JOURNAL OF GEOPHYSICAL RESEARCH: PLANETS, VOL. 118, 190–205, doi:10.1002/jgre.20045, 2013
Normal faulting origin for the Cordillera and Outer Rook Rings
of Orientale Basin, the Moon
Amanda L. Nahm,1,2,3 Teemu Öhman,1,2 and David A. Kring1,2
Received 3 April 2012; revised 1 December 2012; accepted 17 December 2012; published 12 February 2013.
[1] Orientale Basin is the youngest and best-preserved large impact basin on the Moon
with at least four topographic rings contained within the topographic rim marked by the
Cordillera Ring (diameter = 930 km). Its well-exposed interior makes this basin a prime
location to study basin formation processes. Forward mechanical modeling of basin ring
topography shows that the outermost rings, the Cordillera Ring (CR) and Outer Rook Ring
(ORR) are large-scale normal faults with displacements (D) of 0.8 to 5.2 km, fault dip
angles (d) of 54 to 80 , and vertical depth of faulting (T) between 19 and 37 km with most
faults having T = 30 5 km. These faults and the distribution of maria inside the basin
suggest that the transient crater, important for determining many impact-related
characteristics such as projectile size, was contained entirely within the ORR and likely had
a diameter between 500 and 550 km. The difference in crustal thickness between the
western and eastern sides of the basin is not a result of the basin-forming event, which
indicates the formation of the hemispheric crustal thickness asymmetry was likely before
the formation of Orientale Basin 3.68 to 3.85 Ga.
Citation: Nahm, A. L., T. Öhman, and D. A. Kring (2013), Normal faulting origin for the Cordillera and Outer Rook Rings of
Orientale Basin, the Moon, J. Geophys. Res. Planets, 118, 190–205, doi:10.1002/jgre.20045.
1. Introduction
[2] Orientale Basin, with a diameter in excess of 900 km
(Figure 1), is the best-preserved and youngest multi-ring basin
on the Moon [Hartmann and Kuiper, 1962; Head, 1974a;
Moore et al., 1974; McCauley, 1977; Spudis, 1993], with age
estimates of the basin-forming event ranging from 3.85 Ga
[Wilhelms, 1987] to 3.68 Ga [Whitten et al., 2011]. Centered
at approximately 265 E, 20 S, it is located on the Moon’s
western limb [Hartmann and Kuiper, 1962; Head, 1974a;
Spudis et al., 1984] in the transition region between thin
nearside and thick farside crust (Figure 2) [Ishihara et al.,
2009]. Understanding the formation of multi-ring basins is
central to lunar and planetary science because the distribution
of their ejecta is critical to interpreting lunar stratigraphy [Head,
1974a; Moore et al., 1974] and the geologic context of the
Apollo [Head, 1974a] and Luna samples. Orientale Basin is
particularly interesting because its interior remains relatively
unobscured by mare deposits unlike most of the prominent
All Supporting Information may be found in the online version of this
article.
1
Center for Lunar Science and Exploration, USRA - Lunar and
Planetary Institute, 3600 Bay Area Blvd., Houston, TX, 77058, USA.
2
NASA Lunar Science Institute, Moffett Field, California, USA.
3
Department of Geological Sciences, University of Texas at El Paso,
500 W. University Avenue, El Paso, TX, 79968, USA.
Corresponding author: A. L. Nahm, Department of Geological Sciences,
University of Texas at El Paso, 500 W. University Avenue, El Paso, TX
79968, USA. ([email protected])
©2013. American Geophysical Union. All Rights Reserved.
2169-9097/13/10.1002/jgre.20045
nearside basins. The exposed structures can, therefore, be used
to test concepts of basin formation that will be applicable to
other basins on the Moon and elsewhere in the solar system.
[3] The basin interior contains at least four topographically
defined, roughly concentric rings (Figure 1) [Hartmann and
Kuiper, 1962; McCauley, 1977]. The outermost ring scarp,
called the Cordillera Ring (CR), is the most prominent basin
ring both topographically and morphologically and, thus,
defines Orientale’s diameter of 930 km [e.g., Head, 1974a;
Spudis, 1993]. The interior ring structures, the Outer Rook Ring
(ORR) and Inner Rook Ring (IRR), have diameters of 620 and
480 km, respectively [Head, 1974a; McCauley, 1977]. An inner
ring (IR) with a diameter of about 320 km contains a flat, low
albedo unit, known as Mare Orientale [Head, 1974a], which
appears to partially fill the center of Orientale Basin.
[4] The cratering processes that produce multi-ring basins
remain poorly understood, and their resolution is one of the
science priorities targeted by the National Research Council
[2007]. Several classes of competing hypotheses for the
formation of the rings of Orientale and similar structures in
other basins have been suggested. These include the hydrodynamic or tsunami hypothesis [e.g., Baldwin, 1949, 1963,
1972, 1974, 1981], the megaterracing hypothesis [Hartmann
and Yale, 1968; Head, 1974a, 1977; Howard et al., 1974;
McCauley, 1977], the ring tectonic hypothesis [Melosh and
McKinnon, 1978; McKinnon and Melosh, 1980; McKinnon,
1981; Melosh, 1989], and the nested melt cavity hypothesis
[Head, 2010]. This study tests aspects of these diverse models,
and the results may be applied to further studies of impact basin
formation mechanisms. We show that the Cordillera and Outer
Rook Rings are large-scale normal fault scarps, which likely
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Figure 1. Orientale Basin. All parts of the figure are displayed in an orthographic projection, centered
on 265 E, 20 S. CR, Cordillera Ring; ORR, Outer Rook Ring; IRR, Inner Rook Ring; IR, Inner Ring.
Dotted lines indicate uncertainty in ring location. Locations of geographic features, craters, and pyroclastic vent noted for reference. (a) Lunar Reconnaissance Orbiter (LRO) Wide Angle Camera (WAC)
monochrome mosaic, 100 m/px (http://wms.lroc.asu.edu/lroc_browse/view/orientale). (b) LRO Lunar
Orbiter Laser Altimeter (LOLA) topography overlain on shaded relief map. Resolution: 512 px/deg.
(c) Simplified geologic map (modified from Scott et al. [1977]) overlain on shaded relief map derived
from LOLA data in Figure 1b. Cc, Copernican crater material; Ec, Eratosthenian crater material; EIm,
Eratosthenian/Imbrian mare material; EIph, Eratosthenian/Imbrian hilly plateau material; Ic2, Imbrian
crater material (post-Orientale); Ip, Imbrian plains material; Iom, Maunder Formation; Iorm, Montes
Rook Formation massif facies; Iork, Montes Rook Formation knobby facies; Iohi, Hevelius Formation
inner facies; Ioht, Hevelius Formation transverse facies; Ioho, Hevelius Formation outer facies; Iohs,
Hevelius Formation secondary crater facies; Ic1, Imbrian crater material (pre-Orientale); INt, Imbrian/
Nectarian undivided terra material; Nhb, Hertzsprung basin material; Nhsc, Hertzsprung secondary
crater material; Nc, Nectarian crater material; NpNhf, Nectarian hilly and furrowed material; pNc,
Pre-Nectarian crater material. (d) Sketch map of the Orientale Basin showing the inferred location
of the rings and geographic features, craters, and pyroclastic vent for reference.
formed as a result of the collapse of the transient cavity walls to
form the basin equivalent of a complex crater modification zone
during basin formation.
characteristics, impact and post-impact deposits, and crustal
thickness, which provide geologic context for our measurements and models.
2. Orientale Basin
2.1. Basin Topography
[5] Impact basins represent a class of structures characterized
as depressions with widespread concentric-radial lineament systems [Gilbert, 1893; Baldwin, 1942, 1943, 1949; Dietz, 1946;
Hartmann and Wood, 1971] having diameters in excess of
~300 km for the Moon [Wilhelms, 1987; Spudis, 1993]. In the
following, we briefly summarize the basin topography, ring
[6] Figure 1b shows topographic data (512 px/deg) collected
by the Lunar Orbiter Laser Altimeter (LOLA) [Smith et al.,
2010] (onboard the Lunar Reconnaissance Orbiter (LRO))
draped over a shaded relief model. On average, elevations are
~6 km higher on the western side of the basin than on the east.
The four topographic rings of Orientale are prominent as
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NAHM ET AL.: FAULTING IN ORIENTALE BASIN
approximately concentric circles (Figure 1b). In profile, the
basin has a stair-step shape, with the highest elevation on the
western rim at about 6.5 km, stepping down to the lowest
elevation in the basin-interior mare of around 3 km. Low
topographic slopes (~1–3 ) exist between the ORR and CR
with the elevation decreasing toward the CR and then abruptly
increasing at the CR scarp by ~4 km.
Figure 2. Crustal thickness map of the Orientale region.
Data from Ishihara et al. [2009] overlain on a LOLA
DEM. top panel: Regional map of crustal thickness, with
locations of crustal thickness profiles shown in lower panels.
CR, Cordillera Ring; ORR, Outer Rook Ring; IRR, Inner
Rook Ring; IR, Inner Ring. second panel: Profile A showing
transition from thick highland crust to thin nearside crust
north of Orientale. Note slope showing thinning to the east.
third panel: Profile B taken through the center of the Orientale
Basin showing the variation in crustal thickness at the major
outer rings and in the basin center, along with the overall
trend from profiles A and C. fourth panel: Profile C showing
similar thinning of crust as in Figure 2b with eastward slope.
2.2. Rings
[7] The Cordillera Ring (diameter = 930 km) consists of a
generally continuous inward-facing scarp and rises up to 4 km
above the surrounding plains [Head, 1974a]. In places, it takes
on a saw-toothed appearance [Head, 1974a], which has been
attributed to the preexisting “lunar grid” [Fielder, 1961; Head,
1974a, 1974b] where the ring structures may have formed along
preexisting fractures oriented in the northwest–southeast
and northeast–southwest directions. The Outer Rook Ring
(diameter = 620 km) is composed of kilometer-scale
massifs with steep slopes facing the basin interior [Head,
1974a] and is the most continuous and topographically rugged
of the rings [McCauley, 1977]. Linear portions of the ring
parallel those of the Cordillera Ring [Head, 1974a] in the west.
As reported for the Cordillera Ring, the Outer Rook Ring
height is variable [Head, 1974a] and ranges from 1 to 5 km
in elevation. Spectroscopic studies of the ORR suggest a
composition ranging from noritic anorthosite to anorthositic
norite [Spudis et al., 1984], although shocked anorthosite has
also been observed [Bussey and Spudis, 1997, 2000; Hawke
et al., 2003; Pieters et al., 2009].
[8] In the southwest where the concentric trends of the rings
are lost, the Cordillera and Outer Rook Rings are generally
indistinguishable (Figure 1d) [Scott et al., 1977] using imagery
alone. The complicated structures have been suggested to be
related to an oblique basin-forming impact, orientation of
preexisting crustal weaknesses [Scott et al., 1977], the presence
of a pre-Orientale basin [Schultz and Spudis, 1978; Spudis,
1993; Wood and Collins, 2011], or increasing lithospheric
thickness towards the west [Head and Solomon, 1980]. The
Cordillera and Outer Rook Rings do not show any significant
difference in crater density relative to the ejecta blanket
[Hartmann and Wood, 1971], indicating that these rings formed
contemporaneously with the Orientale basin-forming event and
with each other.
[9] The Inner Rook Ring (diameter = 480 km), characterized
by disconnected peaks that range in height from 1 to 3 km and 2
to 10 km in width [Head, 1974a], is well developed on the
northwest, southwest, and southeast sides of the basin
[McCauley, 1977]. The eastern portions of the IRR are
interpreted to be composed of essentially pure anorthosite
of both shocked and nonshocked varieties, in contrast to
the generally more mafic compositions of other Orientale
units [Spudis et al., 1984; Head et al., 1993, 2010; Bussey
and Spudis, 1997, 2000; Hawke et al., 2003; Ohtake et al.,
2009; Pieters et al., 2009], implying that the material that
comprises this ring is likely derived from the upper crust
[e.g., Spudis and Davis, 1986]. The IRR has been
variously interpreted as marking the location of the rim
of the transient cavity [Floran and Dence, 1976], a central
uplift analogous to central peak rings [e.g., Head, 1974a;
Moore et al., 1974; McCauley, 1977; Solomon and Head,
1980; Head et al., 1993], and the surface expression of
the crust-mantle boundary [Wilhelms et al., 1977; Hodges
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NAHM ET AL.: FAULTING IN ORIENTALE BASIN
and Wilhelms, 1978] (meaning that the formation of this
ring is the result of the rheological layering of the target).
[10] The Inner Ring (diameter = 320 km) is discontinuous
and consists of a rounded step-like scarp that separates the
lowest parts of the inner basin from the more rugged terrain
outside [McCauley, 1977]. In the south to southwest, the Inner
Ring scarp may consist of small steep-faced fault segments in
contact with the mare; to the northeast, the ring is expressed
as elongated fractured ridges [McCauley, 1977]. The Inner Ring
may delineate the deepest part of the original cavity [McCauley,
1977] into which basalt was extruded but from which it was
later partially withdrawn [Scott et al., 1977], causing the inner
part of the basin to collapse. Alternatively, the depression bound
by the Inner Ring may be the result of thermal contraction and
subsidence of the impact melt sheet and substrate [Bratt et al.,
1985; Head et al., 1993]. The inward-facing scarp of the Inner
Ring has also been suggested to represent a major strength
discontinuity in the lunar subsurface [Scott et al., 1977].
[11] In addition to the four well-defined topographic rings
(Figure 1), at least two rings external to the Cordillera Ring have
also been suggested, with diameters of 1300 [Hartmann and
Kuiper, 1962] and 1900 km [Pike and Spudis, 1987]. They
appear to have subdued scarp-like morphology [Spudis,
1993], possibly from obscuration due to deposition of ejecta
on top of these structures. A possible exterior ring, with a
diameter of 1460 km, was noted by van Dorn [1969]. A
possible ring interior to the IR may exist as a roughly
basin-concentric wrinkle ridge system in western Mare
Orientale [Scott et al., 1977; Solomon and Head, 1980].
2.3. Impact-related Deposits
[12] The Orientale Group includes all materials produced by
and deposited contemporaneously with the basin-forming
event: the Hevelius, Montes Rook, and Maunder Formations
[McCauley, 1977; Scott et al., 1977]. The Hevelius Formation,
including the inner (Iohi), transverse (Ioht), outer (Ioho), and
secondary crater (Iohs) facies (Figure 1c), lies mostly outside
the Cordillera Ring [McCauley, 1977; Scott et al., 1977]. It
has been interpreted as an early-stage ejecta blanket and consists
of hummocky, lineated, and swirl-textured deposits [McCauley,
1977; Scott et al., 1977; Spudis, 1993]. The Hevelius Formation
extends at least one basin diameter beyond the Cordillera Ring
[Spudis et al., 1984; Head et al., 1993], and its thickness has
been estimated to be ~3.6 km [Moore et al., 1974; Scott et al.,
1977] or ~2.9 0.3 km [Fassett et al., 2011] at or near the
CR. The Cordillera Ring is draped by the Hevelius Formation
over ~80% of its circumference; this has been interpreted to
indicate that the CR formed before the emplacement of the
ejecta [Howard et al., 1974; McCauley, 1977]. Head [1974a],
however, notes that no Orientale ejecta has been deposited
against the Cordillera Ring scarp, which may imply that the
formation of the scarp occurred after and/or contemporaneously
with the deposition of ejecta.
2.4. Mare Deposits
[13] Three major mare units occur within Orientale Basin
[Scott et al., 1977] (Eratosthenian/Imbrian mare material:
EIm; Figure 1c). In decreasing areal extent, these are Mare
Orientale, Lacus Veris, and Lacus Autumni. Basin volcanism
is preferentially concentrated along the interior of basin rings
and in the inner part of the basin [Hartmann and Kuiper,
1962]. Mare Orientale is located in the basin center and
encompassed almost entirely by the Inner Ring. Lacus Veris
is located between the Inner and Outer Rook Rings, and Lacus
Autumni is contained between the Outer Rook and Cordillera
Rings (Figure 1a). Based on crater counting studies [Greeley
et al., 1993; Whitten et al., 2011], mare volcanism occurred significantly later than basin formation, as indicated by the lower
relative crater density of the mare when compared to the basin
ejecta blanket [Hartmann and Wood, 1971]. Mare Orientale is
the oldest identified in Orientale Basin, with model ages of
3.7–3.45 Ga [Greeley et al., 1993; Whitten et al., 2011]. Lacus
Veris and Lacus Autumni have estimated age ranges of
3.69–3.20 and 3.47–1.66 Ga, respectively [Greeley et al.,
1993; Whitten et al., 2011], consistent with previous estimates
of the duration of lunar volcanism [e.g., Hiesinger et al., 2011].
2.5. Crustal Thickness
[14] The Orientale Basin straddles the nearside–farside
crustal thickness asymmetry (Figure 2), with thicker crust
present on the western side of the basin in the lunar highlands
[Ishihara et al., 2009]. Estimates of crustal thickness have been
obtained from inversion of lunar gravity and topography data
[Hikida and Wieczorek, 2007; Ishihara et al., 2009]. Here, the
term crust refers to the upper compositional stratum of the
Moon and is defined in these models by the density of crustal
materials (e.g., anorthosite and basalt).
[15] The thickness of the lunar crust in the vicinity of
Orientale ranges from 20 to 90 km [Hikida and Wieczorek,
2007; Ishihara et al., 2009]. The thickest crust in the Orientale
Basin (~80 km) occurs in the northwestern, western, and
southern sections of the ORR [Ishihara et al., 2009]. The crust
in the basin interior is ~20 km thick [Ishihara et al., 2009],
although it may be as thin as 0.7 km [Hikida and Wieczorek,
2007]. Most of the crust in the Orientale Basin region, however,
is ~70–80 km thick [Ishihara et al., 2009]. The crustal thickness
notably decreases with a gradient of 20 m/km from the west to
the east (Figure 2).
3. Overview of Basin Formation Models
[16] The variety of basin ring morphologies on the terrestrial
planets and icy satellites suggests that several mechanisms may
be responsible for ring formation in impact basins [McKinnon
and Melosh, 1980]. While there is no consensus on the
formation of structures inside the transient cavity, i.e., peak
rings and central peaks, there is general agreement that the
collapse of the central uplift is important [e.g., Melosh, 1989;
Melosh and Ivanov, 1999; O’Keefe and Ahrens, 1999; Morgan
et al., 2000; Collins et al., 2002]. Here, we focus our brief
discussion on the hypotheses for the formation of the rings
outside the transient cavity; for Orientale Basin, we consider
these to be the Cordillera and Outer Rook Rings.
[17] Previous hypotheses for the formation of outer rings in
lunar impact basins can be divided into several classes of
models: the hydrodynamic or tsunami model [e.g., Baldwin,
1949, 1963, 1972, 1974, 1981; van Dorn 1968, 1969], the ring
tectonic model [Melosh and McKinnon, 1978; McKinnon and
Melosh, 1980; McKinnon, 1981; Melosh, 1989], the nested melt
cavity model [Head, 2010], and the megaterracing model
[Hartmann and Yale, 1968; Head, 1974a, 1977; Howard
et al., 1974; McCauley, 1977]. These hypotheses share
the idea that stresses in the Moon, built up due to violent
dissipation of energy imparted to the lunar surface during
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basin formation, result in the formation of concentric ring
structures [Hartmann and Wood, 1971] but differ in the
manner(s) in which these stresses are relieved and predictions on the presence and type of faults occurring at the
basin rings.
[18] Ring formation in the hydrodynamic or tsunami model
as proposed by Baldwin [1949, 1963, 1972, 1974, 1981]
results from impact-induced shock modification of rock,
causing the target material to behave as a fluid-like rubble; as
stresses decay, the rock no longer behaves as a fluid, forming
rings as anticlines and synclines when impact-induced
tsunami-like shock waves “freeze” in place [Baldwin, 1949,
1963, 1972, 1974, 1981] or from repeated rebound and
collapse (i.e., oscillation) of the central uplift [Murray,
1980]. This model gives no predictions on the presence or
absence of faults at basin rings.
[19] In the ring tectonic model, basin rings are normal
fault scarps along which crustal blocks move radially inward
toward the basin center; material on the basinward side of
these scarps is downthrown relative to the pre-impact surface [McKinnon and Melosh, 1980; Melosh, 1989]. This
model predicts that the rings of basins formed on bodies
with thick lithospheres such as the Moon are composed of
antithetic, approximately circumferential normal faults
induced by inward asthenospheric flow (asthenosphere
forming a low-viscosity channel) [Melosh and McKinnon,
1978; McKinnon, 1981; Melosh, 1989]. These normal faults
are assumed to form when the depth of the transient cavity
exceeds the thickness of the lithosphere at the time of the basin
formation [Melosh, 1989; c.f. McKinnon and Melosh, 1980].
[20] The nested melt cavity model [Head, 2010] for the
formation of basin rings is based on the observation that the
proportion of impact-produced melt increases with increasing
transient crater size [e.g., Cintala and Grieve, 1998]. After
impact, a melt cavity forms at the sub-impact point, where
the boundary between the melt zone and the displaced zone
consists of material that experienced high shock pressures
but is still solid [Head, 2010]. The highly shocked rocks of
the displaced zone rebound to form the expanded peak ring,
moving upward and inward, displacing the impact melt in
the central depression, which results in melt coating the
collapsing cavity floor and ponding of melt in the central crater
[Head, 2010]. Deep-seated radially inward listric normal
faulting is initiated along the base of the displaced zone,
propagating outward to the base of the structurally uplifted
rim, which becomes an additional ring with an inward-facing
fault scarp [Head, 2010]. Collapse of the rim forms a megaterrace [Head, 1974a, 1977, 2010], which may be analogous to
the CR in Orientale.
[21] The most commonly proposed mechanism for basin
ring formation is the megaterracing model [Hartmann and
Kuiper, 1962; Hartmann and Yale, 1968; McCauley, 1968,
1977; Mackin, 1969; Hartmann and Wood, 1971; Head,
1974a, 1977; Howard et al., 1974; Scott et al., 1977]. In this
model, outer rings of Orientale and other multi-ring basins
form as large normal faults, where the interior of the basin
is displaced downward relative to the preexisting lunar
surface. In particular, on the basis of the morphology similar
to fault scarps with steep inner faces [Hartmann and Kuiper,
1962], the Cordillera Ring has been interpreted to represent
a large-scale normal fault scarp [e.g., Hartmann and Wood,
1971; Head, 1974a, 1977; Howard et al., 1974].
[22] As most hypotheses for the formation of basin rings
require or predict normal faulting, we will test these by
performing mechanical modeling of Orientale CR and
ORR cross-sectional topography, as described in section 4.
4. Modeling Approach
[23] To test the fault-related models of basin ring formation,
we initially assume that the ORR and CR scarps are normal
faults and adopt a standard technique that utilizes the inversion
of fault-related topography [e.g., Cohen, 1999; Schultz and Lin,
2001] to determine if the rings can be successfully modeled as
normal faults. Models of this type are used to calculate surface
displacements due to underlying faults with prescribed
geometries and displacement magnitudes [Schultz and Lin,
2001]. Forward mechanical modeling has been used
successfully to model the surface displacements from faults
on Mercury [Watters et al., 2002], Earth [e.g., King et al.,
1988; Toda et al., 1998; Cohen, 1999; Muller and Aydin,
2005; Resor, 2008], asteroids [Watters et al., 2011], and
Mars [e.g., Schultz, 2000; Schultz and Lin, 2001; Schultz
and Watters, 2001; Grott et al., 2007]. This approach
provides remarkably good fits to the structural topography above
a fault for a relatively narrow range of parameters [e.g., Cohen,
1999; Schultz and Lin, 2001]. As shown below, a good correspondence between the output model displacements and
observed LOLA topography would suggest that the fault
parameters obtained from modeling are representative of
the characteristics of the putative ring-forming faults [e.g.,
Schultz and Lin, 2001].
4.1. LOLA Topography across the Rings of Orientale
Basin
[24] Several sets of topographic profiles were derived from
the 512 pixel/deg (~60 m/px at the equator) LOLA digital
elevation model (DEM) [Smith et al., 2010] perpendicular
to ring scarps (although not necessarily always basin radial)
in each location of inferred normal faulting (CR and ORR;
Figure 3), with profiles A through J shown in Figures 4 and
5. Additional profiles (K–N in Figure 3) were derived from
IRR and shown in Figure 6 for comparison (see below).
The interior of the basin is to the left in most profiles; those
that have a different orientation are noted in Figure 4. The
mean topography for each set of profiles, calculated by stacking and averaging individual profiles, is shown as heavy
black lines in Figures 4 and 5. The number of profiles
averaged and the distances over which these are spread are
given in the corresponding figure captions. Thin black lines
indicate the spread of one standard deviation. Where a regional
slope of ~1 was apparent (e.g., profiles C, D), it was subtracted from the average profiles [e.g., Grott et al., 2007] prior
to modeling.
[25] The main topographic features evident in the suite of
profiles are the scarp faces, having slopes of 2 to 24 (accounting for vertical exaggeration), and the more gently dipping
footwall back scarps, having slopes of less than 1 to 4 . The
primary criteria for identification of potential faults in the
topographic profiles are steep slopes occurring near or at ring
boundaries, nonlinearly sloping backscarps, elevation offset,
and elevations leveling off away from the scarp.
[26] As ejecta deposition significantly contributes to the
morphology of impact basins and the surrounding regions,
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of structural topography can be neglected when performing
forward mechanical modeling of fault topography.
[28] Shown in Figure 6 are representative profiles across
the IRR for comparison with those taken across the CR
and ORR (Figures 4 and 5). These profiles illustrate that
typical cross-sectional views of the massifs in the IRR do
not show the characteristics of normal faulting (listed above)
as are evident in profiles A–J. Instead they show single
(profiles L and N) or double peaks (profiles K and M) with
sharp to sloping scarps where the profile crosses the mapped
IRR ring boundary as well as on both sides of the IRR
massif peaks. Therefore, this indicates that the IRR did not
form from normal faulting, and accordingly, the IRR profiles
are not included in the modeling.
Figure 3. Locations of profiles used in this study shown on
LOLA DEM. Profiles A through F cross the Cordillera Ring
and are shown in Figure 4. Profiles G through J cross the Outer
Rook Ring and are shown in Figure 5. Profiles K through N
cross the Inner Rook Ring and are shown in Figure 6.
possible effects on the resultant scarp topography must be
estimated. Assuming fault formation occurred prior to ejecta
deposition [Howard et al., 1974; McCauley, 1977; Scott et al.,
1977], the shape of any footwall could have been significantly
modified by the deposition of ejecta. The difference in topography and decrease in ejecta thickness with distance was
calculated at the ORR and CR based on the equation derived
by Fassett et al. [2011]. Fassett et al. [2011] estimate the
thickness of ejecta deposited at the CR to be on the order of
2.9 km and up to ~9 km at the ORR. However, the latter
value is heavily dependent on the presumed diameter of
the transient crater. Although ejecta deposition drastically
changes the topography of the region surrounding the transient crater, to first-order, no significant difference in the
shape of the topographic profiles results from these calculations as the slope and concavity of the decrease in topography with distance due to faulting and ejecta deposition are
roughly similar. Indeed, removal of the ejecta from the topographic profiles slightly decreases the angle of the backslope.
Comparison of the model results using this topography
yields model parameters within ~15% of the reported values.
Thus, the effect of ejecta deposition is generally minor with
respect to the model fits. A caveat to this is that the deposition of ejecta on top of the fault topography would increase
the elevation of the scarps, which may lead to artificially
high values of displacement on each fault determined from
our models. Thus, the results reported here are likely maximum values of displacements along the ring-forming faults.
[27] On the other hand, if the formation of the CR and
ORR scarps occurred after the deposition of ejecta [Head,
1974a], the present-day topography of the footwalls closely
approximates the topography shortly after faulting occurred.
In addition, slopes due to ejecta deposition calculated from
Fassett et al. [2011] are 1 and 0.7 , determined from the
CR out 10 and 100 km, respectively. Thus, to first-order,
complications associated with potential ejecta modification
4.2. Model Details
[29] We use the forward mechanical dislocation modeling
program Coulomb (the software, which requires MATLAB
to run, can be downloaded from http://earthquake.usgs.gov/
research/modeling/coulomb/overview.php) [Lin and Stein,
2004; Toda et al., 2005] to model surface displacements
associated with normal faulting. Stress and material displacement calculations are made in an elastic half-space with
uniform isotropic elastic properties following the equations
derived by Okada [1992].
[30] In our models, a fault is idealized as a rectangular plane
with the sense of slip (i.e., normal, thrust, strike-slip, or
oblique), magnitude of displacement, fault dip angle, depth of
faulting, and fault length specified (Figure 7). We use a constant
value of pure dip-slip displacement (no strike-slip component)
[Schultz and Lin, 2001] in both the strike and dip directions
along each modeled fault. While a nonconstant slip distribution
can be prescribed in Coulomb in both the strike and dip
directions, the way in which the slip is distributed is unknown
and cannot be determined from the data currently available for
the Moon. For example, some studies recommend a triangular
slip distribution [e.g., Cowie and Scholz, 1992; Dawers et al.,
1993; Manighetti et al., 2005], while others suggest an elliptical
slip distribution better approximates natural faults [e.g., Pollard
and Segall, 1987; Bürgmann et al., 1994; Willemse et al.,
1996]. Thus, while a tapered slip distribution is more representative of the way faults behave, imparting a slip distribution to
the faults modeled here would add another layer of uncertainty.
Therefore, we choose to model the faults with as few assumptions as possible for simplicity. Using a tapered slip distribution,
which is more realistic for natural faults, would likely result in
deeper depths of faulting (T) than the method used here.
[31] Good fits to the LOLA topography were obtained by
iteratively adjusting the values of fault dip angle, displacement,
and depth of faulting in the model. The initial displacement
magnitude is estimated from the relief of the scarp and adjusted
based on model output. Final model parameters were
determined based on visual fits to the shape of the footwall
between observed (measured LOLA) and predicted (modeled)
topography [Watters et al., 2002], as the shape of the footwall
uplift is characteristic of normal fault topography [e.g., Jackson
and McKenzie, 1983]. On the Earth, the uplift of extensional rift
flanks has been shown to be the result of isostatic rebound of the
lithosphere following normal faulting; the shape of the uplift
forms at the time of faulting and represents permanent deformation of the lithosphere [Jackson and McKenzie, 1983; Weissel
and Karner, 1989]. Applying this to the Moon and lunar impact
195
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
Figure 4. Average topographic profiles (heavy black lines) derived from LRO LOLA DEM 512 px/deg
with one standard deviation (thin black lines) shown for the Cordillera Ring. Profile locations shown in
Figure 3. Best-fit forward mechanical modeling results shown in red. Average profiles shown above
created from averaging N profiles for each profile set: A, N = 8; B, N = 5; C, N = 8; D, N = 6; E, N = 9;
F, N = 9. The spread of the profiles perpendicular to strike, with the average distance between the profiles
in parentheses, are as follows: A, 12.9 km (1.6 km); B, 6.4 km (1.3 km); C, 2.1 km (0.26 km); D, 2.5 km
(0.42 km); E, 5.7 km (0.62 km); F, 2.5 km (0.27 km). The basin interior is to the left in all profiles, except
for profile C. Best-fit model parameter values listed in Table 1. Faults are numbered for profiles with more
than one fault; numbering corresponds to those in Table 1. Elevations are referenced to a sphere with
radius of 1734 km. Note that all faults dip toward the basin interior with the exception of faults C and
the fault F2.
basin large-scale normal faults, the present-day topography
likely is representative of the original structural topography as
modification of the initial topography by erosion on the Moon
is minor. However, mare deposits have often infilled the bases
of the fault scarps which further emphasizes that the often poor
fits between the topography of the hanging wall and the model
are not a major concern; the footwall preserves the original fault
topography much better than the hanging wall.
[32] It should be noted, however, that flexure of the
lithosphere is a long-term, slow, viscoelastic process, and this
model does not and cannot model this type of behavior.
Given the strain rates associated with impact structure
formation (i.e., orders of magnitude faster than geologic strain
rates [e.g., Key and Schultz, 2011]), brittle elastic processes
are dominant. A sufficiently thick lithosphere, such that it does
not flow and allow stress relaxation on short timescales, like that
for the Moon, would remain rigid, and footwall uplift would be
negligible at the surface during impact basin formation.
Neglecting the effects of viscous deformation that has taken
place since the formation of the basin means that our calculated
displacement values should be regarded as minima.
[33] This modeling approach is not sensitive to the removal
of regional slopes (i.e., detrending) carried out during the
LOLA topographic data processing [e.g., Grott et al., 2007].
Although small variations in the parameters are permitted by
the topographic data, a remarkably good match between the
models and data is obtained for a relatively narrow range of
the fault parameters.
[34] A Young’s modulus E of 83 GPa [Turcotte and
Schubert, 2002], Poisson’s ratio n of 0.25, and coefficient of
196
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
Figure 5. Average topographic profiles (heavy black lines)
derived from LRO LOLA DEM 512 px/deg with one standard
deviation (thin black lines) shown for the Outer Rook Ring.
Profile locations shown in Figure 3. Best-fit forward mechanical modeling results shown in red. Average profiles shown
above created from averaging N profiles for each profile set:
G, N = 8; H, N = 6; I, N = 7; J, N = 7. The spread of the profiles
perpendicular to strike, with the average distance between the
profiles in parentheses, are as follows: G, 2.2 km (0.27 km); H,
3.4 km (0.56 km); I, 3.8 km (0.53 km); J, 4.5 km (0.64 km).
The basin interior is to the left in all profiles. Best-fit model
parameter values shown in Table 2. Faults are numbered for
profiles with more than one fault; values correspond to those
in Table 2. Elevations are referenced to a sphere with radius
of 1734 km.
Figure 6. Topographic profiles derived from DEM shown
for the Inner Rook Ring. Profile locations shown in Figure 3.
The basin interior is to the left in all profiles. Elevations are
referenced to a sphere with radius of 1734 km.
197
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
this model. For example, fault block rotations and changes in
fault dip with depth (i.e., listric fault geometry) usually associated with normal faults are not accounted for in the Coulomb
model. Future work, which applies finite element models of
finite thickness plates, may be able to clarify the roles of fault
block rotation, listric, or otherwise nonplanar fault geometries
and tapered fault slip distribution and their effects on the
determination of important fault parameters. Thus, the work
presented here represents an important first step in determining
the characteristics of the ring-forming faults at depth.
5. Results
Figure 7. Schematic representation of model parameters
and definition of terms used in this study. (a) Block diagram
of single rectangular normal fault showing slip direction and
geometry parameters: fault length L, downdip fault height H,
depth of faulting T, and fault dip angle d. (b) Cross-sectional
view showing the relationship between fault displacement
D, dip angle d, vertical displacement or throw, horizontal
displacement, and structural uplift of the footwall for a
single normal fault. Modified after Schultz and Lin [2001]
and Schultz et al. [2010].
friction m of 0.6 are assumed for the anorthositic rock mass for
the pre-impact surface [Schultz, 1995, 1996; Schultz and Lin,
2001]; these values were held constant throughout the modeling
process as changes in these parameters do not have significant
effects on the resultant fault-related topography. The primary
variables that affect the shape of the modeled displacement
are the fault dip angle d, vertical depth of faulting T, and the
magnitude and sense of displacement D. Briefly, variations in
the fault dip angle affect both the amplitude and shape of the
fault-related topography [Schultz and Lin, 2001] (Figure S1a
in the Supporting Information), increasing the magnitude of displacement increases the structural relief (Figure S1b), and the
shapes of the footwall uplift and hanging wall are affected by
changes in the vertical depth of faulting [Schultz and Lin,
2001] (Figure S1c). A more in-depth discussion of the effects
of changing these parameters is presented in the Supporting
Information.
[35] Modeling using an elastic half-space minimizes the
deformation that would (or did) occur in a plate with a finite
thickness (such as the lithosphere/crust of a planetary body).
Thus, misfits between the model output and the measured
LOLA topography are expected; additional sources of misfits
may come from the simplistic approach to fault mechanics in
[36] Values for the best-fit parameters to the LOLA topography across the CR and ORR scarps are listed in Tables 1 and 2,
respectively. In general, the depth of faulting T ranges from 19
to 37 km, fault dip angle d varies between 54 and 80 , and fault
displacement D varies between 0.8 and 5.2 km. For the CR, the
depth of faulting T ranges from 19 to 37 km, fault dip angle d
varies between 61 and 75 , and fault displacement D varies
between 0.8 and 5.0 km. For the ORR, the depth of faulting T
ranges from 20 to 30 km, fault dip angle d varies between 54
and 80 , and fault displacement D varies between 1.6 and
5.2 km. This relatively constant depth of faulting may correlate
to some sort of mechanical layer thickness (perhaps similar to
the lithosphere) during the impact event. Rheology of geological materials has been shown to affect the ways in which rock
behaves when subjected to different strain rates (i.e., the
strength envelope) [e.g., Brace and Kohlstedt, 1980]. Thus,
under these high strain rate conditions, the rock at depth may
have behaved mechanically homogenously down to about
35 km. Given the inherent uncertainty in the approach of
forward modeling techniques, these results represent the bestfitting solutions to the topography data and are likely not unique
solutions. However, the fits of the model outputs match the
averaged topography well, as shown in Figures 4 and 5, and
thus, the values for the best-fit parameters are significant.
[37] Profiles A, B, C, and E (Figure 4) and G, H, and I
(Figure 5) show individual normal faults, with the uplifted
footwall to the right and the down-dropped hanging wall to
the left. In these cases, best-fit models are chosen based on the
match to the footwall uplift alone, with the exception of profile
I (Figure 5). The reasoning is twofold; the shape of the footwall
Table 1. Best-fit Model Parameters for Faults in Average Profiles
for the Cordillera Ringa
Profile
A
B
C
D
1
1
1
1
2
3
4
1
1
2
E
F
198
Fault
Fault
Depth of
Number Length L Faulting T
(km)
(km)
a
140
120
65
85
85
85
85
36
46
28
34
33
19
30
30
25
30
30
37
27
Fault Displacement Fault
Dip d
D (km)
Heightb H
(km)
(deg)
69
70
70
75
75
70
75
61
66
69
Profile locations shown in Figure 3.
Downdip fault height H = T/sin d (see Figure 7).
b
4.9
5.0
1.5
4.3
1.7
3.4
0.8
5.0
2.5
0.9
36.4
34.1
20.2
31.4
31.2
26.6
31.2
34.3
40.5
28.9
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
uplift is more characteristic of normal fault topography than the
shape of the hanging wall [e.g., Jackson and McKenzie, 1983],
and the model outputs material displacements relative to an
original flat, horizontal datum to which displacements are
referenced. As observed in profiles A, B, C, E, G, and H, this
is reflected by the best-fit model having higher elevations and
plotting above the LOLA topography on the hanging wall side.
The shape of the model may mimic the shape of the hanging
wall as in profile C (Figure 4) or match the hanging wall fairly
well (profile I; Figure 5), indicating that the method can
accurately calculate the hanging wall displacements but not
the elevation offset as a result of fault motion.
[38] The remaining profiles show more complicated normal
fault geometries. Profile F (Figure 4) shows a horst with the
master fault, fault 1, on the basinward side. Profile D (Figure 4)
shows a complicated section of the western Cordillera Ring
structure over 250 km long and encompasses a minimum of
four synthetic normal faults. Profile J (Figure 5) shows at least
two synthetic normal faults dipping toward the basin interior.
The topography of fault 2 is well matched by the best-fit model,
but the topography of fault 1 is not. This may be due to the
presence of a small fault at approximately 60 km. Models that
fit portions of the fault 1 topography had unrealistic fault
parameters (e.g., T = 10 km, D = 5 km), and models with three
faults also did not produce satisfactory fits. The remaining
possibilities to explain the misfit are that the topography of fault
1 was modified during ejecta emplacement after fault formation
or that fault 2 is the only fault transected by profile J. The
remarkable fit of the model to profile D between faults 1 and
2 and 2 and 3 and in profile F between faults 1 and 2 demonstrates the robustness of this method.
[39] Figure 8 shows a map view of the modeled faults in
Orientale. There are likely many other faults present, but
without confirmation by modeling or unambiguous interpretations of imagery and topography, only those faults successfully
modeled are shown. A cross-sectional view of Orientale with
fault locations and dip angles determined from modeling
profiles A and D is shown in Figure 8.
subsurface imaging techniques involving gravity and magnetic
surveys, seismic reflection and refraction studies, and modeled
with hydrocode. Those results show that the final crater rim is
defined, in part, by an inward dipping normal fault scarp [Morgan
and Warner, 1999], which is one of several produced during the
collapse of the ~100 km diameter transient crater [Kring, 1995;
Morgan et al., 1997], forming a modification zone and the final
basin. The central uplift collapsed to form a ~90 km diameter
peak ring whose radially outward movement produced a collision
with and overlapping of the inward collapsing transient cavity
wall [Morgan and Warner, 1999; Collins et al., 2002; Ivanov,
6. Discussion and Implications
6.1. Insights from a Terrestrial Analog: Chicxulub
[40] As the Chicxulub impact structure, located on the Yucatán
Peninsula, Mexico, is the largest (diameter 180 km), youngest
multi-ring impact basin on Earth [e.g., Morgan and Warner,
1999], it represents the best terrestrial analog for large-scale
impact structures on the Moon and the terrestrial planets. The
structure has been probed by multiple boreholes, scanned with
Table 2. Best-fit Model Parameters for Faults in Average Profiles
for the Outer Rook Ringa
Profile
G
H
I
J
Fault
Fault
Depth of
Number Length L Faulting T
(km)
(km)
1
1
1
1
2
a
36
36
77
77
25
20
28
29
30
25
Fault Displacement Fault
Dip d
D (km)
Heightb H
(deg)
(km)
80
80
79
70
54
Profile locations shown in Figure 3.
Downdip fault height H = T/sin d (see Figure 7)
b
4.4
4.4
5.2
4.0
1.6
20.3
28.4
29.5
31.9
30.8
Figure 8. Orientale Basin map view and cross-section.
Upper panel: Map view of the Orientale Basin topography
showing the distribution of the modeled faults in this study.
Letters denote the profile that crosses the mapped faults.
Black boxes occur on the down-dropped side of the fault
(the hanging wall). Heavy black lines denote location of
topographic profile (X–Y–Z) shown at top of lower panel.
Lower panel: Cross-sectional view of Orientale showing
modeled faults from profile D (X to Y), profile A (Y to Z),
and the fault from profile H projected onto the section line
Y to Z (dash-dotted line). LOLA topography along profile
line X–Y–Z shown at top with locations where profile
crosses basin rings (IR, IRR, ORR, and CR) noted. Vertical
exaggeration, 12:1. Approximate depth of excavation of
50 km [Croft, 1980] noted.
199
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
2005; Morgan et al., 2011]. Seismic data reveal a vertical offset
of stratigraphy of ~3 to 6 km along the normal faults at the basin
rim [Morgan et al., 1997; Morgan and Warner, 1999], with
single displacements up to 2.5 km [Morgan and Warner, 1999].
An additional circumferential fault system is characterized by
outward-verging thrust faults occurring beyond the final crater
rim with a diameter of ~250 km [Morgan and Warner, 1999].
The structures in the Chicxulub basin are most similar to those
associated with the megaterracing model [Hartmann and Yale,
1968; Head 1974a, 1977; Howard et al., 1974; McCauley,
1977; Morgan and Warner, 1999], although the outermost ring
may have formed by the process described by the ring tectonic
model [Melosh and McKinnon, 1978; McKinnon and Melosh,
1980; McKinnon, 1981; Morgan and Warner, 1999].
[41] Forward mechanical modeling of the Outer Rook and
Cordillera Ring topography performed in this study indicates
that these rings were formed by large-scale normal faults for
which the vertical extent of faulting T ranges from 19 to
37 km, fault dip angle d varies between 54 and 80 , and fault
displacement D varies between 0.8 and 5.2 km. These results
are consistent with observations at the Chicxulub impact crater.
As such, it has been suggested that the ORR and CR of
Orientale are analogous to the crater rim and outer ring of
Chicxulub, respectively [Morgan and Warner, 1999].
6.2. Transient Crater Location and Dimensions
[42] An integral stage in impact crater formation is the
production of a transient crater, the idealized initial cavity
formed at the end of the excavation stage [Gault et al., 1968].
The size (depth, diameter, volume) of the transient crater is used
to determine many fundamental crater characteristics, including
the size of the projectile, the amount of impact melt produced,
and the depth of the final crater [e.g., Collins et al., 2005 and
references therein]. Currently, there is no consensus on the
location of the rim of the transient crater for Orientale, although
the majority of authors interpret the ORR as at least the
approximate location of the transient crater rim [e.g., Head,
1974a; Moore et al., 1974; Scott et al., 1977; Head et al.,
1993; Fassett et al., 2011; Table 3]. Previous transient crater
diameter estimates range from 100 to 620 km [Head, 1974a;
Table 3], and all four major topographic rings have been
suggested to mark the location of the transient crater rim.
However, there is no physical reason to assume that any of
the observed rings should approximate the transient crater rim
[e.g., Howard et al., 1974; Pike and Spudis, 1987].
[43] Because the ORR is a fault scarp of the type produced
during the modification phase of impact cratering, the transient crater must be interior to it. If the CR represents the final crater diameter, then traditional scaling [Schmidt and
Housen, 1987] suggests the transient crater had a diameter
of ~500 km, well within the ORR. If the transient crater extended to the current location of the ORR, the morphology
of the ring would not be controlled by a deep normal fault
with its characteristic footwall topography as our modeling
implies, but instead brecciated material at the location of
the ORR would have collapsed into the transient cavity
floor, and the terrace-like modification of the transient cavity
would have resulted in faulting beyond the present location
of the ORR.
[44] As described above (sections 2.2 and 4.1; Figure 6), the
IRR has a different morphology than the ORR and CR, being
instead composed of massifs. In the Chicxulub structure, that
type of peak ring structure has been associated with the collapse
of the central uplift onto the margins of the transient cavity and a
collision with material collapsing inward from the modification
zone [e.g., Morgan and Warner, 1999; Collins et al., 2002;
Ivanov, 2005; Kring, 2005; Morgan et al., 2011]. If that analogy
applies to Orientale, then the margin of the transient crater had
an approximate diameter greater than 480 km. The distribution
of post-impact mare fill is therefore consistent with the IRR
and the ORR bounding the transient crater. Sharply defined
circumferential deposits (e.g., Lacus Veris at the base of
the ORR) would be produced from magmas propagating
up well-defined faults, as is common on Earth [e.g., Smith
and Bailey, 1968; Pitcher and Bussell, 1977], whereas the more
distributed Mare Orientale erupted through a fractured floor
beneath and breccia fill within the remnants of the transient
cavity (e.g., Figures 9d and 9e).
[45] Here, we propose that the transient crater lies within the
Outer Rook Ring but interior to Lacus Veris, with a diameter of
Table 3. Selected Transient Crater (TC) Diameter DTC (or Corresponding Ring) Estimates for Orientale Basin
DTC (km)
100
134
320 or 480
352
368
397 10, ~400
<400
<450–500
496
500–620
500–550
600–620
600
Equivalent TC Ring
Notes
Reference(s)
IR/IRR
Between IR and IRR; thermal profile 1
Between IR and IRR
Between IR and IRR
Perhaps TC equals IR or IRR
Between IRR and ORR; thermal profile 2
Between IRR and ORR
Inside Lacus Veris, between IRR and ORR
ORR
ORR
IRR
ORR
CR
CR
CR equals the final rims of complex craters [1974]
or the “true crater” [1981]
200
van Dorn [1968]
van Dorn [1969]
Floran and Dence [1976]
Potter et al. [2012]
Dence [1973]
Wieczorek and Phillips [1999], Hikida and Wieczorek
[2007]
Melosh [1989]
Scott et al. [1977]
Potter et al. [2012]
Spudis et al. [1984], Spudis [1993]
This work
Head [1974a, 1977, 2010], Head et al. [1993, 2010],
Fassett et al. [2011]
Moore et al. [1974]
Schultz and Spudis [1978]
McCauley [1977]
Wilhelms et al. [1977], Hodges and Wilhelms [1978],
Wilhelms [1984, 1987]
Baldwin [1974, 1981]
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
500–550 km. Our interpretation of the transient crater diameter
differs from the commonly held view that the ORR marks the
greatest plausible extent of the transient cavity (Table 3). Of
the transient crater size estimates, our results are most compatible with those of Spudis [1993], Spudis et al. [1984], and the
numerical modeling with thermal profile 2 (mantle not reaching
solidus) of Potter et al. [2012] but are notably larger than, for
example, the diameter estimate of ~400 km based on the
Orientale gravity anomaly [Wieczorek and Phillips, 1999;
Hikida and Wieczorek, 2007], the <400 km estimate by
Melosh [1989], or the 352 km numerical model estimate
based on Potter et al.’s [2012] thermal profile 1 (mantle
reaching solidus).
6.3. Crustal Thickness
[46] The crustal thickness in eastern Orientale, derived
from inversion of gravity and topography data, is ~12 km
thinner than in the west, consistent with the regional gradient
of crustal thickness derived from profile A in Figure 2. This
suggests that the difference in crustal thickness may be a
primary structure, existing prior to the impact event and
did not occur as a result of the impact and basin-forming
processes. Thus, this evidence may provide a constraint on
the timing of the formation of the crustal thickness asymmetry, which had to be in place before the estimated time of
Orientale Basin formation: 3.85 Ga [Wilhelms, 1987] to
3.68 Ga [Whitten et al., 2011].
6.4. Relationship between Normal Faults and Maria
[47] Further evidence for normal faulting of the rings comes
in the form of the distribution of locations and ages of mare
deposits within Orientale Basin. The crust in the basin center
has been significantly thinned and fractured due to material
excavation and impact shock, respectively (Figure 2b), favoring
magma ascent there (Figure 9d); thus, the materials in Mare
Orientale erupted earliest and are the oldest, consistent with
published age estimates [Greeley et al., 1993; Whitten et al.,
2011]. The central mare load may have exerted tension on the
lithosphere, causing dilation of the impact-induced ring faults
and created discrete conduits for magma ascent to the surface
[e.g., Hartmann and Kuiper, 1962; McGovern and Litherland,
2011]. Simultaneously, compressional stresses were induced
below the basin, closing off previously utilized conduits
[McGovern and Litherland, 2011], stopping additional magma
from erupting as lava in Mare Orientale. As lava erupted on
the surface along the ORR, more weight was exerted on
the crust, and the zone for favorable magma ascent and
lava eruption may have been diverted to the nearest outer
structural weakness, the CR. This sequence of faulting and
loading from mare fill may account for the decrease in
mare ages from the basin center outward [Greeley et al.,
1993; Whitten et al., 2011] as well as providing support
for our hypothesis that the extent of the transient crater lies
entirely interior to the Lacus Veris mare ponds. Additionally, the crustal thickness asymmetry (Figure 2) may be a
factor in the preferential location of mare materials on the
eastern side of the basin where the crust is ~10–30 km
thinner than on the west, likely influencing fault and ring
formation and, therefore, locations of discrete conduits
available for magma ascent.
6.5. Formation of Basin Rim Topography
[48] The topographic data and modeled structures indicate
that the presently observed CR is reminiscent of a final complex
crater rim that has slumped in a manner broadly similar to the
terrace formation in complex craters. This can be seen, for
example, in profile D in Figure 3, where several normal faults
are present on the southwestern side of Orientale in the location
where individual rings are not easily discerned. We suggest that
this normal faulting, similar in principle to the formation of terraces in complex craters, is sufficient to cause the modification
of the final crater rim into the observed Cordillera Ring and
Outer Rook Rings with accompanying smaller normal
faults without a need to invoke formation of a single
megaterrace far beyond the uplifted rim [Head, 1974a,
1977]. Nevertheless, further faulting beyond the final
topographic rim is possible and, in the case of Orientale,
may have resulted in the formation of the 1300 km
diameter outer ring [e.g., Hartmann and Kuiper, 1962;
Figure 9. Schematic formation scenario for Orientale
Basin. Not to scale. Ejecta and impact melt are neglected
for clarity and to focus on structural evolution of the basin.
(a) Excavation stage showing transient cavity. (b, c) Formation of the IRR from the collapse of central peak and
transient rim. (d) Faulting from collapse of transient cavity,
forming the CR and ORR rings. (e) Magma ascent along
faults and fractures and eruption of lava at the surface.
201
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
Hartmann and Wood, 1971; Pike and Spudis, 1987].
Testing this hypothesis is, however, beyond the scope of
the present study. As the faults forming both the ORR
and the CR scarps are deep (Figure 8), uplift of the elastic
lithosphere was probably an important part of the structural uplift process forming the topographic rings. Thus,
in addition to ejecta, a structural component is also clearly
required to form the observed topography of the ORR
and CR.
[49] In addition, as we have shown above (Figure 6), the
Inner Rook Ring displays topography unlike the normal
fault-dominated ORR and CR (Figures 4 and 5), implying
that the IRR formed by a different mechanism. We consider
this to support the interpretation that IRR is analogous to
peak rings in peak-ring craters [e.g., Head, 1974a, 1977;
McCauley, 1977; Scott et al., 1977; Head et al., 1993] like
that also seen in Chicxulub [Morgan and Warner, 1999;
Collins et al., 2002; Ivanov, 2005; Kring, 2005; Morgan
et al., 2011]. If the IRR (diameter = 480 km) is a peak ring,
then its size relative to that of the CR (930 km) is similar
to the ~0.5 value expected [e.g., Wood and Head, 1976;
Pike and Spudis, 1987; Baker et al., 2011]. Because peak
rings are composed of material uplifted from depth, this
interpretation is fully consistent with the spectral signature of
partially shocked and uplifted plagioclase (anorthosite) in the
IRR [Spudis et al., 1984; Head et al., 1993, 2010; Bussey
and Spudis, 1997, 2000; Ohtake et al., 2009; Pieters et al.,
2009], which is distinct from the noritic spectral signature of
the outer basin.
[50] To summarize, we suggest the following scenario for
the formation of the three main rings of the Orientale Basin
(Figure 9). The IRR is a peak ring formed by the collapse of
the central uplift, with probable contribution from the collapsing
transient cavity rim material. The CR is analogous to final
complex crater rims, with the Cordillera Ring scarp marking
the location of a deep, inwardly dipping, normal fault and the
ring topography being mainly due to ejecta deposition and uplift
induced by the deep faulting.
[51] The ORR is located outside the transient cavity, and the
scarp marks the location of the second major normal fault in the
Orientale Basin. The topography of the ORR is also the result of
ejecta deposition and uplift induced by the deep faulting, with
probable contribution from the rotation of the block between
the ORR and CR. The collapse of the transient cavity took place
mostly as deep normal faults, with additional shallow normal
faults particularly important in the southwestern portion of the
basin where the morphologic and topographic ORR and CR
scarps are difficult to distinguish.
[52] This formation model is consistent with the observed
ring morphology, topography, spectral properties of the rings,
location of the mare and the lacūs, the origin of the Montes
Rook Formation as the deepest ejecta deposited just outside
the transient crater, peak-ring formation models, IRR/CR
diameter ratio, and the results of forward mechanical modeling
of the ORR and CR scarps.
6.6. Comparison with Other Formation Hypotheses
[53] The observed topography of the ORR and CR matches
that expected from ejecta deposition and deep normal faulting
and is inconsistent with features predicted by the tsunami model
[van Dorn, 1968, 1969]. Thus, we consider Baldwin’s [1974]
statement that the CR appears “to be a ripple, not just a ring
fault” to be invalid. In addition, the predicted locations of faults
by the nested melt cavity model [Head, 2010] do not match
observations and modeling and therefore is not considered to
be a viable formation mechanism for large impact basins on
the Moon.
[54] Compared to the ring tectonic-based approach by
Melosh [1989], our model differs by having a notably larger
transient crater diameter (<400 km versus 500–550 km; see
Table 3). If the IRR is a peak ring [e.g., Head, 1974a; Moore
et al., 1974; McCauley, 1977; Solomon and Head, 1980; Head
et al., 1993], a transient crater diameter of ~400 km or less
[Melosh, 1989; Wieczorek and Phillips, 1999; Hikida and
Wieczorek, 2007; Table 3] is in stark contrast with formation
models in which peak rings result from the outward collapse
of the central peak combined with the inward collapse of
the transient cavity [e.g., Morgan and Warner, 1999;
Morgan et al., 2000, 2011; Collins et al., 2002]. If the
transient crater had a diameter of ~400 km or less [Melosh,
1989; Wieczorek and Phillips, 1999; Hikida and Wieczorek,
2007; Table 3], the IRR (diameter = 480 km) could not be a
peak ring formed by the outwardly collapsing central peak
and inwardly collapsing transient cavity model.
[55] Although in many respects broadly similar, the hypothesis presented in this study differs in two important ways
from the megaterrace hypothesis: (1) The dimensions of the
transient crater are different, and the reason for placing it in
this location is that the magma conduits that facilitated the formation of the lacūs are deep faults caused by the collapse of
the transient cavity, and thus, the transient cavity has to have
been inside the lacūs; and (2) the formation processes of
complex crater rims and the Cordillera Ring do not have to
be drastically different. Rather, there is a continuum from
complex crater terraces to basin rings, as shown by the
highly complex faulting in the southwest part of the basin,
transitioning to other parts of the basin where there
typically is at least one major fault accompanied by several
smaller ones.
7. Conclusions
[56] 1. Forward mechanical modeling of ring topography
reveals that the Outer Rook and Cordillera Rings are largescale normal faults with the vertical depth of faulting T from
19 to 37 km, with most faults having T = 30 5 km, fault dip
angle d between 54 and 80 , and fault displacement D
between 0.8 and 5.2 km.
[57] 2. The collapse of the transient cavity rim took place by
normal faulting, either by several smaller normal faults roughly
similar to complex crater terraces as in the southwest part of
the Orientale Basin (e.g., profile D in Figure 4) or mostly by
two deep normal faults forming the ORR and CR scarps, as
shown by forward mechanical modeling. The Cordillera
Ring is morphologically [e.g., Moore et al., 1974; Head,
1977; Wilhelms et al. 1977; Wilhelms, 1987] and topographically similar to final complex crater rims, suggesting a similar
formation mechanism.
[58] 3. The distribution of mare deposits inside Orientale
Basin suggests that the CR and ORR normal faults were used
as conduits for magma ascent subsequent to their formation
during or slightly after the collapse of the transient cavity.
[59] 4. Our results generally agree with the models that
predict normal faulting, such as megaterracing [Hartmann and
202
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
Yale, 1968; Head, 1974a, 1977; Howard et al., 1974;
McCauley, 1977] and ring tectonics [Melosh and McKinnon,
1978; McKinnon and Melosh, 1980; McKinnon, 1981;
Melosh, 1989], although details of these models compared to
ours may differ.
[60] 5. The Inner Rook Ring does not display characteristic
normal fault topography, implying that the IRR formed by a
different mechanism, supporting the interpretation that IRR is
analogous to peak rings in peak-ring craters.
[61] 6. The crustal thickness difference between the eastern
and western sides of the basin is likely preexisting, attributable
to the global crustal thickness asymmetry, and not the result
of basin formation processes. The formation of the hemispheric crustal thickness asymmetry can then be placed
before the formation of Orientale Basin 3.68 to 3.85 Ga.
[62] 7. The transient crater rim was located inside the Outer
Rook Ring and likely had a diameter of ~500–550 km.
[63] Acknowledgments. The authors thank M. Weller (Rice University)
for discussions about fault modeling using Coulomb, R. Schultz (ConocoPhillips)
for detailed discussions regarding faulting behavior, C. Mercer (LPI and USGS
Denver) and P. McGovern (LPI) for discussions about magma ascent
around impact basins, R. Potter (LPI) for discussions regarding hydrocode
models, Y. Ishihara (National Astronomical Observatory of Japan) for
sharing GIS-compatible crustal thickness data, L. Gaddis (USGS), the
USGS ISIS team, and B. Fessler (LPI) for valuable computer support for
ArcMap and ISIS. We are also grateful for the data collected and provided
by the LRO LOLA team. The authors thank Jeff Andrews-Hanna and
Christian Klimczak for thorough and detailed comments that greatly
improved the manuscript. This research was partially funded by NASA
under the LPI Cooperative Agreement NNX08AC28A, NASA Lunar
Science Institute contract NNA09DB33A (PI David A. Kring), and NASA
Outer Planets Research grant NNX09AP33G (PI B. R. Smith-Konter).
This is LPI contribution 1712.
References
Baker, D. M. H., J. W. Head, C. I. Fassett, S. J. Kadish, D. E. Smith, M. T.
Zuber, and G. A. Neumann (2011), The transition from complex crater to
peak-ring basin on the Moon: New observations from the Lunar Orbiter
Laser Altimeter (LOLA) instrument, Icarus, 214, 377–393.
Baldwin, R. B. (1942), The meteoritic origin of the lunar craters, Pop.
Astr., 50, 365–369.
Baldwin, R. B. (1943), The meteoritic origin of lunar structures, Pop. Astr.,
51, 117–127.
Baldwin, R. B. (1949), The Face of the Moon, Univ. Chicago Press,
Chicago.
Baldwin, R. B. (1963), The Measure of the Moon, Univ. Chicago
Press, Chicago.
Baldwin, R. B. (1972), The tsunami model of the origin of ring structures
concentric with large lunar craters, Phys. Earth Planet. Int., 6, 327–339.
Baldwin, R. B. (1974), On the origin of the mare basins, Proc. Lunar Sci.
Conf., 5, 1–10.
Baldwin, R. B. (1981), On the tsunami theory of the origin of multi-ring
basins, in Multi-ring Basins, edited by P. H. Schultz and R. B. Merrill,
Proc. Lunar Planet. Sci., 12A, pp. 275–288, Lunar and Planetary Institute,
Houston, TX.
Brace, W. F. and D. L. Kohlstedt (1980), Limits on lithospheric stress
imposed by laboratory experiments, J. Geophys. Res., 85, 6248–6252.
Bratt, S. R., S. C. Solomon, and J. W. Head (1985), The evolution of impact
basins: Cooling, subsidence, and thermal stress. J. Geophys. Res., 90,
12,415–12,433.
Bürgmann, R., Pollard, D. D., and S. J. Martel (1994), Slip distributions on
faults: Effects of stress gradients, inelastic deformation, heterogeneous
host-rock stiffness, and fault interaction, J. Struct. Geol., 16, 1675–1690.
Bussey, D. B. J. and P. D. Spudis (1997), Compositional analysis of the
Orientale basin using full resolution Clementine data: Some preliminary
results, Geophys. Res. Lett., 24, 445–448.
Bussey, D. B. J. and P. D. Spudis (2000), Compositional studies of the
Orientale, Humorum, Nectaris, and Crisium lunar basins, J. Geophys.
Res., 105, 4235–4243.
Cintala, M. J. and R. A. F. Grieve (1998), Scaling impact melting and crater
dimensions: Implications for the lunar cratering record, Meteorit. Planet.
Sci., 33, 889–912.
Cohen, S. C. (1999), Numerical models of crustal deformation in seismic
zones, Adv. Geophys., 41, 133–231.
Collins, G. S., H. J. Melosh, J. V. Morgan, and M. R. Warner (2002),
Hydrocode simulations of Chicxulub crater collapse and peak-ring formation, Icarus, 157, 24–33, doi:10.1006/icar.2002.6822.
Collins, G. S., H. J. Melosh, and R. A. Marcus (2005), Earth impact effects
program: A web-based computer program for calculating the regional
environmental consequences of a meteoroid impact on Earth, Meteorit.
Planet. Sci., 40(6), 817–840.
Cowie, P. A. and C. H. Scholz (1992), Displacement-length relationships
for faults: Data synthesis and discussion, J. Struct. Geol., 14, 1149–1156.
Croft, S. K. (1980), Cratering flow fields: Implications for the excavation and transient expansion stages of crater formation, Proc. 11th
Lunar Planet. Sci. Conf., 2347–2378, Lunar and Planetary Institute,
Houston, TX.
Dawers, N. H., Anders, M. H., and C. H. Scholz (1993), Growth of normal
faults: Displacement-length scaling, Geology, 21, 1107–1110.
Dence, M. R. (1973), Dimensional analysis of impact structures,
Meteoritics, 8, 343–344.
Dietz, R. S. (1946), The meteoritic impact origin of the Moon’s surface
features, J. Geol., 54, 359–375.
Fassett, C. I., J. W. Head, D. E. Smith, M. T. Zuber, and G. A. Neumann
(2011), Thickness of proximal ejecta from the Orientale Basin from
Lunar Orbiter Laser Altimeter (LOLA) data: Implications for multiring basin formation, Geophys. Res. Lett., 38, L17201, doi:10.1029/
2011GL048502.
Fielder, G. (1961), The grid system and lattice patterns, in Structure of the
Moon’s Surface, pp. 161–196, Pergamon Press, New York, NY.
Floran, R. J. and M. R. Dence (1976), Morphology of the Manicouagan
ring-structure, Quebec, and some comparisons with lunar basins and
craters, Proc. Lunar Sci. Conf., 7, 2845–2865.
Gault, D. E., W. L. Quaide, and V. R. Oberbeck (1968), Impact cratering mechanics and structures, in Shock Metamorphism of Natural
Materials, edited by B. M. French and N. M. Short, pp. 87–99, Mono
Book Corp., Baltimore, MD.
Gilbert, G. K. (1893), The Moon’s face: A study of the origin of its features,
Bull. Phil. Soc. Wash., 12, 241–292.
Greeley, R., et al. (1993), Galileo imaging observations of lunar maria and
related deposits, J. Geophys. Res., 98, 17,180–17,205.
Grott, M., E. Hauber, S. C. Werner, P. Kronberg, G. Neukum (2007),
Mechanical modeling of thrust faults in the Thaumasia region, Mars,
and implications for the Noachian heat flux, Icarus, 186, 517–526,
doi:10.1016/j.icarus.2006.10.001.
Hartmann, W. K. and G. P. Kuiper (1962), Concentric structures surrounding lunar basins, Comm. Lunar Planet. Lab., 1, 51–66.
Hartmann, W. K. and C. A. Wood (1971), Moon: Origin and evolution of
multi-ring basins, The Moon, 3, 3–78.
Hartmann, W. K. and F. G. Yale (1968), Lunar crater counts IV: Mare
Orientale and its basin system, Comm. Lunar Planet. Lab., 7, 131–137.
Hawke, B. R., C. A. Peterson, D. T. Blewett, D. B. J. Bussey, P. G. Lucey,
G. J. Taylor, and P. D. Spudis (2003), Distribution and modes of occurrence of lunar anorthosite, J. Geophys. Res., 108, 5050, doi:10.1029/
2002JE001890.
Head, J. W. (1974a), Orientale multi-ringed basin interior and implications for the petrogenesis of lunar highland samples, The Moon, 11,
327–356.
Head, J. W. (1974b), Morphology and structure of the Taurus-Littrow
Highlands (Apollo 17): Evidence for their origin and evolution, The
Moon, 9, 355–395.
Head, J. W. (1977), Origin of outer rings in lunar multi-ringed basins—
Evidence from morphology and ring spacing, in Impact and Explosion
Cratering: Planetary and Terrestrial Implications, edited by D. J.
Roddy, R. O. Pepin, and R. B. Merrill, pp. 563–573, Pergamon Press,
Inc., New York.
Head, J. W. (2010), Transition from complex craters to multi-ringed basins
on terrestrial planetary bodies: Scale dependent role of the expanding melt
cavity and progressive interaction with the displaced zone, Geophys. Res.
Lett., 37, L02203, doi:10.1029/2009GL041790.
Head, J. W. and S. C. Solomon (1980), Lunar basin structure: Possible
influence of variations in lithospheric thickness, Lunar Planet. Sci. Conf.
11th, 421–423, Lunar and Planetary Institute, Houston, TX.
Head, J. W., S. Murchie, J. F. Mustard, C. M. Pieters, G. Neukum,
A. McEwen, R. Greeley, E. Nagel, and M. J. S. Belton (1993),
Lunar impact basins: New data for the western limb and far side
(Orientale and South Pole-Aitken Basins) from the First Galileo
Flyby, J. Geophys. Res., 98, 17,149–17,181.
Head, J. W., et al. (2010), The lunar Orientale Basin: Structure and crustal
mineralogy from Chandrayaan-1 Moon Mineralogy Mapper (M3) data,
Lunar Planet. Sci. Conf. 41st, Abs. 1030, Lunar and Planetary Institute,
Houston, TX.
203
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
Hiesinger, H., J. W. Head III, U. Wolf, R. Jaumann, and G. Neukum (2011),
Ages and stratigraphy of lunar mare basalts: A synthesis, in Recent
Advances and Current Research Issues in Lunar Stratigraphy, edited by
W. A. Ambrose and D. A. Williams, Spec. Pap. Geol. Soc. Am., 447,
pp. 1–51, Geol. Soc. Am., Boulder, CO.
Hikida, H. and M. A. Wieczorek (2007), Crustal thickness of the Moon:
New constraints from gravity inversions using polyhedral shape models,
Icarus, 192, 150–166, doi:10.1016/j.icarus.2007.06.015.
Hodges, C. A. and D. E. Wilhelms (1978), Formation of lunar basin rings,
Icarus, 34, 294–323.
Howard, K. A., D. E. Wilhelms, and D. H. Scott (1974), Lunar basin
formation and highland stratigraphy, Rev. Geophys. Space Phys.,
12, 309–327.
Ishihara, Y., S. Goossens, K. Matsumoto, H. Noda, H. Araki, N. Namiki,
H. Hanada, T. Iwata, S. Tazawa, and S. Sasaki (2009), Crustal thickness
of the Moon, Implications for farside basin structures, Geophys. Res.
Lett., 36, L19202, doi:10.1029/2009GL039708.
Ivanov, B. A. (2005), Numerical modeling of the largest terrestrial meteorite
craters, Sol. Syst. Res., 39, 381–409, doi: 10.1007/s11208-005-0051-0.
Jackson, J. and D. McKenzie (1983), The geometrical evolution of normal
fault systems, J. Struct. Geol., 5, 471–482, doi: 10.1016/0191-8141(83)
90053-6.
Key, W. R. O. and R. A. Schultz (2011), Fault formation in porous sedimentary rocks at high strain rates: First results from the Upheaval Dome
impact structure, Utah, USA, Geol. Soc. Am. Bull., 123, 1161–1170,
doi:10.1130/B30087.
King, G. C. P., R. Stein, and J. Rundle (1988), The growth of geological
structures by repeated earthquakes 1. Conceptual framework, J. Geophys.
Res., 93, 13,307–13,318, doi:10.1029/JB093iB11p13307.
Kring, D. A. (1995), The dimensions of the Chicxulub impact crater and
impact melt sheet, J. Geophys. Res., 100, 16,979–16,986, doi:10.1029/
95JE01768.
Kring, D. A. (2005), Hypervelocity collisions into continental crust
composed of sediments and an underlying crystalline basement:
Comparing the Ries (~24 km) and Chicxulub (~180 km) impact
craters, Chem. Erde Geochem., 65, 1–46.
Lin, J. and R. S. Stein (2004), Stress triggering in thrust and subduction
earthquakes and stress interaction between the southern San Andreas
and nearby thrust and strike-slip faults, J. Geophys. Res., 109, B02303,
doi:10.1029/2003JB002607.
Mackin, J. H. (1969), Origin of lunar maria, Bull. Geol. Soc. Am., 80, 735–747.
Manighetti, I., Campillo, M., Sammis, C., Mai, P. M., and G. King (2005),
Evidence for self-similar, triangular slip distributions on earthquakes:
Implications for earthquake and fault mechanics, J. Geophys.
Res., 110, B05302, doi:10.1029/2004JB003174.
McCauley, J. F. (1968), Geologic results from the lunar precursor probes,
Am. Inst. Aeronaut. Astronaut. J., 6, 1991–1996.
McCauley, J. F. (1977), Orientale and Caloris, Phys. Earth Planet. Inter.,
15, 220–250.
McGovern, P. J. and M. M. Litherland (2011), Lithospheric stress and basaltic
magma ascent on the Moon, with implications for large volcanic provinces
and edifices, Lunar Planet. Sci. Conf. 42nd, Abs. 2587, The Woodlands, TX,
March 7–11, Lunar and Planetary Institute, Houston, TX.
McKinnon, W. B. (1981), Application of ring tectonic theory to Mercury
and other solar system bodies, in Multi-ring Basins, edited by P. H.
Schultz and R. B. Merrill, Proc. Lunar Planet. Sci., 12A, pp. 259–273,
Lunar and Planetary Institute, Houston, TX.
McKinnon, W. B. and H. J. Melosh (1980), Evolution of planetary
lithospheres: Evidence from multiringed structures on Ganymede and
Callisto, Icarus, 44, 454–471.
Melosh, H. J. (1989), Impact cratering: A geologic process, Oxford University
Press, New York.
Melosh, H. J. and W. B. McKinnon (1978), The mechanics of ringed basin
formation, Geophys. Res. Lett., 5, 985–988.
Melosh, H. J. and B. A. Ivanov (1999), Impact crater collapse, Annu. Rev.
Earth Planet. Sci., 27, 385–415.
Moore, H. J., C. A. Hodges, and D. H. Scott (1974), Multiringed
basins–illustrated by Orientale and associated features, Proc. 5th
Lunar Conf., 71–100, Lunar and Planetary Institute, Houston, TX.
Morgan, J. and M. Warner (1999), Chicxulub: The third dimension of a
multi-ring impact basin, Geology, 27, 407–410.
Morgan, J. V., et al. (1997), Size and morphology of the Chicxulub impact
crater, Nature, 390, 472–476.
Morgan, J. V., M. R. Warner, G. S. Collins, H. J. Melosh, and G. L. Christeson
(2000), Peak ring formation in large impact craters, Earth Planet. Sci.
Lett., 183, 247–354.
Morgan, J. V., M. R. Warner, G. S. Collins, R. A. F. Grieve, G. L. Christeson,
S. P. S. Gulick, and P. J. Barton (2011), Full waveform tomographic images
of the peak ring at the Chicxulub impact crater, J. Geophys. Res., 116,
B06303, doi:10.1029/2010JB008015.
Muller, J. R. and A. Aydin (2005), Using mechanical modeling to constrain
fault geometries proposed for the northern Marmara Sea, J. Geophys.
Res., 110, B03407, doi:10.1029/2004JB003226.
Murray, J. B. (1980), Oscillating peak model of basin and crater formation,
Moon Planet., 22, 269–291.
National Research Council (2007), The Scientific Context for the Exploration of
the Moon, National Academies Press, Washington, DC.
Ohtake, M., et al. (2009), The global distribution of pure anorthosite on the
Moon, Nature, 461, 236–240, doi:10.1038/nature08317.
O’Keefe, J. D. and T. J. Ahrens (1999), Planetary cratering mechanics,
J. Geophys. Res., 98, 17,001–17,028.
Okada, Y. (1992), Internal deformation due to shear and tensile faults in a
half-space, Bull. Seismol. Soc. Am., 82, 1018–1040.
Pieters, C. M., et al. (2009), Mineralogy of the lunar crust in spatial context:
First results from the Moon Mineralogy Mapper (M3), Lunar Planet. Sci.
Conf. 40th, Abs. 2052, Lunar and Planetary Institute, Houston, TX.
Pike, R. J. and P. D. Spudis (1987), Basin-ring spacing on the Moon,
Mercury, and Mars, Earth Moon Planets, 39, 129–194.
Pitcher, W. S. and M. A. Bussell (1977), Structural control of batholith
emplacement in Peru: A review, J. Geol. Soc. London, 133, 249–256.
Pollard, D. D. and P. Segall (1987), Theoretical displacements and
stresses near fractures in rock: With applications to faults, joints, veins,
dikes, and solution surfaces, in Fracture Mechanics of Rocks, edited by
B. K. Atkinson, Academic Press Geology Series, Academic Press,
London, U. K.
Potter, R. W. K., Kring, D. A., Collins, G. S., Kiefer, W. S., and McGovern,
P. J. (2012), Estimating transient crater size using the crustal annular
bulge: Insights from numerical modeling of lunar basin-scale impacts,
Geophys. Res. Lett. 39, L18203, doi: 10.1029/2012GL052981.
Resor, P. G. (2008), Deformation associated with a continental normal fault system, western Grand Canyon, Arizona, Geol. Soc. Am. Bull., 120, 414–430.
Schmidt, R. M. and K. R. Housen (1987), Some recent advances in
the scaling of impact and explosion cratering, Int. J. Impact Eng.,
5, 543–560.
Schultz, P. H. and P. D. Spudis (1978), The dark ring of Orientale: Implications
for pre-basin mare volcanism and a clue to the identification of the transient
cavity rim, Lunar Planet. Sci. Conf. 9, 1033–1035, Lunar and Planetary
Institute, Houston, TX.
Schultz, R. A. (1995), Limits on strength and deformation properties of
jointed basaltic rock masses, Rock Mech. Rock Eng., 28, 1–15,
doi:10.1007/BF01024770.
Schultz, R. A. (1996), Relative scale and the strength and deformability of
rock masses, J. Struct. Geol., 18, 1139–1149, doi:10.1016/0191-8141
(96)00045-4.
Schultz, R. A. (2000), Localization of bedding plane slip and backthrust faults
above blind thrust faults: Keys to wrinkle ridge structure, J. Geophys. Res.,
105, 12,035–12,052, doi:10.1029/1999JE001212.
Schultz, R. A. and J. Lin (2001), Three-dimensional normal faulting models
of the Valles Marineris, Mars, and geodynamic implications, J. Geophys.
Res., 106, 16,549–16,566, doi:10.1029/2001JB000378.
Schultz, R. A. and T. R. Watters (2001), Forward mechanical modeling of
the Amenthes Rupes thrust fault on Mars, Geophys. Res. Lett.,
28, 4659–4662.
Schultz, R. A., R. Soliva, C. H. Okubo, and D. Mège (2010), Fault populations, in Planetary Tectonics, edited by T. R. Watters and R. A. Schultz,
pp. 457–510, Cambridge University Press, Cambridge, U. K.
Scott, D. H., J. F. McCauley, and M. N. West (1977), Geologic map of the
west side of the Moon, scale 1:5,000,000, Misc. Inv. Ser. Map I–1034,
U. S. Geol. Surv., Denver, CO.
Smith, R. L. and R. A. Bailey (1968), Resurgent cauldrons, Geol. Soc. Am.
Mem., 116, 613–662.
Smith, D. E., et al. (2010), The Lunar Orbiter Laser Altimeter investigation on the Lunar Reconnaissance Orbiter mission, Space Sci.
Rev., 150, 209–241.
Solomon, S. C. and J. W. Head (1980), Lunar mascon basins: Lava filling,
tectonics, and evolution of the lithosphere, Rev. Geophys. Space Phys.,
18, 107–141.
Spudis, P. D. (1993), The geology of multi-ring impact basins: The Moon
and other planets, Cambridge Univ. Press, Cambridge, U. K.
Spudis, P. D. and P. A. Davis (1986), A chemical and petrological
model of the lunar crust and implications for lunar crustal origin,
Proc. 17th Lunar Planet. Sci. Conf., Part 1, J. Geophys. Res.,
91, suppl., E84–E90.
Spudis, P. D., B. R. Hawke, and P. Lucey (1984), Composition of Orientale
basin deposits and implications for the lunar basin-forming process, Proc.
15th Lunar Planet. Sci. Conf., J. Geophys. Res., 89, C197–C210.
Toda, S., R. Stein, P. Reasenberg, J. Dieterich, and A. Yoshida (1998),
Stress transferred by the 1995 Mo = 6.9 Kobe, Japan, shock: Effect on
aftershocks and future earthquake probabilities, J. Geophys. Res.,
103, 24,543–24,565, doi:10.1029/98JB00765.
204
NAHM ET AL.: FAULTING IN ORIENTALE BASIN
Toda, S., R. S. Stein, K. Richards-Dinger, and S. B. Bozkurt (2005),
Forecasting the evolution of seismicity in southern California: Animations built on earthquake stress transfer, J. Geophys. Res., 110, B05S16,
doi:10.1029/2004JB003415.
Turcotte, D. L. and G. Schubert (2002), Geodynamics, Cambridge Univ.
Press, Cambridge, U. K.
van Dorn, W. G. (1968), Tsunamis on the Moon?, Nature, 220, 1102–1107.
van Dorn, W. G. (1969), Lunar maria: Structure and evolution, Science,
165, 693–695.
Watters, T. R., R. A. Schultz, M. S. Robinson, and A. C. Cook (2002), The
mechanical and thermal structure of Mercury’s early lithosphere,
Geophys. Res. Lett., 29(11), 1542, doi:10.1029/2001GL014308.
Watters, T. R., P. C. Thomas, and M. S. Robinson (2011), Thrust faults and
the near-surface strength of asteroid 433 Eros, Geophys. Res. Lett.,
38, L02202, doi:10.1029/2010GL045302.
Weissel, J. K. and G. D. Karner (1989), Flexural uplift of rift flanks due to
mechanical unloading of the lithosphere during extension, J. Geophys.
Res., 94, 13,919–13,950.
Whitten, J., J. W. Head, M. Staid, C. M. Pieters, J. Mustard, R. Clark, J.
Nettles, R. L. Klima, and L. Taylor (2011), Lunar mare deposits
associated with the Orientale impact basin: New insights into mineralogy, history, mode of emplacement, and relation to Orientale
Basin evolution from Moon Mineralogy Mapper (M3) data from
Chandrayaan-1, J. Geophys. Res., 116, E00G9, doi:10.1029/
2010JE003736.
Wieczorek, M. A. and R. J. Phillips (1999), Lunar multiring basins and the
cratering process, Icarus, 139, 246–259.
Wilhelms, D. E. (1984), Moon, in The Geology of the Terrestrial Planets,
NASA SP-469, edited by M. H. Carr, pp. 107–205, Washington, DC.
Wilhelms, D. E. (1987), The geologic history of the Moon, U.S. Geol. Surv.
Prof. Pap., 1348.
Wilhelms, D. E., C. A. Hodges, and R. J. Pike (1977), Nested-crater model
of lunar ringed basins, in Impact and Explosion Cratering: Planetary and
Terrestrial Implications, edited by D. J. Roddy, R. O. Pepin, and R. B.
Merrill, pp. 539–562, Pergamon Press, Inc., New York.
Willemse, E. J. M., Pollard, D. D., and A. Aydin (1996), Three-dimensional
analyses of slip distributions on normal fault arrays with consequences for
fault scaling, J. Struct. Geol., 18, 295–309.
Wood, C. A. and J. W. Head (1976), Comparison of impact basins
on Mercury, Mars and the Moon, Proc. Lunar Planet. Sci. Conf.
7 th , 3629–3651, Lunar and Planetary Institute, Houston, TX.
Wood, C. A. and M. J. S. Collins (2011), New light on old basins, Lunar
Planet. Sci. Conf. 42nd, Abs. 1314, The Woodlands, TX, March 7–11,
Lunar and Planetary Institute, Houston, TX.
205