Structural geology of impact craters

Abstract

The formation of impact craters is a highly dynamic and complex process that subjects the impacted target rocks to numerous types of deformation mechanisms. Understanding and interpreting these styles of micro-, meso- and macroscale deformation has proved itself challenging for the field of structural geology. In this paper, we give an overview of the structural inventory found in craters of all size ranges on Earth, and look into the structures of craters on other planetary bodies. Structural features are discussed here that are caused by i) extremely high pressures and temperatures that occur during the initial passage of the shock wave through the target rock and projectile, ii) the resulting flow field in the target that excavates and ejects rock materials, and iii) the gravitationally induced modification of the crater cavity into the final crater form. A special focus is put on the effects that low-angle impacting bodies have on crater formation. We hope that this review will help both planetary scientists and structural geologists understand the deformation processes and resulting structures generated by meteorite impact.

Full text

Review article
Structural geology of impact craters
Thomas Kenkmann
*
, Michael H. Poelchau, Gerwin Wulf
Institut für Geo- und Umweltnaturwissenschaften eGeologie, Albert-Ludwigs-Universität Freiburg, Albertstraße 23-B, D-79104 Freiburg, Germany
article info
Article history:
Received 22 May 2013
Received in revised form
9 January 2014
Accepted 25 January 2014
Available online 13 February 2014
Keywords:
Impact cratering
Simple crater
Complex crater
Shock metamorphism
Oblique impact
abstract
The formation of impact craters is a highly dynamic and complex process that subjects the impacted
target rocks to numerous types of deformation mechanisms. Understanding and interpreting these styles
of micro-, meso- and macroscale deformation has proved itself challenging for the field of structural
geology. In this paper, we give an overview of the structural inventory found in craters of all size ranges
on Earth, and look into the structures of craters on other planetary bodies. Structural features are dis-
cussed here that are caused by i) extremely high pressures and temperatures that occur during the initial
passage of the shock wave through the target rock and projectile, ii) the resulting flow field in the target
that excavates and ejects rock materials, and iii) the gravitationally induced modification of the crater
cavity into the final crater form. A special focus is put on the effects that low-angle impacting bodies have
on crater formation. We hope that this review will help both planetary scientists and structural geolo-
gists understand the deformation processes and resulting structures generated by meteorite impact.
Ó2014 Published by Elsevier Ltd.
1. Introduction
The heavily cratered surfaces of almost all solid planetary
bodies in the solar system emphasize the important role
hypervelocity impacts have played in the formation and subse-
quent evolution of planets and satellites. The present structure of
planetary crusts has been influenced by collision processes.
Hypervelocity impacts also pose a threat to human civilization.
The 15-m meteorite impact crater, formed 2007 in Carancas, Peru,
(Kenkmann et al., 2009) or the recent encounter of a 19 m bolide
that exploded in the atmosphere near Chelyabinsk, Russia on Feb.
15, 2013 (Popova et al., 2013) together with the near flyby of
asteroid 2012 DA14 on the same day caused a worldwide sensa-
tion, and indicate the continued threat we face from impact
events.
The number of impact craters discovered on Earth so far con-
tinues to increase; currently 184 are known (Earth impact data
base:http://www.passc.net/EarthImpactDatabase). It is assumed
that hundreds to thousands of impact craters are still undetected
due to their poor state of morphological preservation and their
destruction via erosion. Together with craters that have been
identified, these buried or partly eroded crater structures may
significantly contribute to the overall structure of the earth’s crust.
The recognition of these hidden impact structures requires
knowledge of the structural inventory of impact craters in general.
To this day, only a few structural geologists have focused on this
fascinating topic that also bears a strong economic potential. This
article is aimed at stimulating the exchange of knowledge between
impact crater research and planetology on one hand, and structural
geology on the other hand. In this review article we present
deformation features and structures that are characteristic for
impact craters of various sizes, with a focus on the macro-scale.
Specific attention is drawn to structures that result from oblique
incidences. Many structural aspects presented here are only briefly
discussed.
The understanding of impact processes has significantly
improved over the last several decades. A strong contribution came
from numerical simulations in combination with advancements in
computational power. The hydrocodes in use today are capable of
accounting for 3D-geometries, multi-material targets, strength,
strain localization, fragmentation and particleeatmosphere in-
teractions, to name some examples (e.g., Wünnemann and Ivanov,
2003; Collins et al., 2004; Pierazzo et al., 2008; Elbeshausen et al.,
2009; Senft and Stewart, 2009; Artemieva and Pierazzo, 2011).
Likewise, achievements in experimental techniques and the real-
time recording of high-speed phenomena that provide ground
truth data (e.g., Hugoniot data, equations of state, and constitutive
properties) are important for numerical modeling of impact phe-
nomena (e.g., Holsapple, 1993; Anderson et al., 2003; Kraus et al.,
2010; Kenkmann et al., 2011a). Research on impact crater forma-
tion has also stimulated the structural, microstructural, and
*Corresponding author.
E-mail address: Thomas.kenkmann@geologie.uni-freiburg.de (T. Kenkmann).
Contents lists available at ScienceDirect
Journal of Structural Geology
journal homepage: www.elsevier.com/locate/jsg
http://dx.doi.org/10.1016/j.jsg.2014.01.015
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Journal of Structural Geology 62 (2014) 156e182

petrographical analysis of terrestrial craters (e.g., Grieve, 1991;
Osinski and Spray, 2005; French and Koeberl, 2010; Kenkmann
et al., 2011b). Complementary to the study of terrestrial craters,
there are numerous new spacecraft missions, for example, to Mars
or the Moon, and to the icy satellites of Jupiter and Saturn that
continue to deliver high-resolution remote-sensing images of
impact craters (e.g., McEwen et al., 2007; Robinson et al., 2010;
Wulf et al., 2012; Mouginis-Mark and Boyce, 2012). These data
provide new constraints on crater formation for a variety of target
materials and other boundary conditions, such as surface gravity.
Impact craters display a large variety of deformation features
that are formed under very different and rapidly changing
boundary conditions (Fig. 1). Maximum stresses and strain rates are
reached at the very beginning of an impact process when a shock
wave passes through the target rocks. During the collision of a
kilometer-sized asteroid, pressures may vary over eight orders of
magnitude, from several hundreds of gigapascals to a few kilo-
pascals during late stage movements (O’Keefe and Ahrens, 1975;
Collins et al., 2005). Temperatures range over five to six orders,
from several ten-thousand degrees near the point of impact to
room temperature, within a short period of time (Collins et al.,
2005). Excluding long term relaxations of impact craters, strain
rates may vary from over 10
þ6
s
1
during shock loading to 10
1
s
1
(Huffman and Reimold, 1996). The duration of deformation ranges
from microseconds during shock loading to minutes when gravity-
driven mass movements take place (Fig. 1). Because of the magni-
tudes of strain rates and the duration of deformation, deformation
mechanisms are generally ruled out in which time-dependent
diffusion or solution/precipitation processes are involved, e.g.,
pressure solution creep or diffusion and dislocation creep. Likewise,
multi-incremental crack-seal processes typical for upper crustal
tectonic shear zones do not occur in the context of impact cratering.
Early deformation in impact cratersistheresultofrapidloading
and unloading of rocks by shock waves, causing shock-metamorphic
effects. The volume of rock that is altered by shock metamorphism is
concentrated in the center of the structure and comprises a fraction of
the total crater volume that depends primarily on the impact energy.
However, a large portion of rocks in impact craters shows no signs of
shock metamorphism and plastic deformation, but is deformed at the
sub-shock level by brittle mechanisms. Equally important to the
deformation that is related to the passage of the shock wave is the
subsequent deformation that results from rapid movements associ-
ated with the outward ejection of rock and the gravity-driven collapse
of the crater cavity.
In its organization this article follows the common subdivision
of the impact process into different stages, (i) the contact and
compression stage, (ii) the excavation stage, and (iii) the modifi-
cation stage (Gault et al., 1968). Although the formation of an
impact crater is a continuous and very rapid process, this subdivi-
sion into three stages has proven very useful to distinguish different
processes of cratering that occur in succession. After briefly out-
lining the physical background we assign deformation features to
the different stages with a focus on the modification stage of cra-
tering. Special attention is given to the effects of oblique impacts in
each phase.
2. The contact and compression stage
2.1. Physical background
The contact and compression stage is the first and most brief
phase of the three stages of impact cratering and comprises the
period from the initial contact of an impacting body (referred to as
the “projectile”) with the ground (referred to as “target”) until the
compression ends. During the contact and compression phase, the
kinetic energy of the impacting projectile is instantaneously par-
titioned into internal energies of projectile and target and
remaining kinetic energies of both the projectile and target. This
conversion results in the generation of shock waves that develop at
the interface of the projectile and target. Shock waves are charac-
terized by an abrupt, nearly discontinuous change in pressure,
temperature and density. They travel through media at a higher
Fig. 1. A comparison of endogenic tectonic processes and impact cratering. A) Strain
rate is plotted against the duration of tectonic and cratering processes. Highest strain
rates occur during impact within the shock wave but have a duration of less than 1 s
even for large-scale impacts. The total cratering process, including the gravitationally
induced formation of the central uplift, is typically over within a few tens of seconds to
minutes. B) PeT diagram comparing regional metamorphism and shock meta-
morphism. During shock compression, rocks experience much higher pressures of up
to several hundreds of gigapascals and temperatures of over 10,000 C near the point
of impact. Hugoniot curves are based on the shock behavior of single-crystal quartz
and sandstone (Modified from Stöffler and Langenhorst, 1994). Note that for a given
shock pressure, temperatures are much higher in porous rocks. The lack of overlap
between regional metamorphism and shock metamorphism allows the unambiguous
identification of impact structures.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 157

T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182158

speed than an elastic wave, induce a flow of the media they have
traversed, cause irreversible deformation within the media that is
known as shock metamorphism, and heat the material. Mass,
momentum, and energy are conserved during the transition from
an unshocked to a shocked state (HugonioteRankine equations),
but not entropy. An equation of state relates pressure, specific
volume, and the internal energy for each material.
The amplitude of the shock decays with the distance traveled by
the shock wave due to geometric spreading and consumption of
energy and usually follows a power-law decay curve, as shown by
dimensionless analysis by Holsapple (1987). The degree of atten-
uation is material-dependent, e.g., in porous rocks the shock wave
magnitude decays faster than in low-porosity rocks. Behind the
shock wave, a rarefaction wave generated by reflection of the shock
wave at the free surface (i.e., the rear of the projectile and the air-
target interface), releases the compressed material from its high-
pressure state and finishes the first phase of impact cratering.
Shock wave compression is irreversible, whereas decompression is
reversible; hence, the passage of the shock wave results in a net
increase in temperature (internal energy and entropy) and a re-
sidual particle velocity in the rocks (Melosh, 1989). The duration of
the contact and compression stage depends on the diameter of the
projectile and the shock wave velocity. Even for mega-impacts such
as the Chicxulub event this period may last only for several tenths
of a second (Fig. 1).
2.2. Shock metamorphism
The Hugoniot elastic limit (HEL) represents the maximum stress
a rock can withstand under shock compression without plastic
deformation and internal rearrangement. Above this limit rocks
undergo a shock-metamorphic overprint; below this limit brittle
deformation dominates (Fig. 2A). Shock-deformation effects in
minerals and rocks allow the unambiguous identification of impact
craters even if the craters are morphologically degraded by erosion,
as there is no other natural process capable of producing certain
types of shock features. Strongly shocked minerals and rocks
(>25 GPa) are affected by the subsequent excavation flow in such a
way that they are displaced and found as allochthonous material
either as a part of the ejecta blanket outside the crater, as crater fill
material, or as injected material in the crater floor. The highest
shock levels recorded beneath the crater floor usually do not exceed
25 GPa and occur in the central parts of the crater. In case of larger,
complex impact craters this region is usually uplifted (“central
uplift”).
Except for shatter cones (Fig. 2B), all other unambiguous in-
dicators for shock metamorphism are visible only at the micro-
scopic scale. Shock effects in major rock-forming minerals quartz
(e.g., Short, 1968; Hörz, 1968; Stöffler, 1972; Langenhorst and
Deutsch, 1994; Langenhorst, 2002), feldspar (Ostertag, 1983;
Dressler, 1990), and olivine (Reimold and Stöffler, 1978; Stöffler
et al., 1991; Schmitt, 2000) have been systematically shock-
calibrated by means of planar shock recovery experiments and
can be used as piezometers to reconstruct shock pressure isobars if
rocks are devoid of pore space (Table 1). Low shock pressures are
indicated by shatter cones (w1e10 GPa, Fig 2B), planar fractures
(PFs, Fig. 2C) in quartz (w5e10 GPa), and feather features (FFs,
Fig. 2C) (Poelchau and Kenkmann, 2011), which all form through
dynamic fragmentation processes.
Planar deformation features (PDFs, Fig. 2D and E) are probably
the most important shock criteria. They are narrow-spaced parallel
sets of amorphous lamellae oriented along specific crystallographic
directions. In quartz they develop at a shock pressure interval
ranging from 5-10 GPae35 GPa (Hörz, 1968; Stöffler and
Langenhorst, 1994). The width of the lamellae increases with the
shock magnitude so that above 35 GPa entire grains are trans-
formed into diaplectic glass (Fig. 2F). Diaplectic glasses are the
result of solid-state amorphization. They form pseudomorphs after
their original crystals and do not show indications of liquid flow.
They have slightly higher refractivity indices than standard glass.
Diaplectic quartz glass develops in a shock interval of 35e60 GPa,
while diaplectic feldspar glass, named maskelynite, occurs in the
shock pressure interval 35e45 GPa (Table 1).
Phase transformations in SiO
2
are usually closely linked with the
formation of diaplectic glasses. Stishovite nuclei grow within
amorphous PDF lamellae above w13 GPa and coesite forms glob-
ular clusters (Fig. 2F) within diaplectic quartz glass at 35 GPa
(Stöffler, 1972). At even higher shock pressures true melting occurs
(Fig. 2G) with the development of liquid flow features like schlieren
textures (lechatelierite) (Fig. 2H) occuring upon pressure release.
Vaporization of rock forming minerals occurs roughly above
w100 GPa ( Table 1). A detailed description of shock metamorphism
is beyond the scope of this article. For comprehensive reviews of
the progressive stages of shock metamorphism we refer to Stöffler
and Langenhorst (1994), French (1998), and French and Koeberl
(2010). The shock pressure calibration applies for dense materials
but not for porous rocks. Porosity leads to pore space crushing and
localized shock pressure excursions with magnitudes as high as
four times as the average shock pressure (Güldemeister et al., 2013).
For a shock classification scheme in sandstone we refer to Kieffer
et al. (1976).
2.3. Effects of oblique impact angles on shock distribution and
deformation
The volume of shocked rocks in relation to the crater volume
depends predominantly on impact energy and the target lithology.
The impact angle is of further importance, which also affects the
distribution of shocked rocks and the peak pressure. Peak pressure
and cratering efficiency (the ratio of excavated crater mass to the
initial projectile mass) scale with the vertical component of the
impact velocity and display a sinusoidal decrease with increasing
impact obliquity (Fig. 3b; Chapman and McKinnon, 1986;
Elbeshausen et al., 2009; Davison et al., 2012) if the friction coef-
ficient of the target material is not significantly lower than 0.7.
Three-dimensional numerical simulations of crater formation
(Pierazzo and Melosh, 2000) show that for oblique impacts the
shock isobars concentrate in the downrange portion of the target,
which represents the sector of the crater which is located along
the projection of the projectile’s trajectory into the ground
Fig. 2. Shock-metamorphic features and their corresponding shock stages. A) Close-spaced fracture network in quartzite of the Matt Wilson crater, Australia. Such fracture net-
works, often displaying stair-stepping displacements, form at pressures below shock metamorphism and are typical for craters but provide no evidence for an impact. B) Shatter
cone from the Steinheim crater, Germany. C) Micrograph of quartz from a sandstone in the Matt Wilson structure, Australia, exhibiting a set of planar fractures and feather features.
D) Micrograph of shocked quartz in a granitic clast from suevite of the Nördlinger Ries crater, Germany. Multiple sets of planar deformation features (PDFs) along with a planar
fracture and feather features are visible. Crossed polarizers. E) TEM image of shocked quartz from a shock-recovery experiment into quartzite. The quartz grain shows amorphous
PDF lamellae. F) Coesite crystals and diaplectic glass in a suevite sample from the Nördlinger Ries, Germany. G) EDX element mapping superimposed on SEM-secondary electron
image. The image shows the intensive intermingling of melted iron projectile with welded and comminuted sandstone particles (image courtesy of Tobias Salge). The sample is from
a cratering experiment with a 1 cm steel projectile impacting Seeberger sandstone at 5300 ms
1
(Kenkmann et al., 2011). H) Melted quartz (lechatelierite) with schlieren textures
and vesicles. Suevite sample from Seelbronn, Ries.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 159

(Fig. 3A). These numerical results are supported by cratering ex-
periments that demonstrate that shock-induced damage beneath
craters formed by oblique impacts is stronger in the downrange
direction than in up range direction (Ai and Ahrens, 2005). Like-
wise, the uneven distribution of impact melt in impact structures
like the Ries crater (Stöffler et al., 2002) was interpreted to be an
effect of an oblique impact. Stress wave measurements in oblique
impact experiments showed that the magnitude of peak stress is
about twice as large in the target in the downrange direction (Dahl
and Schultz, 2001). Elevated peak stresses also suggest enhanced
damage in the downrange region. For very low-angle impacts
frictional shear heating is strongly enhanced along the projectile/
target interface and increases the amount of vaporization (Schultz,
1996).
3. The excavation stage
3.1. Physical background
The excavation stage includes both the expansion and dissipa-
tion of the shock wave and the opening of the crater cavity. As the
shock wave propagates hemispherically from the initial contact
area, the engulfed material is compressed and accelerated (Fig. 4).
Upon unloading the material is not completely decelerated and a
residual velocity remains (Melosh, 1989). This residual velocity
plays a crucial role in impact crater excavation and is the most
important aspect that distinguishes hypervelocity impacts from
low velocity impacts. The value of this final particle velocity
component typically corresponds to between one-third to one-fifth
Table 1
Overview of shock-metamorphic features and the shock classification scheme for non-porous rocks. The table data are from Stöffler (1984);Stöffler and Langenhorst (1994) and
French (1998).
Shock stage Shock pressure [GPa] Postshock temperature [

C] Physical effects Features
00e8 Fracturing Narrow spaced fracture networks
Shatter cones, planar fractures
Breccia formation
Twinning Basal twinning in qtz
Ia 8e20 >100 Fracturing Shatter cones, planar fractures
Localized amorphization Planar deformation features
Growth of high pressure polymorphs Stishovite
Ib 20e35 >170 Localized amorphization Planar deformation features
Bulk reduction of refractive indices and birefringence
Growth of high pressure polymorphs Coesite
II 35e45 >300 Complete amorphization Diaplectic qtz and fsp glasses
III 45e60 >900 Complete amorphization Diaplectic qtz glasses
Melting Vesiculated fsp melt
IV 60e80 >1500 Bulk melting Melted qtz and fsp
V>80e100 >2500 Bulk vaporization Condensed glass
For non-porous rocks. Data from Stöffler (1984), Stöffler and Langenhorst (1994) and French (1998).
Fig. 3. A) Numerical simulations of a 10 km dunite asteroid striking a granitic target at 20 km/s at different angles. Regions of high shock pressures are located downrange relative to
the point of impact. A general reduction of the volume of shocked material occurs at lower impact angles. B) Peak shock pressures calculated from the simulations in A) decrease
with decreasing impact angle, following a simple sinusoidal dependence. (Modified from Pierazzo and Melosh, 2000).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182160

of the peak particle velocity during shock wave compression
(Melosh, 1989). Finally the shock wave decays to a stress wave that
travels with the bulk sound speed in the rock.
A kinematic model of the excavation flow, the so-called Z-
model, was proposed by Maxwell (1977). The streamlines of the
excavation flow near the point of impact are directed downward
and outward, away from the point of impact (Fig. 4). Streamlines
radiating outward gradually bend upward and outward due to the
reduced pressure gradient towards the surface, caused by the
reflection of the shock wave off the free surface. The flow field leads
to the opening of a cavity that has a parabolic shape in cross-section
(Fig. 4). It is important to note that the ejecta do not include ma-
terial excavated from the full depth of the cavity. Target material
from the deeper parts is displaced downward into the ground and
does not leave the cavity. It is only the material from the upper one-
third to half that is entrained in the excavation process and is
deposited as allochthonous ejecta (Fig. 4). On airless planetary
bodies the debris ejected from the growing crater follows ballistic
trajectories. The innermost ejecta are launched first, become ejec-
ted to high altitudes, and travel fastest. Ejecta originating farther
from the center are launched later, move more slowly and become
deposited in the ejecta blanket nearer to the crater rim (Melosh,
1989). This combination leads to the formation of an ejecta cur-
tain that has the shape of an inverted cone. The cone expands
outwards and the ejecta debris of this cone is deposited on the
target surface at increasing distances from the crater. The excava-
tion flow is halted and the cavity stops growing when the
remaining kinetic energy is insufficient to displace the target
against its own weight (gravity-dominated cratering) or to over-
come the cohesive strength of the target material (strength-
dominated cratering) (e.g., Kenkmann et al., 2012). The resulting
unstable cavity at the end of the excavation stage is called the
transient cavity, and on planetary scales its diameter can range
between a few projectile diameters to over one hundred projectile
diameters, depending on projectile velocity and size, and gravity
(Holsapple, 1993). Experimental observation of transient craters in
particulate targets yields depth-diameter ratios of roughly 1/3
(Barnouin et al., 2011).
3.2. Macroscopic structures related to the excavation stage
3.2.1. Allochthonous, excavated lithologies
The lithological products of the excavation stage are different
types of allochthonous breccias, megablocks, diamictites, and
impact melt rocks that are observed in the ejecta blanket and the
crater walls and floor (Fig. 5). The debris of the ejecta curtain forms
a blanket around the transient crater cavity. These ejecta blankets
are continuous to approximately 2e3 crater radii, but become
discontinuous with increasing radial range. Monomict lithic brec-
cias (Fig. 6A) often develop either in the proximal ejecta blanket
near the edge of the transient cavity, where displacements remain
moderate and mixing of lithologies is subdued, or in the frag-
mented and faulted basement rocks of the crater. Allochthonous
breccias are usually polymict in nature (Fig. 6BeD) and may contain
a variable amount of shocked fragments. As the streamlines of the
excavation flow cross the shock isobars, the polymict breccias of
ejecta blankets contain fragments of various degrees of deforma-
tion and shock metamorphism. If melted particles are visible with
the naked eye and are embedded in a fine-grained clastic matrix,
these polymict lithic breccias are named “suevite”(Fig. 6C). If the
matrix itself is melted the rock is named impact melt rock or impact
melt rock breccias (Fig. 6D) (Stöffler and Grieve, 2007).
Beside the deposits of the ejecta blanket, the crater cavity itself
is filled with various, often melt-bearing types of allochthonous
breccias deposited from a presumably very hot ejecta plume above
the cavity (Fig. 5). Such deposits often closely intermingle with
diamictites from slumps that emanate from the steep cavity walls.
Fig. 4. Schematic diagram of the initial impact, and subsequent formation of the transient crater cavity during the excavation stage. Upon impact, a hemispherical shock wave
passes through the target rocks. Close to the point of impact, rocks are melted or turned to vapor. With increasing distance, the pressure of the shock wave decays and deforms the
rock to different stages of shock metamorphism (SIVeS0). For details of these shock stages see Table 1. The passage of the shock wave sets the target material in motion, following
specific particle paths that either lead to the excavation and ejection of the material in the top third of the transient crater, or to the displacement and downward deflection of target
rocks in the bottom two thirds of the transient crater.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 161

Fig. 5. Schematic cross-section of a mid-sized impact crater with lithologies formed through the impact cratering process. Near-surface or ejected lithologies include lithic breccias
that can be either monomict or polymict. If a polymict breccia with a clastic matrix contains melt fragments, it is referred to as a suevite. If the matrix itself is composed of “melt”,it
is referred to as an impact melt rock. In the crater subsurface, monomict breccias commonly are generated through fracturing, faulting and the formation of megablocks. Clastic
breccia dikes are found injected into the crater walls, and melt veins with clastic components, i.e., pseudotachylites, occur in many larger impact structures.
Fig. 6. Impact breccia lithologies. A) Monomict quartzite breccia of the central uplift of the Matt Wilson crater, Australia. B) Polymict lithic breccia (Bunte breccia) of the continuous
ejecta blanket of the Ries crater at Gundelfingen, Germany. C) Suevite containing SIV-type melt lumps (black), SIIeIII crystalline fragments, and a lithic groundmass, Ries crater,
Seelbronn, Germany. D) Impact melt rock consisting of a variety of variously shocked crystalline fragments embedded in a melt matrix from the Ries crater, Polsingen, Germany.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182162

The formation of breccias requires the fragmentation of rock
masses involved in the excavation flow, which occurs through
several subsequent processes: (i) tensile failure during shock
pressure release, (ii) shear and tensile failure plus abrasion during
the dense material stream flow within the transient cavity, in
particular near the cavity floor (Fig. 7), (iii) tensile break-up during
ballistic flight by aerodynamic drag forces and particle collision and
by the release of stored elastic strain, (iv) fragmentation during the
initial deposition of the ejecta blanket, and (v) further fragment size
reduction by the entrainment in the radially outward moving ejecta
blanket debris flow (Fig. 7).
3.2.2. Parautochthonous target structures
Rocks underneath the crater cavity which are not directly
involved in the excavation flow also show signs of outward
movement. Shear sense indicators often preserve and document
the outward and upward directed excavation flow in the crater’s
sub-surface. For instance, concentrically striking drag folds or
asymmetric folds with inclined fold planes and top-outward ver-
gencies are indicative for this second stage of cratering (Roddy,
1977; Shoemaker et al., 2005). Parallel to the steeply dipping
crater walls concentrically striking reverse faults can form, while
reverse faults that dip away from the crater center can occur near
the surface further outwards (Pilon et al., 1991). The degree of
fragmentation varies: zones of monomict brecciation alternate
with intact blocks and faulted as well as folded rock units.
During the excavation flow a considerable amount of rock be-
comes injected into the walls of the crater subsurface (Fig. 7).
Injection can either occur via dike formation or by the emplace-
ment of intact rock masses. Dikes usually consist of intensively
cataclastically deformed and fluidized rock debris with or without
contribution of melted material (e.g., Wittmann et al., 2004). Such
dikes have sharp contacts with the host rock, and display numerous
branching points with blind terminations. A second type of injec-
tion can be observed, e.g., at Meteor Crater (Barringer Crater), AZ,
USA, where coherent blocks are injected into the cavity walls
(“interthrust wedges”;Poelchau et al., 2009). The emplacement of
dikes and interthrust wedges is a mechanism of thickening and
uplift of the transient crater rim, and leads to an anticlinal doming
of the rock layers above interthrust wedges (Figs. 7 and 8C).
The total elevation of transient crater rims is the sum of uplift
induced by the injection of material into the crater walls and the
deposition of ejecta at the edge of the transient crater. This ejecta is
only moderately displaced and therefore often forms coherent and
intact masses with inverted stratigraphy or monomict breccias. The
proximal ejecta forms an overturned flap (Fig. 7,Fig. 8A and B). In
structural terms this overturned flap is the upper limb of an iso-
clinal recumbent fold with a circumferential fold hinge.
The excavation flow field is affected by pre-existing target het-
erogeneities such as joints. For instance the rectangular joint and
fissure pattern at Barringer Crater, AZ, USA, resulted in a more
efficient excavation flow along the fissures (Eppler et al., 1983;
Poelchau et al., 2009; Watters et al., 2011) than between them
and consequently increased the radius of the cavity along this di-
rection. This led to a strong deviation in circularity and resulted in a
more rectangular planform of crater cavities (Eppler et al., 1983;
Fig. 7. Schematic profile of a crater rim showing macroscopic structures formed during the excavation stage. The excavation flow field is directed outwards and upwards and leads
to the ejection of target material, which forms an ejecta curtain that is ballistically deposited outside of the crater at progressively further distances. At the crater rim, the most
proximal ejecta forms an overturned flap of coherent to semi-coherent overturned layers, while further outwards the ejecta is sheared and brecciated upon initial deposition due to
the horizontal component of momentum of the ejecta curtain. In the crater wall, material of the excavation flow can be injected into the target rock either in the form of dike
breccias, melt or as coherent blocks, termed “interthrust wedges”, that lead to an uplift of the target surface. These injections can exploit weaknesses in the target rock that were
formed through spallation effects caused by the interaction of the shock wave with the target surface. Further displacements are observed in top outwards and bottom outwards
thrusting.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 163

Öhman et al., 2010). Pre-existing joints in the corners of such
polygonal craters are called “radial corner faults”(Fig. 8A and D)
(Poelchau et al., 2009). The main component of movement dis-
played by these faults is vertical along with a rotational component
or “scissors-type of displacement”(Shoemaker, 1960). The kine-
matics of the radial corner faults is comparable to mode III shear
fractures (cf. Twiss and Moores, 1992) with a shear displacement
parallel to the edge of the fault. At Meteor Crater, radial corner
faults displace rock beds vertically by up to 50 m (Shoemaker and
Kieffer, 1974; Poelchau et al., 2009).
Another important process associated with the excavation
process is spallation. It occurs when the expanding shock wave or
pressure wave approaches free surfaces and interferes with ten-
sile stress waves reflected from these surfaces. Spallation domi-
nantly occurs near the target surface when the local tensile
strength of the target rock is exceeded (Melosh, 1984). The zone
of spallation may have a wider lateral extent than the transient
crater cavity and is thus also recognizable in the periphery of the
crater cavity. Spallation is an important mechanism enabling dike
injection into the cavity walls as it promotes the short-term uplift
of rock (Fig. 7). In bedded target strata, spallation effects can be
observed several tens of meters beneath the surface in the pe-
riphery of craters up to 1.8 crater radii, e.g., in the form of sub-
horizontal shear planes (detachments) (Fig. 9). The best place to
study these effects is the Ries crater in Germany (Kenkmann and
Ivanov, 2006; Kenkmann and Schönian, 2006). Weak spallation
outside the transient crater leads to a short-term uplift and
decoupling, and causes mode I tensile fracturing of the target
material along the spall surfaces. The subsequent arrival of the
ejecta curtain at the target surface delivers radially outward
directed horizontal momentum to such decoupled uppermost
target layers and can result in a re-activation of these spallation
planes as shear planes (Fig. 9). Striations on these detachment
planes and offsets of markers indicate top-outward shearing with
radial slip vectors and displacements ranging from meters to
decameters. At the outer front of these spallation-induced
detachment planes reverse faults may develop (Kenkmann and
Schönian, 2006).
3.3. Fracturing during the early stages of cratering
The response of minerals and rocks to high loading rates and
loads well beyond the point of failure is critical for the under-
standing of fracturing and fragmentation during the early phase of
an impact event. The importance of inherent flaws as sites of
weakness for the nucleation and coalescence of fractures is
described in the Griffith theory (Griffith, 1920). Unlike quasi-static
fracture mechanics, a material-specific rate dependency exists for
high strain rates that controls the fragmentation process. The
fracture stress and fragment size depend on the loading strain rate
(e.g., Grady and Kipp, 1993) in such a way that the fragment size
decreases and the fracture stress increases with increasing strain
rate. Material tests over a large range of strain rates show that for
both uniaxial compressive strength and tensile strength, after a
certain “transition strain rate”(w10
1
e10
3
s
1
) the material’s
Fig. 8. Structural features of the excavation stage observed at Barringer Crater, Arizona,
USA (1.2 km diameter). A) Along the crater wall, autochthonous beds of the Kaibab and
Moenkopi Formation are overlain by the allochthonous ejecta, composed mainly of
overturned Kaibab beds. The rock layers are displaced by a radial corner fault by
several tens of meters. Talus covers the lower parts of the crater wall. B) Close up of the
overturned flap, showing uplifted, outwards dipping Moenkopi beds (red color) and
overturned Moenkopi and Kaibab units that were folded along a hinge that strikes
parallel to the crater wall. The crater center is to the left. C) Competent blocks were
thrust into the crater as an interthrust wedge, causing anticlinal doming and localized
uplift in the crater rim. D) Close up of a radial corner fault. The shear displacement
occurs along sub-vertical fault planes that lead to local drag folds of the layered
bedrock. Note the step-wise displacement of the Kaibab-Moenkopi contact in the
anastomosing system of faults.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182164

strength begins to rapidly increase (Fig. 10). As a cause for this
behavior, it has been suggested that at high strain rates, material
inertia begins to affect the nucleation and propagation of cracks
(e.g., Kipp et al., 1980). Relative to their quasi-static strengths,
brittle materials show an up to four-fold increase in compressive
strength and an eight-fold increase in tensile strength, within the
range of experimentally generated strain rates (e.g., Kimberley
et al., 2013; Zhang and Zhao, 2013).
An inverse dependence on strain rate also exists for the length of
a fracture or fault. Melosh (2005) showed that faults cannot be
longer than the distance that sounds travels in the time it takes to
exceed the yield stress. Thus, fractures which develop early in the
cratering process, when strain rates are very high, tend to be short
and very closely spaced. Fractures associated with the passage of
the shock wave usually have lengths of less than a centimeter and
displacements of less than a millimeter. Often irregular to parallel
fracture networks develop (Fig. 2A) that are pervasive and trans-
granular. During the excavation flow fractures coalesce and
become longer. Faults of hundreds of meters to even kilometers
length with single-slip off-sets of up to several kilometers (Spray,
1997) are always formed during the final stage of cratering, the
modification stage.
Fig. 9. Typical exposure of the continuous ejecta blanket of the Ries crater on top of target rocks outside the crater. A) Radial striations are observed outside of the craterrim on the
target surface, caused by the outward flow of the ejecta blanket. In the sub-surface, top-outwards directed shearing displaces karst cleavages at the Gundelsheim quarry 7 km
outside the Ries crater. The detachment exploits weaknesses of the rock layers that were opened during spallation processes. B) Numerical simulation of the ejecta curtain of the
Ries crater shows a strong horizontal velocity component of up to 240 m/s at 13 km distance from the crater center. (Impact crater is to the left) C) After a first pulse of motion
induced by spallation the ejecta curtain drag in the simulation induces horizontal velocities of w10 m/s in the target rocks at 75 m depth after 50 s (Modified from Kenkmann and
Ivanov, 2006).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 165

The development from narrow-spaced fracture networks under
high strain rate conditions during the early impact phase to more
localized large fault zones during the modification stage indicates
that the degree of strain localization also shifts from a relatively
homogeneous bulk strain behavior towards a strongly heteroge-
neous strain distribution at the end of the cratering process.
The geometry of fractures formed in the very early stage of
cratering is also affected by the propagation speed of the fracture
tip. The theoretical speed limit for tensile fracture propagation is
the Raleigh wave speed. Sagy et al. (2006) showed that fractures
that propagate at high speed, reaching about half of the Raleigh
wave velocity, tend to become unstable in brittle materials and
bifurcate when velocity excursions occur, e.g., induced by local
asperities of the fracture plane. Networks of hierarchical bi-
furcations of fractures characterize, for instance, the surfaces of
shatter cones (Sagy et al., 2002, 2004)(Fig. 2B).
Brittle failure of rock is related to a reduction of particle size. It
has been shown that resulting particle sizes are properly described
by a fractal size distribution (e.g., Marone and Scholz,1989; Sammis
and Biegel, 1989) and can be quantified by a resulting power-law, in
which the exponent is referred to as the “fractal dimension”or D-
value. Recent experimental analysis of the particle size distribution
in impact cratering experiments also suggests a fractal particle-size
distribution for impact-loaded rocks (Buhl et al., 2013). In these
experiments, D-values decrease with increasing distance from the
point of impact, reflecting a decrease in damage.
3.4. Effects of oblique impacts on the excavation flow and related
structures
The most probable impact angle for all planetary bodies is 45

(Gilbert, 1893; Shoemaker, 1962) regardless of the magnitude of
the gravitational field. The probability of incidence angles follows
a Gaussian distribution (Gilbert, 1893; Shoemaker, 1962). In spite
of the prevalence of oblique impacts, the crater shape remains
circular for impact angles above 10e15

from the target surface
(Gault and Wedekind, 1978; Bottke et al., 2000) and thus can
rarely give implications for the impact direction or angle. The
reason for this apparently puzzling circumstance is that the
asymmetric region controlled by the transfer of momentum and
energy in oblique impacts may become an insignificant fraction of
the final crater, which is eventually about twenty times larger in
diameter than the impactor. However, the distribution of the
ejecta blanket is a good indicator for oblique impacts. It loses its
radial symmetry at angles below 45-35

(Fig. 11A) with a preferred
downrange ejection and at lower angles forms “forbidden”zones
in the uprange sector (Fig. 11B). With the expansion of the
forbidden zone and the preferred distribution of ejecta in
Fig. 10. Compressive strength plotted against strain rate for various brittle materials.
Compressive strength is normalized by quasistatic compressive strength. An abrupt
increase in strength occurs above a specific transition strain rate. (Modified after
Kimberley and Ramesh, 2011).
Fig. 11. Examples of Martian impact craters that illustrate the influence of obliquity on ejecta emplacement. A) Unnamed non-oblique impact crater (17.79N 313.56E) with a
symmetric ejecta pattern (mosaic of THEMIS VIS and CTX data). B) Unnamed oblique impact crater (2.29N 64.83E) with an asymmetric ejecta pattern including an uprange
forbidden zone that indicates an impact from NE (black arrow) (CTX mosaic). C) Highly oblique impact crater from NW (black arrow) with a crossrange concentration of ejecta and
additional uprange and downrange forbidden zones forming a butterfly ejecta pattern (CTX mosaic).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182166

crossrange directions, bilaterally symmetric “butterfly”patterns
(Fig. 11C) develop at very shallow incidences of less than 20

(Gault and Wedekind, 1978; Herrick and Hessen, 2006). The po-
sition of the symmetry axis gives the trajectory of the impact. The
asymmetric ejecta flow in oblique craters also involves deviations
from a pure radial flow. When extrapolated backward into the
crater, striations on the ejecta blankets do not meet in the crater
center but focus along a line running from the uprange section of
the crater to its center, suggesting non-radial, outward flow from a
moving source of ejection, as also suggested from experiments
(e.g., Anderson et al., 2003; Schultz et al., 2007).
Unfortunately, on Earth ejecta blankets of impact craters are
usually eroded and cannot be used to determine the impact di-
rection for terrestrial craters. Hence, other potential indicators for
an oblique impact are needed if the impact trajectory is to be
determined in terrestrial craters. The asymmetric and bilateral
symmetric excavation flow field can be measured applying tech-
niques of structural geology (Poelchau and Kenkmann, 2008;
Poelchau et al., 2009). Deviation from a non-radial excavation
flow leads to a measurable deviation in strike of layered rock units
from a concentric direction. This can be measured in the hinge zone
of the crater rim and within the overturned flap, where strata are
often coherent over large distances. In these settings it can be
assumed that strata strike is perpendicular to the excavation flow
direction (for originally horizontal bedding). The expected pattern
of strike should be bilaterally symmetric to the direction of impact
and, on the basis of the analysis of Tooting Crater on Mars and
Wolfe Creek, Australia, have two ‘‘corners’’ between the uprange
and crossrange sector, in which an abrupt change in strike orien-
tation of layered rock units occurs (Poelchau and Kenkmann, 2008;
Poelchau et al., 2009).
4. The modification stage
4.1. Physical background
The excavation flow describes the motion of target material
away from the impact center. In the modification stage, the direc-
tion of flow is effectively reversed and acts to modify and collapse
the transient cavity. The modification stage begins when the
excavation flowcomes to a halt and the transient cavity reaches its
largest horizontal extent at the level of the target surface. It is
important to mention here that the excavation flow does not stop
simultaneously in all parts of the transient cavity before reversing
its direction; e.g., excavation flow is still active in the rim after the
transient crater has reached its final depth (Turtle et al., 2005). This
overlap of excavation in the rim with simultaneous modification of
the crater floor increases the more the impact process is governed
by gravity rather than the strength of the target, i.e., for larger
impacts.
Gravity is the principal force that drives the collapse of the
transient cavity. Depending on the degree of modification, the final
crater is classified into either a simple or complex morphology
(Dence, 1965). Simple and complex impact craters have funda-
mental differences in morphology and structure; both being a
function of size. Simple craters generally have a bowl-shaped
morphology (Fig. 12a), while complex craters have terraced rims
and an uplifted central portion of the crater floor (Fig. 12b). With
increasing size, complex craters are further subdivided into central
peak, peak-ring craters, and multi-ring basins based on the
morphology of the uplifted target rocks.
The end of the modification stage is reached when all signifi-
cant motion of the target comes to rest. The duration of crater
Fig. 12. Comparison of simple and complex crater morphologies. a) Unnamed simple crater on Mars (38.7N/316.1E) displaying an elevated crater rim and steeply dipping upper
cavity walls. The mid and lower part of the wall is covered by talus deposits (HiRISE image). b) The complex impact crater Aristarchus on the Moon, showing a central peak, a flat
crater floor and an extensive slump terrace zone (Kaguya/SELENE image). Note the different scale bars in the two images.
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 167

modification depends on impact energy, gravity, and strength
properties of the target material; its duration increases with
increasing impact energy, decreasing gravity or decreasing target
strength and can take up to a few minutes for a 25 km sized
impact crater like the Ries crater in southern Germany (Fig. 13).
More generally speaking, the duration of collapse is proportional
to (D/g)
1/2
, which is approximately the period of a gravity wave of
a wavelength equal to the crater diameter Dat a surface gravity g
(Melosh, 1989).
Long-term lithostatic crustal relaxations that are triggered by
mass deficits of very large impact craters can persist over millions
of years and are accommodated by plastic flow at mid to lower
crustal levels or even in the underlying mantle, as was observed in
the case of the 180 km Chicxulub crater (Morgan et al., 2000).
Structural crater modifications may also occur due to variations in
the post-impact sediment loading (Tsikalas and Faleide, 2007). This
has been demonstrated, for example, at the Mjølnir impact struc-
ture, Barents Sea, Norway, where differential compaction during
long-term subsidence laterally varies within the crater and caused
the formation of a very prominent central high. These processes are
not further detailed here. Subsequent geologic processes that are
unrelated to the impact process itself, such as a regional tectonic
overprint or erosion, are not part of the modification stage but
represent the post-impact history.
4.2. The crater flow field during collapse
The transient cavity is negatively buoyant and creates an up-
ward force on the crust due to the displaced mass in the crater. This
gravitational force attempting to close the crater is counteracted by
the strength of the target. If the buoyancy force is too small,
strength prohibits the target material from moving upwards and
inwards to modify the cavity and a simple crater is formed (Collins
et al., 2004; Kenkmann et al., 2012). If the size of the transient
cavity exceeds a certain threshold, buoyancy forces exceed the
target’s strength and the entire cavity can completely or partly
collapse. This starts at the deepest point of the cavity in the very
center, where the upward-directed buoyancy force is strongest and
a complex crater is formed (Fig. 13). The relative balance between
both mechanisms, the buoyancy acting to close the crater, and
material strength acting to stabilize the cavity, determines whether
a simple or complex crater is formed.
4.2.1. Simple craters
The principal shape of a simple crater consists of a bowl-shaped
depression and a raised crater rim and is thus similar in shape to
the transient cavity (Figs. 12a and 14A). Modification is mainly
restricted to mass movements along the steep transient cavity rim
and the presence of brecciated material that fills the cavity. The
visible floor of the crater is underlain by a lens of allochthonous
unshocked and shocked target-rock breccias (Grieve,1987). Slumps
initiate along the steepest parts of the crater walls near the slope
top end and lead to an increase in cavity diameter by 10e20%. The
lower part of the autochthonous crater wall is usually covered by
talus deposits (Fig. 8A). The increase in diameter and the infill of the
cavity explains the decrease of the depth/diameter ratio from 1/3
for the transient cavity to about 1/5 for the simple crater. A prime
example of a young, well-preserved and well-documented simple
impact crater on Earth is the 1.2 km diameter Barringer Crater, AZ,
USA (Fig. 14A).
4.2.2. Complex craters
Complex impact craters are generally larger than simple craters
and differ significantly from the shape of their transient cavity. The
strong modification is the result of extensive gravity-driven
collapse. The collapse occurs first at the deepest point of the tran-
sient cavity. The cavity floor starts to rise upward and inward,
causing a rotational flow field underneath the cavity. The greatest
total uplift and uplift rate exist in the center, pushing the cavity
floor upward (Fig. 13). The upward and inward flow creates a mass
deficit in the subsurface beneath the cavity rim, which ultimately
Fig. 13. Numerical simulation of crater modification using the SALE-2D hydrocode
(Image courtesy of B. A. Ivanov). Note that the initial 12 km diameter cavity transforms
into a 24e26 km wide crater structure with a central uplift in its center. Model pa-
rameters are given in Kenkmann et al. (2000a).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182168

results in a down-sagging of the steep crater walls and causes an
enlargement of the rim of the cavity involved into the modification
flow (e.g., Kenkmann et al., 2012). Near the cavity surface, material
directly moves down slope towards the center in a manner similar
to that described for simple craters.
With increasing diameter, the gravitationally-driven central
uplift can grow fast and high enough that inertial flow leads to an
overshooting of the equilibrium height of the central uplift (Fig. 13).
The central uplift itself becomes gravitationally unstable and col-
lapses under its own weight. In this case, a downward and outward
directed collapse flow occurs. This flow interacts with and super-
poses the material that is simultaneously moving from the cavity
rim towards the crater center. Craters showing this type of exten-
sive modification are called peak-ring craters.
Complex craters have a depth-to-diameter ratio in the order of
1/10. The most remarkable morphological difference between
simple and complex impact craters is that complex impact craters
contain an uplifted crater floor of deformed and shocked rocks that
either forms a central uplift (Fig. 14B), a patchy distribution of hills
and hummocks, a peak ring, or a flat crater floor (Kenkmann et al.,
2012). The stratigraphic uplift of the central crater floor systemat-
ically increases with increasing final crater size. The crater floor that
surrounds the central uplift typically appears flat (Figs. 12b and
14b). In this region the autochthonous crater floor is overlain by
allochthonous breccias with various shock stages and impact melt
rock. The crater rim region of complex impact craters is subdivided
into terraces separated by steep scarps (Fig. 12b). Terrace widths are
narrow close to the crater center, but increase as the rim is
approached. The widest, best-defined, and last-formed terrace with
the largest off-set normally occurs just below the crater rim (Pearce
and Melosh, 1986; Leith and McKinnon, 1991). The diameter of
complex craters is enlarged by a factor of 1.5e2.0 with respect to
the transient cavity (Grieve et al., 1981). In pristine craters the
crater rim can be delineated by a morphological crest line, in
eroded craters on Earth the outermost circumferential fault that
delimits the outer crater rim terrace is used for defining the crater
size (Turtle et al., 2005).
4.2.3. The simple-to-complex transition
The concept of classifying crater morphology into simple and
complex applies for all planetary bodies in the solar system. The
crater diameter at which the simple-to-complex transition occurs
varies between planetary bodies and is inversely proportional to the
surface gravity (Pike, 1988) indicating that gravity is the main
driving force for crater modification. On the Moon (surface gravity:
1.6 2 ms
2
) the largest impact craters with simple geometries have
diameters of 16 km, on Mars (3.69 ms
2
), simple craters reach
maximum diameters of w8 km on average, and on Earth
(9.81 ms
2
), the largest simple craters were formed in crystalline
targets and have diameters of up to 4 km (Brent, Canada). The size-
morphology progression is also controlled by strength as the me-
chanical property of the target material working against the
modification of the transient crater. For instance, in sedimentary
targets the size limit for simple craters on Earth is about 2 km
diameter. On icy bodies (i.e., Europa, Ganymede, Callisto) the tran-
sition lies at 2e3 km diameter and is nearly an order of magnitude
smaller than the transition diameter of the Moon, despite similar
gravitational fields (Schenk, 2002; Barnouin et al., 2012).
4.3. Macroscopic structures related to the modification stage
The deformation inventory formed during the modification
stage is in some respect comparable to that of landslides and also to
certain tectonic environments. Major differences between gravity-
driven collapse of large impact craters and upper crustal tectonics
occur in the slip behavior and the particle trajectory field. As a first
approximation, particle paths are radially symmetric during inward
flow with respect to the impact center, which results in the con-
ditions for plane-strain deformation not being fulfilled (Kenkmann,
2002). Shear displacements occur as single-slip events, with
Fig. 14. Typical terrestrial examples of simple and complex impact craters. A) Panorama view of Meteor (Barringer) Crater, AZ, USA, B) Panorama view of the recently discovered
Jebel Waqf as Suwwan crater, Jordan. (Modified from Kenkmann et al., 2012).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 169

displacements ranging from centimeters to probably kilometers in
very large craters as can be derived e.g., from the vertical off-set of
terraces of lunar or Martian craters. As the structures formed in a
known period of time, terrestrial impact craters can be used as
large-scale laboratories for structural investigations. The macro-
scopic structures described are ordered according to their occur-
rence from the rim towards the center.
4.3.1. The crater rim
The crater rim of pristine craters usually shows a weak
morphological elevation with a scarp on the inner side. The
prominent escarpment inside corresponds to the outermost fault
visible on the uneroded target surface and usually forms one of the
major terrace steps in the crater rim region that often appears
stepped (Figs. 12 and 15). On Earth, where the original morphology
of craters is often barely visible or strongly modified, the outermost
continuous concentric normal fault usually defines the final crater
diameter of a complex impact crater (Fig. 15). Turtle et al. (2005)
stated that the outermost normal faults visible in eroded craters
can lie further outwards than the main escarpments of uneroded
structures. They suggest the terms “rim diameter”for uneroded
craters and “apparent diameter”for eroded structures. The rela-
tionship of these two diameters is complex and not completely
clarified. Crater rim faults typically undergo unconstrained (free-
surface) dip-slip (Spray, 1997). The main faults are often associated
with synthetic or antithetic faults. Pre-existing faults and joints can
be reactivated during crater modification. Such craters often appear
as polygonal craters with straight rim segments that run along the
pre-existing joints (Eppler et al., 1983).
Very deeply eroded impact structures are typically not defined
by concentric normal faults. Instead circumferential monoclines or
a combination of inward dipping normal faults and monoclines are
common, particularly if the target is a sedimentary and stratified
one (Fig. 15). The inner limb of a crater rim monocline usually dips
downward towards the crater, and the crater rim can be defined by
the trace of the monocline’s hinge (Kenkmann et al., 2012). Ex-
amples for this type of crater rim are present at Upheaval Dome, UT,
U.S.A. (Kriens et al., 1999), or Matt Wilson, NT, Australia (Kenkmann
and Poelchau, 2009). If the craters formed underwater, resurging
water degrades and modifies the crater rim area (e.g., Ormö and
Lindström, 2000).
4.3.2. The crater moat
The moat between the crater rim and the central uplift is often
termed ring syncline (Figs. 14Band15). Ring synclines are mostly
asymmetric in radial cross section, with a steeply dipping or even
Fig. 15. Schematic block diagram illustrating the structural inventory and the locations of certain structures in the sub-surface of a complex impact crater. Note that the letters AeF
also correspond to Fig. 16. A) Low-angle normal fault and detachment within the ring syncline. B) Lateral thrust ramps. C) Radial transpression ridge/positive flower structure. D)
Radial folding with outward plunging fold hinges. E) Radial syncline with vertical to overturned plunging fold axis within the central uplift. F) Imbrication of blocks thrust onto each
other in the core of the central uplift. (Sketch modified from Kenkmann et al., 2012).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182170

overturned inner limb (Fig. 13) that transitions into the central
uplift and a more gently dipping outer limb that is often
segmented by normal faulting. In all known terrestrial complex
impact craters, the ring syncline is not a simple synform. It is
subdivided radially and concentrically into numerous fault-
bounded segments or disintegrated into blocks (if fault zones
completely frame and isolate a certain rock volume). Between the
crater rim and the axis of the ring syncline, normal faults of more
or less concentric strike are frequent. Normal faulting along non-
planar faults is commonly associated with antithetic or synthetic
rotation of the hanging-wall unit. These faults often develop listric
shapes. In stratified target rocks they can merge into low-angle
detachments at depth (Figs. 15 and 16A) to compensate for the
inward movement of material during crater collapse (Kriens et al.,
1999; Kenkmann and von Dalwigk, 2000). The presence of low-
angle faults or detachments is favored by the large-scale rota-
tional flow field that exists during crater collapse that comprises
uplift in the center and associated down-sagging in the periphery
(Fig. 13). Bedding planes of the stratified sediments are often used
as glide planes. Displacements related to the modification stage
commonly indicate inward and downward motion within the ring
syncline. Due to the formation of the central uplift and passive
Fig. 16. Field observations of structural deformation in complex craters. A) A low-angle normal fault merges into a bedding-parallel detachment within the ring syncline of the
Upheaval Dome crater, UT, USA. Note that the section is parallel to the fault strike. B) Lateral thrust ramps in the periphery of the central uplift of Matt Wilson crater, NT, Australia. C)
Radial transpression ridge (positive flower structure) in the periphery of the central uplift of Matt Wilson crater, NT, Australia. D) Radial syncline with steeply outward plunging fold
hinges within the central uplift of the Serra da Cangalha crater, Brazil. The fold axes are bent over toalmost vertical plunge with increasing elevation. E) Open radial folds with gently
outward dipping axes characterize the periphery of the Upheaval Dome crater, UT, USA). F) Imbrication of blocks thrust onto each other in the core of the central uplift (Matt Wilson,
NT, Australia).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 171

rotation, the low-angle normal faults may transform into outward
dipping thrust faults and reverse faults in the inner limb of the
ring syncline (Jahn and Riller, 2009).
The structural complexity of the ring syncline increases towards
the center because the amount of lateral constriction increases by
the motion of rock towards the center. The convergent particle
trajectories during inward flow can be compensated either by a
bulk thickening of inward sliding masses (tight folding, stacking of
rock units along reverse faults, plastic flow) or by the formation of
radial transpression zones (Fig.15). They develop at the edges of the
obliquely converging rock masses during inward flow (Kenkmann
and von Dalwigk, 2000). Within such radial transpression ridges
material is uplifted to the free surface to accommodate the
convergent mass flow. Different modes of uplift are possible,
including lateral over-thrusting (Fig.16B), the formation of positive
flower structures (Fig. 16C), radial folding (Fig. 16D and E)
(Kenkmann et al., 2012), or complete brecciation of the contact
zone.
4.3.3. Central peaks
The intensity of deformation attributed to the modification
stage of cratering culminates within the central uplift. The defor-
mation inventory of central uplifts is extremely complex and could
be unraveled only for a few terrestrial impact craters on Earth with
a sedimentary target, e.g., Decaturville (Offield and Pohn, 1979),
Haughton (Osinski and Spray, 2005), Upheaval Dome (Kenkmann
et al., 2005), Matt Wilson (Kenkmann and Poelchau, 2009), Jebel
Waqf as Suwwan (Kenkmann et al., 2010) and Serra da Cangalha
(Kenkmann et al., 2011b). In the absence of appropriate marker
beds, impacts into crystalline rock targets often do not allow the
motions to be reconstructed in detail. Faulting, however, dominates
over folding in these non-layered, mechanically isotropic rocks.
Field analysis of a number of impact craters eroded to different
levels and numerical modeling have proven that the centrally
uplifted area becomes broader with depth while the observable
stratigraphic uplift decreases. Simultaneously, the diameter of the
crater shrinks with depth. Both circumstances cause the increase of
the ratio of the central uplift diameter to the apparent crater
diameter with increasing depth of erosion. Consequently, the cen-
tral uplift is by far the dominant structural feature for deeply
eroded craters; the moat and crater rim become faint features at
depth.
Complex terrestrial impact craters up to 10e20 km formed in
sedimentary targets show common features in the central uplift
structure: anticlines and synclines with radial fold axes are typical
for the periphery of central uplifts and the inner part of the ring
syncline (Figs. 15 and 16D and E). They result from constriction
caused by the convergent mass flow. Radial fold axes usually
plunge outward (Fig. 17) and cause the serrated appearance of
central uplifts in geological maps. The hinge line of these radially
striking folds is often bent and plunges more steeply with
increasing proximity to the core of the central uplift. Near the
center, vertical folds or folds with overturned hinges can develop
(Figs. 16Eand17). Steeply inclined, vertical or even overturned
beds of the central uplift result from this. Moreover, a gradational
transition in fold tightness, wavelength, and amplitude may be
detected, with open symmetrical anticlines of the central uplift
periphery changing into isoclinal and overturned folds towards
the center. Spatial incompatibilities of the folds increase with
increasing fold tightness. This leads to the initiation of reverse
faults in the core of these folds and their rapid propagation into
the limbs to finally offset one of the fold limbs from the other. As a
consequence, fold limbs become detached into sheet-like blocks,
bounded by reverse faults, which “stack up”against the core of
the central uplift in an imbricated fashion (Fig. 15). This process
can be compared to the closure of an iris diaphragm of a camera
lens.
Fault-bounded blocks usually build up the core of a central uplift
(Fig. 17). In sedimentary targets, the stratigraphic context can be
completely broken up. The occurrence of brecciation and breccia
bodies is at first limited to the edges of thrust units and blocks as
fault breccias. They become the dominant rock type in the core of
the central uplift (Fig. 17). Due to the immense stratigraphic uplift,
very large complex impact structures on Earth, such as the Vre-
defort Dome, South Africa, show an increase in pre-impact meta-
morphic grade toward the center of the structure. In such cases, an
eroded central uplift provides a condensed section through the
upper part of the crust (Gibson and Reimold, 2000).
4.3.4. Peak ring craters
At a certain threshold diameter, central uplifts become gravi-
tationally unstable and start to collapse under their own weight to
form a morphological ring of peaks in pristine craters. For
terrestrial craters this critical size is reached when the final crater
size exceeds w13e23 km, depending on the involved lithologies
(Pike, 1985). Craters near the transition diameter show vertical
flanks of their central uplifts. Localized kinking and buckle folding
of the vertically uplifted strata indicate the onset of collapse.
Widespread overturning of strata in the central uplift periphery
(Morgan et al., 2000; Lana et al., 2003; Jahn and Riller, 2009)is
one mechanism which enables the outward collapse. The over-
turning strata collide with the inward moving blocks of the sur-
rounding ring syncline and form a complex interference zone
(Kenkmann et al., 2000a; Morgan et al., 2000). The gravitational
collapse also induces the development of normal faults dipping
radially outward and offsetting the uplifted strata (Osinski and
Spray, 2005). The hummocky morphological appearance of fresh
lunar peak-ring structures suggests that the intensity of collapse
varies sector-wise and the formation of radial transtension troughs
may contribute to this selective collapse (Kenkmann and von
Dalwigk, 2000). Eventually, the collapsing central uplift flows
outward, thereby overthrusting the downfaulted rocks of the
surrounding ring syncline (e.g., Grieve et al., 2008). Craters of this
size contain a melt sheet of considerable volume. The Sudbury
Igneous Complex of the w200 km diameter Sudbury impact
structure is generally interpreted as the melt sheet of this crater
differentiated into layers that contain world-class CueNi-PGE
mineralizations (Grieve et al., 2008). From thickness variations of
the melt sheet layers the topography of the crater floor was
inferred. According to Dreuse et al. (2010) the topography of the
final crater floor at Sudbury was characterized by variations of up
to 400 m over distances of hundreds of meters to a few kilometers,
and variations of up to 1500 m over a distance of about 25 km.
Similar variations in the magnitude of crater floor topography are
known from the Manicouagan impact structure, Quebec (Spray
and Thompson, 2008). Melt from the melt pool may permeate
or even be sucked into the crater basement via tensile fractures
that can be active for a long period of time due to crustal re-
laxations. The offset dikes at Sudbury may represent such conduits
(Riller et al., 2010).
The 65 Ma Chicxulub impact crater in Mexico is the most
prominent example of a peak ring crater. The collapse of the central
uplift led to the formation of a rugged peak ring of 40 km radius
that stands several hundred meters above the otherwise relatively
flat crater basin floor (Brittan et al., 1999). The existence of slumped
blocks of the annular trough beneath the peak ring was inferred
through reflection seismic studies, which demonstrate inward
dipping reflectors, which have been interpreted to indicate a
simultaneous outward collapse of the central uplift and inward
collapse of the transient crater (Brittan et al., 1999; Grieve et al.,
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182172

2008). Hydrocode modeling has reproduced this dynamic behavior,
showing that the overturned flap of the transient crater rim moved
into the cavity and is located beneath the peak ring (Morgan et al.,
2000).
4.3.5. Multi-ring basins
Multi-ring formation is also related to the collapse of transient
crater cavities. Multi-ring basins are known from Mercury (Caloris
basin, 1300 km in diameter), the Moon (e.g., Orientale, 900 km in
Fig. 17. Structure map of the central uplift of the Jebel Waqf as Suwwan impact structure, Jordan. The oldest rocks exposed in the core of the uplift (Cenomanian) are strongly
brecciated. Limestone and marl beds (purple) have Campanian age. Chert beds (blue) of Maastrichtian/Paleogene age are strongly folded and segmented into w100 m blocks. Note
that synclines with overturned fold axes occur in the northeast, whereas the southwest is characterized by outward-plunging anticlines and synclines. The symmetry axis of the
central uplift trends SWeNE, and suggests an impact trajectory from the SW. For more details the reader is referred to Kenkmann et al. (2010).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 173

diameter), and Mars (e.g., Argyre, 1800 km in diameter) (Ivanov
et al., 2010). The Chicxulub crater has been interpreted as a
multi-ring structure by Sharpton et al. (1993) but others classify
Chicxulub as a peak-ring structure. The largest multi-ring structure
in the solar system is Valhalla on the icy Jovian satellite Callisto, and
extends over 3800 km in diameter (Schenk, 1995).
It is believed that multiple ring structures develop when the
depth of the transient cavity is comparable to the thickness of the
lithosphere. McKinnon (1981) showed that for a given impact en-
ergy the number of rings primarily depends on the thickness and
strength of the lithosphere and the viscosity of the underlying
asthenosphere. For a very thin, weak lithosphere and underlying
asthenosphere of sufficiently low viscosity (almost a liquid), the
basin formation is followed by multiple oscillations of the cavity
and outward propagation of gravity waves which disrupt the entire
lithosphere (ripple ring basins). For more realistic asthenosphere
viscosities, the asthenosphere will flood the cavity but dampen the
propagating gravity waves. Still, brittle concentric ring fractures
will permeate the entire thickness of the lithosphere. Thicker
lithospheres restrict the number of rings that can form. Each ring
represents scarps that result from conjugate systems of normal
faults (graben) with circumferential strike on average. Ring for-
mation is suppressed if the lithosphere is too thick. An unusual
multi-ringed structure, whose impact origin is likely, but not yet
proven, is Silverpit in the North Sea, which measures only 20 km
diameter (Stewart and Allen, 2002).
4.4. Brittle deformation during crater modification
Geological observations show that the target underneath the
crater floor is disintegrated into blocks, in particular in the central
uplift (Fig. 17). These blocks are commonly internally deformed
(bent or folded) at the millimeter to decameter scale, rather than
being entirely rigid and bounded on all sides by faults. An average
block size of w100 m was determined from the Vorotilovskaya
Deep Borehole (5374 m), drilled through the central uplift of the
40-km diameter Puchezh-Katunki impact crater in Russia (Ivanov
et al., 1996). Mapping at Upheaval Dome, UT., U.S.A., (7 km
diameter) (Kenkmann et al., 2006) and Waqf as Suwwan, Jordan,
(6 km diameter) (Kenkmann et al., 2010)(Fig. 17) revealed block
sizes of 50e100 m, with evidence of both a lithological control on
block size (smaller blocks were observed in limestone relative to
chert) and an increase in block size as a function of distance from
the crater center, which is in accordance to theoretical models of
rock fragmentation (Grady and Kipp, 1987).
The intensity of impactdeformation increases from therim to the
center. Numerical modeling shows that growth and collapse of the
transient cavity leads to an accumulated strain of w1inthematerial
underneath the crater, decreasing to w0.25 at the crater rim (Collins
et al., 2004; Kenkmann et al., 2012). This is associated with a tran-
sition from localized brittle faulting to a more pervasive cataclasis
and granular flow. Thus, in the crater rim zone, large-scale dis-
placements occur on localized fault planes, with blocks that are large
and internally only weakly damaged. The inner crater shows an in-
crease in brittle deformation and blocks that are smaller or more
internally damaged. Pervasive cataclasis and granular flow down to
the grain scale is present between the blocks (Kenkmann, 2003). At
Upheaval Dome the innermost strata indicate an almost complete
loss of internal coherence during deformation and display extreme
thickness variations, blind terminations and frequent embranch-
ments. Microstructural analysis of these dike networks revealed that
the macroscopically ductile appearance is achieved by distributed
cataclastic flow that was initiated by grain crushing, collapse of pore
space, and subsequent inter-granular shear. The damaged rocks
subsequently flowed as a granular medium (Kenkmann, 2003).
4.5. Impact-induced pseudotachylites and impact melt
Melt is a common characteristic of impact structures, in
particular of large impact craters. Two processes are responsible for
melt generation: (i) shock-related melting and (ii) friction-
controlled melting. A long-lasting controversial debate exists on
the discrimination of both and their significance for the mechanics
of the cratering process.
(i) Shock melting of rock-forming minerals such as quartz or
feldspar occurs during shock unloading of strong shock
waves with pressures exceeding 45e60 GPa. The products of
impact melting at terrestrial impact structures range from
small glass spherules, over melt lumps within suevitic
breccias (Fig. 6C) to thick sheets of coherent impact melt rock
(Grieve et al., 1977). Relative to the volume of the transient
cavity, the volume of impact melt rock (Fig. 6D) increases
with crater size (Grieve et al., 1977). In small craters, melt
volume is a tiny fraction of the transient crater volume that
forms an unevenly distributed and relatively thin sheet lin-
ing the final crater floor. In contrast, large impact craters such
as the 100 km Popigai crater, Russia, the 180 km Chicxulub
crater, Mexico, or the 200 km Sudbury structure, Canada,
contain kilometer thick sheets of impact melt rock forming a
pool in the crater center.
(ii) Frictional melting of rocks can occur tectonically during co-
seismic faulting, gravitational sliding, and during impact
cratering as a result of high rates of deformation and a nearly
complete transformation of kinetic energy to heat. The type
locality for pseudotachylites, Parys, South Africa (Shand,
1916), is situated in the Vredefort impact structure. These
pseudotachylites resulted from the cratering process. In the
context of impact craters, the origin of pseudotachylites is
ambiguous and one has to distinguish between millimeter-
sized veinlets and large occurrences of meters to hundreds
of meters extent.
At least two generations of pseudotachylites form during an
impact and have to be considered separately (e.g., Lambert, 1981;
Spray, 1998). The first generation has been referred to as A-type
pseudotachylites (Martini, 1978, 1991), or S-type (shock-related)
pseudotachylites (Spray, 1998). These pseudotachylites are believed
to form during shock compression underneath the crater floor of
the transient crater. They are typically thin (<2 mm) and are
randomly distributed in the crater floor. Localized melt veins in
meteorites (Stöffler et al., 1991) containing high-pressure poly-
morphs are morphologically very similar to these S-type pseudo-
tachylites. S-type pseudotachylites were successfully produced by
shock compression experiments and particularly develop along
heterogeneities where shock impedance contrasts exist (Fiske et al.,
1995; Langenhorst et al., 2002; Kenkmann et al., 2000b).
Second-generation pseudotachylites (Fig. 18) are believed to
form during the modification stage as frictional melts and were
termed B-type (Martini, 1991)orE-type (endogenic related) pseu-
dotachylites (Spray, 1998). These pseudotachylites may reach a
considerable thickness in the order of meters.
Reimold and Gibson (2005) introduced the term pseudotachy-
litic breccias for breccias containing a melted rock matrix that re-
sembles pseudotachylite but where the genetic origin of the melt is
unclear. Several potential mechanisms for producing the melt in
these breccias have been suggested and include friction melting
(Spray and Thompson, 1995), shock melting (Fiske et al., 1995),
decompression melting (Reimold and Gibson, 2005), a combination
of shock and decompression melting (Mohr-Westheide and
Reimold, 2011), or drainage from the initially superheated impact
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182174

melt pool above into tensional fractures during central uplift for-
mation (Lieger et al., 2009; Riller et al., 2010). Pseudotachylitic
breccias are abundant in the central uplift of the Vredefort structure
and are particularly pervasive in the central core rocks, which were
uplifted furthest and from the greatest depths (Reimold and
Gibson, 2006).
4.6. Rheological considerations of target weakening
The extensive target deformation that occurs during the collapse
of the transient crater cavity and the formation of the central uplift is
inconsistent with standard strength properties of rocks. The paradox
of crater collapse readily becomes obvious by a simple analysis of
stresses surrounding a hemispherical cavity. For an uplift of the
transientcavity floor tooccur, stresses must exceed the strength of the
target material Y.ForaconstantY, large craters will collapse if their
transientcrater depth exceeds the ratio 15Y/
r
g,wheregis the planet’s
gravity and
r
is the target density (Melosh, 1977). Hence, this simple
model explains both the presence of a transition from simple crater
formationto complex crater formation and the 1/gdependence of the
crater size at which this transition occurs on different planetary
bodies (e.g., Kenkmann et al., 2012). Rigorous static analysis of cavity
slumpinghas shown that, for substantial rim collapseor floor uplift to
occur, the actualeffective strength mustbe less than w3MPa(Melosh,
1977), with little or no internal friction (McKinnon, 1978). Moreover,
the same lowstrength and friction are required by dynamic models of
crater formation (e.g., O’Keefe and Ahrens, 1999; Wünnemann and
Ivanov, 2003). Consequently the effective strength and friction
coefficient of the target rocks underneath the crater floor must
somehow be temporarily reduced. Providing a physical explanation
for the apparent transitory low strength of the target is an enduring
problem in impact cratering mechanics (Melosh, 1989; Melosh and
Ivanov, 1999; Senft and Stewart, 2009). In the following we briefly
outline possible weakening mechanisms operating during crater
modification:
4.6.1. Acoustic fluidization
An obvious weakening mechanism is fracturing and fragmen-
tation induced by the passage of the shock wave. A well-developed
theory for temporarily reducing friction both within a rock mass
and along fault zones is acoustic fluidization (Melosh, 1979, 1996).
The shock wave that passes through the target rocks generates
scattered seismic vibrations within the fractured rock mass
beneath and surrounding the crater and along the narrow fault
zones. These vibrations result in pressure fluctuations of the
ambient overburden pressure. During periods of low pressure,
frictional resistance is diminished, leading to slip events in low-
pressure zones. The time- and space-averaged effect of this pro-
cess is that the rock mass behaves rheologically as a viscous fluid
with a certain yield strength. Hence, this mechanism elegantly
explains the distributed brittle deformation and formation of
discrete blocks that are orders of magnitude smaller than the size of
the transient cavity. It also explains the apparent crater-scale
continuous deformation, as observed in many terrestrial craters.
The only pre-requisites are that the target is pervasively fractured
by the expanding shockwave, and that the scattered pressure-wave
Fig. 18. Pseudotachylite vein indicating a weak right-lateral offset with an extensional component (Vaal river banks, 6 km W of Parys, Vredefort crater, South Africa).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 175

field behind the shock has sufficient amplitude and longevity to
facilitate slip for the duration of transient cavity collapse. Slip
movements may generate additional seismic vibrations that pro-
vide a positive feedback for the weakening mechanism, and thus
enable large displacements along localized fault zones (Melosh,
1996). Simple parameterizations of the macro-scale effect of
acoustic fluidization have been used in numerical impact models
for many years (e.g., Melosh and Ivanov, 1999), producing dynamic
crater formation models that are in good agreement with obser-
vational constraints on cavity collapse, and reproducing the general
size-morphology progression of craters on crystalline planetary
surfaces (e.g., Wünnemann and Ivanov, 2003). Field evidence for
acoustic fluidization is documented by the presence of narrow-
spaced fault networks, blocks and mega-blocks, in particular if
the central uplift is composed of sedimentary strata. Oscillating
movements could only occasionally be documented (Kenkmann
et al., 2000a). If the central uplift is composed of crystalline rocks
indications for acoustic fluidization are sparse. At Manicouagan, the
anorthositic central uplift is very much coherent, with the original
pre-impact metamorphic foliation preserved over kilometers. Only
the presence of localized veins hints towards relative movements of
rocks (Spray, 2010; Biren and Spray, 2011).
4.6.2. Shock heating, frictional heating and rate-dependent
softening
The strength of rock substantially drops as their temperature
approaches the melting point (e.g., Stesky,1974). Shock heating and
the heat that remains in the rocks after shock-loading (post-shock
heat; Table 1) are included in all numerical models in use today. The
circumstance that almost all rocks that suffered strong shock
(>25 GPa) are involved in the excavation flow and become ejected
strongly reduces the potential of a substantial temperature increase
in the target. Thus, only in very large impacts like the 200-km
diameter Vredefort impact structure is shock and post-shock
heating an important weakening mechanism (e.g., O’Keefe and
Ahrens, 1999). At Vredefort shock-induced thermal softening is
restricted to the central uplift core that, in addition, was pre-heated
due to its pre-impact position at a deep crustal level.
Melting during the excavation and modification flow might
provide sufficient lubrication to lower the strength of block con-
tacts during the later stages of movement (Dence et al., 1977; Spray,
2010). Indeed, aligned melt networks of presumably frictional
origin (Spray and Thompson, 1995; Spray, 1997) were identified at
several large craters. In other cases, however, particularly in small
to mid-sized complex craters, pseudotachylitic veins are not found.
Moreover, although the total displacements and slip rates neces-
sary to generate melt along a fault are easily achieved during crater
formation, the volume of friction melt expected is very small (less
than a few volume percent) compared to the volume of collapsing
material (Melosh, 2005). Hence, whether or not sufficient friction
melt can be formed to lubricate crater collapse remains uncertain.
Dynamic friction shows a state- and rate-dependency. Slip
weakening and strain-rate weakening has been well documented
experimentally for different lithologies (e.g., Spray, 1988; Scholz,
2002; Di Toro et al., 2004). Strain-rate weakening dominantly oc-
curs at relatively low temperature and pressure under abrasive
wear conditions. During crater collapse very rapid, long-distance,
high strain motions occur under relatively low normal load, thus
these weakening mechanisms might play a relevant role. Senft and
Stewart (2009) implemented a parametric strain-rate weakening
model in a numerical hydrocode of crater formation that reduces
the friction in damaged cells that exceed specified minimum total
strain and strain-rate criteria. Despite a dependence of the detail of
the results on resolution, the model’s success in matching observed
features in large craters suggests that the temporary strain-rate
weakening of fault zones is sufficient to explain complex crater
collapse. However, the physical mechanism for this weakening is
uncertain and may be different in different target lithologies. Po-
tential explanations include: friction melting, pore-fluid pressuri-
zation, granular flow of fault gouge material, silica gel formation in
quartz-rich rocks, acoustic fluidization of fault gouge and flash
heating along asperities (Senft and Stewart, 2009 and references
therein).
To conclude, the low abundance of melt in most terrestrial
impact craters and the predominance of brittle deformation
suggest that the weakening mechanisms that enable crater collapse
are dominated by some type of dynamic frictional interaction of
cold rocks. This might be the internal friction in a breccia, the
friction between large blocks or the frictional resistance to slip
along discrete fault zones. Most impact researchers favor this
concept of acoustic fluidization to explain the temporary target
weakening during crater modification, although proving the theory
of acoustic fluidization is difficult in the field due to its transient
nature. A future and promising avenue is to implement the rate-
dependency of friction into models of acoustic fluidization.
4.7. Effects of oblique impact incidences on crater modification
The shape of the crater rim is largely insensitive to the impact
trajectory and remains circular with the exception of highly oblique
impacts (<10

with respect to the target surface) (Fig. 11). In fact,
there is only one confirmed crater on Earth that shows an elliptical
outline, namely Matt Wilson, NT, Australia (Kenkmann and
Poelchau, 2009). Still, a number of morphological crater features
have been cited as diagnostic of oblique impacts, such as a
depressed rim with a steepened inner slope uprange, a large central
uplift diameter relative to crater diameter, or an uprange offset of
the central peak (e.g., Schultz and Anderson, 1996). The latter indi-
cator, however, has been disputed by statistical analysis of Venusian
(Ekholm and Melosh, 2001) and lunar craters (Goeritz et al., 2009).
Morphological criteria are of limited use for the analysis of
terrestrial craters that are very rarely pristine. In contrast, the
eroded sub-surface beneath crater floors is often accessible and
enables the study of the cratering kinematics. Systematic de-
viations from axial symmetry were observed in numerous eroded
central uplifts that stimulated the hypothesis that they are formed
as a result of oblique incidences. It was recently possible to prove
this assumption through independent methods and to further the
understanding of these observations (Wulf et al., 2012).
4.7.1. Terrestrial case studies
The subsurface structures of some terrestrial impact craters
formed in originally flat-lying stratified bedrocks show a preferred
strike and dip orientation of strata and fault planes, and an
arrangement of folds that systematically deviate from axial sym-
metry and imply a preferred transport direction during the crater
modification process. Such deviations occur at Upheaval Dome, UT,
USA (Kenkmann et al., 2005), Spider, WA, Australia (Fig. 19)
(Shoemaker and Shoemaker, 1985), Gosses Bluff, NT, Australia
(Scherler et al., 2006), Matt Wilson, NT, Australia (Kenkmann and
Poelchau, 2009) and Jebel Waqf as Suwwan, Jordan (Kenkmann
et al., 2010)(Fig. 17)(Table 2). In the case of Matt Wilson, the
elliptical outline of the eroded impact structure independently
restricts the impact trajectory to two possible directions and, hence,
enables a correlation of the abundant structural features to the
impact trajectory. The structural features indicative for an oblique
incidence are: (i) dominance of strata dip in uprange direction and
strata strike perpendicular to the impact direction (Fig. 19A) (ii)
dominance of a thrust direction within the central uplift in
downrange direction leading to a stacking sequence (Fig. 19A), (iii)
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182176

Fig. 19. A) The structural arrangement of imbricate thrust stacks in the central uplift of Spider crater, WA, Australia, shows a bilateral symmetry and indicates a transport in
southerly direction. The thrust stacks surround the core of the central uplift. B) The arrangement of thrust faults is explained by combining a pure radially converging flow field
(idealized collapse flow field for vertical impacts) with parallel trajectories that transport material from uprange to downrange (as suggested for an oblique impact).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 177

bilateral symmetry of the central uplift, with an axis of symmetry
corresponding to the impact trajectory (Fig. 17;Fig. 19A), (iv)
occurrence of anticlines and synclines parallel to the symmetry axis
(Fig. 17), (v) normal dipping strata uprange and overturned strata
downrange, and (vi) normally plunging fold axes uprange and
overturned plunging axes downrange (Table 2)(Fig. 17). The
bilaterally symmetric arrangement of thrust stacks at the Spider
crater, Australia, (Fig. 19A) which are bent around the core of the
central uplift, deserves a special emphasis. To explain the change in
orientation of the thrust slices from downrange to uprange a very
simplified, in fact oversimplified, approach of vector summation in
the horizontal plane might be helpful. The idealized flow field
during crater collapse of a vertical impact is a radially converging
vector field. If this flow field is combined with a parallel flow field
with transport from uprange to downrange which results from
momentum transfer from an obliquely impacting projectile
(Fig. 19B), the resulting flow field displays curved trajectories
whose orientation fit to the arrangement of the thrust slices at
Spider. The deformation features studied in the craters listed above
suggest a downrange transport of rock and a central uplift that
initiates uprange and migrates downrange as the central uplift and
crater grows to its final size. This is in agreement with flow fields
inferred from sophisticated three-dimensional numerical models
of oblique impact cratering (Shuvalov and Dypvik, 2004;
Elbeshausen et al., 2009)(Fig. 20). Layered sedimentary rocks
with much less resistance to horizontal movement than to vertical
movement seem to be particularly susceptible to this type of
deformation.
4.7.2. Martian case studies
The unprecedented quality, resolution and availability of remote
sensing data of the Martian surface, in particular High Resolution
Table 2
Compilation of structural features observed in terrestrial and Martian impact craters that are characteristic for oblique impacts.
Structural criteria to infer the impact
trajectory in complex impact craters
Martian craters Terrestrial craters
Martin Bacau Arima Matt Wilson Upheaval Dome Spider Waqf as Suwwan
Ejecta pattern Asymmetry of the ejecta
blanket
xxx
Crater shape Elliptical crater shape x
Central uplift Strata strike preferentially
perpendicular to trajectory
xxxx x x
Central uplift Strata dip preferentially
up range
exx x x x
Central uplift Bilateral symmetry along
trajectory
xxxxxx
Central uplift Faults parallel to trajectory xxxx x x
Central uplift Stacking x x x x x
Central uplift Fold axes preferentially
parallel to trajectory
xxx x
Central uplift Overturned beds
downrange
x ? ? (x) x
References Wulf et al.,
2012. Icarus
Wulf et al.,
2012. Icarus
Wulf et al.,
2012. Icarus
Kenkmann
and Poelchau
2009. Geology
Kenkmann et al., 2005.
GSA-SP.; Scherler et al.,
2006. EPSL
Shoemaker and
Shoemaker, 1985.
Meteoritics
Kenkmann et al.,
2010. GSA-SP
Fig. 20. Snapshot of a three-dimensional numerical model of a 45oblique impact. (Image courtesy of Dirk Elbeshausen, Museum of Natural History, Berlin). Modeling shows that
the crater collapse starts uprange and progressively shifts downrange. This leads to the formation of an asymmetrical central uplift that displays a downrange vergency and contains
an imbricated inner structure. For further details concerning model parameters the reader is referred to Elbeshausen et al. (2009).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182178

Imaging Science Experiment (HiRISE) images up to 25 cm/pixel,
(McEwen et al., 2007), permit the application of structural geology
methods to Martian craters with a high measure of precision
(Fig. 21)(Table2), in particular if the rocks display a stratification.
For the study of oblique impacts, Martian impact craters offer an
ideal opportunity to verify terrestrial observations due to the fact
that structural trends in central uplifts can be combined with an
independent indicator of the impact direction, i.e., the asymmetric
ejecta blanket or an elliptical crater shape (Fig. 11), which is usually
lacking on Earth.
Structural analyses of oblique Martian impact craters show that
the (i) the strike of bedrock in the central uplift is predominantly
perpendicular to the impact trajectory (Fig. 21b) and (ii) the ma-
jority of faults trend parallel to the impact trajectory. The impact
direction could be clearly determined using the ejecta pattern as an
unequivocal indicator (Wulf et al., 2012). Formation of central up-
lifts during the modification stage of the impact cratering process is
the result of an inward and upward movement of rock towards the
crater center and may result, at least under ideal, non-oblique
conditions, in either concentric strike of the tilted bedrock in
more distal parts of the central uplift or in radial strike in the
central portion of the uplift due to constrictional faulting and
folding (Kenkmann, 2002). The preferred orientation of the origi-
nally horizontal to sub-horizontal layered bedrock of volcanic
origin in the central uplifts of the oblique Martian impact craters
clearly deviates from this idealized structure, indicating a preferred
transport direction during the crater formation process. The results
of the analyzed Martian impact craters are in good agreement
with observations from terrestrial impact craters and confirm
their classification as oblique impacts due to the specific internal
structure of their central uplifts, even without an available ejecta
blanket (Wulf et al., 2012).
4.7.3. Future research in the structural analysis of craters
A rigorous structural analysis and mapping of the deformation
inventory of terrestrial impact craters is the basis for understanding
the kinematics and dynamics of crater formation. High-quality maps
are still lacking for many craterson Earth. The increasingly extensive
data set of high-resolution remotesensing imagery of extraterrestrial
impact craters now enables comparative studies of terrestrial and
planetary craters and generally expands the field of activity for
structural geologists, not only in the context of impact cratering.
The expertise of structural geologists is particularly necessary and
helpful for the characterization of the micro-mechanisms of shock-
metamorphosed minerals and rocks. Relatively little attention has
been drawn to the effects of weak shock waves on mineral defor-
mation and on the rate-dependencies of brittle failure at high strain
rates. A better understanding of the rheological behavior of tran-
siently loaded rocks based on microstructural analysis would help to
narrow the gap between observation and modeling of crater
formation.
Acknowledgments
We very much appreciate the invitation by the editor-in-chief,
Prof. Dr. T. Horscroft, to prepare a review article on the structural
inventory of impact craters for the Journal of Structural Geology.
We hope that this contribution may stimulate the mutual
communication between impact and planetary researchers and
structural geologists. The work would not have been possible
Fig. 21. Example of a Martian impact crater showing the strong influence of obliquity on the internal structure of the central uplift. A) Topographical overview of Martin crater
(21.38S 290.73E) in Thaumasia Planum (superposed HRSC DTM over HRSC nadir). B) The mean strike trend of all measured layers of the central uplift is NWeSE (red line) and thus
nearly perpendicular to the assumed impact direction (modified after Wulf et al., 2012). C) Perspective view of the central part of the central upflift showing imbrication of layered
bedrock perpendicular to the impact direction (vertical exaggeration is 1.5, superposed HiRISE DTM (1 m/px) over HiRISE image (25 cm/px)).
T. Kenkmann et al. / Journal of Structural Geology 62 (2014) 156e182 179

without the sustained and ongoing support by the German
Research Foundation. The study of processes of crater collapse was
supported by the grants KE 732/6, KE 732/8, KE 732/19 and KE 732/
20. For the study of oblique craters we acknowledge grant KE 732/
11. Experimental studies on crater formation were supported by
grants FOR-887 and KE 732/16, KE 732/17 and KE 732/18. TK would
like to express special thanks to K. Wünnemann, D. Stöffler and
W.U. Reimold (Museum of Nat. History Berlin), A. Deutsch (Uni-
versity Münster) B.A. Ivanov and N.A. Artemieva (Russ. Acad. of Sci.
Moscow) G.S. Collins (Imperial College London), K. Thoma and F.
Schäfer (Ernst-Mach Inst Freiburg), numerous involved master and
Ph.D. students, and many other international collaborators for the
intensive and inspiring cooperation over the last decade. The paper
strongly benefited from the reviews of John Spray and Ulrich Riller
and the editorial handling by Cees Passchier.
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