, brittle crust/mantle lithosphere above and a hot, ductile partially molten mantle asthenosphere beneath. During excavation, this rheological layering leads to outward plastic flow in the asthenosphere (high plastic strain below 120 km depth) and elastic upwarping of the lithosphere above (low plastic strain above 120 km depth). During the subsequent uplift and overshoot of the crater floor and inward collapse of the crater rim, the asthenospheric flow reverses, leading to permanent inward displacement of the mantle beneath 120 km (Figures 6b–6c). At the same time, the overlying, predominantly undamaged, lithosphere bends downward, and subtle zones of localized strain propagate down from the surface, which might indicate extensional faulting (discussed presently). Thermal profile 2 produces a similar, though less dramatic rheological contrast, but at a much greater depth so that it has no influence on crater modification. Nevertheless, in Video S6 at time 13–19 min, when the rate of central uplift and inward rim collapse is greatest, localized strain is observed at a radial distance of ∼700 km. If this feature is real and not an instability in the model, then it would be indicative of external normal faulting (see also Figure 7). It could also be analogous to the suggested 1300 km diameter outer ring of Orientale [Pike and Spudis, 1987]. The features seen in our models are similar to those described in the ring tectonic theory for multi-ring basin formation [Melosh and McKinnon, 1978; Melosh, 1989], though ring tectonics relies on the inward asthenospheric flow dragging the overlying crust inward, rather than allowing it to bend down.

Figure 7.

Post basin-formation accumulated plastic strain distribution within the crust and upper mantle for each thermal profile's best-fit impact. Overlaid on the strain plots is the Lagrangian tracer grid. Tracers are spaced every 10 km in Figures 7a and 7c, and every 6 km in Figures 7b and 7d. The boxed regions in Figures 7a and 7c are expanded in Figures 7b and 7d to highlight possible zones of faulting. Fault location and motion is inferred from offsets in the tracer grid and highlighted by yellow lines and arrows. Downward arrows indicate possible normal faulting; upward arrows indicate possible thrust faulting. The observed locations of Orientale's ring structures are also plotted: CR (Cordillera Ring), ORR (Outer Rook Ring), IRR (Inner Rook Ring), and IR (Inner Ring). rtc shows the transient crater radius for each best-fit impact.

[33] Strain and deformation within the basins, post-impact, is illustrated in greater detail in Figure 7. Strain is greatest around the basin center, as would be expected, and is accommodated by incompressible ductile flow; this distribution is similar to modeled deformation within the Chicxulub impact structure [Collins et al., 2008]. Strain decays outward from the basin center but remains high near the surface to radial distances beyond 800 km and within the crustal annular bulge. The former is due to both the shock/free-surface interaction and the deposition of high-energy ejecta. High strain within the cooler, brittle rocks of the crustal annular bulge implies extensive communition and fracturing, which will create porosity and reduce the density of the rocks. Hence, the annular bulge is likely to be lower in density than the crust outside the basin. Consistent with this result, GRAIL gravity observations show an annulus of low bulk density surrounding the Orientale basin, which has been interpreted as a region of higher than average porosity [Wieczorek et al., 2013].

[34] Offsets in the tracer grid in Figure 7 can be used to infer possible zones of faulting within the target; these are highlighted by the yellow lines in Figures 7b, 7c, and 7d. Motion on the inferred faults is indicated by the yellow arrows: downward arrows indicate normal faulting, and upward arrows indicate thrust faulting. Though target material may lose cohesion during the basin-forming process, the final structure would contain attributes mappable as faults. In the best-fit TP1 impact, the sense of deformation underneath the observed IR location is consistent with thrust faulting (Figure 7b). Farther out, tracers are downthrown toward the basin center, consistent with normal faulting, at the ORR location. In the best-fit TP2 impact, deformation patterns are consistent with thrusting beneath the IR and IRR locations and normal faulting beneath the CR (Figure 7d). Deformation consistent with thrust faulting at locations equivalent to the IRR and normal faulting at the ORR and CR in our models agrees with the interpretations of Nahm et al. [2013]. Neither best-fit model, however, shows evidence of faulting at each ring structure. This could be due to the models not being able to resolve smaller-scale deformation. Alternatively, this could imply additional faulting occurred after the main basin-forming process.

[35] Localization of strain is inherently sensitive to numerical instability, and additional higher-resolution simulations are required to verify the robustness of these results. The observed patterns of strain are, however, suggestive of numerical models being close to simulating multiple-ring features.

4.1.3 The Transient Crater

[36] As stated previously, the transient crater, though an ephemeral feature, is an important part of the cratering process, as it can be used to infer impact energy and momentum, impact melt volume, and the maximum depth and volume of excavated material. The transient crater (approximately) marks the division between the stages of excavation and modification. It is often defined by its diameter; past studies of Orientale have provided a broad range of diameter estimates, from 620 km, the ORR [e.g., Head, 1974; Head et al., 1993; Fassett et al., 2011], to <400 km [Melosh, 1989], well within the IRR. That range in transient crater diameters has produced an unsatisfactory range in calculated impact energy and other parameters for Orientale. We also note that in basin-forming events, the distinction between the excavation and modification stages is blurred; hydrocode simulations [e.g., Turtle et al., 2005] indicate modification of the basin floor (uplift) begins before the maximum transient crater diameter is reached.

[37] In this work, the transient crater was defined as forming once the expanding transient cavity reached its maximum volume (similar to the approach of Elbeshausen et al. [2009]) and measured at the preimpact surface. Figure 8 plots the transient crater sizes, against impactor radius, for all our modeled impact scenarios. For the best-fit models, the transient craters had diameters of 324 km (TP1) and 466 km (TP2). The former is similar in size to Orientale's IR (320 km); the latter is similar to Orientale's IRR (480 km). These estimates imply the ORR and CR are unsuitable transient crater approximations—they are too large. Nahm et al. [2013] also believe the transient crater must be interior to the ORR as this ring (as with the CR) was formed by large-scale normal faulting; we have already pointed out deformation patterns consistent with normal faulting at radial distances equivalent to the ORR and CR in our best-fit models. The two best-fit estimates also place the transient crater within the IRR. Our models suggest the IRR was formed by the collapse of the central peak, with some contribution from the collapsing transient crater rim (Figures 3 and 6), agreeing with Nahm et al. [2013]. The impact using TP1 produced a transient crater similar in size to the IR, which has also been suggested to have formed as a result of thermal contraction and subsidence of the impact melt sheet and substrate [Bratt et al., 1985; Head et al., 1993].

Figure 8.

Transient crater radius against impactor radius for all modeled TP1 (red icons) and TP2 (blue icons) impacts. Models are also subdivided by impact velocity (10 km/s: triangles, 15 km/s: inverted triangles, and 20 km/s: diamonds). The best-fit impact for each thermal profile is highlighted by a solid color icon. The radius of Orientale's Outer Rook Ring (ORR), Inner Rook Ring (IRR), and Inner Ring (IR) are plotted for reference. The model results show transient crater radius increases with increasing impactor radius and impact velocity.

4.1.4 Material Excavation

[38] In this work, the maximum depth of excavation was defined as the deepest preimpact material that was above the target surface when the transient cavity was as its greatest volume (i.e., the transient crater). Figure 9 plots excavation depth against impactor radius for all our models. In the best-fit models, mantle material was not excavated; excavation depths were within the lower crust: 42 km for the impact into TP1 and 55 km for the impact into TP2. These results agree with spectroscopic data sets [Yamamoto et al., 2010] which suggest no mantle material was excavated during Orientale's formation as no mantle-like signatures have been found within the basin. Mantle-like signatures have, however, been found within basins both smaller (e.g., Crisium) and larger (e.g., Serenitatis) than Orientale. The lack of mantle-like signatures around Orientale is probably due to the regional crustal thickness. This is estimated to be greater at Orientale (∼40–60 km) [Wieczorek et al., 2013] than many other basins, including Crisium and Serenitatis (both ≲40 km) [Wieczorek et al., 2013], and would therefore require more energetic impacts to excavate mantle material.

Figure 9.

Excavation depth against impactor radius for all modeled TP1 (red icons) and TP2 (blue icons) impacts. Models are also subdivided by impact velocity (10 km/s: triangles, 15 km/s: inverted triangles, and 20 km/s: diamonds). Crustal thickness was 60 km for all models based on gravity-inferred crustal thickness around Orientale. The best-fit impact for each thermal profile is highlighted by a solid color icon. The model results show excavation depth increases with increasing impactor radius and impact velocity.

4.2 Basin Structure

4.2.1 Crustal Distribution

[39] Figure 10 illustrates the distribution of crustal and mantle (discussed later) material following basin formation for the two best-fit models. Crustal material is divided to highlight the distribution of crust originally at different depths: upper (0–20 km), middle (20–40 km), and lower (40–60 km). This segregation also highlights the crustal overturn with originally deeper-seated crustal material overlying shallow crustal material within the crustal annular bulge. On the surface at radial distances equivalent to the CR, a thin (few kilometers) layer of midcrustal material overlays upper crust in the TP1 impact; in the TP2 impact, lower crust is present overlaying a thin mixed unit of upper, middle, and lower crust. At radial distances equivalent to the ORR, for the TP1 impact, a thin (few kilometer) overturned lower-crustal layer overlays midcrustal material; for TP2, material is from the lower crust. At radial distances equivalent to the IRR, the basin floor was composed of, again, lower-crustal material in the TP1 best-fit, while the basin floor for the TP2 best-fit was composed of a thin crustal layer (∼1 km thick) and partially molten mantle melt. At radial distances equivalent to the IR, the basin floor was composed of a mixture of a (thin) crustal layer and partially molten mantle (melt fraction >0.5) in the TP1 best-fit model; the TP2 best-fit model had a similar composition. Inside of the IR, a thin (few kilometers) crust overlaid (partially) molten mantle in both best-fit impacts, though this layer was discontinuous in the TP2 best-fit impact.

Figure 10.

Crust and mantle distribution for each thermal profile's best-fit impact. Crustal material is subdivided into material originally at depths of 0–20 km, 20–40 km, and 40–60 km. Mantle material is subdivided into solid, unmolten material (gray) and (partially) molten material. Partially molten material is represented by the melt fraction ranging from >0 (the mantle solidus) to 1 (the mantle liquidus). The observed locations of Orientale's ring structures are also plotted: CR (Cordillera Ring), ORR (Outer Rook Ring), IRR (Inner Rook Ring), and IR (Inner Ring).

[40] Observations of Orientale suggest the Hevelius Formation, beyond the CR, is highly feldspathic impact ejecta [Spudis et al., 1984]. If this material belongs to the upper crust, then our best-fit model using TP1 matches this observation well as beyond the location of the CR only upper-crustal material is present on the surface, some of which is ejecta from the basin center. For the TP2 impact, beyond the location of the CR, surface materials gradually grade from lower crust to upper crust; upper-crustal materials are not continuous on the surface until a radial distance of ∼700 km.

[41] The Montes Rook Formation, primarily between the ORR and CR has a weak mafic signature [Head et al., 1993] suggesting a relatively deeper crustal composition. Our best-fit models are consistent with this interpretation, as surface materials are generally midcrustal in the TP1 impact and more lower-crustal in the TP2 impact.

[42] Hawke et al. [2003] suggested, based on the composition of the IRR, that a pure anorthosite layer was overlain by a mixed anorthositic-norite/noritic-anorthosite layer within the lunar crust. In our models, surface crustal material at locations equivalent to the IRR originally came from the lower (>40 km) crust. If our model, and the suggestion of Hawke et al. [2003], is correct, then this pure anorthosite layer could be deep (>40 km) within the crust at Orientale.

[43] It should be noted, however, that comparison of spectral data with numerical model data is nontrivial, as the former detects composition in the upper few meters of the surface, whereas (our) numerical models resolve features inline image1 km. Additionally, material was removed from our simulations if its density was <300 kg/m3. Though the volume removed was small (∼1% of the mass of material excavated/displaced by the transient crater), this material would not be accounted for in the final basin structure.

[44] Our numerical models also use a simplified, pristine preimpact target composition. The stratigraphy of the Orientale site preimpact is likely to have been complex and heavily affected by ejecta from the older, far larger, and relatively nearby, South Pole-Aitken basin impact. This complex preimpact stratigraphy could explain why a pure anorthosite layer is possibly overlaid by a mixed anorthositic-norite/noritic-anorthosite layer within the lunar crust.

4.2.2 Mantle Distribution and Melt Volume

[45] Our best-fit models, unlike those derived from gravity data, produce a large centralized body of mantle material (Figure 10). Some crustal material is still present at the basin surface and entrained within the mantle melt, but relative to the mantle volume, this volume is minor (<5%). This centralized body of mantle is (partially) molten, and its presence agrees with previous modeling of basin-scale impacts [Ivanov et al., 2010; Stewart, 2011; Potter et al., 2012a].

[46] The distribution of mantle material, both unmolten (gray) and partially molten (whites, blues, and pinks), is illustrated in Figure 10. Melt fraction is used to quantify the relative volume of melt and varies linearly from zero at the mantle solidus to one at the mantle liquidus (completely molten). The melt pools for both best-fit impacts extend to a depth of ∼300 km and radial distances of 150 km (TP1) and 280 km (TP2). However, very little mantle material, less than 1% of the total melt volume for each best-fit impact, is completely molten (melt fraction  = 1). Material that is completely molten is found discontinously at, or just below, the lunar surface out to distances of 120 km and 250 km for the impacts into TP1 and TP2, respectively. Material with a melt fraction >0.5 was found to depths of 20 km (TP1) and 70 km (TP2) and out to radial distances of 50 km (TP1) and 150 km (TP2). Vaughan et al. [2013], who suggested a 175 km radius central topographic depression within Orientale represents the solidified impact melt pool, estimated the melt pool was 12.5–16 km deep at its center, assuming all subsidence in the central depression was due to melt cooling and solidification. The radial extent of this central topographic depression compares favorably with the radial extent of mantle in the TP1 impact (which is also similar to the IR radius).

[47] Although our models imply mantle-dominated impact melt compositions, no spectral evidence of ultramafic compositions has thus far been found within lunar basins [Yamamoto et al., 2010]. As described above, both models in Figure 10 include substantial volumes of mantle material with large melt fractions. The TP1 best-fit model has >0.50 melt fraction to a depth of 20 km, and the TP2 best-fit impact has >0.75 melt fraction to a depth of 50 km. These melt layers are sufficiently thick and would have a low enough viscosity that significant differentiation of these melt zones may occur prior to solidification. Morrison[1998] and more recently, Vaughan et al. [2013] and Hurwitz and Kring [2013] have suggested that differentiation of such a melt sheet could produce a composition similar to crustal material and therefore mask any mantle-like signatures. The numerical model results in this paper can be used as the initial condition for a melt sheet differentiation model, but such models are beyond the scope of the current paper.

[48] Budney et al. [1998] used a melt differentiation model and compared iron and titanium signatures to those seen in crater ejecta within Orientale suggesting melting was confined to the upper crust and that differentiation of the Orientale melt pool did not take place. They did not, however, examine ejecta from craters around the center of Orientale (the proposed melt pool zone) and, therefore, ignored the 55 km diameter complex crater Maunder, which may have uplifted differentiated material. Recent spectral observations of Maunder, using M-cubed data, suggest its walls and central peak are noritic [Whitten et al., 2011]. It is not known whether this represents original lower crust material or differentiated impact melt material.

[49] Figure 11 plots the total impact-generated mantle melt volume against impactor radius for our suite of impact scenarios. The total mantle melt volumes for the best-fit impacts into TP1 and TP2 are 1.1  × 106 km3and 2.4  × 106 km3, respectively. The larger melt volume for the impact into TP2 is due to the impactor's total energy, which is 4 times greater than the TP1 impactor, and so more than compensates for the somewhat cooler initial state in the 100–300 km depth range in TP2. These values for the total impact-generated melt volume compare favorably with other Orientale estimates: equation (8) of Cintala and Grieve [1998] predicts, for impact velocities of 10–20 km/s, melt volumes of 2.9–3.8  × 106 km3, assuming a chondrite impactor (density: 3580 kg/m3), anorthosite target (density: 2734 kg/m3) and, using equation (9) of Croft [1985], a 509 km transient crater diameter. Vaughan et al. [2013] calculated a melt volume of 106 km3by assuming a 175 km radius central topographic depression within Orientale represents the solidified impact melt pool and that all subsidence within this central depression was due to cooling and solidification of the melt. Budney et al. [1998], who modeled the composition and size of the melt sheet, also suggested a melt volume of 106 km3. An alternative melt volume estimate can be found using the analytical equations in Abramov et al. [2012]. Their equations also allow for targets with a thermal gradient to be considered. Using a near-surface thermal gradient equivalent to that in this study (10 K/km), the Abramov et al. [2012] equations predict an Orientale melt volume of 7.5  × 105 km3(using our best-fit impactor energy for TP1) and 3.0  × 106 km3(using our best-fit impactor energy for TP2).

Figure 11.

Impact-generated mantle melt volume (Vm) against impactor radius (ri) for all modeled TP1 (red icons) and TP2 (blue icons) impacts. Models are also subdivided by impact velocity (10 km/s: triangles, 15 km/s: inverted triangles, and 20 km/s: diamonds). The individual best-fit impact for each thermal profile is highlighted by a solid color icon. Fits to specific data sets are also shown—TP1 20 km/s: inline image; TP2 10 km/s: inline image.

4.2.3 Comparisons: Orientale Best-Fit Models and Chicxulub

[50] Chicxulub is the Earth's best preserved large-scale impact structure, making it the ideal terrestrial analog for large-scale impact structures on silicate bodies, such as the Moon. Chicxulub has a crater rim diameter of 180 km [Hildebrand et al., 1991] defined, in part, by a topographically high, normally faulted scarp which represents the head of a terraced modification zone caused by the collapse of a ∼100 km diameter transient crater [Kring, 1995; Morgan et al., 1997]. The current working hypothesis is that during the cratering event, a central uplift rose from the transient crater floor and subsequently collapsed outward, thrusting over the inwardly collapsing transient crater wall, to produce a ∼90 km diameter peak ring [Collins et al., 2002]. The crust beneath Chicxulub's crater center is thinned by ∼1–2 km [Morgan et al., 2000; Christeson et al., 2009] and bound by a slight (∼0.5–1 km thicker) [Christeson et al., 2009] crustal annular bulge with a diameter approximately equal to that of the peak ring. Beyond the crater rim, there is a ∼250 km diameter ring defined by a thrust fault.

[51] The same process that is hypothesized to have produced the peak ring of Chicxulub is evident in both of the TP1 and TP2 best-fit models, producing deformation consistent with outward thrust faulting at locations equivalent to Orientale's IR and IRR, respectively. That type of origin is consistent with the massif-style morphological characteristics of the IRR and lack of normal faulting at that ring [Nahm et al., 2013]. According to numerical simulations, peak rings should be composed of lithologies uplifted from approximately 0.1–0.15 times the transient crater diameter. In the case of Chicxulub, it is expected that granitic crust was uplifted and thrust over carbonate platform sedimentary rocks. In our best-fit model for TP1, we find lower-to-middle crust at the locations of the IR and IRR and mostly mantle material at these locations in the TP2 models. Because anorthosite is seen in the IRR [Hawke et al., 2003; Ohtake et al., 2009; Cheek et al., 2012], the TP1 models may be a better fit. The IR region has a distinctly different composition [e.g., Head et al., 1993; Cheek et al., 2012], implying a slightly different origin. Bratt et al. [1985] suggested the IR could have formed due to thermal contraction and subsidence from the Orientale melt sheet (in our best-fit models, the IR is within the melt sheet in TP2 and on its edge in TP1).

[52] An important difference between the models of Orientale basin formation presented here and previous simulations of Chicxulub [e.g., Ivanov, 2005; Collins et al., 2008] is that central uplift formation and collapse is more pronounced for Orientale because of its greater scale. The enhanced collapse of the Orientale central uplift dominates over transient crater rim collapse (Figures 3 and 6) resulting in the peak ring forming outside the estimated transient crater. A number of lunar basins (including Imbrium, Crisium, and Serenitatis) have more than one ring structure inside of their inferred basin rim [Spudis, 1993]. Our models of Orientale suggest that it is possible that a second peak ring may form internal to the first during a second, much subdued stage of central uplift and collapse involving highly mobile melt-rich material.

[53] If the IRR of Orientale is the collapsed peak ring, then the ORR could correspond to a section within the modification zone. That assignment of the ORR is consistent with the normal faulting observed at the margins of the ring [Nahm et al., 2013] and the deformation we see at the ORR in our TP1 best-fit impact. In this case, the CR could be comparable to Chicxulub's crater rim—this correlation is also consistent with both being the topographically highest structures. Alternatively, the ORR could correspond to Chicxulub's final crater rim with Orientale's CR corresponding to the ring structure beyond Chicxulub's final crater rim. The latter seems unlikely, however, because the CR is characterized by a massive normal fault [Nahm et al., 2013], whereas the exterior ring at Chicxulub is characterized by a thrust fault.

[54] Based on a comparison of our model results with Chicxulub, an impact into the TP1 target appears a more suitable scenario for the Orientale-forming impact as: (1) the area around the location equivalent to the CR is locally disrupted in TP1, whereas it is heavily modified in TP2, (2) TP1 shows signs of normal faulting at the ORR (more so than TP2), (3) observations of Orientale suggest a highly feldspathic, therefore upper crustal, surface composition beyond the CR; our best-fit TP1 impact agrees with this observation, (4) the material distribution in our best-fit models at radial distances similar to the ORR is comparable to observations (i.e., deeper, more mafic crustal material is observed at the Montes Rook Formation where we predict lower crustal material to be exposed; see Figure 10), and (5) the elevation and radial location of the topographic high for TP1 (445 km) is similar to that observed (465 km), whereas the elevation and radial location for TP2 topographic high is well beyond the CR.

[55] If an impact into a TP1 target is the best scenario, then that favors an energy of 2.5  × 1025 J for the Orientale basin-forming impact, into a target where upper mantle temperatures were at, or close to, the solidus. This produces a transient crater ∼320 km in diameter, similar to the IR diameter, and a melt volume of 1.1  × 106 km3. Thermal conditions similar to TP2 cannot, however, be discounted.

5 Conclusions

[56] In this work, we have numerically modeled the formation of Orientale basin with the aim of resolving uncertainties in this basin's attributes. By primarily constraining model results to the location and thickness of Orientale's crustal annular bulge (as seen in the gravity-derived crustal structure), we have estimated a number of its features.

[57] The thermal state of the Moon was found to greatly influence basin formation and structure producing two different Orientale energy estimates for the two thermal profiles investigated (2.5 × 1025 and 9.9 × 1025 J). These best-fit impacts, therefore, produced two alternative estimates for Orientale's transient crater: one comparable in size to Orientale's Inner Ring, the other comparable to the Inner Rook Ring. This implies the Outer Rook Ring is an unsuitable approximation for Orientale's transient crater. Analysis of crustal distribution, deformation, and a comparison to the Chicxulub impact structure suggests Orientale's Inner Rook Ring is comparable to Chicxulub's peak ring, the Outer Rook Ring to somewhere within Chicxulub's modification zone, and the Cordillera Ring to Chicxulub's crater rim.

[58] Mantle material was not excavated in the modeled impacts (maximum excavation depths were within the lower crust), agreeing with spectroscopic data. Impact-generated melt volumes for the best-fit impacts in both thermal profiles matched well with other estimates, suggesting an impact melt volume on the order of 106 km3. This volume of melt could potentially differentiate, producing a mineralogically evolved upper layer and, therefore, mask any mantle-like signatures around the basin center.

[59] Based on analysis of material distribution and deformation, the best-fit TP1 impact appears a closer match to Orientale's observed structure than the TP2 impact, though an impact similar to the TP2 best-fit cannot be ruled out. It is possible, however, that alternative models may provide as good a fit for thermal profiles between our two scenarios. Nevertheless, our best-fits imply a near-surface thermal gradient of ∼10 K/km (similar to TP1 and TP2) and an upper mantle temperature close to, if not at, the mantle solidus, at the time of Orientale's formation at the end of the lunar basin-forming epoch.

Acknowledgments

[60] We thank Boris Ivanov, Jay Melosh, Kai Wünnemann and Dirk Elbeshausen for their work developing iSALE. This work was partially supported by NASA Lunar Science Institute contract NNA09DB33A (PI David A. Kring). We also acknowledge STFC grant ST/J001260/1 and NERC grant NE/E013589/1. The authors thank Boris Ivanov and an anonymous reviewer for their constructive, and overall positive, reviews of this paper.

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