Christian KOEBERL, Ao. Univ. Professor Dr.

(after: Koeberl, C., 1997, Impact cratering: The mineralogical and geochemical evidence. In: Proceedings, "The Ames Structure and Similar Features", ed. K. Johnson and J. Campbell, Oklahoma Geological Survey Circular 100, 30-54)
Abstract Shock Waves in Rocks - Hugoniot Equations
Introduction Shock Metamorphism
General Characteristics of Impact Craters Geochemistry and Detection of Meteoritic Components in Impactites
Recognition of Impact Structures Conclusions
Formation of Impact Craters Acknowledgements
_ References cited
Over the last 15 years, new studies related to the events that caused the extinction of the majority of life on earth at the end of the Cretaceous period have led to the hypothesis that a large-scale asteroid or comet impact occurred 65 Ma ago. In the past, impact cratering as a geological process has not been much appreciated by the general geological community, despite the fact that, on all other planets and satellites with a solid surface, impact cratering is the most important process that alters the surface at the present time, and during most of the history of the solar system. Detailed studies, mainly since the 1960s, have led to the recognition of about 150 impact structures on earth. Here, some fundamental mineralogical and geochemical properties of impact-derived rocks that are used to recognize impact craters are reviewed. The formation conditions of impact craters lead to pressure and temperature conditions in the target rocks that are significantly different from those reached during any internal terrestrial process. Among the most characteristic changes induced by the impact-generated shock waves are irreversible changes in the crystal structure of rock-forming minerals, such as quartz and feldspar. These shock metamorphic effects are characteristic of impact and do not occur in natural materials formed by any other process. In addition, geochemical methods are used to find traces of the meteoritic projectile in impact melt rocks and glasses. A complete and diligent mineralogical, petrological, and geochemical study is necessary, before any conclusions regarding an impact origin of geological structures can be reached.
During the 1980s and early 1990s, a lively debate was held in the geological community regarding the cause of the mass extinction that marks the end of the Cretaceous period, at the Cretaceous-Tertiary (K-T) boundary. Interest in the events at the K-T boundary was renewed by a publication by Alvarez et al. (1980), who found that the concentrations of the rare platinum group elements (PGEs; Ru, Rh, Pb, Os, Ir, and Pt) and other siderophile elements (e.g., Co, Ni) are enriched by up to 4 orders of magnitude in the thin clay layer marking the K-T boundary compared to those of normal terrestrial crustal rocks. These observations were interpreted by Alvarez et al. (1980) as the result of a large asteroid or comet impact, which caused extreme environmental stress. This hypothesis was later strongly supported by the finding of shocked minerals in the K-T boundary layer by Bohor et al. (1984, 1987). It turned out that one of the main problems impeding the acceptance of the theory that a large impact took place 65 Ma ago was a lack of detailed knowledge of impact cratering and shock metamorphic processes in the general geological community. Similar debates - regarding impact versus internal origin - have been held in discussing the origin of a variety of "unusual" structures around the world, including the Ames structure in Oklahoma. Thus, it seems useful to briefly review our basic knowledge of terrestrial impact craters and shock metamorphism. The discussion of general properties of impact craters is the topic of the paper by Grieve (this volume), while here I will review mainly mineralogical and geochemical aspects of impact structures.
Historically, the concept of impact cratering on earth has not been much accentuated in classical geological studies. The concept of classical Huttonian and Lyellian geology is that slow, endogenic processes lead to gradual changes in our geological record. In this uniformitarian view, internal forces are preferred over seemingly more exotic processes to explain geological phenomena that often give the impression of occurring over very long periods of time. In contrast, impact appears as an exogenic, relatively rare, violent, and unpredictable event, which violates every tenet of uniformitarianism. The explanation of craters on the moon or on earth as being of impact origin has been opposed by many geologists over much of this century. It is almost ironical that it was Alfred Wegener who published a little-known study (Wegener, 1922), in which he concluded that the craters on the moon are of meteorite impact origin. The history of study and acceptance of impact cratering over this century is somewhat similar to the record of the acceptance of plate tectonics (see, e.g., Mark, 1987; Hoyt, 1987; Marvin, 1990; Glen, 1994, for a historical account of impact crater studies).
Planetary exploration and extensive lunar research in the second half of our century led to the conclusion that essentially all craters visible on the moon (and many on Mercury, Venus, and Mars) are of impact origin. Therefore, it has to be concluded that, over its history, the earth was subjected to a larger number of impact events than the moon. Part of the reason why this conclusion was not widely accepted among geologists may be that terrestrial processes (weathering, plate tectonics, etc.) effectively work to obliterate the surface expression of these structures on earth. Through studies of the orbits of asteroids and comets, astronomers have a relatively good understanding of the rate with which these objects strike the earth (e.g., Shoemaker et al., 1990, Weissman, 1990). For example, minor objects in the solar system with diameters >=1 km (mainly asteroids) collide with the earth at a frequency of about 4.3 impacts per million years (Shoemaker et al., 1990), and each such impact forms a crater >=10 km in diameter. Impactors about 2 km in diameter collide with the earth about every 1 - 2.106 years. Impact of earth-orbit crossing asteroids dominate the formation of craters on earth that are smaller than about 30 km in diameter, while comet impact probably forms the majority of craters that are larger than about 50 km in diameter (Shoemaker et al., 1990). However, the orbits of asteroids are better known than those of comets, because many of the latter have such long periodicities that no appearance has yet been observed during the time of human civilization.
In an important historical and sociological evaluation of the K-T boundary debates, Glen (1994, p. 52) found that "resistance to the [impact] hypothesis seemed inverse to familiarity with impacting studies." Thus, planetary scientists, astronomers, and meteoriticists, have grown accustomed to view "large-body impact as a normal geological phenomenon - something to be expected throughout earth history - but another group, the paleontologists, is confounded by what appears to be an ad hoc theory about a nonexistent phenomenon" (Raup in Glen, 1994, p. 147). Thus, we may conclude that one scientist's uniformitarianism is another scientist's deus ex machina.
However, it may be important to consider the time scales involved in this discussion. What geologists have called "uniformitarianism" is the result of integrating individual catastrophes of various magnitudes over a sufficiently long time span. Earthquakes, volcanic eruptions, landslides, etc., are locally devastating if time spans of maybe 20-100 years are concerned, but if the whole world and longer time spans are concerned, these "catastrophes" become part of the "uniformitarian" process of explosive volcanism, earthquake history, or erosion. The bias in what is considered uniformitarian is related to the life span of humans and the human civilization. As large meteorite impacts have not been observed during the last few millennia (with rare exceptions, such as the Tunguska event, which occurred in a remote tundra location of Siberia in 1908, but even this event was too small to produce a crater), such events tend to be neglected when constructing the "uniformitarian" history of the earth. The falls of small meteorites have been observed quite frequently. There is no real conflict between uniformitarianism and meteorite impact. We just have to learn to apply the same principle that is being used for extrapolating the frequency of volcanic eruptions and earthquakes also to the scaling of meteorite impacts - the large and devastating ones occur less often than the small events.
About 150 impact structures are currently known on earth (e.g., Grieve and Shoemaker, 1994; Grieve, this volume). However, it is somewhat embarrassing that almost two thirds of the confirmed or probable impact craters in the USA have only been studied superficially (see Koeberl and Anderson, 1996). Considering that some impact events severely affected the geological and biological evolution on earth and that even small impacts can disrupt the biosphere and lead to local devastation (Chapman and Morrison, 1994), the understanding of impact structures and the processes by which they form should be of interest not only to earth scientists, but also to society in general.
General Characteristics of Impact Craters
As no large impact event has been observed by humans over the last several thousand years (which is, of course, not a geologically long period of time), impact experiments and the detailed study of impact craters on earth are essential to understand these features. During an impact event, the geological structure of the target area is changed in a characteristic way, which can be used to help distinguish volcanic structures from meteorite impact craters. Meteorite impact craters are circular surfical features without deep roots, while in volcanic structures the disturbances continue to (or, rather, emerge from) great depth. Impact craters are practically always circular, with only very few exceptions that result either from highly oblique impacts (see, e.g., the Rio Cuarto structures in Argentina - Schultz et al., 1994), or from post-formational distortion due to, for example, tectonism or erosion (e.g., the Sudbury structure in Canada - e.g., Stöffler et al., 1994). It is useful to distinguish between the impact crater, i.e., the feature that results from the impact, and the impact structure, which is what we observe today, long after formation and modification of the crater.
Impact craters (before erosion) occur in two distinctly different morphological forms, namely as small bowl-shaped craters (<=4 km diameter), and large (>=4 km diameter) complex craters with a central uplift. All craters have an outer rim and some crater infill (e.g., brecciated and/or fractured rocks, impact melt rocks), while the central structural uplift in complex craters consists of a central peak or of one or more peak ring(s) and exposes rocks that are uplifted from considerable depth. The diameters of impact craters on Earth show a variation, which is, however, the result of biased processes, chiefly different effects of age and differential erosion of large and small craters. The erosional processes that obliterate small (0.5-10 km diameter) craters after a few million years create a severe deficit of these craters, compared to the number that is expected from the number of larger craters and astronomical observations (Grieve and Shoemaker, 1994). This also explains why most small craters are young. Older craters of larger initial diameter also suffer erosion degradation leading to the destruction of the original topographical expression, or to burial of the structures under post-impact sediments. For details on crater morphology, see Grieve (this volume).
Recognition of Impact Structures
As a consequence of the obliteration, burial, or destruction of impact craters on earth, they can be difficult to recognize, requiring the development of diagnostic criteria for the identification and confirmation of impact structures. The most important of these characteristics are: a) evidence for shock metamorphism, b) crater morphology, c) geophysical anomalies, and d) the presence of meteorites or geochemical discovery of traces of the meteoritic projectile. Of these, only the presence of diagnostic shock metamorphic effects and, in some cases, the discovery of meteorites, or traces thereof, can provide unambiguous evidence for an impact origin.
However, morphological and geophysical observations are important in providing supplementary - but not confirming - evidence. Geophysical methods are also useful in identifying candidate sites for further studies. It should be noted that in complex craters the central uplift usually contains severely shocked material and is often more resistant to erosion than the rest of the crater. In old eroded structures the central uplift may be the only remnant of the crater that can be identified. Geophysical characteristics of impact craters that have been investigated include gravity, magnetic properties, reflection and refraction seismics, electrical resistivity, and others (see Pilkington and Grieve, 1992, for a review). In general, simple craters have negative gravity anomalies due to the lower density of the brecciated rocks compared to the unbrecciated target rocks, whereas complex craters often have a positive gravity anomaly associated with the central uplift that is surrounded by an annular negative anomaly. Magnetic anomalies can be more varied than gravity anomalies, but seismic data show the loss of seismic coherence due to structural disturbance, slumping, and brecciation. Such geophysical surveys are important for the recognition of anomalous subsurface structural features, which may be deeply eroded craters or simply covered by post- impact sediments (e.g., in the U.S.: Ames, Avak, Chesapeake Bay, Manson, Newporte, Red Wing Creek - see Koeberl and Anderson, 1996; Koeberl and Reimold, 1995a,b; Koeberl et al., 1995b, 1996b, 1996c). However, to better appreciate the other criteria for identification of impact structures, we need to briefly consider some physical processes that operate during crater formation.
Formation of Impact Craters
The formation of a crater by hypervelocity impact is - not only in geological terms - a very rapid process that is usually divided into three stages: 1) contact/compression stage, 2) excavation stage, and 3) post-impact crater modification stage. Crater formation processes have been studied for many decades, but space limitations require that the reader is referred to the literature (see, e.g., Gault et al., 1968; Roddy et al., 1977; Melosh, 1989; and references therein) for a detailed discussion of the physical principles of impact crater formation. Here, only a few key ideas can be mentioned.
During the impact of a large meteorite, asteroid, or comet, large amounts of kinetic energy (equal to 1/2mv2, m = mass, v = velocity) are released. Earlier in the century, the amount of energy was largely underestimated, because the velocities with which extraterrestrial bodies hit the earth had not been known or assessed properly. However, any body that is not slowed down by the atmosphere will hit the earth with a velocity between about 11 and 72 km/s. These velocities are determined by celestial mechanics. Thus, a 250-m-diameter iron or stony meteorite has a kinetic energy roughly equivalent to about 1000 megatons of TNT, which would lead to the formation of a crater about 5 km in diameter. The relatively small Meteor (or Barringer) crater in Arizona (1.2 km diameter) was produced by an iron meteorite of about 30-50 m in diameter. Many of the characteristics of an impact crater are the consequence of the enormous kinetic energy that is released almost instantaneously during the impact. This energy can be compared to that of "normal" terrestrial processes, such as volcanic eruptions or earthquakes. During a small impact event, which may lead to craters of 5-10 km in diameter, about 1024-25 ergs (1017-18 J) are released, while during formation of larger craters (50-200 km diameter) about 1028-30 ergs (1021-23 J) are liberated (e.g., French, 1968; Kring, 1993). On the other hand, about 6·1023 ergs (6.1016 J) were released over several months during the 1980 eruption of Mount St. Helens, and 1024 ergs (1017 J) in the big San Francisco earthquake in 1906. It may also be surprising that the total annual energy release from the earth (including heat flow, which is by far the largest component, as well as volcanism and earthquakes) is about 1.3.1028 ergs (1.3.1021 J/y) (French, 1968; Sclater et al., 1980; Morgan, 1989). The latter amount of energy is comparable to the energy that is released almost instantaneously during large impact events. It is also important to realize that the energy that is liberated during an impact is concentrated at almost a point on the Earth's surface, leading to an enormous local energy density.
Shock Waves in Rocks - Hugoniot Equations
Structural modifications and phase changes in the target rocks occur during the compression stage, while the morphology of a crater is defined in the second and third stage. For a more detailed description of crater formation, see, e.g., Grieve (1987, 1991), Melosh (1989), and references therein. During this early impact phase, the impacting body is stopped after about 2 projectile radii and the kinetic energy (1/2mv2) is transformed into heat and shock waves that penetrate into the projectile and target. The most important phenomenon, which is characteristic of impact, is the generation of a supersonic shock wave that is propagated into the target rock. The effects of shock waves on matter are well understood from decades of experimental evidence. The following discussion is based mainly on information from Melosh (1989). Matter is being accelerated very rapidly and, as a consequence of the decrease of compressibility with increasing pressure, the resulting stress wave will become a shock wave moving at supersonic speed (up to about 2/3 of the impact velocity). Shock waves are inherently nonlinear and shock fronts are abrupt. They can be mathematically represented as a discontinuous jump of pressure, density, particle velocity, and internal energy. In reality, shock waves have a finite thickness, which is, however, very limited. For example, the widths of shock waves in gas are limited to about 10 micrometers, which is roughly equal to one molecular mean free path, but shock waves in solids are wider, up to a few meters in rocks, depending on their porosities.
The shock wave leads to compression of the target rocks at pressures far above a material property called the Hugoniot elastic limit. The Hugoniot elastic limit (HEL) can generally be described as the maximum stress in an elastic wave that a material can be subjected to without permanent deformation. Above this limit plastic, or irreversible, distortions occur in the solid medium through which the compressive wave travels (see, e. g., compilations by Roddy et al., 1977; Melosh, 1989; and references therein). The value of the HEL is about 5-10 GPa for most minerals and whole rocks. For example, single crystals of quartz have HELs ranging from 4.5 to 14.5 GPa (depending on the crystal orientation), for feldspar the HEL is at 3 GPa, and for olivine at 9 GPa. For rocks, the HEL of dolomite is 0.3 GPa, for granite 3 GPa, and for granodiorite, 4.5 GPa. The only known process that produces shock pressures exceeding the HELs of most crustal rocks and minerals in nature is impact cratering. Volcanic processes are not known to exceed 0.5 to 1 GPa. In addition to structural changes, phase changes may occur as well.
For a thermodynamical treatment of shock fronts travelling through matter, the so-called Hugoniot equations are used (see Melosh, 1989). These equations link the pressure P, internal energy E, and density r in front of a shock wave (uncompressed: P0, E0, r0) to values after the shock front (compressed: P, E, r). The density is also expressed as the specific volumes V = 1/r and V0 = 1/r0 for the compressed and uncompressed cases, respectively. Initial pressure, energy, and density before the shock are known values, while the respective values after the shock are unknown quantities, as are the shock velocity U and particle velocity up behind the shock front. The Hugoniot equations are then written as:
r(U - up) = r0U
P - P0 = r0upU
E - E0 = (P + P0)(V0 - V)/2

These equations express the conservation of mass, momentum, and energy across the shock front to reduce the number of unknown variables from five to two. For a derivation of the Hugoniot equations, see Appendix 1 in Melosh (1989), and Boslough and Asay (1993). In the uncompressed material, the initial particle velocity should be zero, and the initial pressure P0 can be neglected, yielding the approximation E - E0 = up2/2. In addition to the three equations mentioned above, a forth one, the equation of state, is necessary to specify conditions on either side of the shock front. This equation links pressure, specific volume (density), and internal energy: P = (V,E). Equations of state have been determined experimentally for a large number of different materials (e.g., Marsh, 1980).
The shock wave equation of state data can be plotted in pressure versus specific volume (Fig. 1) or shock velocity versus particle velocity diagrams. The curves in these diagrams are not equivalent to conventional equilibrium in thermodynamical P,V diagrams, but represent loci of several individual shock events, i.e., each point on a curve is the result of one particular shock wave compression event. The HEL appears as a kink in the shock curve, indicating yielding at the maximum stress of the elastic wave (Fig. 1).
After the shock wave passes, the high pressure is released by a so-called rarefaction, or release, wave, which trails the shock front. The rarefaction wave is a pressure, not shock, wave and travels the speed of sound in the shocked material. As it moves faster than the shock wave, it gradually overtakes the shock front and causes a decrease in pressure with increasing distance of propagation. While the pressure behind a rarefaction wave may drop to near zero, the residual particle velocity actually accelerates material, leading to impact crater excavation. In addition, the rarefaction wave does not only conserve mass, energy, and momentum (as the shock wave does), but also entropy. Thus, rarefaction is a thermodynamically reversible adiabatic process, while shock compression is thermodynamically irreversible. During shock compression, a large amount of energy is being introduced into a rock. Upon decompression, the material follows a release adiabat in a pressure versus specific volume diagram. The release adiabat is located close to the Hugoniot curve, but usually at generally somewhat higher P and V values, leading to excess heat appearing in the decompressed material, which may result in phase changes (e.g., melting or vaporization). The effects of the phenomena described above can be observed in various forms in shocked minerals and rocks.
Shock Metamorphism    
In late January 2005 a sampling meeting (called a “sampling party”) was held at the ICDP headquarters in Potsdam, Germany. Over a dozen research tems sampled the two dep drill cores. Samples were distributed in the spring of 2005 and first scientific results became available in early 2006.
As some recent literature indicates, there is still some incomplete understanding in the geological community about the precise nature of shock metamorphism (for a discussion, see, e.g., French, 1990; Sharpton and Grieve, 1990). In contrast to some assertions (e.g., Lyons et al., 1993), the existence of definite shock metamorphic features in volcanic rocks has never been substantiated (see, e.g., de Silva et al., 1990; Gratz et al., 1992b). Static compression, as well as volcanic or tectonic processes, yield different products because of lower peak pressures and strain rates that are different by more than 11 orders of magnitude. It should be reaffirmed that the study of the response of materials to shock is not a recent development, but has been the subject of thorough investigations over several decades, in part stimulated by military research. Numerous shock recovery experiments (i.e., controlled shock wave experiments, which allow the collection of the shocked samples for further studies), using various techniques, have been performed in the last three to four decades. These experiments have led to a good understanding of the conditions for formation of shock metamorphic products and a pressure-temperature calibration of the effects of shock pressures up to about 100 GPa (see, e.g., Hörz, 1968; French and Short, 1968; Stöffler, 1972, 1974; Gratz et al., 1992a,b; Huffman et al., 1993; Stöffler and Langenhorst, 1994; and references therein).


Fig. 1. Idealized representation of a Hugoniot equation of state curve. The Hugoniot curve does not represent a continuum of states as in thermodynamical diagrams, but the loci of individual shock compression events. The yielding of the material at the Hugoniot Elastic Limit is indicated. See text for detailed discussion.


Fig. 2. Comparison of pressure-temperature fields of endogenic metamorphism and shock metamorphism. Also indicated are the onset pressures of various irreversible structural changes in the rocks due to shock metamorphism. The curve on the right side of the diagram shows the relation between pressure and post-shock temperature for shock metamorphism of granitic rocks. (After Grieve, 1987, and B. M. French, personal communication, 1995)
Table 1 lists the most characteristic products of shock metamorphism, as well as the associated diagnostic features. The best diagnostic indicators for shock metamorphism are features that can be studied easily by using the polarizing microscope. They include planar microdeformation features, optical mosaicism, changes in refractive index, birefringence, and optical axis angle, isotropization, and phase changes.

Before discussing the various shock metamorphic features, the type and location of impactite lithologies should be mentioned (Fig. 3). In an impact crater, shocked minerals, impact melts, and impact glasses are commonly found in various impact-derived breccias. Well-preserved ejecta at the crater rim may display a stratigraphic sequence that is inverted compared to the normal stratigraphy in the area. The impact process leads to the formation of various breccia types (e.g., Fig. 4), which are found within and around the resulting crater (see also Stöffler and Grieve, 1994, and Koeberl et al., 1996a). The three main breccia types include monomict or polymict breccias consisting of 1) cataclastic (fragmental), 2) suevitic (fragmental with a melt fragment component), or 3) impact melt (melt breccia - i.e., melt in the matrix- with a clastic component) breccias. The breccias can be allochthonous or autochthonous. In addition, dikes of injected or locally formed fragmental or pseudotachylitic breccias (Reimold, 1995), which contain evidence of melting, can be found in the basement rocks. The schematic distribution of breccias, melt, and breccia dikes at a simple crater is shown in Fig. 3. Whether these various breccia types are indeed present in a crater depends on factors including the size of the crater, the composition, and the porosity of the target area (e.g., Kieffer and Simonds, 1980), and the level of erosion (see, e.g., Roddy et al., 1977; Hörz, 1982; Hörz et al., 1983; Grieve, 1987; and references therein).
Shatter Cones
The occurrence of shatter cones has long been discussed as a good macroscopic indicator of shock effects, and a variety of structures were proposed to be of impact origin on the basis of shatter cone occurrences (e.g., Dietz, 1968; Milton, 1977). Such cones have also been formed in (chemical) explosion crater experiments (see, e.g., Milton, 1977). Their formation is dependent on the type of target rock (i.e., they are better developed in certain lithologies than in others) and has been estimated to take place at pressures in the range of 2 to 30 GPa. In general, shatter cones are cones with regular thin grooves (striae) that radiate from the top (the apex) of a cone. They can range in size from less than one centimeter to more than one meter (Fig. 5). Shatter cones occur mostly in the outer and lower parts of a crater and may be preserved even if a structure is deeply eroded.

Unfortunately, conclusive criteria for the recognition of "true" shatter cones have not yet been defined. If they are strongly eroded, it is possible to confuse concussion features, pressure-solution features (cone-in-cone structure), or abraded or otherwise striated features with shatter cones. It would be important to arrive at some generally accepted criteria for the correct identification of shatter cones, as some impact craters have been identified almost exclusively by the occurrence of shatter cones (see, e,g., compilation by Koeberl and Anderson, 1996; cf. Koeberl et al., 1996b). However, shatter cones are important as potential macroscopic shock indicators, as they are developed in large volumes of rock, and are useful as a guide for the presence of more definitive shock indicators, such as shocked minerals (see below).
Mosaicism is a microscopic effect of shock metamorphism and appears as an irregular mottled optical extinction pattern (Fig. 6a), which is distinctly different from undulatory extinction that occurs in tectonically deformed quartz. Mosaicism can be measured in the optical microscope by determining the scatter of optical axes in different regions of crystals showing mosaicism. Mosaicism can be semiquantitatively defined by X-ray diffraction studies of the asterism of single crystal grains, where it shows up as characteristic increase (with increasing shock) of the width of individual lattice diffraction spots in diffraction patterns. Highly shocked quartz crystals show a diffraction pattern that becomes similar to a powder pattern, because of shock-induced polycrystallinity. Many shocked quartz grains that show planar microstructures also show mosaicism. In addition, it should be noted that the crystal lattice of shocked quartz shows lattice expansion above shock pressures of 25 GPa, leading to an expansion of the cell volume by <=3% (Langenhorst, 1994).
Planar Microstructures
Two types of planar microstructures are apparent in shocked minerals: planar fractures (PFs) and planar deformation features (PDFs). Their characteristics are summarized in Table 2. PDFs in rock-forming minerals (e.g., quartz, feldspar, or olivine) are generally accepted to be diagnostic evidence for shock deformation (see, e.g., French and Short, 1968; Stöffler, 1972, 1974; Alexopoulos et al., 1988; Sharpton and Grieve, 1990; Stöffler and Langenhorst, 1994). PFs, in contrast to irregular, non-planar fractures (which are caused by rarefaction waves), are thin fissures, spaced about 20 micrometers or more apart, which are parallel to rational crystallographic planes with low Miller indices, such as (0001) or {1011} (e.g., Engelhardt and Bertsch, 1969). PFs form at lower pressures than PDFs, and may not provide conclusive evidence of shock metamorphism, but can act as guide to other more characteristic shock deformation effects.

PDFs, together with the somewhat less definitive planar fractures (PFs), are well developed in quartz (Stöffler and Langenhorst, 1994). PDFs are parallel zones with a thickness of <=1-3 mm and are spaced about 2-10 micrometers apart (see Figs. 6a- e). The degree of planarity of the individual sets of PDFs is an important parameter for the correct identification of bona fide PDFs and allow their distinction from (sub-)planar features that are produced at lower strain rates, e.g., in tectonically deformed quartz. It was demonstrated in Transmission Electron Microscopy (TEM) studies (see, e.g., Goltrant et al., 1991) that PDFs consist of amorphous silica. The structural state of the glassy lamellae is, however, slightly different from that of regular silica glass (Goltrant et al., 1991). The fact that the PDF lamellae are filled by glass allows them to be preferentially etched by, e.g., hydrofluoric acid, emphasizing the planar deformation features (see Fig. 6c). Figures 6a-e show various appearances of PDFs in natural samples from impact structures in the United States.

Fig. 6. Shocked quartz and feldspar. a) Quartzitic clast in impact melt rock from the Manson crater, Iowa, showing two prominent sets of PDFs and shock mosaicism; crossed polars, width of image 2.2 mm (see Koeberl et al., 1996a); b) Close-up of K-feldspar grain from the Ames structure, Oklahoma, 18-4 Nicor Chestnut core, sample 9011.0 (from 2747 m depth), showing incipient brecciation in the feldspar grain, which contains three sets of PDFs and shows the closely spaced nature of the lamellae; width of image 900 micrometers, crossed polars (courtesy W.U. Reimold, Univ. of the Witwatersrand). c) SEM image of quartz grain from the K-T boundary layer at DSDP Site 596 (Southwest Pacific), after brief etching with HF, showing three different sets of PDFs (courtesy B. Bohor, U.S. Geological Survey); d) shocked quartz grain from a drill core into the Red Wing Creek impact structure, North Dakota, depth 2301 m, within brecciated Kibbey sandstone, with PDFs of 2 different orientations; width of image 375 micrometers, crossed polars (see Koeberl and Reimold, 1995a, Koeberl et al., 1996b); e) quartz grain from the Newporte crater, North Dakota (Koeberl and Reimold, 1995b), with three sets of PDFs, in granitic clast from granitic fragmental breccia D9462.2, parallel polars, 355 micrometers wide.
Engelhardt and Bertsch (1969) have classified PDFs into four groups: (a) non-decorated PDFs (extremely fine lamellae, cannot be resolved in the optical microscope), (b) decorated PDFs (the lamellae are lined by, or replaced with, small spherical or elliptical bubbles, often representing fluid inclusions), (c) homogeneous lamellae (thicker lamellae that can be resolved in the microscope), and (d) filled PDFs (where the lamellae are filled with very small fine-grained crystals). Types (a) and (b) are the most common ones.
In addition, TEM studies have shown that there is a second type of PDFs, which consists of very thin multiple lamellae of Brazil twins. Brazil twins have been observed in hydrothermally grown quartz, but always parallel to the {1011} plane, while the impact-derived Brazil twins form at pressures >8 GPa, are of mechanical origin, and are exclusively parallel to the (0001) plane (Goltrant et al., 1991, Leroux et al., 1994). These latter authors have documented that such Brazil twins, from the Vredefort impact structure in South Africa, were formed due to annealing of the shocked rocks.
Most rock-forming minerals, as well as accessory minerals, such as zircon (Fig. 7), develop PDFs. The occurrence of diagnostic shock features is by far the most important criterion for evaluating the impact origin of a crater, particularly when several of the features that are typical of progressive shock metamorphism, as listed in Table 1, have been found. The occurrence of PDFs and PFs can be used, together with other shock effects, to determine the maximum shock pressure in impactites (Fig. 8). Most commonly, quartz is used to study these shock effects, as it is the simplest, best studied, and most widely distributed rock-forming mineral that develops PDFs.
PDFs occur in planes corresponding to specific rational crystallographic orientations. In quartz, the most abundant mineral that develops distinctive PDFs, the (0001) or c (basal), {1013} or w, and {1012} or p orientations are the most common ones. In addition, PDFs often occur in more than one crystallographic orientation per grain. With increasing shock pressure, the distances between the planes decrease, and the PDFs become more closely spaced and more homogeneously distributed over the grain, until at about >=35 GPa complete isotropization has been achieved. Depending on the peak pressure, PDFs are observed in 2 to 10 (maximum 18) orientations per grain. To properly characterize PDFs, it is necessary to measure their crystallographic orientations by using either a universal stage (Reinhard, 1931; Emmons, 1943) or a spindle stage (Medenbach, 1985), or by transmission electron microscopy (TEM; see, e.g., Goltrant et al., 1991; Gratz et al., 1992a; Leroux et al., 1994).
It is possible to use the relative frequencies of the crystallographic orientations of PDFs to calibrate shock pressure regimes, as given in Table 3 (see, e.g., Robertson et al., 1968; Hörz, 1968; Stöffler and Langenhorst, 1994). For example, at 5 to 10 GPa, PDFs with (0001) and {1011} orientations are formed, while PDFs with {1013} orientations start to form between about 10 and 12 GPa. Such studies are done by measuring the angles of the c-axis and of the PDFs in individual quartz grains with a universal stage. In a stereographic projection (Fig. 9), the optical axis (c-axis) is rotated into the center of projection, the locations of the poles of PDFs are plotted, and then those positions are compared with the stereographic projection of the rational crystallographic planes in quartz (as listed in Fig. 9).

Fig. 9. Standard stereographic projection (lower hemisphere) of rational crystallographic planes in alpha-quartz, which is used to index crystallographic planes of PDFs based on universal-stage measurements. The arrows indicate the three a- axes of quartz, and the c-axis (the (0001) plane) is in the center of the projection. Also indicated are the low Miller indices in a part of the diagram (other indices can be derived from crystal symmetry). The circles are about 5º in diameter and indicate the accuracy of the U-stage measurements (see, e,g., Engelhardt and Bertsch, 1969).
The measured angles that fall within 5º of the theoretical polar angle of the plane are considered valid and can be indexed. Figure 10 shows the results of this procedure in the form of a histographic plot of indexed PDFs. Such plots are used to identify the relative frequencies, in which the various shock- characteristic crystallographic orientations occur.
The pre-shock temperature of a target rock also influences the formation and distribution of PDFs. Reimold (1988) and Huffman et al. (1993) presented the results of shock experiments with quartzite at room temperature (25ºC), and preheated to 450ºC and 750ºC. They noticed a slight difference in the relative distribution of the {1013} and {1012} orientations and a large difference in the number of PDF sets per grain (Fig. 11). Langenhorst (1993) compared PDF orientations in shocked quartz single crystals preheated to a higher temperature than Huffman et al. (1993) and found a distinct change in the relative frequencies of the {1013} and {1012} orientations.
Fig. 10. Crystallographic orientation of PDFs in quartz from the Newporte (North Dakota) impact structure, shown as a histogram giving the frequency of indexed PDFs versus angle between c-axis and poles of PDFs, without plotting unindexed planes (see Grieve and Therriault, 1995); the shock-characteristic orientations (0001), {1013}, {1012}, {1122}, {1011}, {0111}, and {1121} (c, w, p, x, r,z, and s, respectively), are dominating. (After Koeberl and Reimold, 1995b)
Bulk Optical and other Properties
Recent experimental evidence shows that there is a decrease of the density of shocked quartz with increasing shock pressure (Langenhorst, 1993). At shock pressures up to about 25 GPa, only a slight decrease is noticeable, followed by a significant drop in density between 25 and 35 GPa, depending on the direction of the shock wave relative to the c-axis of the quartz crystal, and the pre-shock temperature (Fig. 12). Optical properties, such as the birefringence of quartz and the refractive index, show also an inverse relationship with shock pressure in the 25 to 35 GPa range (Fig. 13). At 35 GPa, isotropization (formation of diaplectic quartz glass) occurs. Figure 13 also indicates that with increasing shock pressure the birefringence (nw - ne) decreases. Still other properties of shocked minerals can be used to either confirm a shock history or calibrate shock pressures. For example, intensity and wavelength of infrared absorption bands, the electron paramagnetic resonance, and peak width in a 29Si magic angle spinning nuclear magnetic resonance (NMR) spectrum all depend in a quantitative way on the shock pressure (e.g., Boslough et al., 1995; references in Stöffler and Langenhorst, 1994).
Fig. 11. Histogram with crystallographic orientation of PDFs in quartz from Hospital Hill quartzite, showing the dependency of the orientations on the pre-shock temperature (after Huffman et al., 1993). a) Pre-shock temperature 25ºC, shock pressure 28 GPa. b) Pre-shock temperature 440ºC, shock pressure 28 GPa. The main difference between the sets is that about half of the quartz grains in the high-temperature experiment remain unshocked, while in the low-temperature experiment, almost all quartz grains are shocked.
Bulk Optical and other Properties
Recent experimental evidence shows that there is a decrease of the density of shocked quartz with increasing shock pressure (Langenhorst, 1993). At shock pressures up to about 25 GPa, only a slight decrease is noticeable, followed by a significant drop in density between 25 and 35 GPa, depending on the direction of the shock wave relative to the c-axis of the quartz crystal, and the pre-shock temperature (Fig. 12). Optical properties, such as the birefringence of quartz and the refractive index, show also an inverse relationship with shock pressure in the 25 to 35 GPa range (Fig. 13). At 35 GPa, isotropization (formation of diaplectic quartz glass) occurs. Figure 13 also indicates that with increasing shock pressure the birefringence (nw - ne) decreases. Still other properties of shocked minerals can be used to either confirm a shock history or calibrate shock pressures. For example, intensity and wavelength of infrared absorption bands, the electron paramagnetic resonance, and peak width in a 29Si magic angle spinning nuclear magnetic resonance (NMR) spectrum all depend in a quantitative way on the shock pressure (e.g., Boslough et al., 1995; references in Stöffler and Langenhorst, 1994).

Diaplectic Glass
At shock pressures in excess of about 35 GPa, diaplectic glass is formed (Table 1), which has been found at numerous impact craters. It is an isotropic phase that preserves the crystal habit, original crystal defects, and, in some cases, planar features. It forms without melting by solid-state transformation and has been described as a phase "intermediate between crystalline and normal glassy phases" (Stöffler and Hornemann, 1972). For example, maskelynite forms from feldspar. Diaplectic glass has a refractive index that is slightly lower, and a density that is slightly higher, than that of synthetic quartz glass. At pressures that exceed about 50 GPa, lechatelierite, a mineral melt, forms by fusion of quartz. Other minerals also undergo melting (fusion) at similar pressures. This complete melting is not the same process that results in the formation of diaplectic glass. The distinction between diaplectic glass and lechatelierite (both after quartz) was described by Stöffler and Hornemann (1972) and Stöffler and Langenhorst (1994).
High-Pressure Polymorphs
Another form of shock deformation are phase transitions to high-pressure polymorphs of minerals in a solid state transformation process. Such transformation can be predicted from Hugoniot data. Many minerals form metastable high- pressure phases (Stöffler, 1972). These include (density in g/cm3 is given in parentheses): stishovite (4.23) and coesite (2.93) from quartz (2.65); jadeite (3.24) from plagioclase (2.63-2.76), majorite (3.67) from pyroxene (3.20-3.52); and ringwoodite (3.90) from olivine (3.22-4.34). In contrast to expectations from the equilibrium phase diagram of quartz, stishovite forms at lower pressures than coesite, probably because stishovite forms directly during shock compression, while coesite crystallizes during pressure release.
The formation probabilities and conditions for these phases are strongly dependent of the porosity of the target rocks. While stishovite has never been found in any natural non-impact related rocks, there are rare findings of coesite within metamorphic rocks or in kimberlites. However, coesite within metamorphic rocks occurs as large single crystals within, or associated with, high-pressure minerals of metamorphic or volcanic origin, but never associated with quartz. On the other hand, impact-derived coesite occurs as fine-grained, colorless to brownish, polycrystalline aggregates of up to 200 micrometers in size, which are usually embedded in diaplectic quartz, or, rarely, in nearly isotropic shocked quartz. In addition to morphological differences, shock-produced coesite occurs in a disequilibrium assemblage of quartz-coesite-stishovite-glass (see also Grieve et al., 1996).
In addition to high-pressure phases of rock-forming minerals, impact-derived diamonds (the high pressure modification of carbon) have also been found at various craters. These diamonds form from carbon in the target rocks, mainly graphite- bearing (e.g., graphitic gneiss) or coal-bearing rocks (Koeberl et al., 1995a). Impact diamonds commonly preserve the crystal habit of their precursor material, which is mostly graphite (some are coal-derived). The diamonds that formed after graphite are called "apographitic" diamonds. Many of them were found to contain up to several 10 vol% lonsdaleite, the rare hexagonal diamond modification.
Mineral and Rock Melts
At pressures in excess of about 60 GPa, rocks undergo complete (bulk) melting to form impact melts (see Table 1). The melts can reach very high temperatures due to the passage of shock waves that generate temperatures far beyond those commonly encountered in normal crustal processes or in volcanic eruptions. This is shown by the presence of inclusions of high-temperature minerals, such as lechatelierite, which forms from pure quartz at temperatures >1700ºC (see above), or baddeleyite, which is the thermal decomposition product of zircon, forming at a temperature of about 1900ºC. Impact melts may also undergo a phase of superheating (i.e., staying liquid even though the vaporization temperature has been exceeded) at temperatures of 10,000ºC or higher (e.g., Jakes et al., 1992). Depending on the initial temperature, the location within the crater, the composition of the melt, and the speed of cooling, impact melts either form impact glasses (if they cool fast enough), or, more commonly, (mostly) fine-grained impact melt rocks (if they cool slower). Impact melt rocks are also found in suevitic breccias in the form of melt clasts. Impact melt rocks contain clasts of shocked minerals or lithic clasts (Fig. 14a-c).

Fig. 14. Microphotographs of impact melt rocks. a) Largely melted quartzitic clast in flow-banded, extremely fine-grained, melt matrix, in sample 277.8 from drill core M8, Manson impact structure, Iowa (see Koeberl et al., 1996a); note the fine feathery recrystallization texture; 2.2 mm wide, parallel polars; b) Impact melt breccia 9018.1 from the Ames crater, Oklahoma, 18-4 Nicor Chestnut core, depth 2748.7 m, with fractured and shocked mineral grains set in a finer-grained matrix, showing feathery spherulitic devitrification texture in center and upper right, and a diplectic quartz glass grain on the upper left, 3.4 mm wide, crossed polars (courtesy W.U. Reimold); c) Aphanitic impact melt breccia with K-feldspar clasts set in a fine-grained matrix, sample 1341.5 from the Exmore drill core, Chesapeake Bay impact structure, Virginia (see Poag et al., 1994; Koeberl et al., 1995b, 1996c); 3.4 mm wide, crossed polars.
As glasses are metastable supercooled liquids, impact glasses slowly recrystallize (if dissolution is not acting faster), at a rate that depends on the composition of the glass and post-impact environmental conditions. Therefore, impact glasses are more commonly found at young impact craters than at old impact structures. Very fine-grained recrystallization textures are often characteristic for devitrified impact glasses (Fig. 14a,b). Impact glasses have chemical and isotopic compositions that are very similar to those of individual target rocks or mixtures of several rock types. For example, it is possible to use the rare earth element (REE) distribution patterns, or the Rb-Sr isotopic composition, which are identical to those of the (often sedimentary or metasedimentary) target rocks, to distinguish the impact melt rocks from intrusive or volcanic rocks (e.g., Blum and Chamberlain, 1992; Blum et al., 1993). Furthermore, impact glasses have much lower water contents (about 0.001 - 0.05 wt%) than volcanic or other natural glasses (e.g., Koeberl, 1992b). Detailed descriptions of impact melts and glasses and their characteristics and compositions are discussed by, for example, El Goresy et al. (1968), Dence (1971), Stöffler (1984), Koeberl (1986, 1992a,b), and references therein.
Impact melts and glasses (or minerals that have recrystallized from the melt; e.g., Krogh et al., 1993; Izett et al., 1994) have another important use, as they often are the most suitable material for the dating of an impact structure. The methods most commonly used for dating of impact melt rocks or glasses include the K-Ar, 40Ar-39Ar, fission track, Rb-Sr, Sm-Nd, or U- Th-Pb isotope methods. However, dating impact craters is complicated and tedious and, if not done with utmost care, can easily lead to erroneous results (see, e.g., Bottomley et al., 1990, and Deutsch and Schärer, 1994, for reviews of impact crater dating).
Geochemistry and Detection of Meteoritic Components in Impactites
No meteorites have been found at most meteorite impact craters. This may seem a contradiction, but it follows as a logical consequence of the physics of an impact event. A shock wave, similar to the one that penetrates through the target, also passes through the meteoritic impactor and, within fractions of a second, vaporizes most or all of the projectile. Only during the impact of small objects (less than about 40 m in diameter, depending on impact angle and velocity), because of spallation during entry into the atmosphere or due to lower impact velocity resulting from atmospheric drag, a small fraction of the initial mass of the meteorite may survive. The cut-off diameter of impact craters at which some fraction of meteoritic material may be preserved is about 1 - 1.5 km. Thus, even under optimistic conditions, meteoritic fragments are only preserved at very young and small craters. The absence of meteorite fragments can, therefore, not be used as evidence against an impact origin of a crater structure.
A more generally applicable impact-diagnostic method is the detection of traces of the meteoritic projectile in target rocks. This allows to establish the impact origin for a crater structure. The meteoritic projectile undergoes vaporization in the early phases of crater formation. A small amount of the meteoritic vapor is incorporated with the much larger quantity of target rock vapor and melt, which later forms impact melt rocks, melt breccias, or glass. In most cases, the contribution of meteoritic matter to these impactite lithologies is very small (commonly <<1%), leading to only slight chemical changes in the resulting impactites. Only elements that have high abundances in meteorites, but low abundances in terrestrial crustal rocks, can be used to detect such a meteoritic component. During the last two decades, studies of the abundances and interelement ratios of the siderophile elements, such as Cr, Co, Ni, and, especially, the platinum group elements (PGEs) have been used for these investigations (see, e.g., Morgan et al., 1975; Palme, 1982; Evans et al., 1993; and references therein). However, the use of elemental abundances does not necessarily lead to unambiguous results, as ultramafic rocks or ore minerals may be present among the target rocks, resulting in elevated PGE abundances. Another complication is the possible fractionation of the siderophile elements in the impact melt while it is still molten. This effect may be significant in larger craters, because there the melt can stay hot for many thousand years. Different mineral phases, such as sulfides or oxides (e.g., magnetite, chromite), may take up various proportions of the PGEs or other siderophile elements, leading to an irregular distribution of these elements and possibly fractionated interelement ratios and patterns. Such irregular distribution of siderophiles is known from, for example, the East and West Clearwater impact structures (Palme et al., 1979), or the Chicxulub impact structure (Koeberl et al., 1994c; Schuraytz et al., 1996). Hydrothermal processes associated with the hot impact melt may also change PGE abundances.
The use of the Re-Os isotopic system has numerous advantages over the use of elemental abundances of the PGEs. The Re-Os isotope method is superior with respect to detection limit and selectivity, as discussed by Koeberl and Shirey (1993, 1996) and Koeberl et al. (1994a,b). In principle, the abundances of Re and Os and the 188Os/187Os isotopic ratios, which are measured by very sensitive mass spectrometric techniques, allow to distinguish the isotopic signatures of meteoritic and terrestrial Os. Meteorites (and the terrestrial mantle) have much higher (by factors of 104 - 105) PGE contents than terrestrial crustal rocks. In addition, meteorites have relatively low Re and high Os abundances, resulting in Re/Os ratios less or equal to 0.1, while the Re/Os ratio of terrestrial crustal rocks is usually no less than 10. More importantly even, the 188Os/187Os isotopic ratios for meteorites and terrestrial crustal rocks are significantly different.
187Os is formed from the ß-decay of 187Re (with a half-life of 42.3±1.3 Ga). Thus, due to the high Re and low Os concentrations in old crustal rocks, their 187Os/188Os ratio increases rapidly with time. The present day 187Os/188Os ratio of mantle rocks is about 0.13. Meteorites also have low 187Os/188Os ratios of about 0.11 to 0.18. Osmium is much more abundant in meteorites than Re, leading to only small changes in the meteoritic 187Os/188Os ratio with time. Because of the high Os abundances in meteorites, the addition of a minute meteoritical contribution to the crustal target rocks leads to an almost complete change of the Os isotopic signature of the resulting impact melt or breccia (see Fig. 15 for an example). For details about this method, see Koeberl and Shirey (1993; 1996; 1997) and Koeberl et al. (1994a,b). Similar to studies of shock metamorphism, Re-Os isotopic measurements of target rocks and impactites may provide good evidence for an impact origin.
Impact cratering still remains one of the least appreciated geological processes, even though, over the past three decades, researchers have studied impact cratering and craters in nature, in the laboratory, and by computer modelling. Identification of further impact structures on earth can only be achieved with diligent and careful investigations. Impact crater research is an excellent example to illustrate the necessity - and success - of interdisciplinary studies. This paper was aimed at describing how mineralogical and geochemical studies should be applied to the identification and characterization of impact craters and impact-derived rocks. As the discussions regarding the impact origin of the Ames structure, or in relation to the K-T boundary, have illustrated, there are still lots of misconceptions and a lack of understanding of the mineralogical and geochemical characteristics of shocked rocks. Thus, it is essential that the proper methods for identifying impact craters are understood and used before drawing any conclusions regarding the impact origin of a geological structure.
Fig. 15. Ratios of 187Os/188Os versus 187Re/188Os for target rocks (shale and sandstone) of the Kalkkop impact crater, South Africa, in comparison with data for four impact breccias (solid triangles) and the data array for chondritic and iron meteorites (small solid dots). The sample depths in the drill core are given as well (after Koeberl et al., 1994b). The dotted area marks the mixing field between target rocks and meteorites. All impact breccias fall within this mixing field; breccia-3 plots very close to the meteorite data array, indicating a larger meteoritic contamination than any other breccia, which is in good agreement with the high Os content of this sample (about 0.2 ppb). In addition to distinct differences in the Os isotopic composition, the Os content of the target rocks, at about 0.02 ppb, is ca. 10 times lower than the Os content of the breccias.
Many colleagues have helped with information, references, or have provided various other support; I am especially grateful to B. Bohor (U.S. Geological Survey, Denver), B. M. French (Smithsonian Institution, Washington), and R.A.F. Grieve (Geological Survey of Canada, Ottawa). I am also grateful to D. Jalufka for drafting the figures, and B. M. French and W. U. Reimold for detailed comments on the manuscript. I thank K. Johnson and J. Campbell for the invitation to contribute to the symposium, for which this paper is part of the proceedings. Parts of this study were supported by the Austrian Fonds zur Förderung der wissenschaftlichen Forschung, Project P08794-GEO (to C.K.).
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