Untitled

Bulletin of the Czech Geological Survey, Vol. 77, No. 4, 243–252, 2002 

© Czech Geological Survey, ISSN 1210-3527 

Asteroids: Their composition and impact threat 

THOMAS H. BURBINE 

Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, DC 20560-0119, USA; e-mail: 

burbine.tom@nmnh.si.edu 

Abstract. Impacts by near-Earth asteroids are serious threats to life as we know it. The energy of the impact will be a function of the mass of 

the asteroid and its impact velocity. The mass of an asteroid is very difficult to determine from Earth. One way to derive a near-Earth object’s mass is 

by estimating the object’s density from its surface composition. Reflectance spectra are the best way to determine an object’s composition since many 

minerals (e.g. olivine, pyroxene, hydrated silicates) have characteristic absorption features. However, metallic iron does not have characteristic absorption 

bands and is very hard to identify from Earth. For a particular size, asteroids with compositions similar to iron meteorites pose the biggest 

impact threat since they have the highest densities, but they are expected to be only a few percent of the impacting population. Knowing an asteroid’s 

composition is also vital for understanding how best to divert an incoming asteroid. 

Key words: Earth, asteroids, impact features, meteorites, mineral composition, geologic hazards 

Introduction 

As we enter a new millennium, we are constantly being 

bombarded with news of close encounters with near-Earth 

objects (NEOs). Observers have discovered over 2,000 near- 

Earth asteroids (NEAs); luckily none are known to be on a 

collision course with the Earth. Comets are also serious impact 

threats as shown by the collision of Shoemaker-Levy-9 

with Jupiter in 1994. Since the discovery of the iridium 

anomaly at the Cretaceous-Tertiary (K-T) boundary layer by 

Alvarez et al. (1980), there has been a considerable discussion 

of the possibility and consequences of such an impact 

(e.g. Chapman and Morrison 1994, Adushkin and Nemchinov 

1994, Toon et al. 1997, Garshnek et al. 2000). 

As we discover more Earth-approaching asteroids, we 

are also learning more about their compositions and structure. 

Charge-coupled devices (CCDs) now allow us to obtain 

visible and near-infrared telescopic spectra of 

near-Earth asteroids that are a few hundred meters or 

smaller in diameter (e.g. Binzel et al. 2001a). Technological 

improvements to the radio telescope at the Arecibo Observatory 

have allowed similarly sized objects to be 

characterized by radar (e.g. Ostro et al. 2002). Spacecraft 

missions now show asteroids to be geologic bodies with a 

variety of morphologic features. More meteorites are discovered 

every year and are being extensively studied in the 

laboratory with more precise analytical techniques. 

However as we continue our research on asteroids, a 

number of questions should be asked. “How do asteroid 

compositions affect the impact threat?” “How well can we 

determine asteroid compositions from Earth?” This paper 

will review what we currently know about asteroid compositions 

and how it affects the impact threat. Since we 

have not yet returned any samples from any asteroids, our 

knowledge of asteroid compositions is derived from analyses 

of meteorites, remote sensing observations from Earth, 

and spacecraft missions. 

Impact threat 

I will first briefly review what we know about the number 

of near-Earth objects and the energy and effects of impacts. 

The near-Earth object population is defined as small 

bodies with perihelion distance less than 1.3 AU (astronomical 

units) and aphelion distance greater than 0.983 

AU (Morbidelli et al. 2002). These objects are primarily 

thought to be asteroids ejected from the main belt; however, 

a few extinct comets probably exist in the population. 

Recent estimates of the numbers of NEAs larger than one 

kilometer in diameter vary from 855 (±110) (Morbidelli et 

al. 2002) to 1227 (uncertainties of +170 and –90) (Stuart 

2001). 

The kinetic energy (E) of an incoming asteroid is 

(1/2)mv 2 where m is the mass and v is the velocity. Mass is 

a function of the density (ρ) and volume (V) of the object. 

Since the energy of an impact is usually given as megatons 

of TNT, the kinetic energy equation can be written (Morbidelli 

et al. 2002) as 

E = 62.5 ρd 3 v 2 

where E is in megatons, ρ is in g/cm 3 , d is the diameter of 

the impactor in kilometers, and v is in km/s. The average 

impact velocity for asteroids with Earth are ~ 20 km/s (e.g. 

Hughes, 1998). Comets have lower densities (estimated to 

be around ~ 1 g/cm 3 ), but some long-period comets have 

much higher average impact velocities (~ 55 km/s) (Marsden 

and Steel 1994). Since water covers approximately 

three-fourths of the Earth’s surface, “large” impacts are 

likely to cause tsunamis (e.g. Paine 1999), giant tidal 

waves created by sudden disturbances. 

The best-characterized localized catastrophe is the impact 

at Tunguska where an object exploded 5–10 km in the 

air over an uninhabited region of Siberia (e.g. Chyba et al. 

1993, Vasilyev 1998). The energy was estimated to be 

~ 10–20 megatons and devastated an area of ~ 2000 km 2 

of forest area. In comparison, the energy is ~ 1,000 times 

243

Thomas H. Burbine 

larger than the bomb dropped on Hiroshima and the area is 

~ 4 times larger than New York City. 

The best-characterized “large” impact is the Chicxulub 

crater found in the Yucatan peninsula (e.g. Hildebrand et 

al. 1998). Hildebrand et al. (1998) argues that the crater is 

180 km in diameter, but other researchers (e.g. Morgan et 

al. 1998) argue that the actual crater size could be as large 

as 270 km. The energy of the impactor for producing the 

Chicxulub crater is estimated to be ~ 10 8 megatons (e.g. 

Toon et al. 1997). This impact occurred ~ 65 million years 

ago, which is the same age as the K-T extinction. The K- 

T boundary marks the demise of the dinosaurs as the dominant 

animal species on Earth (e.g. Milner 1998). 

We currently have evidence for over 160 impact craters 

on Earth over the last ~ 2 billion years that range from 15 

meters to 300 kilometers in diameter (Earth Impact 

Database 2002). Many more impacts have occurred over 

this time with evidence for these impacts wiped out due to 

erosion by water and wind. Cratering rates for the Earth 

(e.g. Neukum and Ivanov, 1994) can be estimated from determinations 

of the NEO population, the lunar cratering 

rate (where the effects of the Earth’s atmosphere needed to 

be added), and crater counts on terrestrial cratons. 

Neukum and Ivanov (1994) estimate that craters one kilometer 

in size or greater form on the Earth every ~ 1600 

years and craters 100 kilometers in size or greater form 

every 27 million years. 

Morbidelli et al. (2002) have done a recent study of 

collision probabilities of NEOs with the Earth. They estimate 

that a 1000 megaton impact, which would produce 

large-scale regional damage and a crater ~ 5 km in size 

(e.g. Hughes 1998), should occur every 63,000 ± 8,000 

years. They estimate that known NEOs carry only ~ 18% 

of the overall collision probability. 

The consequences of any impact on Earth will be a 

function of the mass (density times volume) and impact 

velocity plus the location, date, and time of the impact on 

Earth. Astrometric observations, which determine the 

location of an object in the sky, can help determine an asteroid’s 

orbit with a high degree of certainty. These observations 

can be used to predict the probability that an object 

will strike the Earth. If an object is on a collision course 

these astrometric observations can be used to predict the 

impact velocity and the location of the impact. Radar observations 

(e.g. Ostro et al. 2002) can image an asteroid (if 

it is large enough and close enough to Earth) and then determine 

its diameter, which can be used to calculate its 

volume. 

However, the determination of an asteroid’s mass is 

much more difficult from Earth. For the largest asteroids, 

masses can be determined by asteroid-asteroid interactions 

or their perturbations on Mars (Britt et al. 2002). To determine 

a near-Earth asteroid’s mass, the object needs to have 

a natural body revolving around it (e.g. Margot et al. 2002) 

or a spacecraft encounter (e.g. Veverka et al. 1999, Yeomans 

et al. 2000). 

However in the absence of such a mass determination, 

the best way to estimate a near-Earth asteroid’s mass 

would be to determine its surface composition, which 

would give insight into the object’s density. (Since near- 

Earth asteroids are fragments of much larger asteroids, we 

assume that the surface composition is representative of 

the object as a whole.) The mass could then be estimated 

by multiplying the inferred density by the object’s volume. 

Such an estimate would only give an upper limit on an object’s 

mass, since many asteroids are thought to be rubble 

piles with significant amounts of macroporosity (e.g. Britt 

and Consolmagno 2001, Britt et al. 2002). The estimated 

macroporosities for ~ 20 asteroids range from 0 to ~ 80% 

(Britt et al. 2002). 

Meteorites 

Except for ~ 50 samples from the Moon and Mars, meteorites 

appear to be fragments of sub-planetary sized bodies 

(asteroids) that formed ~ 4.56 billion years ago. 

Meteorites can basically be broken into two types: those that 

experienced heating but not melting (chondrites) (e.g. 

Brearley and Jones 1998) and those that experienced melting 

and differentiation (achondrites, stony-irons, irons) (e.g. 

Mittlefehldt et al. 1998). Silicate-rich meteorites are often 

referred to as stony and would include all chondrites and 

achondrites. Meteorites that are similar in terms of petrologic, 

mineralogical, bulk chemical, and isotopic properties 

are separated into groups (Table 1). In general, groups contain 

five or more members to allow for the compositional 

characteristics of the group to be adequately characterized. 

Currently, 13 groups (Table 1) of chondritic meteorites 

have been defined. Chondritic groups are subdivided according 

to petrologic type (1–6) with 1 being the most 

aqueously altered, 3.0 being the least altered, and 6 being 

heated to the highest temperatures. Mineralogically, meteorites 

of petrologic type 1 and 2 (CI, CM, CR) have compositions 

dominated by phyllosilicates. Visually, these 

meteorites are extremely dark. The other types of carbonaceous 

chondrites (CH, CV, CO, CK) tend to have 

mineralogies dominated by mafic silicates. The R chondrites 

are dominated by olivine. Ordinary chondrites (H, 

L, LL) are mixtures of mafic silicates and metallic iron, 

while enstatite chondrites are composed of enstatite (virtually 

FeO-free pyroxene) and metallic iron. In addition to 

these 13 well-defined groups, ~ 14 chondritic grouplets or 

unique meteorites (Meibom and Clark 1999, Weisberg et 

al. 2001, Mittlefehldt, 2002) have been recognized. 

The differentiated meteorites range from those that experienced 

only limited differentiation (primitive achondrites) 

to those (differentiated achondrites, stony-irons, 

irons) that were produced by extensive melting, melt migration, 

and fractional crystallization. These processes 

produce a wide variety of lithologies. 

Many of the differentiated meteorites are thought to 

sample the crusts (howardites, eucrites, and diogenites or 

HEDs; angrites), core-mantle boundaries (pallasites), and 

cores (irons) of differentiated bodies. Eucrites contain primarily 

plagioclase and both low-Ca and high-Ca pyrox- 

244


enes, while diogenites are predominantly magnesian orthopyroxene. 

Howardites are breccias of eucritic and diogenitic 

material. Pallasites are mixtures of metallic iron 

and olivine. We do not appear to be sampling the mantles 

of these differentiated bodies (e.g. Burbine et al. 1996); 

however, we do appear to be sampling the mantles of veryreduced 

differentiated bodies in the form of aubrites. 

Iron meteorites are composed of metallic iron with 

5–20% Ni plus accessory phases such as sulfides, 

schreibersite, and silicate inclusions. Iron meteorites are 

classified according to siderophile (“iron-loving”) element 

(Ga, Ge, Ir, Ni) concentrations. Of the thirteen groups of 

iron meteorites, ten (IC, IIAB, IIC, IID, IIF, IIIAB, IIIE, 

IIIF, IVA, IVB) have fractional crystallization trends suggestive 

of prolonged cooling as expected from the cores of 

differentiated bodies. Three of the groups (IAB, IIE, and 

IIICD) do not display well-developed fractional crystallization 

trends and contain abundant silicate inclusions, 

which argues that they are not core fragments. Approximately 

90 irons (Grady 2000) are not classified as members 

of the 13 groups and are labeled anomalous. These 

ungrouped irons are believed to require 50-70 distinct parent 

bodies (Wasson 1995, Burbine et al. 2002b). 

Other types of differentiated meteorites have undergone 

varying amounts of melting. These include a number 

of primitive achondritic meteorites (acapulcoite-lodranites, 

winonaites), which are samples of partially differentiated 

asteroids. These meteorites are mixtures of olivine, 

pyroxene, and metallic iron. Mesosiderites are breccias 

composed of HED-like clasts of basaltic material mixed 

with metallic clasts. One possible scenario for the formation 

of the mesosiderites is the disruption of an asteroid 

with a molten core (Scott et al. 2001). Rounding out the 

differentiated meteorites are the olivine-dominated brachinites 

and the carbon-rich ureilites, whose origins are 

still being debated (e.g. Mittlefehldt et al. 1998). 

Meteorite densities and strengths 

Consolmagno and Britt (1998), Flynn et al. (1999), and 

Wilkinson and Robinson (2000) have recently done studies 

of meteorite bulk densities. The only CI chondrite 

measured had a bulk density of 1.58 g/cm 3 while the bulk 

densities of CM chondrites were 2.08–2.37 g/cm 3 . CO 

chondrites (2.36–2.98 g/cm 3 ) and CV chondrites 

(2.60–3.18 g/cm 3 ) tended to have slightly higher bulk densities. 

The only enstatite chondrite measured had a bulk 

density of 3.36 g/cm 3 . On average, ordinary chondrites 

(3.05–3.75 g/cm 3 for the most reliable measurements) 

have the highest bulk densities of the chondrites. 

HEDs have bulk densities of 2.99–3.29 g/cm 3 . Stonyirons 

such as mesosiderites (4.16–4.22 g/cm 3 ) and pallasites 

(4.82–4.97 g/cm 3 ) have higher bulk densities due to 

the presence of significant amounts of metallic iron. As expected, 

iron meteorites tend to have the highest bulk densities 

of all meteorite types (6.99–7.59 g/cm 3 for relatively 

unweathered specimens). 

Table 1. Meteorite groups 

Groups Composition* Fall percentages # (%) 

Carbonaceous chondrites 

CI phy, mag 0.5 

CM phy, toch, ol, 1.7 

CR phy, px, ol, met 0.3 

CO ol, px, CAIs, met 0.5 

CH px, met, ol, Only finds 

CV ol, px, CAIs 0.6 

CK ol, CAIs 0.3 

Enstatite chondrites 

EH enst, met, sul, plag, ± ol 0.8 

EL enst, met, sul, plag 0.7 

Ordinary chondrites 

H ol, px, met, plag, sul 34.1 

L ol, px, plag, met, sul 38.0 

LL ol, px, plag, met, sul 7.9 

R chondrites ol, px, plag, sul 0.1 

Primitive achondrites 

Acapulcoites a px, ol, plag, met, sul 0.1 

Lodranites a px, ol, met, ± plag, ± sul 0.1 

Winonaites ol, px, plag, met 0.1 

Differentiated achondrites 

Angrites TiO 2-rich aug, ol, plag 0.1 

Aubrites enst, sul 0.1 

Brachinites ol, cpx, ± plag Only finds 

Diogenites b opx 1.2 

Eucrites b pig, plag 2.7 

Howardites b eucritic-diogenitic breccia 2.1 

Ureilites ol, px, graph 0.5 

Stony-irons 

Mesosiderites basalt-met breccia 0.7 

Pallasites ol, met ol, met 0.5 

Irons c met, sul, schreib 4.2 

Table is revised from tables found in Burbine et al. (2002b). 

* Minerals or components are listed in decreasing order of average abundance. 

Abbreviations: ol – olivine, px – pyroxene, opx – orthopyroxene, 

pig – pigeonite, enst –enstatite, aug – augite, cpx – clinopyroxene, plag 

– plagioclase, mag – magnetite, met – metallic iron, sul – sulfides, phy – 

phyllosilicates, toch – tochilinite, graph – graphite, CAIs – Ca-Al-rich 

refractory inclusions, schreib – schreibersite, ± – may be present 

# Fall percentages are calculated from the 942 classified falls that are listed 

in Grady (2000), Grossman (2000), and Grossman and Zipfel (2001). 

a Acapulcoites and lodranites appear to come from the same parent body 

(e.g. Mittlefehldt et al. 1998). 

b Howardites, eucrites, and diogenites (HEDs) appear to come from the 

same parent body (e.g. Mittlefehldt et al. 1998). 

c There are 13 iron meteorite groups plus ~100 ungrouped irons. 

In terms of physical strength, most chondritic meteorites 

can be easily crushed into powders with a mortar 

and pestle. Meteorites containing phyllosilicates tend to be 

the most fragile with the CI chondrites being the weakest 

of these objects with laboratory crushing strengths of 1–10 

bars (e.g. Lewis 2000). What cannot be easily pulverized 

is metallic iron. Metallic iron is ductile and tends to flatten 

and elongate while being crushed at room temperatures. It 

is unclear how ductile metallic iron is at the colder temperatures 

found in the asteroid belt. 

Iron meteorites are extremely strong with strengths of 

approximately 3.5 kbars (e.g. Lewis, 2000). The much 

stronger physical strength of irons allows them to survive 

in space much longer than stony bodies as seen by their 

longer cosmic ray exposure ages. Cosmic ray exposure 

ages record the time an object has spent as a meter-sized 

245


Normalized Reflectance 

6 

5 

4 

3 

2 

1 

0 

0.2 0.6 1 1.4 1.8 2.2 2.6 

Wavelength (µm) 

or less body in space or within a few meters of the surface. 

Irons have cosmic ray exposure ages ranging from hundreds 

of millions to a few billion years (e.g. Voshage and 

Feldman 1979) while stony meteorites tend to have much 

shorter exposure ages (


Asteroid observations 

This section will detail what we currently think we understand 

about asteroid compositions. Asteroid spectral 

surveys (e.g. Zellner et al. 1985, Bus and Binzel 2002a) 

have primarily been done in the visible due to the peaking 

of the illuminating solar flux and the relative transparency 

of the atmosphere at these wavelengths. Near-infrared observations 

(1-3.5 µm) have now become easier to obtain 

due to the advent of SpeX (an infrared spectrograph at the 

Infrared Telescope Facility on Mauna Kea) (e.g. Binzel et 

al. 2001b). 

Asteroids are generally grouped into classes (Table 2) 

based on their visible spectra (~ 0.4 to ~ 0.9–1.1 µm) and 

visual albedo (when available). The most widely used taxonomy 

(Tholen 1984) classifies objects observed in the 

eight-color asteroid survey (ECAS) (Zellner et al. 1985). 

Bus and Binzel (2002a) did a CCD spectral survey of over 

1300 objects and developed an expanded taxonomy (Bus 

and Binzel 2002b) with many more classes and subclasses 

to represent the diversity of spectral properties seen in 

higher-resolution spectra. Representative spectra of a 

number of Bus and Binzel (2002a, 2002b) asteroid classes 

are shown in Figure 2. 

As stated earlier, to accurately determine an asteroid’s 

mineralogy, observations are needed in the near-infrared 

since many minerals have diagnostic features in this wavelength 

region. Near-infrared asteroid spectra (Gaffey et al. 

1993, Rivkin et al. 2000, Burbine and Binzel 2002) at a variety 

of wavelengths have shown that each class tends to 

contain a wide variety of mineralogies. Only qualitative 

mineralogical descriptions will be given for asteroid classes 

discussed in this paper; quantitative analyses of asteroid 

compositions (determining the proportion and composition 

of different mineral species) are very difficult (e.g. 

Clark, 1995) since an asteroid’s reflectance spectrum is a 

function of a number of factors such as mineralogy, mineral 

chemistry, particle size, and temperature. 

Asteroid taxonomy is based on astronomically observed 

parameters without regard to composition. However, many 

asteroid classes have been given letter designations that 

“imply” specific compositions. This includes the letter “M” 

for metallic, “S” for siliceous (or stony or stony-iron), and 

“C” for carbonaceous. (Bus and Binzel (2002b) do not differentiate 

between E, M, and P asteroids and call them all X 

objects since there spectra are similar and albedo is not used 

in their taxonomy.) Since each asteroid class only groups asteroids 

with specific spectral characteristics, not all objects 

in a class have to have similar surface mineralogies. For example, 

it is unknown how many M-class asteroids actually 

have surfaces similar to iron meteorites. For example, enstatite 

chondrites (primarily mixtures of metallic iron and 

enstatite) also have relatively featureless spectra (Gaffey 

1976) and similar albedos to metallic iron. Also, a subgroup 

of M-class asteroids (called the W class) have distinctive 3 

µm features (e.g. Rivkin et al. 2000), which apparently indicate 

hydrated minerals and surface compositions inconsistent 

with metallic iron. 

Radar observations have been used to estimate the metal 

contents of asteroids (Ostro et al., 2002). Radar experiments 

of asteroids usually measure the distribution of echo 

power reflected from an object in time delay and Doppler 

Table 2. Asteroid Classes 

Class characteristics 

A Distinctive olivine absorption features 

B Weak UV feature, blue past 0.4 µm, subclass of C types; low albedo (generally less than 0.1) 

C a Weak UV feature, flat to reddish past 0.4 µm; low albedo (generally less than 0.10) 

D Very red spectrum; low albedo (usually around 0.05 or less) 

E Flat to slightly red, featureless spectrum; high albedo (> 0.30); usually associated with aubrites 

F Very weak UV feature, flat to bluish past 0.4 µm, subclass of C types; low albedo (< 0.10) 

G Strong UV feature, flat past 0.4 µm, subclass of C class; usually have strong 3 µm features; low albedo (< 0.10) 

J Stronger 1 µm feature than V types; appear to have compositions similar to the HEDs 

K Spectrum intermediate between S and C asteroids; usually associated with CO3/CV3 chondrites 

L b Very strong UV feature and then becoming approximately flat 

M Flat to slightly red, featureless spectrum; moderate albedo (0.10-0.30) 

O Weak UV feature out to 0.44 µm, very strong 1 µm feature; type spectrum is 3628 Božněmcová 

P Flat to slightly red, featureless spectrum; low albedo (usually around 0.05 or less) 

Q Strong UV and 1 µm feature; spectrum similar to ordinary chondrites 

R Strong UV and 1 µm feature; type spectrum is 349 Dembowska 

S c Strong UV feature; usually has 1 µm feature, indicating olivine and/or pyroxene; moderate albedo (0.10–0.30) 

T Weak UV feature, reddish past 0.4 µm; low albedo (< 0.10) 

V Distinctive 1 and 2 µm features due to pyroxene; appear to have compositions similar to the HEDs 

W M-class visible spectrum with a 3 µm absorption feature 

Spectrum similar to E, M, or P types, but no albedo information 

X d 

This table is revised from tables in Wetherill and Chapman (1988), Pieters and McFadden (1994), and Bus and Binzel (2002b). The term “red” refers 

to reflectances increasing with increasing wavelength and the term “blue” refers to reflectances decreasing with increasing wavelength. 

a Bus and Binzel (2001b) subdivided the C-class objects into the C, Cb, Cg, Ch, and Cgh subclasses based on CCD spectra. Bus and Binzel (2001b) 

do not define the F and G classes. 

b Bus and Binzel (2001b) subdivided the L-class objects into the L and Ld classes based on CCD spectra. 

c Gaffey et al. (1993) subdivided the S-class objects into the S(I) to S(VII) subclasses on the basis of both visible and near-infrared spectra. Bus and 

Binzel (2002b) have subdivided the S class into the S, Sa, Sk, Sl, Sq, and Sr subclasses based on CCD spectra. 

d Bus and Binzel (2002b) subdivided the X-class class objects into the X, Xc, Xe, and Xk subclasses based on CCD spectra. 

247



2.8 

2.3 

1.8 

1.3 

1929 Kollaa (V) 

1542 Schalen (D) 

6 Hebe (S) 

221 Eos (K) 

64 Angelina (Xe) 

19 Fortuna (Ch) 

16 Psyche (X) 

0.8 

0.4 0.5 0.6 0.7 0.8 0.9 1 


Fig. 2. Reflectance spectra of a number of Bus and Binzel (2002a, 

2002b) asteroid classes. All spectra are normalized to unity at 0.55 µm 

and offset in reflectance from each other. Error bars are one sigma. All 

spectra are available at the website http://smass.mit.edu. 

frequency in the opposite sense (OC) of circular polarization 

and the same sense. OC radar albedos are equal to the 

OC radar cross section divided by the target’s projected 

area. For homogeneous and particulate surfaces, radar 

albedos are functions of the near-surface bulk density, 

which is related to both the solid-rock density and the surface 

porosity of the object. Increasing the solid-rock density 

or decreasing the surface porosity would increase the 

radar albedo of an object. Without knowing the porosity of 

the surface, it is impossible to conclusively determine the 

surface assemblage of an object. M-asteroids with the 

highest radar albedos (0.6–0.7) could have surfaces of 

metallic iron with “lunar-like” porosities (35-55%) or solid 

enstatite-chondritic material with little to no porosity. It 

is unclear if such solid surfaces of enstatite chondrite material 

could exist on the surface of an asteroid. M asteroids 

tend to have higher radar albedos than C or S asteroids 

(Magri et al. 1999), apparently implying that they are richer 

in metallic iron. 

S asteroids tend to have 1 µm absorption features, 

which indicates assemblages containing Fe 2+ -bearing silicates 

such as olivine and/or pyroxene. On the basis of 

high-resolution CCD spectra, Bus and Binzel (2002b) subdivide 

the S asteroids into a number of subtypes (S, Sa, Sk, 

Sl, Sq, and Sr). The subscript represents that the objects is 

intermediate in spectral properties between that class (A, 

K, L, Q, and R) and the S class; however, it is currently 

unclear if each of these subclasses is grouping objects with 

similar compositions. From near-infrared spectra, Gaffey 

et al. (1993) finds that S asteroids tended to have compositions 

that ranged from olivine-rich (which he defined as 

S(I)) to pyroxene-rich (S(VII)). To classify these objects, 

they use the band area ratio (ratio of the area of Band I to 

the area of Band II) and the Band I minimum, which are 

both function of the olivine/pyroxene abundance (Cloutis 

et al., 1986). The band parameters of S(IV)-objects appear 

consistent with ordinary chondrites, but also consistent 

with other types of meteorites such as ureilites or acapulcoites/lodranites. 

S asteroids are the most abundant type of classified asteroid 

since they tend to be found in the inner main belt 

(closer to the Sun) and have higher albedos than C-types, 

making them brighter and easier to discover and observe. 

The biggest question concerning S asteroids is what fraction 

is compositionally similar to ordinary chondrites. S 

asteroids are spectrally redder than ordinary chondrites 

and tend to have weaker absorption bands (Figure 3). It 

has long been argued (e.g. Wetherill and Chapman 1988) 

whether these spectral differences are due to inherent compositional 

differences or simply due to an alteration 

process that can redden ordinary chondrite material. Analyses 

of lunar regolith (e.g. Pieters et al. 2000) and alteration 

experiments (e.g. Sasaki et al. 2001) appear to show 

that this reddening on asteroidal surfaces could be due to 

surface alteration processes (e.g. micrometeorite impacts, 

solar wind sputtering) that produce vapor-deposited coatings 

of nanophase iron. 

C-type asteroids (including the B, C, F, G, and P classes) 

tend to have relatively featureless spectra in low-resolution 

(photometric) surveys, which was consistent with 

carbonaceous meteorites. Higher-resolution spectral surveys 

(e.g. Bus and Binzel 2002b) have shown that almost 

half of observed C-type asteroids have a 0.7 µm feature. 

Observations (Jones et al. 1990) indicate that approximately 

two-thirds of all observed C-type asteroids have 

3 µm features. 

A-class asteroids tend to have strong UV and 1 µm features 

that appear similar to those of olivine. Near-infrared 

spectra (Figure 3) of these objects (Bell et al. 1988) confirm 

that these surfaces contain significant amounts of 

olivine as seen by the three distinctive olivine bands that 

make up the 1 µm feature. Two types of meteorites (brachinites 

and pallasites) have silicate mineralogies dominated 

by olivine and have been postulated to have 

compositions similar to the A asteroids. 

V- and J-type asteroids are the asteroids whose mineralogies 

appear to be the best-determined by remote sensing. 

These objects (including the 500-km diameter 4 Vesta 

and much smaller objects with diameters of 10 km or less) 

have visible and near-infrared spectra (Figure 3) (e.g. 

McCord et al. 1970, Binzel and Xu 1993, Burbine et al. 

2001b) similar to the HEDs, which have very distinctive 

spectral features due to pyroxene. The smaller V- and J-objects 

(called Vestoids) have been found in the Vesta family 

and between Vesta and the 3:1 and the ν 6 resonances. The 

presence of only one “large” body (4 Vesta) in the main 

belt with a spectrum similar to HEDs argues that Vesta is 

248


the parent body of the HEDs (e.g. Consolmagno and 

Drake, 1977). 

E asteroids, due to their relatively featureless spectra and 

high albedos (greater than 0.3), have been interpreted as 

having surfaces composed predominately of an essentially 

iron-free silicate (e.g. Zellner et al. 1977). The most obvious 

meteoritic analog is the aubrites (enstatite achondrites), 

which are igneous meteorites composed primarily of essentially 

iron-free enstatite (Watters and Prinz 1979). However, 

the identification of an absorption feature in the 3 µm wavelength 

region of a number of main-belt E-asteroid spectra 

(Rivkin et al., 1995) has been interpreted as indicating hydrated 

minerals on the surfaces of some of these objects, 

which is inconsistent with an igneous origin. A feature centered 

at ~ 0.5 µm has also been identified in a number of E- 

class asteroids (Bus 1999, Fornasier and Lazzarin 2001) and 

these objects have been classified as Xe by Bus and Binzel 

(2002b) (Figure 2). This feature is believed to be due to a 

sulfide (Burbine et al. 2002a), which is commonly found in 

aubrites (Watters and Prinz 1979). 

The surface mineralogies of many other asteroid classes 

are thought to be somewhat understood. The K-class asteroids 

tend to have visible and near-infrared spectral 

properties similar to CO3/CV3 chondrites (Bell 1988, 

Burbine et al. 2001a, Burbine and Binzel 2002). Q asteroids 

(most notably 1862 Apollo) tend to have spectral 

properties similar to ordinary chondrites. D and P objects, 

which tend to be found in the outer main belt, are thought 

to have very primitive, organic-rich surfaces due to their 

relatively featureless and red spectra (e.g. Vilas and Gaffey, 

1989). R asteroids (most notably 349 Dembowska) 

appear to be mixtures of olivine and pyroxene (Gaffey et 

al., 1989). T asteroids tend to be very dark and featureless 

and Hiroi and Hasegawa (2002) have noted the spectral 

similarity of unusual carbonaceous chondrite Tagish Lake 

to T asteroids. 

The surface compositions of a few asteroid classes are 

unclear, but do appear to contain silicates. O-asteroid 3628 

Božněmcová has a 1 µm feature unlike any known meteorite 

(Burbine and Binzel 2002). L class objects have been 

newly defined by Bus and Binzel (2002b) and appear intermediate 

in spectral properties in the visible between the 

K and some S asteroids. 

Spacecraft missions 

The first dedicated asteroidal mission was the NEAR- 

Shoemaker spacecraft rendezvous with S-asteroid 433 

Eros (e.g. McCoy et al. 2002) and its flyby of C-asteroid 

253 Mathilde (e.g. Veverka et al. 1999). NEAR-Shoemaker 

Spacecraft obtained high-resolution images, reflectance 

spectra of different lithologic units, bulk density, magnetic 

field measurements, and bulk elemental compositions of 

433 Eros. Eros had a density of ~ 2.7 g/cm 3 , consistent 

with a silicate-rich assemblage while Mathilde had an extremely 

low density of 1.3 g/cm 3 . 

Eros was classified as an S(IV) object (Murchie and 


3.5 

3 

2.5 

2 

1.5 

1 

Bouvante (eucrite) 

1929 Kollaa (V) 

Ehole (H5) 

6 Hebe (S) 

289 Neneta (A) 

0.5 

0.4 0.8 1.2 1.6 2 2.4 


Fig. 3. Reflectance spectra of S-asteroid 6 Hebe versus H5 chondrite 

Ehole, A-asteroid 289 Nenetta, and 1929 Kollaa versus eucrite Bouvante. 

The 6 Hebe and 289 Nenetaa spectra are a combination of data from 

Binzel and Bus (2002a) and Bell et al. (1988) while the 1929 Kollaa 

spectrum is a combination of data from Binzel and Bus (2002a) and Burbine 

and Binzel (2002). All spectra are normalized to unity at 0.55 µm 

and offset in reflectance from each other. Error bars are one sigma. All 

the Binzel and Bus (2002a) and the Burbine and Binzel (2002) spectra 

are available at the website http://smass.mit.edu. The Bouvante spectrum 

is from Burbine et al. (2001b) and was taken at Brown University’s 

RELAB facility. 

Pieters 1996) and it was hoped that the NEAR-Shoemaker 

data would help “solve” the S-asteroid/ordinary chondrite 

question. The average olivine to pyroxene 

composition derived from band area ratios (McFadden et 

al. 2001) and elemental ratios (Mg/Si, Fe/Si, Al/Si, and 

Ca/Si) derived from X-ray data (Nittler et al. 2001) of Eros 

are consistent (McCoy et al. 2001) with ordinary chondrite 

compositions. However, the S/Si ratio derived from X-ray 

data (Nittler et al. 2001) and the Fe/O and Fe/Si ratios derived 

from gamma-ray data (Evans et al. 2001) are significantly 

depleted relative to ordinary chondrites. McCoy et 

al. (2001) believe that the best meteoritic analogs to Eros 

are an ordinary chondrite, whose surface mineralogy has 

been altered by the depletion of metallic iron and sulfides, 

or a primitive achondrite, derived from a precursor assemblage 

of the same mineralogy as one of the ordinary chondrite 

groups. 

What is the percentage of near-Earth asteroids 

with iron meteorite compositions? 

What is extremely difficult from Earth to determine is 

if an object has a composition similar to iron meteorites. 

This is a problem since for a particular size, the biggest 

249


devastation among asteroid impacts will occur for these 

objects. Metallic iron has no distinctive absorption features 

and radar observations are often inconclusive since 

we do not know the surface porosity. There is some evidence 

that these objects do exist in the near-Earth asteroid 

population. M-type near-Earth asteroid (6178 1986 DA) 

(Tedesco and Gradie 1987), has one of the highest radar 

albedos (~ 0.6) of any observed asteroid, strongly implying 

a metallic iron surface (Ostro et al. 1991). 

Fall percentages give a probability of 4% of an iron 

meteorite being seen to land on the Earth and a 94% probability 

of seeing a stony or stony-iron meteorite fall. It is 

unclear if these fall statistics can be extrapolated to hundreds 

of meters to tens of kilometer-sized bodies in the 

near-earth population; however, compositions interpreted 

from the classifications derived from NEA spectral surveys 

roughly correspond to these percentages. Binzel et al. 

(2001a) have published the spectra and classifications of 

48 near-Earth asteroids. Approximately 90% of the objects 

(S, B, C, L, Q, V, O, K) had compositions consistent with 

silicate-dominated mineralogies if our compositional interpretations 

of these asteroid classes are correct. Approximately 

10% of the objects were classified as X types. As 

stated before, only some percentage of the X types are 

probably similar in composition to iron meteorites. 

This estimated percentage of iron projectiles is roughly 

consistent with numbers derived from impact craters. 

Koerbel (1998) lists 41 impact craters with diameters 

greater than 100 meters where the compositional characteristics 

of the impactor have been tentatively identified. 

He listed 14 (34%) as being due to iron meteorite projectiles 

with 3 impacts (7%) being due to either chondritic or 

iron-meteorite assemblages. The rest of the impacts appeared 

to be due either to chondritic, achondritic, stone, or 

stony-iron objects. However, the impacts due to iron meteorite 

assemblages are predominately associated with the 

smallest craters. Ten of the eleven craters smaller than 1.2 

kilometers in diameter are all thought to be due to iron meteorite 

projectiles. This preponderance of iron meteorite 

impacts among small craters is due to the atmospheric selection 

effect (e.g. Melosh 1989) that allows smaller iron 

meteorite projectiles (since they are denser compared to 

chondritic ones) to reach the ground and produce hypervelocity 

impacts. For the 30 craters larger than 2 km in diameter, 

~ 25% could be due to iron (or stony-iron) 

projectiles. However, only one of the eleven craters greater 

than 24 kilometers in diameter listed by Koeberl (1998) is 

believed to be due to an asteroid with a composition similar 

to iron meteorites. 

The meteorite fall statistics, classification of NEAs, 

and the characterization of the impactors that produced 

large craters all tend to argue that iron meteorite projectiles 

are a small percentage of the impacting population. 

All these lines of evidence argue that iron meteorite projectiles 

are a few percent and certainly less than 10% of 

near-Earth asteroids. Silicate-dominated assemblages are 

much more likely to strike the Earth. 

Is it important to know the composition 

of an impacting asteroid? 

To estimate the effects of an impact, the mass needs to 

be known. Since most meteorites that land on the Earth 

have bulk densities between 2 and 4 g/cm 3 , we can estimate 

pretty well an upper limit on an incoming object’s 

mass by using a density of 4 g/cm 3 and multiplying it by 

the object’s volume computed from its diameter. Also 

making this mass estimate an upper limit is the fact that asteroids 

may have significant macroporosity. 

But in a real-world situation where a sizable asteroid 

(e.g. a kilometer in diameter) that will be extremely destructive 

(~ 5,000 MT) is known to be a collision course 

with Earth, we will care more about diverting the object 

than knowing exactly how destructive it will be. Compositional 

information will be vital for determining how best to 

divert an object (e.g. Ahrens and Harris 1994). Deflection 

techniques such as impacting an asteroid with a spacecraft 

or a nuclear explosion will work best with knowledge of 

the surface composition. We need to know how the surface 

will be affected by an impact or blast and this information 

can only be derived from compositional studies. For example, 

an asteroid with a composition similar to a carbonaceous 

chondrite would be expected to fracture much 

more easily than an object with a composition similar to 

iron meteorites. 

Conclusions 

It is certain that the Earth will be hit in the future by an 

asteroid. The only question is “When?” For this impacting 

object, compositional studies will be vital for trying to determine 

how destructive the impact will be and for diverting 

the object. 

Acknowledgments. The author would like to thank Clark Chapman 

and Guy Consolmagno for very thoughtful reviews and the editorial work 

of Roman Skála. Almost all of the meteorite spectra in this paper were 

measured at Brown University’s Keck/NASA Reflectance Experiment 

Laboratory (RELAB), which is a multi-user facility supported by NASA 

grant NAG5-3871. 

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Handling editor: Roman Skála 

252



The recognition of terrestrial impact structures 

ANN M. THERRIAULT – RICHARD A. F. GRIEVE – MARK PILKINGTON 

Natural Resources Canada, Booth Street, Ottawa, Ontario, KIA 0ES Canada; e-mail: ATherria@NRCan.gc.ca 

Abstract. The Earth is the most endogenically active of the terrestrial planets and, thus, has retained the poorest sample of impacts that have 

occurred throughout geological time. The current known sample consists of approximately 160 impact structures or crater fields. Approximately 30% 

of known impact structures are buried and were initially detected as geophysical anomalies and subsequently drilled to provide geologic samples. The 

recognition of terrestrial impact structures may, or may not, come from the discovery of an anomalous quasi-circular topographic, geologic or geophysical 

feature. In the geologically active terrestrial environment, anomalous quasi-circular features, however, do not automatically equate with an 

impact origin. Specific samples must be acquired and the occurrence of shock metamorphism, or, in the case of small craters, meteoritic fragments, 

must be demonstrated before an impact origin can be confirmed. Shock metamorphism is defined by a progressive destruction of the original rock and 

mineral structure with increasing shock pressure. Peak shock pressures and temperatures produced by an impact event may reach several hundreds of 

gigaPascals and several thousand degrees Kelvin, which are far outside the range of endogenic metamorphism. In addition, the application of shockwave 

pressures is both sudden and brief. Shock metamorphic effects result from high strain rates, well above the rates of normal tectonic processes. 

The well-characterized and documented shock effects in quartz are unequivocal indicators and are the most frequently used indicator for terrestrial impact 

structures and lithologies. 

Key words: Earth, impact structures, shock metamorphism, melting, glasses, shatter cones, geophysical anomaly 

Introduction 

On Earth, compared to the other terrestrial planets, the 

very active geologic environment tends to modify and destroy 

the impact crater record. Approximately 160 impact 

structures or crater fields are currently known on Earth. 

Impact involves the transfer of massive amounts of energy 

to a spatially limited area of the Earth’s surface, in an extremely 

short time interval. As a consequence, local geology 

of the target area is of secondary importance. The 

effects of impact are, however, scale-dependent and show 

progressive changes with increasing energy of the impact 

event. The net result is that, impacts of similar scale produce 

similar first-order geological and geophysical effects. 

Thus, general observations can be derived with respect to 

the appearance and geological and geophysical signatures 

of terrestrial impact structures, in specific size ranges. 

Terrestrial impact structures were first recognized by 

their bowl-like shape and meteorite fragments found in their 

vicinity or within them (the classic example being Meteor 

or Barringer Crater, Arizona). In the 1960’s, petrographic 

studies of rocks from impact structures defined a series of 

unique characteristics produced by a style of deformation 

called shock metamorphism (e.g. French and Short 1968). 

Shock metamorphic effects include shatter cones (e.g. Dietz 

1947, Milton 1977), the only macroscopic diagnostic shock 

effect observed at terrestrial impact structures, a number of 

microscopic effects in minerals, some of which are diagnostic 

of shock, and impact melting. 

The aim of this paper is to summarize the morphology 

and geoscientific aspects of terrestrial impact structures 

and provide a general description of shock-metamorphic 

effects. 

Morphology 

On most planetary bodies, well-preserved impact 

structures are recognized by their characteristic morphology 

and morphometry. The basic shape of an impact structure 

is a depression with an upraised rim. Detailed 

appearance, however, varies with crater diameter. With increasing 

diameter, impact structures become proportionately 

shallower and develop more complicated rims and 

floors, including the appearance of central peaks and interior 

rings. Impact craters are divided into three basic morphologic 

subdivisions: simple craters, complex craters, 

and basins (Dence 1972, Wood and Head 1976). 

Small impact structures have the form of a bowl-shaped 

depression with an upraised rim and are known as simple 

craters (Fig. 1). The exposed rim, walls, and floor define the 

so-called apparent crater. At the rim, there is an overturned 

flap of ejected target materials, which displays inverted 

stratigraphy, with respect to the original target materials. 

Beneath the floor is a lens of brecciated target material that 

is roughly parabolic in cross-section (Fig. 2). This breccia 

lens is allochthonous and polymict, with fractured blocks of 

various target materials. In places, near the top and the base, 

the breccia lens may contain highly shocked, and possibly 

melted, target materials. Beneath the breccia lens, parautochthonous, 

fractured rocks define the walls and floor 

of what is known as the true crater. In the case of terrestrial 

simple craters, the depth to the base of the breccia lens (i.e., 

the base of the true crater) is roughly twice that of the depth 

to the top of the breccia lens (i.e., the base of the apparent 

crater, Fig. 2). Shocked rocks in the parautochthonous materials 

of the true crater floor are confined to a small central 

volume at the base of the true crater. 

253

Ann M. Therriault – Richard A. F. Grieve – Mark Pilkington 

a 

Figure 1. (a) Oblique aerial view of 1.2 km diameter, 50,000 years old simple crater, Meteor or Barringer Crater, Arizona, U.S.A. (b) Vertical aerial 

view of 3.8 km diameter, 450 ± 30 million years old, Brent Crater, Ontario, Canada. Note how this ancient crater has no rim, has been filled by sediments 

and lakes and is a generally subtle topographic feature. 

b 

With increasing diameter, simple craters show increasing 

evidence of wall and rim collapse and evolve into complex 

craters (Fig. 3). The transition diameter varies 

between planetary bodies and is, to a first approximation, 

an inverse function of planetary gravity (Pike 1980). 

Other variables, such as target material and possibly projectile 

type and velocity, play a lesser role, so that the transition 

diameter varies over a small range. The most 

obvious effect of secondary variables appears on Earth, 

where there are major areas of both sedimentary and crystalline 

rocks at the surface. Complex craters on Earth first 

occur at diameters greater than 2 km in layered sedimentary 

target rocks but not until diameters of 4 km or greater 

in stronger, more coherent, igneous or metamorphic, crystalline 

target rocks (Dence 1972). 

With a central topographic peak or peaks, a broad, flat 

floor, and terraced, inwardly slumped rim areas (Fig. 4), 

complex craters are a highly modified craterform compared 

to simple craters. The rim of a typical complex 

D 

 

yyyyyyyyy 

a D t 

 

yyyyyyyyy 

 

yyyyyyyyy Slump 

Autochthonous 

target rocks 

Simple crater – final form 

D 

Basal melt pool 

(minor fall back) 

breccia fill 

Figure 2. Schematic cross-section of a simple crater. D is the diameter 

and d a and d t are the depths of the apparent and true crater, respectively. 

See text for details. 

crater is a structural feature corresponding to a series of 

fault terraces. Interior to the rim lays an annular trough, 

which is partially filled by a sheet of impact-melt rock 

and/or polymict allochthonous breccia (Fig. 4). Only in 

the central area of the crater is there evidence of substantial 

excavation of target materials. This region is structurally 

complex and, in large part, occupied by a central 

peak, which is the topographic manifestation of a much 

broader and extensive area of uplifted rocks that occurs 

beneath the center of complex craters. Readers interested 

in the details of cratering mechanics at simple and complex 

structures are referred to Melosh (1989) and references 

therein. 

With increasing diameter, a fragmentary ring of interior 

peaks appears, marking the transition from craters to 

basins. While a single interior ring is required to define a 

basin, basins have been subdivided, with increasing diameter, 

on other planetary bodies, into central-peak basins, 

with both a peak and ring; peak ring basins, with only a 

ring; and multi-ring basins, with two or more interior rings 

(Wood and Head 1976). There have been claims that the 

largest known terrestrial impact structures have multi-ring 

forms, e.g. Chicxulub, Mexico (Sharpton et al. 1993), 

Sudbury, Canada (Stöffler et al. 1994, Spray and Thompson 

1995) and Vredefort, South Africa (Therriault et al. 

1997). Although certain of their geological and geophysical 

attributes form annuli, it is not clear that these correspond, 

or are related in origin, to the obvious 

topographical rings observed, for example, in lunar multiring 

basins (Spudis 1993, Grieve and Therriault 2000). 

Most terrestrial impact structures are affected by erosion. 

In extreme cases, the craterform has been completely 

removed. In such cases, recognition of structural and 

254


a 

Figure 3. (a) Oblique aerial photograph of the Gosses Bluff impact structure, Australia. 

Note that all that is visible of this originally 22 km, 142.5 ± 0.8 million 

years old structure is a 5 km annulus of hills, representing the eroded remains of 

a central uplift. See text for details. (b) Shuttle photograph of the Manicouagan 

impact structure, Canada, 100 km in diameter and 214 ± 1 million years old. Note 

that the annular trough (with a diameter of ~ 65 km) is filled by water. 

b 

geological effects of impact in the target rocks is essential 

to the identification of an impact structure rather than the 

presence of a characteristic craterform. For example, 

Gosses Bluff, Australia has a positive topographical form 

consisting of an annular ring of hills, approximately 5 km 

in diameter (Fig. 3). The ring consists of erosionally resistant 

beds from within the original central uplifted area 

of a complex impact structure. The original craterform, 

which has an estimated diameter of approximately 22 km 

(Milton et al. 1996), has been removed by erosion. There 

are several other impact structures, which have some form 

of rings, e.g. Manicouagan, Canada (Floran and Dence 

1976), Haughton, Canada (Robertson and Sweeney 1983), 

but it is not clear whether these are primary forms or secondary 

features, with some relation to primary structural 

features (Grieve and Head 1983). 

There are also other subtleties to the character of 

craterforms in the terrestrial record that do not appear on 

the other terrestrial planets. A number of relatively young, 

and, therefore, only slightly eroded, complex impact structures 

(e.g. Haughton, Canada; Ries Germany; Zhamanshin, 

Kazakhstan) do not have an emergent central peak or 

other interior topographical expression of a central uplift 

(Garvin and Schnetzler 1994). These structures are in 

mixed targets of platform sediments overlying crystalline 

basement. Although there are no known comparably 

young complex structures entirely in crystalline targets, 

the buried and well-preserved Boltysh structure, Ukraine, 

which is of comparable size, has a central peak (emergent 

from the crater-fill), similar to the appearance of lunar central 

peak craters. This difference in form is probably a target 

rock effect but it has not been studied in detail. 

The morphology of impact craters formed in marine 

environment is also quite distinct. These impact structures 

are characterized by a broad and shallow brim at the periphery 

of the crater, extensive infilling, and prominent 

fault blocks floored by apparent low-angle décollement 

surfaces at the periphery of the crater (e.g. Tsikalas et al. 

1999, Ormö and Lindström 2000). The extensive infilling 

is most likely due to large amounts of ejecta and crater 

wall material transported into the excavated crater by the 

collapse of the impact-induced water cavity and the subsequent 

rapid surge of sea water (Tsikalas et al. 1999, Ormö 

and Lindström 2000). The 40-km-diameter Mjølnir submarine 

impact structure in the Barents Sea, for example, 

consists of a central region of deep excavation surrounded 

by a shallow excavated shelf, without a raised crater rim 

(Tsikalas et al. 1998, 1999). This morphology is also observed 

at the 13.5-km-diameter Lockne impact structure, 

Sweden (Lindström et al. 1996). 

Attempts to define morphometric relations, particularly 

depth-diameter relations, for terrestrial impact structures 

have had limited success, because of the effects of 

erosion and, to a lesser degree, post-impact sedimentation. 

Unlike depth, the variation of stratigraphic uplift (SU, Fig. 

4) with diameter at complex impact structures is fairly 

d t 

d a 

Complex structure – final form 

Central Uplift Area 

Dcp 

D 

Melt/allochthonous 

breccia sheet 

"Autochthonous" 

crater floor 

SU 

Figure 4. Schematic cross-section of complex impact structure. Notation 

as in Figure 2, with SU corresponding to structural uplift and D cp to the 

diameter of the central uplift. Note preservation of beds in outer annular 

trough of the structure, with excavation limited to the central area. See 

text for details. 

255


well constrained with SU = 0.86D 1.03 (n = 24), where n is 

the number of data points (Grieve and Pilkington 1996). 

Similarly, the diameter of the central uplift area (D cp , Fig. 

4), at its maximum radial expression, is constrained by D cp 

= 0.31 D 1.02 (n = 44) (Therriault et al. 1997). 

Geophysics of impact structures 

Geophysical anomalies over terrestrial impact structures 

vary in their character and, in isolation, do not provide 

definitive evidence for an impact origin. About 30 per 

cent of known terrestrial impact structures are buried by 

post-impact sediments. Geophysical methods resulted in 

their initial discovery and subsequent drilling provided 

geologic samples, which confirmed their impact origin. 

Interpretation of a single geophysical data set over a suspected 

impact structure can be ambiguous (for example, 

Hildebrand et al. 1998, Sharpton et al. 1993). When combined, 

however, with complementary geophysical methods 

and the existing database over other known impact structures, 

a more definite assessment can be made (e.g. Ormö 

et al. 1999). 

Since potential-field data are available over large areas, 

with almost continuous coverage, gravity and magnetic 

observations have been the primary geophysical indicators 

used for evaluating the occurrence of possible terrestrial 

impact structures. Seismic data, although providing much 

better spatial resolution of subsurface structure, is used 

less, because it is less generally available. Electrical methods 

have been used even less (e.g. Henkel 1992). Given 

space limitations and some lack of specificity of the geophysical 

attributes of terrestrial impact craters, they are 

generally discussed here and the reader is referred to the 

most recent synthesis in Grieve and Pilkington (1996). 

Gravity signature 

The most notable geophysical signature associated 

with terrestrial impact structures is a negative gravity 

anomaly. These gravity lows are generally circular, extending 

to, or slightly beyond, the crater rim, and are due 

to lithological and physical changes associated with the 

impact process. In well-preserved impact structures, lowdensity 

sedimentary infill of the topographic depression of 

the crater contributes to the gravity low. In complex impact 

structures, relatively lower density impact-melt sheets 

also contribute to the negative gravity effect. However, 

such lithological effects are minor compared to density 

contrasts induced by fracturing and brecciation of the target 

rocks. 

In general, the amplitude of the maximum negative 

gravity anomaly associated with impact structures increases 

with the final crater diameter (Dabizha and Fedynsky 

1975, Dabizha and Feldman 1982). Over simple craters, a 

circular bowl-shaped negative anomaly is observed; 

whereas most of larger complex impact structures, greater 

than 30 km in diameter, tend to exhibit a central gravity 

high. Based on data from 58 terrestrial impact structures, 

Pilkington and Grieve (1992) showed that erosional level 

has only a secondary effect on gravity anomaly size. 

It is important to note that due to differences in target 

lithologies, large variations in gravity signature are observed 

between structures of similar sizes. In general, 

structures formed in sedimentary lithologies produce 

smaller anomalies than similar sized ones formed in crystalline 

rocks. Structures formed in unconsolidated sediments 

in continental shelf areas may not produce 

detectable negative gravity anomalies but are marked only 

by a central gravity high. 

Magnetic signature 

In general, magnetic anomalies associated with terrestrial 

impact structures are more complex than gravity 

anomalies. This observation reflects the greater variation 

possible in the magnetic properties of rocks. The dominant 

effect over impact structures is a magnetic low or subdued 

zone ranging in amplitude from tens to a few hundred nanotesla 

that is commonly manifested as a truncation of the 

regional magnetic fabric (Dabizha and Fedynsky 1975, 

Clark 1983). Magnetic lows are best defined over simple 

and some small complex craters, where the anomaly is 

smooth and simple; whereas at larger impact structures, 

the magnetic low can be modified by the presence of 

shorter-wavelength, large-amplitude, localized anomalies 

that usually occur at or near the centre of the structure. 

No correspondence exists between the magnetic anomaly 

character and crater morphology of impact structures. 

Moreover, the presence of a central gravity high does not 

imply the existence of a central magnetic anomaly. There 

are several structures with no obvious magnetic signature. 

Shock effects, thermal effects or chemical effects may 

cause magnetic anomalies related to impact. Shock effects 

in impact structures can serve to increase or decrease magnetization 

levels. Thermal effects may result in the production 

of non-magnetic impact glasses (Pohl 1971) or in 

resetting magnetic minerals through thermoremanent 

magnetization in the direction of the Earth’s magnetic field 

at the time of impact. Chemical effects may result in the 

production of new magnetic phases, through elevated 

residual temperatures and hydrothermal alteration, leading 

to the acquisition of a chemical remanent magnetization in 

the direction of the ambient field. 

Seismic signature 

Reflection seismic surveys allow for detailed imaging 

of impact structure morphology and delineating zones of 

incoherent reflections that are characteristic of brecciation 

and fracturing. The disturbance of coherent subsurface reflectors 

is most prominent in the central uplift of complex 

structures and decreases outward and downward from this 

zone (Brenan et al. 1975). Reflection data can provide es- 

256


timates of such morphological parameters as the dimensions 

of the central uplift, annular trough and faulted 

blocks at the structural rim of complex structures (e.g. 

Morgan et al. 2002). The depth to horizontal reflectors that 

exist below the crater floor can be used to determine the 

amount of structural uplift. 

Electrical signature 

The presence of fluids in impact-induced fractures and 

pore spaces leads to decreased resistivity levels that can be 

mapped effectively by various electrical methods. The 

conductivity of rocks is heavily dependent on their water 

content: < 1% change in water content can produce more 

than an order of magnitude change in conductivity. The 

degree of fragmentation determines the amount and distribution 

of fluids within the rock and hence, its electrical 

properties. 

Where a distinct contrast exists between the allochthonous 

breccia deposits and the underlying autochthonous 

target rocks, electrical profiling using resistivity sounding 

can map the structure of the true crater floor (e.g. Vishnevsky 

and Lagutenko 1986). In order to determine the 

deeper electrical structure associated with impact, magnetotelluric 

surveys have been carried out (e.g. Zhang et al. 

1988, Campos-Enriquez et al. 1997). 

Geology of impact structures 

Although an anomalous circular topographic, structural, 

or geological feature may indicate the presence of an 

impact structure, there are other endogenic geological 

processes that can produce similar features in the terrestrial 

environment. An obvious craterform is an excellent indicator 

of a possible impact origin; particularly, if it has 

the appropriate morphometry, but as noted, such features 

are rare and short-lived in the terrestrial environment. The 

burden of proof for an impact origin generally lies with the 

documentation of the occurrence of shock-metamorphic 

effects. 

Few structures preserve physical evidence of the impacting 

body. Such structures are limited to small, young, 

simple structures, where the impacting body (or, more 

commonly, fragments of it) has been slowed by atmospheric 

deceleration and impacts at less than cosmic velocity. 

These are restricted generally to the impact of iron or 

stony-iron meteorites. Stony meteorites are weaker than 

their iron-bearing counterparts and small stones are generally 

crushed as a result of atmospheric interaction (Melosh 

1981). Larger impacting bodies (>100–150 m in diameter) 

survive atmospheric passage with undiminished impact 

velocity. Consequently, the peak shock pressures upon impact 

are sufficient, in most cases, to result in the melting 

and vaporisation of the impacting body, destroying it as a 

physical entity. 

On impact, the bulk of the impacting body’s kinetic 

Temperature (°C) 

10 000 

1 000 

100 

P-T field of 

endogenic 

metamorphism 

Shatter 

cones 

Rock 

melting 

Fused 

glasses 

Diaplectic 

glasses 

Planar 

features 

Vaporization 

Pressure post-shock 

temperature curve for 

shock metamorphism 

of granitic rocks 

1 10 100 1 000 

Pressure, GPa 

Figure 5. Temperature and pressure range of shock metamorphic effects 

compared to that of endogenic metamorphism. Planar features include 

planar deformation features (PDFs) and planar fractures (PFs). Scale is 

log-log. See text for details. 

energy is transferred to the target by means of a shock 

wave. This shock wave imparts kinetic energy to the target 

materials, which leads to the formation of a crater. It also 

increases the internal energy of the target materials, which 

leads to the formation of so-called shock-metamorphic effects. 

The details of the physics of impact and shockwave 

behavior can also be found in Melosh (1989), and references 

therein. 

Shock metamorphism is the progressive breakdown in 

the structural order of minerals and rocks due to the passage 

of a high-pressure shock wave and requires pressures 

and temperatures well above the pressure-temperature 

field of endogenic terrestrial metamorphism (Fig. 5). The 

dependence on high pressures for the formation of shockmetamorphic 

effects has been shown by their duplication 

in nuclear and chemical explosion craters, and in laboratory 

shock recovery experiments (e.g. Hörz 1968, Müller 

and Hornemann 1969, Borg 1972). Minimum shock pressures 

required for the production of diagnostic shockmetamorphic 

effects are 5–10 GPa for most silicate 

minerals. Strain rates produced by impact cratering 

process are of the order of 10 6 s -1 to 10 9 s -1 (Stöffler and 

Langenhorst 1994), many orders of magnitude higher than 

typical tectonic strain rates (10 -12 s -1 to 10 -15 s -1 ; e.g. Twiss 

and Moores 1992), and shock-pressure duration is measured 

in seconds, or less, in even the largest impact events 

(Melosh 1989). These physical conditions are not reproduced 

by endogenic geologic processes. They are unique 

to impact and, unlike endogenic terrestrial metamorphism, 

disequilibrium and metastability are common phenomena 

in shock metamorphism. 

The extreme pressures and high strain rates of shock 

deformation are fundamental differences from normal endogenic 

causes of compression (Ashworth and Schneider 

1985, Goltrant et al. 1991, 1992, Langenhorst 1994). A 

shock wave passing through a heterogeneous rock mass 

undergoes numerous modifications, as it interacts with 

grain boundaries, fractures, foliations, and different mineral 

species with different shock impedances within the 

257


Figure 6. Outcrop (~ 80 m high) of coherent impact melt rock at the Mistastin 

complex impact structure, Canada. 

rock. There is, thus, local variations in shock pressure. Petrographic 

study indicates that shock pressures may vary by 

a factor of two or more over distances ranging from millimeters 

to meters in outcrops (Grady 1977). Hence, each 

individual mineral grain experiences its own particular 

shock history based upon its physical properties and its relationship 

to both the adjacent grains and the overall structural 

character of the rock. A maximum shock effect in 

grains of a particular mineral species in a hand specimen 

may, thus, be a means of measuring relative deformation 

intensities throughout an impact structure. For example, 

shock pressures of at least ~ 5 GPa are required to produce 

PFs in quartz and greater than 10 GPa to produce PDFs in 

quartz or feldspars. This variation of shock deformation of 

important rock-forming minerals of the target rocks with 

increasing shock pressures have been used to delineate 

zones of shock metamorphism in the floor of a number of 

impact structures, e.g. Charlevoix, Canada (Robertson 

1968), Brent, Canada (Dence 1968), Ries, Germany (von 

Engelhardt and Stöffler 1968), and Manicouagan, Canada 

(Dressler 1990), with the intensity of deformation decreasing 

from the center outwards. 

The exact physical conditions on impact are a function 

of the specific impact parameters. The density of the impacting 

body and the target, and the impact velocity determine 

the peak pressure on impact. The shock wave 

attenuates with distance from the impact point with the kinetic 

energy of the impact event determining the absolute 

radial distance in the target at which a specific shock pressure 

is achieved and, thus, which specific shock-metamorphic 

effects occur. Shock-metamorphic effects are well 

described in papers by Chao (1967), Bunch (1968), Stöffler 

(1971, 1972, 1974), Stöffler and Langenhorst (1994), 

Grieve et al. (1996), French (1998), Langenhorst and 

Deutsch (1998), and Langenhorst (this volume). They are 

discussed here only in general terms. 

Impact melting 

During compression, considerable pressure-volume 

work is done and the pressure release occurs adiabatically. 

Heating of the target rocks, thus, occurs as not all this 

pressure-volume work is recovered upon pressure release 

and results in irreversible waste heat. Above 60 GPa, the 

waste heat is sufficient to cause whole-rock melting and, 

and at higher pressures, vaporisation of a certain volume 

of target rocks (Melosh 1989). This volume is a function 

of the impact velocity, physical properties of the impacting 

body and target, and, most importantly, the size of the impacting 

body (Grieve and Cintala 1992). 

Impact melt lithologies may occur as glass bombs in 

crater ejecta (von Engelhardt 1990), as dykes within the 

crater floor and walls, as glassy to crystalline pools and 

lenses within the breccia lenses of simple craters, or as coherent 

annular sheets (Fig. 6) lining the floor of complex 

craters and stratigraphically located immediately above 

breccias and/or brecciated basement rocks and overlain by 

breccias. 

When crystallized, impact-melt sheets have igneous 

textures, but tend to be heavily charged with clastic debris 

a 

b 

Figure 7. Photomicrographs of far-from-equilibrium textures examples in impact melts: (a) plagioclase crystals with swallow-tail texture, Boltysh impact 

melt sheet, Ukraine, plane light, field-of-view = 2.28 mm; (b) pyroxene-plagioclase spherulitic texture, Vredefort Granophyre impact melt dyke, 

South Africa, plane light, field-of-view = 5 mm. 

258


Figure 8. Photomicrograph of fused glass (lechatelierite), Ries, Germany, 

plane light, field-of-view = 2.5 mm. 

towards their lower and upper contacts. They may, therefore, 

have a textural resemblance to endogenic igneous 

rocks. Impact melts are superheated, reaching thousands 

of degrees Kelvin. Temperature differences with host 

rocks may result in rapid cooling of the melt leading to farfrom-equilibrium 

textures (Fig. 7). Grain-size in thick impact-melt 

sheets increases inwards from the contacts, but, 

in general, impact-melt rocks are usually fine-grained to 

glassy. An important textural property of impact-melt 

rocks is the presence of mineral and rock fragments, which 

have undergone shock metamorphism of different degrees, 

and have been variously reworked by the melt. The size of 

such fragments ranges from millimeters to several hundreds 

of meters, and gradational changes in inclusion content 

are observed in thick melt sheets, varying from one to 

several tens of percent (e.g. von Engelhardt 1984), with 

highest concentrations towards their lower and upper contacts. 

Impact-melt rocks can have an unusual chemistry compared 

with endogenic volcanic rocks, as their composition 

depends on the wholesale melting of a mix of target rocks, 

as opposed to partial melting and/or fractional crystallization 

relationships for endogenous igneous rocks. The composition 

of impact-melt rocks is characteristic of the target 

rocks and may be reproduced by a mixture of the various 

country rock types in their appropriate geological proportions. 

Such parameters as 87 Sr/ 86 Sr and 143 Nd/ 144 Nd ratios 

may also reflect the pre-existing target rocks within the 

impact-melt rocks composition (Jahn et al. 1978, Faggart 

et al. 1985). In general, unlike endogenous magmatic rock 

masses of comparable size (up to a few hundred meters 

thick), even relatively thick impact-melt sheets are chemically 

homogeneous over distances of millimeters to kilometers. 

In cases where the target rocks are not 

homogeneously distributed, this observation may not hold 

true, such as for Manicouagan, Canada (Grieve and Floran 

1978), Chicxulub (Kettrup et al. 2000) and Popigai (Kettrup 

et al. 2002). Differentiation is not a characteristic of 

relatively thick coherent impact-melt sheets (with the exception 

of the extremely thick, ~ 2.5 km, Sudbury Igneous 

Complex, Sudbury Structure, Canada; Ostermann 1996, 

Ariskin et al. 1999, Therriault et al. 2002). 

Enrichments above target rock levels in siderophile elements 

and Cr have been identified in some impact-melt 

rocks. These are due to an admixture of up to a few percent 

of meteoritic material from the impacting body. In 

some melt rocks, the relative abundances of the various 

siderophiles have constrained the composition of the impacting 

body to the level of meteorite class, (e.g. East 

Clearwater, Canada, was formed by a C1 chondrite, Palme 

et al. 1979). In other melt rocks, no siderophile anomaly 

has been identified. This may be due to the inhomogeneous 

distribution of meteoritic material within the impact-melt 

rocks and sampling variations (Palme et al. 

1981) or to differentiated, and, therefore, relatively nonsiderophile-enriched 

impacting bodies, such as basaltic 

achondrites. More recently, high precision osmium-isotopic 

analyses have been used to detect a meteoritic signature 

at terrestrial impact structures (e.g. Koeberl et al. 

1994). Unfortunately, Re-Os systematics are, in themselves, 

not an effective discriminator between meteorite 

classes. 

Fused glasses and diaplectic glasses 

In general, shock fused minerals are characterized 

morphologically by flow structures and vesiculation (Fig. 

8). Peak pressures required for shock melting of single 

crystals are in the order of 40 to 60 GPa (Stöffler 1972, 

1974), for which postshock temperatures (> 1000 °C) exceed 

the melting points of typical rock-forming minerals 

(Fig. 5). At these conditions, the minerals in the rock will 

melt immediately and independently after the passage of 

the shock wave. This melt has approximately the same 

composition as the original mineral before any flow or 

mixing takes place, and the melt regions are initially distributed 

through the rock in the same manner as the original 

mineral grains (French 1998). Melting is mineral 

selective, producing unusual textures in which one or more 

minerals show typical melting features; whereas, others, 

even juxtaposed ones, do not. One of the most common 

fused glasses observed at terrestrial impact structures is 

that of quartz, i.e. lechatelierite (e.g. Fig. 8). 

Conversion of minerals to an isotropic, dense, glassy 

phase at peak pressures of 30 to 50 GPa (Fig. 5) and temperatures 

well below their normal melting point is a shock 

metamorphic effect unique to framework silicates. These 

phases are called diaplectic (from the Greek “destroyed by 

striking”) glasses, which are produced by breakdown of 

long-range order of the crystal lattice without fusion. 

Although diaplectic forms may occur as the direct result of 

compression by the shock wave, they are probably more 

commonly produced by inversion from a high-pressure 

crystalline phase, which is unstable in the postshock P-T 

environment (Robertson 1973). Based on shock recovery 

experiments, the formation of diaplectic glass occurs between 

30 and 45 GPa for feldspar and 35 to 50 GPa for 

quartz (e.g. Stöffler and Hornemann 1972). The morphology 

of the diaplectic glass is the same as the original mineral 

crystal and shows no evidence of fluid textures (e.g. 

Grieve et al. 1996). Diaplectic glasses have densities low- 

259


Figure 9. Photomicrograph of partial conversion to maskelynite of plagioclase 

feldspar crystals, Manicouagan, Canada, cross-polarized light, 

field-of-view = 5 mm. 

er than the crystalline form from which they are derived, 

but higher than thermally melted glasses of equivalent 

composition (e.g. Stöffler and Hornemann 1972, Langenhorst 

and Deutsch 1994). With increasing pressure, the 

bulk density of diaplectic glass decreases. This decrease is 

due in part to progressively greater portions of the mineral 

having been converted to low density, disordered phases, 

but also to the fact that diaplectic phases exist in a 

sequence of intermediate structural states, whose refractive 

index and density decrease with increasing pressure 

and temperature (Stöffler and Hornemann 1972). The refractive 

index of diaplectic glasses is also generally higher 

than for synthetic, or thermally melted, glasses of 

equivalent composition (e.g. Robertson 1973, Grieve et al. 

1996). However, in the case of K-feldspar, its diaplectic 

glass has a slightly lower refractive index than the fused 

feldspar glass (Stöffler and Hornemann 1972). Maskelynite, 

the diaplectic form of plagioclase (Fig. 9), is the 

most common example from terrestrial rocks; diaplectic 

glasses of quartz (Chao 1967) and of alkali feldspar 

(Bunch 1968) are also reported but in lesser abundance. 

Diaplectic glasses of different minerals can exist adjacent 

to one another without mixing (e.g. Robertson 1973). 

High-pressure polymorphs 

Shock can result in the formation of metastable polymorphs, 

such as stishovite and coesite from quartz (Chao 

et al. 1962, Langenhorst this volume) and diamond and 

lonsdaleite from graphite (Grieve and Masaitis 1996, Masaitis 

1998, Langenhorst this volume). Coesite and diamond 

are also products of endogenic terrestrial geological 

processes, including high-grade metamorphism, but the 

paragenesis and, more importantly, the geological setting 

are completely different from that in impact events. 

Under high pressure, the mineral lattice is unstable and 

is converted to a more stable configuration. Such transformation 

begins at ~ 11.5 GPa for K-feldspars (Robertson 

1973) and at ~ 12 GPa for quartz (De Carli and Milton 

1965). With increasing pressure, a greater proportion of 

the mineral is converted to a high-pressure polymorph until 

complete transformation is achieved at ~ 30 GPa for 

feldspars (Ahrens et al. 1969) and ~ 35 GPa for quartz 

(Stöffler and Langenhorst 1994). Neither the high-pressure 

phase of K-feldspar, thought to be the dense hollandite-type 

structure with Al and Si in octahedral 

co-ordination, nor an equivalent plagioclase polymorph 

have been recovered from shock experiments or identified 

in non-impact terrestrial rocks (Robertson 1973). It would 

appear that these phases are very unstable in postshock environments 

and, more likely, invert to more disordered, 

metastable phases. The high-pressure polymorphs of 

quartz (i.e. stishovite and coesite) have only rarely been 

produced by laboratory shock recovery experiments (cf. 

Stöffler and Langenhorst 1994). Contrary to what is expected 

from equilibrium phase diagram, stishovite is 

formed at lower pressures (12–30 GPa) than coesite 

(30–50 GPa; Stöffler and Langenhorst 1994) in impact 

events. This is mainly due to the fact that stishovite is 

formed during shock compression, whereas, coesite crystallizes 

during pressure release. In terrestrial impact structures, 

these polymorphs occur in small or trace amounts as 

very fine-grained aggregates and are formed by partial 

transformation of the host quartz. In crystalline or dense 

rocks, coesite is found in quartz with planar deformation 

features (PDFs) and strongly lowered refractive index and, 

more commonly, in diaplectic glass; whereas, in porous 

sandstone, coesite co-exists with > 80% of quartz displaying 

planar fractures (PFs) and diaplectic quartz glass 

(Grieve et al. 1996). Stishovite occurs most commonly in 

quartz with PDFs and less frequently in diaplectic glass 

(Stöffler 1971). For details on the characteristics of coesite 

and stishovite, the reader is referred to Stöffler and Langenhorst 

(1994) and references therein. 

Planar microstructures 

The most common documented shock-metamorphic 

effect is the occurrence of planar microstructures in tectosilicates, 

particularly quartz (Hörz 1968). The utility of 

planar microstructures in quartz reflects the ubiquitous nature 

of the mineral and its stability, including the stability 

of the microstructures themselves, in the terrestrial environment, 

and the relative ease with which they can be documented. 

For details, the reader is referred to the 

accompanying paper by Langenhorst. Recent reviews of 

the nature of the shock metamorphism of quartz, with an 

emphasis on the nature and origin or planar microstructures 

in experimental and natural impacts, can be found in 

Stöffler and Langenhorst (1994) and Grieve et al. (1996). 

Planar deformation features (PDFs) in minerals are 

produced under pressures of ~ 10 to ~ 35 GPa (Fig. 5). 

Planar fractures (PFs) form under shock pressures ranging 

from ~ 5 GPa up to ~ 35 GPa (Stöffler 1972, Stöffler and 

Langenhorst 1994). 

260


Shatter cones 

The only known diagnostic shock effect that is megascopic 

in scale is the occurrence of shatter cones (Dietz 

1968). Shatter cones are unusual, striated, and horse-tailed 

conical fractures ranging from millimeters to meters in 

length produced in rocks by the passage of a shock wave 

(e.g. Sagy et al. 2002). The striated surfaces of shatter 

cones are positive/negative features and the striations are 

directional, i.e., they appear to branch and radiate along 

the surface of the cone. The acute angle of this distinctive 

pattern points toward the apex of the cone and the shatter 

cones themselves generally point upward with their axes 

lying at any angle to the original bedding. Once the host 

rocks are graphically restored to their original impact position, 

shatter cones indicate the point of impact. 

Shatter cones are initiated most frequently in rocks that 

experienced moderately low shock pressures, 2–6 GPa 

(Fig. 5), but have been observed in rocks that experienced 

~25 GPa (Milton 1977). These conical striated fracture 

surfaces are best developed in fine-grained, structurally 

isotropic lithologies, such as carbonates and quartzites. 

They do occur in coarse-grained crystalline rocks but are 

less common and poorly developed. They are generally 

found as individual or composite groups of partial to complete 

cones (Fig. 10) in place in the rocks below the crater 

floor, especially in the central uplifts of complex impact 

structures, and rarely in isolated rock fragments in breccia 

units. Shatter cones are used as a diagnostic field criterion 

to identify impact structures (e.g. Dietz 1947, Milton 

1977). 

Conclusion 

The detailed study of impact events on Earth is a relatively 

recent addition to the spectrum of studies engaged in 

by the geological sciences. More than anything, it was 

preparations for and, ultimately, the results of the lunar 

and the planetary exploration program that provided the 

initial impetus and rationale for their study. Some recent 

discoveries have resulted from the occurrence or re-examination 

of unusual lithologies, rather than an obvious circular 

geological or topographic feature. For example, 

unusual breccias at Gardnos, Norway and Lockne, Sweden 

had been known for some time, but their shock-metamorphic 

effects were documented only recently, and they 

are now associated with the remnants of impact structures 

(French et al. 1997, Lindström and Sturkell 1992). 

The level of knowledge concerning individual terrestrial 

impact structures is highly variable. In some cases, it 

is limited to the original discovery publication. In terms of 

understanding the terrestrial record, this is compensated, 

to some degree, by the fact that impact structures with similar 

dimensions and target rocks have the same major characteristics. 

Nevertheless, there is still much to be learned 

about impact processes from terrestrial impact structures, 

Figure 10. Complete shatter cone in limestone, Cap de la Corneille, 

Charlevoix, Canada. 

particularly with respect to details of the third dimension. 

This is the property that is unobtainable from impact structures 

on other bodies in the solar system, where it must be 

studied by remote-sensing methodologies. 

Apart from increasing our understanding of impact 

processes, the study of terrestrial impact structures has influenced 

the siting of significant economic deposits 

(Grieve and Masaitis 1994, Donofrio 1997, Grieve 1997). 

In addition, the documentation of the terrestrial impact 

record provides a direct measure of the cratering rate on 

Earth and, thus, a constraint on the hazard that impact 

presents to human civilization (Gehrels 1994). The K/T 

impact may have resulted in the demise of the dinosaurs as 

the dominant land-life form and, thus, permitted the ascendancy 

of mammals and, ultimately, humans. It is, however, 

inevitable that human civilization, if it persists long 

enough, will be subjected to an impact-induced environmental 

crisis of potentially immense proportions. 

Acknowledgements. We would like to thank J. Ormö and 

A. Deutsch for their critical reviews. This paper is GSC contribution 

No. 2002141. 

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263



Shock metamorphism of some minerals: Basic introduction and microstructural observations 

FALKO LANGENHORST 

Bayerisches Geoinstitut, University of Bayreuth, Universitätsstr. 30, D-95447 Bayreuth, Germany; e-mail: Falko.Langenhorst@uni-bayreuth.de 

Abstract. Minerals show a unique behaviour when subjected to shock waves. The ultradynamic loading to high pressures and temperatures 

causes deformation, transformation and decomposition phenomena in minerals that are unequivocal indicators of impact events. This paper introduces 

into the basics of shock compression, required to understand the formation and experimental calibration of these shock effects in minerals, and particularly 

focuses on the recent advances in the field of shock metamorphism achieved by the application of transmission electron microscopy (TEM). 

TEM studies underline that the way minerals respond to shock compression largely depends on their crystal structures and chemical compositions, as 

is illustrated here on the basis of four minerals: quartz, olivine, graphite and calcite. 

The crystal structure of a mineral exerts an important control on the shock-induced deformation phenomena, comprising one- to two-dimensional 

lattice defects, such as dislocations, mechanical twins, planar fractures, and amorphous planar deformation lamellae. For example, dislocations cannot 

be activated in quartz due to the strong covalent bonding, whereas the island silicate olivine easily deforms by dislocation glide. 

Transformation phenomena include phase transitions to (diaplectic) glass and/or high-pressure polymorphs. TEM studies reveal that high-pressure 

polymorphs such as coesite, stishovite and ringwoodite are liquidus phases, which form upon decompression by crystallization from high-pressure 

melts. The graphite-to-diamond transition is however a rare example for a solid-state transformation, taking place by a martensitic shear mechanism. 

Shock-induced decomposition reactions are typical of volatile-bearing minerals and liberate toxic gases that, in case of large impacts, may affect 

Earth’s climate. Shock experiments show that degassing of calcite does not take place under high pressure but can massively occur after decompression 

if the post-shock temperature is sufficiently high. A recombination reaction happens however if CaO and CO 2 are not physically separated. 

Key words: impact features, shock metamorphism, shock waves, minerals, crystal structure, defects, experimental study 

Introduction 

Number of craters (>1 km) per km 2 

0.04 

0.03 

0.02 

0.01 

0.00 

4 3 2 1 0 

Age (Ga) 

Apollo Luna Earth 

Fig. 1. Cratering statistics of the Moon-Earth system for the last 4 Ga 

years [data from Stöffler and Ryder (2001) and regression curve after 

Neukum et al. (2001)]. 

Collisions of solid bodies played a crucial role in the 

formation of Earth and its subsequent evolution. The Proto-Earth 

accreted by collisions of planetesimals (Wetherill 

1984) and underwent, in Archean times, a very heavy 

bombardment (Melosh 1989). The high collision rate at 

the beginning of Earth was simply the consequence of the 

chaotic states in the orbital movements of early solid bodies. 

Subsequently, this early collision rate distinctly declined 

until about 2.5 Ga, and since then we have an 

essentially constant flux of bodies colliding with Earth 

(Fig. 1). 

This knowledge of the collision history of Earth is 

however not exclusively based on observations of terrestrial 

impact craters, as one would possibly expect, but it 

largely relies on the cratering history of the Moon (Chapman 

and Morrison 1994, Neukum et al. 2001, Stöffler and 

Ryder 2001). Due to the absence of an atmosphere (and 

hence weathering) and the early cessation of volcanic activity 

on Moon, impact craters accumulated, in nearly 

undisturbed fashion, particularly in the lunar highlands, 

where crater densities are highest. On the geologically 

very active Earth, the morphologic landforms of impact 

craters easily disappear beyond recognition due to erosion, 

plate tectonics and other exogenic and endogenic processes. 

Although four Archean impact layers have been reported 

(Byerly et al. 2002), we have no knowledge of an 

Archean crater (Fig. 2). According to the cratering statistics 

(Fig. 1), one would however expect that most craters 

formed during this early episode of Earth (Grieve and 

Shoemaker 1994). 

265

Falko Langenhorst 

Identified craters per 1 Ma 

1.0 

0.5 

0.0 

Cenozoic (0–65 Ma) 

Mesozoic (65–250 Ma) 

Impact events are however not only expressed at the 

large scale in the form of circular craters but they manifest 

also at microscopic scales in minerals. This article focuses 

on these microscopic traces in minerals termed shock or 

shock-metamorphic effects. Shock effects in minerals are 

an unequivocal fingerprint and, often, the last remnants of 

impact events. Even if the crater is completely erased, 

shocked minerals can be preserved in distal or global ejecta 

horizons, still providing evidence of the impact and its 

age. In case of the Cretaceous/Tertiary extinction (Alvarez 

et al. 1980), the discovery of shocked minerals in the KT 

boundary layer increased, for example, the efforts to find 

the corresponding impact structure (Bohor et al. 1984 and 

1987), subsequently identified as the Chicxulub crater 

(Hildebrand et al. 1991). 

In the last 10–15 years we have obtained new insights 

and a better understanding of the shock behaviour of minerals, 

mainly due to the increased use of transmission electron 

microscopy (TEM). In terms of spatial resolution, this 

technique is well superior to optical microscopy. Therefore, 

it became possible to decipher the nature of shock 

phenomena already known from optical studies (e.g. planar 

deformation features (PDFs)), to discover so-far unknown 

effects, and to understand their mechanisms of 

Paleozoic (250–570 Ma) 

Geological Era 

Fig. 2. The number of identified impact craters per Ma plotted for different 

eras in the Earth history. Note that despite the high cratering activity 

in the Archean we have no knowledge of an impact crater from this 

time period (cf. Fig. 1, data from http://gdcinfo.agg.nrcan.gc.ca/crater/ 

world_craters_e.html). 

Proterozoic (570–2500 Ma) 

Archean (2500–4550 Ma) 

formation and subsequent alteration. This article does not 

aim to give a comprehensive review on shocked minerals; 

for a more detailed compilation the reader is referred to: 

e.g., French and Short (1968), Stöffler (1972, 1974), Stöffler 

and Langenhorst (1994), French (1998), Deutsch and 

Langenhorst (1998), Langenhorst and Deutsch (1998). It 

is merely an attempt to utilize instructive examples to explain 

the specific response of minerals to shock compression, 

which is distinctly different from the transformation 

and deformation behaviour of minerals in other geologic 

processes, principally of lower strain rate. 

Principles of shock waves 

From a basic understanding of shock waves it will be 

immediately clear why minerals respond differently to impact 

than to other geologic processes. It will also help to 

understand the mechanics of the large-scale crater-forming 

process (Roddy et al. 1977, Melosh 1989). A shock wave 

is an extreme compression wave that propagates with supersonic 

velocity, abruptly compresses, heats, and plastically 

deforms solid matter. In this respect, shock waves are 

fundamentally different from seismic (elastic) waves (e.g. 

Duvall and Fowles 1963). 

A shock wave is produced by the impact of a high-velocity 

projectile or the detonation of an explosive (Roddy 

et al. 1977, Asay and Shahinpoor 1993). In these processes, 

the time of load is extremely short and, therefore, the 

initial stress wave steepens immediately to an almost 

atomically sharp discontinuity, which separates highly 

compressed from uncompressed material. This so-called 

“shock front” represents consequently an infinite discontinuity 

in all state parameters (Duvall and Fowles 1963, 

Melosh 1989): pressure P, temperature T, specific volume 

V, and energy E. Minerals and rocks undergo some kind of 

a physical “shock”, because they have to instantly adapt to 

extreme P,T conditions with strain rates on the order of 10 6 

to 10 9 s -1 . Additionally, a shock wave is very short-lived 

with a typical pulse duration of about 1 second for a natural 

impact of a 10 km diameter projectile. It is due to these 

two time parameters, high strain rates and short shock duration, 

that minerals cannot respond by equilibrium reactions 

but rather by the activation of fast deformation and 

transformation mechanisms. 

The unusual properties of shock waves can be illustrated 

by a simple train model (Fig. 3). The cartoon shows 

the collision of a moving train (projectile) with a standing 

train (target). The collision leads to deformation and compression 

of the hitting locomotive and the wagons in two 

opposite directions. The two boundaries between compressed 

and uncompressed wagons are the shock fronts, 

which propagate with the shock velocity U. An additional 

effect of the compression is material flow in the direction 

of the target (see the displacement of the boundary between 

the hitting locomotive and the last target wagon). 

The material moves behind the shock wave into the target, 

266

time 


at a particle velocity u p , which is distinctly smaller than 

the shock velocity U. As the process proceeds, more and 

more wagons are engulfed by the compression until the 

shock wave in the projectile is reflected as rarefaction 

wave at the free rear side of the projectile (Fig. 3). This 

leads to a backward acceleration of wagons and represents 

the ejection of material in natural impact cratering. As the 

rarefaction wave is propagating in already compressed 

material, it is faster than the shock wave in the target. Consequently, 

the shock wave will be overtaken at some depth 

by the rarefaction wave, i.e. it will decay. 

Both shock U and particle u p velocities characterise the 

state of material under shock compression and are related 

to pressure P, density ρ, and energy E via the Hugoniot 

equations (Duvall and Fowles 1963, Melosh 1989): 

ρ 0 U = ρ (U – u p ) (1) 

P – P 0 = ρ 0 U u p (2) 

P u p = ρ 0 U (u p2 /2) + ρ 0 U (E – E 0 ) (3) 

These equations express the conservation of mass, momentum, 

and energy across the shock front. By combining 

equations (1) to (3) the velocities can be eliminated to give 

the so-called Rankine-Hugoniot equation: 

E – E 0 = (P – P 0 ) (V 0 – V)/2 with V = 1/ρ (4) 

This is an equation of state, describing the physical 

states that can be achieved by shock waves of variable intensity 

in any solid. Commonly, the shock equation of state 

of a particular material is experimentally determined by 

measuring particle u p and shock U velocities. Since the zero-pressure 

densities ρ 0 of minerals are usually well known, 

pressures P and densities ρ (or specific volumes V) can be 

calculated with the above mentioned Hugoniot equations. 

As is usual for all equations of state, the shock data of 

a particular material are displayed in a pressure-specific 

volume plot, defining the so-called Hugoniot curve. The 

Hugoniot curves of rock-forming minerals are commonly 

characterised by kinks, i.e., discontinuous changes in the 

curve slope. One discontinuity is typically at 5–10 GPa, 

the so-called Hugoniot elastic limit, which is the yield 

strength of the material. Discontinuities at higher pressures 

were generally assumed to result from phase transformations 

to high-pressure polymorphs (e.g. Ahrens and 

Rosenberg 1968, Jackson and Ahrens 1979, Ahrens et al. 

1976, Melosh 1989). This interpretation is certainly correct 

for materials such as iron and graphite, which undergo 

martensitic (displacive) transformations, fast enough to 

take place under shock compression. It is however an inappropriate 

assumption for silicate materials such as 

quartz, feldspars and olivine. These minerals can only be 

converted into high-pressure polymorphs via reconstructive 

mechanisms. These mechanisms are however too 

time-consuming to occur under shock compression. 

Therefore, shock experiments on these minerals usually 

result in the formation of dense (diaplectic) glass with possibly 

five- and/or six-fold coordinated silicon or finely recrystallised 

low-pressure phase. 

U 

V 0 , E 0 

P 0 , T 0 

U 

P 1 , T 1 

V 1 , E 1 

shock front 

u p 

material 

front 

u p 

Experimental simulation 

shock front 

rarefaction front 

ejection 

Fig. 3. A train model illustrating the formation and propagation of shock 

and rarefaction waves and the associated material movement (from Langenhorst 

2000). 

Shock experiments are indispensable for understanding 

the formation of shock effects in minerals, because 

they are performed under controlled laboratory conditions 

and because they provide shocked minerals in their initial, 

unmodified state. This has two basic advantages. First, 

shock effects can be calibrated as function of pressure (or 

other factors) and the resultant barometers can then, with 

some care, be applied to nature (Stöffler and Langenhorst 

1994). Secondly, the study of pristine shock effects helps 

to understand possible post-shock modifications (annealing, 

chemical alteration etc.) in the natural impact environment 

(Grieve et al. 1996). 

The major disadvantage of shock experiments is their 

short pulse duration (< 1µs), which is at least 6 orders of 

magnitude shorter than the pressure pulse in a natural 

bolide impact (~ 1 s, Langenhorst et al. 2002a). The short 

pulse duration in an experiment is simply the result of the 

small size of the projectile. The pressure duration corresponds 

approximately to the time that a shock wave needs 

to travel through the projectile and back (i.e. 2 projectile 

diameters, Fig. 3). 

It may be some surprise, but pioneering, experimental 

work on impact processes started with Alfred Wegener 

(1921), the founder of plate tectonics. He simulated impact 

events by throwing cement powder in a tablespoon 

with his bare hand onto a target, which was also composed 

of cement powder. He was able to produce circular craters 

with central uplifts, resembling in shape and proportions 

those observed on the Moon. Since then, a large variety of 

sophisticated shock techniques has been developed to 

simulate cratering mechanics and shock metamorphism of 

267


a 

b 

booster 

laser 

d 

D 

high explosive 

flyer plate 

spacing ring 

cover layer 

specimen 

Fe cylinder 

vacuum 

chamber 

lens 

mirrors 

sample 

steel blocks 

Fig. 4. Two different experimental designs to produce shock waves: (a) High-explosive device used at the Ernst-Mach-Institut, Efringen-Kirchen, Germany 

(modified after Langenhorst and Deutsch 1994) and (b) a laser irradiation setup with the sample being clamped into an Al block. The laser is focused 

on a thin Al foil, acting as flyer plate (modified after Langenhorst et al. 1999a and 2002a). 

minerals (e.g. French and Short 1968, Roddy et al. 1977, 

Asay and Shahinpoor 1993, Davison et al. 2002). 

Most experiments related to cratering mechanics employ 

spherical projectiles that produce and excavate some 

crater in an infinite half space medium via a spherically 

expanding shock wave. In contrast, most experiments related 

to shock metamorphism employ flat-plate projectiles 

that drive a planar shock wave through a similarly planar 

target; the thickness of the latter is typically less than projectile 

thickness to avoid measurable pressure decay 

across the sample (Barker et al. 1993). The target is usually 

a metallic container, encapsulating the sample in a fashion 

to allow its partial or complete recovery. The 

experiments may differ widely in the type of accelerating 

system (Fig. 4): powder and light-gas guns, high-explosive 

charges, electric discharge guns, and laser irradiation techniques 

have been used and tested to successfully reproduce 

shock effects in minerals (e.g. Milton and DeCarli 

1963, Müller and Hornemann 1969, Gratz et al. 1992, 

Stöffler and Langenhorst 1994, Langenhorst et al. 1999a, 

2002a). The principle of the electric gun is to vaporise a 

thin metal foil by rapid electric discharge of a capacitor. 

The high electrical current leads to the instantaneous vaporisation 

of the foil and the production of a shock wave 

(Langenhorst et al. 2002a). Laser irradiation experiments 

can be performed either with or without a projectile. In the 

latter case, the beam is directly focused on the sample surface 

(Langenhorst et al. 2002a), whereby the absorption of 

the laser energy generates rapidly exploding plasma that 

subsequently induces a shock wave. Such plasma techniques 

are capable to produce the highest shock pressures 

with an unbelievable world record of 750 Mbar (Cauble et 

al. 1993). However, higher shock pressures are commonly 

at the expense of shorter shock durations and smaller sample 

volumes. This is because higher impact velocities and 

hence higher shock pressures can only be achieved by reducing 

the size and weight of the projectile. Typical pressure 

pulses in laser irradiation, electric discharge and highexplosive 

experiments are approximately 1 ns, 10 ns, and 

1 µs with shocked sample volumes on the order of 0.1, 1, 

and 100 mm 3 , respectively (Langenhorst et al. 2002a). 

To determine pressures in shock experiments it is necessary 

to measure the velocities associated with the shock 

waves (projectile v, particle u and shock U velocities), using 

electrical pin contact, optical interferometry (VISAR) 

or similar techniques (Hornemann and Müller 1971, Barker 

et al. 1993). It is then possible to calculate the pressures 

temperature (°C) 

3000 

2500 

2000 

1500 

1000 

500 

quartz 

eclogite 

liquid 

coesite 

release 

paths 

stishovite 

porous sandstone 

0 

1 10 

100 

shock pressure (GPa) 

Fig. 5. Phase diagram depicting the pressure-temperature conditions 

reached in quartz, olivine (solid line), and porous sandstone (long dashed 

line) by shock compression (data after Wackerle 1962, Kieffer et al. 

1976, and Holland and Ahrens 1997). The release paths hold for porous 

quartz, which first melts on loading and then solidifies on cooling as coesite 

or stishovite. The equilibrium phase boundaries between quartz, coesite, 

stishovite and liquid are drawn as dashed lines. The grey 

pressure-temperature field represents the conditions reached in regional 

metamorphism; the solid line within this field depicts the pressure-temperature 

path of a diamond-bearing gneiss (Stöckhert et al. 2001). 

quartz 

olivine 

268


by aid of the known Hugoniot relations of the materials involved 

(Marsh 1980). Thus, the determination of shock 

pressure is fundamentally different from the approach 

used in static compression techniques (e.g. piston cylinder, 

multi-anvil or diamond anvil cell experiments), in which 

the pressures are calibrated via properties and/or phase 

transformations of reference materials (e.g. Rubie 1999). 

The determination of temperatures is a less precise and 

more difficult undertaking. So far, pyrospectrometric techniques 

have been used to measure both shock and postshock 

temperatures (e.g. Raikes and Ahrens 1979, Holland 

and Ahrens 1997; Fig. 5). The knowledge of temperatures 

is however of fundamental importance for the determination 

of the melting temperatures of materials at very high 

pressures. For example, the precise measurement of the 

melting curve of iron enables the temperature at the 

boundary between Earth’s inner (solid) and outer (liquid) 

core to be determined (Brown and McQueen 1986), an 

important fix point of the geotherm. 

calcite graphite olivine quartz 

shock effects 

mechanical twins 

PF 

PDF 

mosaicism 

diaplectic glass \ lechatelierite 

stishovite \ coesite 

dislocations 

PF and PDF 

mosaicism 

recrystallisation and staining 

ringwoodite 

kink bands 

diamond 

mechanical twins 

dislocations 

decomposition and recombination products 

shock pressure (GPa) 

0 10 20 30 40 50 60 

0 10 20 30 40 50 60 

Shock effects in minerals 

Fig. 6. Compilation of the pressure intervals over which certain shock effects 

are formed in quartz, olivine, graphite and calcite (modified after 

Stöffler and Langenhorst 1994 and Langenhorst and Deutsch 1998). The 

diagram is based on shock experiments with non-porous samples. 

Shock-induced physical and chemical changes in minerals 

are collectively called shock effects or shock-metamorphic 

effects. This term is relatively broad and covers 

any type of shock-induced change, such as formation of 

lattice defects, phase transformations, decomposition reactions 

and resultant changes in physical and chemical properties. 

A great diversity of natural shock effects has been 

described on the basis of spectroscopic techniques, X-ray 

diffraction, and optical microscopy (Hörz and Quaide 

1973, Stöffler 1972 and 1974, Schneider 1978, Boslough 

et al. 1989). The physical nature of some of the effects was 

not completely clear and others were not even known until 

TEM was used to characterize the mineralogical effects 

at the nanometer scale. Based on TEM observations, the 

following simple subdivision of shock effects and processes 

occurring during shock compression and decompression 

has been proposed (Langenhorst and Deutsch 1998): 

(1) Deformation: formation of dislocations, planar microstructures 

(planar fractures and planar deformation 

features), mechanical twins, kink bands, and mosaicism 

(2) Phase transformations into high-pressure phases and 

diaplectic glass 

(3) Decomposition into a solid residue and a gaseous 

phase 

(4) Melting and vaporisation of entire mineral (subsequently 

quenched as shock-fused glass or polycrystalline 

aggregates) 

This simple list does not contain the post-shock thermal 

effects forming distinctly after the impact. In a wide 

sense, these effects may also be regarded as shock indicators, 

although they are not primary effects of shock compression 

and decompression and simply result from the 

high temperatures prevailing after an impact event. Examples 

of diagnostic post-shock effects are the formation of 

Ballen quartz and checkerboard feldspars in impact melts 

(Carstens 1975, Bischoff and Stöffler 1984) or the crystallization 

of highly oxidised Ni-rich spinels (magnesioferrites) 

in microtektites (Robin et al. 1992). 

We will focus here exclusively on the shock effects 

listed above. The deformation and transformation effects 

largely form during the compression phase of shock 

waves, whereas decomposition, melting and vaporisation 

are temperature dominated processes, which take place 

during the decompression phase when pressure declines to 

a larger extent than temperature. There is no single mineral 

that shows all of these effects (Fig. 6). The response of 

a mineral to shock compression largely depends on its 

crystal structure and composition. A mineral with strong 

three-dimensional covalent bonding between polyhedra in 

the crystal structure can, for example, not respond by dislocation 

glide. Also, minerals without volatile components 

cannot respond by decomposition reactions such as dehydration 

and decarbonation. To highlight these aspects and 

to illustrate the various types of shock effects in minerals 

we will concentrate on four minerals, differing largely in 

terms of crystal structure and composition. 

Quartz 

Among the rock-forming minerals, quartz (SiO 2 ) displays 

the widest variety of shock effects. It possesses a 

crystal structure, consisting of three-dimensionally linked, 

corner-shared SiO 4 -tetrahedra with strong covalent bonding. 

In the low-pressure regime (< 30 GPa), quartz can 

269


(101 — 1) 

a 

50 µm 

0.2 µm 

b 

(0001) 

c 

0.5 µm 

100 nm 

d 

Fig. 7. Shock effects in quartz: (a) Optical micrograph of planar deformation features in shocked quartz in a garnet gneiss from the Popigai crater, 

Siberia, (b) bright-field TEM image of fresh, amorphous PDFs in shocked quartz from the Massignano outcrop, Italy (see Langenhorst 1996), (c) Darkfield 

TEM image of a mechanical Brazil twin in shocked quartz from the Mjølnir crater, Barents Sea and (d) dark-field scanning TEM image of numerous 

stishovite needles embedded in a glassy shock vein of the Zagami meteorite (see Langenhorst and Poirier 2000). 

therefore not react by dislocation glide, particularly because 

the activation and glide of dislocations is controlled 

by the diffusion of water-related defects (see McLaren 

1991 for “hydrolytic weakening”), which is a rather slow 

process compared to the short time frame of shock compression. 

Instead, quartz develops so-called planar microstructures 

(Grieve et al. 1990), which comprise planar fractures 

(PF), planar deformation features (PDFs), and mechanical 

twins (Stöffler and Langenhorst 1994). The mechanical 

twins were previously also regarded as PDFs, because a distinction 

on the basis of optical microscopy is impossible. 

PFs can basically be regarded as high-pressure cleavage 

planes that are only activated by dynamic shock loading; 

they are relatively widely spaced (> 5–20 µm). On the 

contrary, quartz has no cleavage at ambient conditions and 

fails by a glassy-like parting. 

PDFs were previously also called planar elements (Engelhardt 

and Bertsch 1969) and are thin (< 200–300 nm) 

amorphous lamellae with the same composition as the host 

crystal (Fig. 7b); at the TEM scale, their spacing (< 1 µm) 

is much smaller than that of PFs (Müller 1969, Ashworth 

and Schneider 1985, Gratz et al. 1992, Langenhorst 1994). 

PDF orientations are crystallographically controlled and 

show a pressure-dependent variation. Most PDFs are oriented 

parallel to rhombohedral planes such as {101 — 3} and 

{101 — 2}; other less abundant PDF orientations are given in 

Table 1 (see also Stöffler and Langenhorst 1994). In the 

low-pressure regime (< 30 GPa), PDF orientations are regarded 

as the most robust barometer, because even postshock 

annealing and alteration merely converts the PDFs 

into the “decorated” type (French 1969; Fig. 7a), but PDF 

orientations remain unchanged. The decoration of PDFs in 

naturally shocked quartz is due to recrystallization of the 

270


Table 1. Compilation of the most abundant PDF orientations in shocked quartz. 

Miller indices {h k i l} Azimuth angle Pole distance ρ Symbol Form 

(0001) – 0° c basal pinacoid 

{101 — 3}, {011 — 3} p, n 30° 22.95° ω, ω’ rhombohedron 

{101 — 2}, {011 — 2} p, n 30° 32.42° π, π’ rhombohedron 

{101 — 1}, {011 — 1} p, n 30° 51.79° r, z rhombohedron 

{101 — 0} 30° 90° m hexagonal prism 

{404 — 1v, {044 — 1} p, n 30° 78.87° t rhombohedron 

{516 — 0}, {61 — 5 — 0} r, l 40° 90° k ditrigonal prism 

{516 — 1}, {61 — 5 — 1} p, n, r, l 40° 82.07° x trigonal 

{61 — 5 — 1}, {156 — 1} trapezoedron 

{314 — 1}, {43 — 1 — 1} p, n, r, l 45° 77.91° – trigonal 

{41 — 3 — 1}, {134 — 1} trapezoedron 

{213 — 1}, {32 — 1 — 1} p, n, r, l 50° 73.71° – trigonal 

{31 — 2 — 1}, {123 — 1} trapezoedron 

{112 — 2}, {21 — 1 — 2} r, l 60° 47.73° ξ trigonal dipyramid 

{112 — 1}, {21 — 1 — 1} r, l 60° 65.56° s trigonal dipyramid 

{112 — 0}, {21 — 1 — 0} r, l 60° 90.0° a trigonal prism 

{224 — 1}, {42 — 2 — 1} r, l 60° 77.20° trigonal dipyramid 

p = positive, n = negative, r = right, l = left 

amorphous lamellae and resultant exsolution 

of water. The exsolution of water can 

be attributed to the different solubility of 

water in silica glass and crystalline quartz. 

Therefore, the fluids mobilized in impact 

events can initially be dissolved in the 

amorphous PDFs but subsequent recrystallization 

of the glass expels the water in 

form of tiny voids (Goltrant et al. 1991, 

1992, Leroux and Doukhan 1996, Grieve et 

al. 1996, Langenhorst and Deutsch 1998). 

Sub-planar features of tectonic origin 

such as the so-called Böhm lamellae have 

been erroneously assigned as PDFs (Ernstson 

et al. 1985, Vrána 1987; Fig. 8a). TEM 

studies have deciphered the nature of these 

tectonic features, as being subgrain boundaries 

(Cordier et al. 1994, Langenhorst and 

Deutsch 1996, Joreau et al. 1997a). Subgrain 

boundaries are dislocation arrays, 

separating a crystal into sub-grains that are 

slightly tilted (< 5°) with respect to each other. They are 

the result of slow plastic deformation and recovery of deformed 

quartz in a tectonic environment. The deformation 

proceeds by the activation and emission of dislocations 

coupled with simultaneous or subsequent climb and recovery 

of dislocations into sub-grain boundaries (Fig. 8b; 

Poirier 1985). Water can easily penetrate along these internal 

boundaries leading to a decoration with water bubbles. 

A careful optical inspection of the suspected planar 

features in quartz will immediately reveal whether they are 

of endogenic or exogenic origin. Sub-grain boundaries 

show a sub-parallel arrangement but are not perfectly planar, 

unlike shock-produced PDFs. The spacing of tectonic 

features (usually > 5–10 µm) is larger than that of shockproduced 

PDFs (< 1 µm). Under crossed Nicols, the tectonically 

deformed quartz grains show undulatory extinction 

and can exhibit an anomalous optic axial angle of up 

to 10°. In contrast, shocked quartz usually shows a patchy 

extinction pattern, with extinct areas in different parts of 

the crystal. This behaviour is the so-called mosaicism. In a 

tectonic source rock, not all of the deformed quartz grains 

contain sub-grain boundaries in a sub-planar arrangement; 

many quartz grains may even be devoid of sub-grain 

boundaries. On the other hand, it would be rather unusual 

for an impact rock that only one or few quartz grains contain 

PDFs. If the quartz grains have experienced the same 

pressure-temperature conditions, then at least most of 

them should contain PDFs (Hörz 1968, Engelhardt and 

Bertsch 1969). The size and orientation of quartz grains 

with respect to the shock front has however an influence 

on the development of PDFs (Walzebuck and Engelhardt 

a 

50 µm 3 µm 

b 

Fig. 8. (a) Optical micrograph and (b) corresponding bright-field TEM image of subgrain boundaries in a tectonically deformed quartz from Azuara, 

Spain (see Langenhorst and Deutsch 1996). 

271


1979). These effects may therefore be responsible for a 

heterogeneous PDF distribution throughout quartzose 

rocks, particularly at pressures required for incipient PDF 

formation (~ 10 GPa). 

Another important outcome of TEM studies was the 

discovery of mechanical Brazil twins (Leroux et al. 1994), 

which are exclusively oriented parallel to the basal plane 

(0001). Numerous partial dislocations in the twin boundaries 

indicate the mechanical nature of the twins (Fig. 7c). 

In contrast, the commonly known Brazil twins are hydrothermally 

grown twins, which lack partial dislocations 

and are always oriented parallel to (101 — 1) (McLaren and 

Pithkethly 1982, Langenhorst and Poirier 2002). Static deformation 

experiments have shown that a high shear stress 

of 3 to 4 GPa has to be applied to generate the mechanical 

Brazil twins parallel to (0001) (McLaren et al. 1967). In 

crustal rocks, such conditions are only met by an impact 

event. At the optical scale, the mechanical twins become 

only visible if they are decorated with water bubbles. The 

mechanical Brazil twins are more resistant to post-shock 

annealing and alteration. Long-term regional metamorphism 

with complete recrystallization of the shocked rock 

or melting is probably the only way to erase them. At the 

Vredefort structure, most of the PDFs in shocked quartz 

are lost by post-shock overprint but the Brazil twins are 

still present. This discovery could explain why the PDF 

orientations indicate an apparent decrease in shock pressure 

toward the center of the crater (Schreyer 1983, Nicolaysen 

and Reimold 1990). 

In the high-pressure regime (> 25–30 GPa), quartz is 

dominated by phase transformations either to diaplectic 

glass or to high-pressure polymorphs. Diaplectic glass is a 

densified glass, preserving the shape and sometimes even 

internal features of precursor grains (Engelhardt et al. 

1967, Stöffler 1984). X-ray studies demonstrate that the 

transition from crystalline quartz to diaplectic glass is 

marked by increasing mosaicism coupled with decreasing 

domain size (Dachille et al. 1968, Hörz and Quaide 1973, 

Hanns et al. 1978). The transition is also characterized by 

a decrease of refractive indices and densities but diaplectic 

glass has a density about 5% higher than that of synthetic 

silica glass (Langenhorst and Deutsch 1994). It is 

still not fully understood on whether the transformation 

represents a solid-state collapse (Stöffler 1984) or quenching 

of a liquid under high pressure (Langenhorst 1994). 

The latter seems to be more likely because diaplectic glass 

is often associated with coesite that crystallized from highpressure 

melt (see below). 

Solid-state phase transformations of quartz into the 

high-pressure polymorphs coesite and stishovite are reconstructive. 

This means that time is needed for the cooperative 

movement and diffusion of atoms. Consequently, 

the short duration of shock compression impedes a direct 

solid-state transformation of quartz into the high-pressure 

polymorphs. Instead, quartz has to be melted under high 

pressure and the high-pressure phases can then readily 

crystallize during the decompression phase (see release 

paths in Fig. 5). The latter means that coesite and 

stishovite form in their stability fields at pressures, which 

are smaller than the maximum shock pressure achieved 

during the compression phase. The very high temperatures 

needed for melting are reached locally within shocked 

rocks, e.g. at pores or along shock veins, which result from 

shear-induced frictional melting (Kieffer et al. 1976, Langenhorst 

and Poirier 2000). The crystallization of highpressure 

polymorphs from melt is much faster than a 

reconstructive solid-state transformation because the 

liquidus temperatures at high pressures are well above 

2000°C, where kinetics is very fast. Furthermore, it is 

known from NMR studies that compressed silicate melts 

contain five- and six-fold coordinated silicon (Xue et al. 

1989, Stebbins and Poe 1999), which should facilitate per 

se the crystallization of high-pressure polymorphs from 

such dense melts. Indeed, coesite and stishovite have always 

been found in conjunction with silica or rock glass 

(Kieffer et al. 1976, White 1993, Leroux et al. 1994, Langenhorst 

and Poirier 2000). Polycrystalline coesite aggregates 

occur in diaplectic silica glass (Hörz 1965, Stöffler 

1971), e.g. at the Ries and Popigai craters. The rapid crystallization 

led to the formation of numerous (100) rotation 

twins with the composition plane (010) (Leroux et al. 

1994, Grieve et al. 1996). These twins are known to represent 

growth defects (Bourret et al. 1986). Stishovite has 

been identified in the porous Coconino sandstone from the 

Barringer crater (Kieffer et al. 1976), in thin pseudotachylite 

veins in shocked basement rocks of the Vredefort 

structure, South Africa (Martini 1991, White 1993), as 

well as in shock veins in the Martian meteorite Zagami 

(Fig. 7d, Langenhorst and Poirier 2000). Recently, the discovery 

of post-stishovite polymorphs has been reported, as 

well (ElGoresy et al. 2000), but the high beam sensitivity 

prevented a thorough characterization of the phases (Sharp 

et al. 1999). 

Olivine 

Olivine is an island silicate with isolated SiO 4 tetrahedra 

that are joined by divalent cations. As a consequence, 

it is easier than in quartz to break bonds and to activate and 

move dislocations in this crystal structure. In the low-pressure 

regime, olivine deforms thus by dislocation glide. The 

Burgers vector b is always [001] (= c) but depending on 

the orientation of the olivine to the shock front, various 

slip planes can be activated, all sharing the c direction as 

zone axis (Fig. 9a): (010), (100), (110), (hk0) and symmetrically 

equivalent planes (Ashworth and Barber 1975, 

Langenhorst et al. 1995, Leroux et al. 1996, Joreau et al. 

1997b, Langenhorst and Greshake 1999). The dislocation 

densities can be as high as 2 × 10 14 m –2 (Madon and Poirier 

1983, Langenhorst et al. 1995, Leroux 2001), even in experimentally 

shocked olivine (Langenhorst et al. 1999a 

and 2002b). Experiments have also demonstrated that dislocations 

can propagate at approximately half the sound 

velocity (~ 3 km/s, Langenhorst et al. 1999a) and are 

272


0.5 µm 0.2 µm 

a b 

Fig. 9. (a) Dark-field TEM image of numerous c dislocations and a planar fracture in a shocked olivine grain from the Tenham chondrite, (b) Brightfield 

TEM image of ringwoodite in a shock vein of the ordinary chondrite Acfer 90072. Note the numerous stacking faults parallel to {110} planes. 

probably hampered to reach this maximum speed due to 

interactions between themselves. The sources of dislocations 

seem to be the tips of forming planar fractures. When 

the planar fractures are formed, a high stress field is created 

at their tips, from which the dislocations are first emitted 

as loops. Since the edge component is distinctly faster 

than the screw component, the dislocations become very 

straight during further propagation, with the long dislocation 

line representing the screw segment. 

The PFs in olivine are indicative of shock as they are 

oriented parallel to rational crystallographic planes that 

are not known as normal cleavage planes of olivine. The 

PFs are typically parallel to low index planes (Müller and 

Hornemann 1969, Snee and Ahrens 1975, Reimold and 

Stöffler 1978, Bauer 1979, Langenhorst et al. 1995, Stöffler 

et al. 1991): (100), (010), (001), (130), and (110), belonging 

to pinacoidal and prismatic forms. Both, the 

internal fragmentation of olivine by fracturing and the 

high density of dislocations contribute to the patchy extinction 

behaviour under crossed Nicols, known as mosaicism. 

Observations on naturally and experimentally shocked 

olivine reveal that mosaicism increases distinctly in the 

high-pressure regime (> 25–30 GPa; Carter et al. 1968, 

Müller and Hornemann 1969, Snee and Ahrens 1975, 

Reimold and Stöffler 1978, Bauer 1979, Schmitt 2000). 

Additional shock phenomena in this pressure regime are 

staining, recrystallization, and transformation into the 

high-pressure polymorphs wadsleyite and ringwoodite 

(Fig. 9b). Although the production of glass has been reported 

in one experimental study (Jeanloz et al. 1977), diaplectic 

or shock-fused glasses are generally unknown for 

olivine. 

Brownish staining of olivine has been described for experimentally 

and naturally shocked olivine (Stöffler et al. 

1991, Schmitt 2000) but the reason for the colorization is 

unknown; a careful microstructural characterization is certainly 

required to unravel the nature of the staining effect. 

The term recrystallisation decribes the formation of 

fine-grained olivine (1–2 µm) aggregates, mostly in the 

vicinity of shock veins. Shock-induced recrystallisation of 

olivine is generally assumed to be a solid-state process 

(Carter et al. 1968, Ashworth and Barber 1975, Bauer 

1979, Stöffler et al. 1991). This interpretation actually 

makes sense, because recovery and polygonization of dislocations 

are expected to occur at elevated post-shock temperatures, 

resulting in fine-grained strain-free olivine 

aggregates with lattice-preferred orientation of subgrains 

(Lally et al. 1976). 

In analogy to quartz, the high-pressure transformations 

require first the production of an olivine melt, which has 

then to rapidly crystallize upon decompression as highpressure 

polymorphs. Such rapid crystallization is so far 

only manifested in shock veins of ordinary chondrites, i.e. 

in thin localized melt zones that are first formed by shearinduced 

frictional heating (Stöffler et al. 1991) and are 

then rapidly quenched by the surrounding cold host rock. 

The shock veins contain aggregates of tiny (~1 µm) 

wadsleyite and ringwoodite grains, together with other 

high-pressure phases (Binns et al. 1969, Putnis and Price 

1979, Madon and Poirier 1983, Langenhorst et al. 1995, 

Chen et al. 1996). Olivine melt crystallizing at pressures 

beyond the ringwoodite stability field yields for example 

an assemblage of silicate perovskite and magnesiowüstite 

(Sharp et al. 1997, Tomioka and Fujino 1997), the expected 

mineral assemblage in the Earth’s lower mantle. The 

coexistence of high-pressure phases that should not coexist 

under equilibrium indicates that the crystallization of 

shock veins probably happened during decompression, i.e. 

the phases crystallized progressively from the melt while 

the pressure declined. 

In ordinary chondrites, wadsleyite and ringwoodite are 

both characterized by numerous stacking faults. Wadsleyite 

develops faults parallel to the (010) plane (Madon 

and Poirier 1983) and ringwoodite parallel to {110} planes 

(Fig. 9b; Langenhorst et al. 1995, Chen et al. 1996). The 

273


a 

Graphite 

b 

Diamond 

A 

C 

A 

c 

B 

B 

 

C 

C 

A 

Fig. 10. Crystal structures of (a) hexagonal (ABAB…) graphite and (b) cubic (ABCABC…) diamond showing the different stacking sequences in the 

[0001] and [111] directions, respectively. 

stacking faults are regarded as growth defects, i.e. they are 

an inevitable result of the short crystallization times. The 

time for crystallisation of shock veins depends primarily 

on the thickness of the vein and the initial temperature difference 

between vein and adjacent cold host rock. Recent 

calculations suggest that the times for crystallisation of 

high-pressure polymorphs are much shorter than the shock 

duration (Langenhorst and Poirier 2000). 

In terrestrial impact rocks, the high-pressure polymorphs 

of olivine have not been found yet, probably because 

of the lack of well-preserved, olivine-rich target 

rocks with thin shock veins. 

Graphite 

Graphite is composed of pure carbon and is an example 

of a mineral with a pronounced sheet structure (Fig. 

10a). Within the sheets, carbon atoms form hexagonal 

rings with very strong covalent bonds. Weak van-der- 

Waals bonding prevails between the sheets. For such a layered 

structure, it is characteristic to react to shock 

compression by kink or twin operations. In the low-pressure 

regime (< 25–30 GPa), graphite is usually assumed to 

kink but there are no exact measurements of rotation angles 

between different parts of shock-folded graphite crystals 

to exclude the possibility of twin operations (Stöffler 

1972). 

Graphite is of most interest for its phase transformation 

to diamond. In the context of impact events, the formation 

of diamond is a rare example for a solid-state transformation. 

Another mineralogical example for such a transformation 

is the conversion of zircon into the scheelite 

structured high-pressure polymorph reidite (Glass et al. 

2002). Two atomic operations are necessary to convert 

graphite into diamond (Fig. 10). The hexagonal carbon 

layers have to be brought together by compression along 

the c axis and the stacking sequence has to be changed 

from a hexagonal to a cubic array by shearing of the 

hexagonal carbon layers in their a-a plane. Neglecting the 

van-der-Waals bonds, it is not necessary to break any 

strong bonds within the layers. Such shear-induced transitions 

are called martensitic transformations. Even in short 

shock experiments it is possible to produce this martensitic 

transformation; this was first demonstrated by DeCarli 

and Jamieson (1961). 

In light of this first shock synthesis, the long-known diamonds 

in iron meteorites and ureilites (Erofeev and 

Lachinov 1888, Foote 1891) were reinterpreted as shock 

products (Lipschutz 1964). On Earth, a large number of 

impact craters and the K/T boundary are yet known as find 

locations of impact diamonds (Masaitis et al. 1990 and 

2000, Valter et al. 1992, Koeberl et al. 1997, Hough et al. 

1997). They were first discovered in the 100 km sized 

Popigai structure, Siberia (Masaitis et al. 1972), which is 

regarded as largest diamond deposit on Earth (Deutsch et 

al. 2000). In the Ries, impact diamonds were already 

found in 1978 by Rost et al., but it was not before 1995 

that this finding was noticed by a broader scientific community 

(Masaitis et al. 1995, Hough et al. 1995, El Goresy 

et al. 2001). 

The source rocks of diamonds from terrestrial impact 

craters are usually graphite-bearing gneisses or other crystalline 

rocks (Masaitis et al. 1990, ElGoresy et al. 2001). 

Graphite in these rocks was transformed within the very 

short time of shock compression. As a consequence of the 

short transformation time, impact diamonds are very defect-rich 

and inherited some features of the precursor 

graphite. They are birefringent (Fig. 11a), retain the tabular 

hexagonal and sometimes even preserve spectacular 

growth twins rotated about the c axis of graphite (Fig. 

11b). Therefore impact diamonds are regarded as para- or 

pseudomorphs after graphite and are called apographitic 

diamonds (Masaitis et al. 1990). At the TEM scale, these 

impact diamonds contain numerous kink or twin bands 

parallel to (h h 2 — h — l) planes of precursor graphite (Fig. 11c; 

Langenhorst et al. 1999b). The bands were generated 

274


b 

0.1 mm a 

c 

Fig. 11. (a) Optical micrograph of a tabular impact diamond from the 

Ries crater, Germany. The anomalous interference colours are due to internal 

strain; crossed Nicols; (b) secondary electron scanning image of 

impact diamond from the Popigai impact crater, showing inherited twins 

of the precursor graphite, (c) Dark-field TEM image of an impact diamond 

from Popigai crater, showing mechanical twin bands that are inherited 

from precursor graphite. 

when the shock wave was transmitted into the graphite, i.e. 

directly before the phase transformation (Langenhorst 

2000). TEM studies failed, so far, to detect lonsdaleite although 

X-ray diffraction techniques indicate the presence 

of this high-pressure polymorph with a hexagonal stacking 

sequence (Frondel and Marvin 1967, Hannemann et al. 

1967, Masaitis et al. 1990). Instead, one observes that the 

diamonds are very disordered and contain numerous 

stacking faults, changing locally the cubic into a hexagonal 

stacking sequence (Langenhorst 2000). Diffuse X-ray 

scattering on these stacking faults is a way to explain the 

extra-peaks in X-ray diffraction patterns. 

Strongly corroded impact diamonds that are intergrown 

with moissanite (SiC) have been found at the Ries 

impact crater, Germany. Based on this finding, Hough et 

al. (1995) concluded that the diamonds formed by vapour 

condensation (so-called chemical vapour deposition 

(CVD) mechanism). Meanwhile, this idea has been discarded, 

because the moissanite may have formed by reaction 

of diamond with silica-rich impact melt (Langenhorst 

et al. 1999b). This reaction also forms the basis for the industrial 

production of SiC (Mehrwald 1992). Incorporation 

of impact diamonds into hot impact melt additionally 

provides an elegant explanation for the corroded surfaces. 

Calcite 

Calcite possesses a crystal structure with planar CO 3 

groups that are bridged via Ca atoms. The CO 3 groups are 

arranged in layers normal to the c-axis. The bonding between 

Ca cations and CO 3 groups is rather weak, allowing 

this structure to cleave as rhombohedra and to deform by 

mechanical twinning and dislocation glide (Nicolas and 

Poirier 1976). Static deformation experiments have shown 

that calcite can basically develop three types of mechanical 

twins (Barber and Wenk 1979a): e = {011 — 8}, r = {101 — 4} 

and f = {011 — 2} using the hexagonal unit cell setting with 

a = 4.99 Å and c = 17.06 Å. Little is known about the 

shock deformation of calcite but these twin laws apparently 

operate also under low to moderate shock pressures 

(Fig. 12a; Barber and Wenk 1979b, Langenhorst et al. 

2002a). At slightly higher pressures, the microstructure of 

shocked calcite becomes more dominated by dislocations 

(Fig. 12b; Barber and Wenk 1976, Langenhorst et al. 

2002a), occurring in a high density of up to 10 14 m -2 . 

X-ray line broadening observed for naturally shocked calcite 

might be due to such high dislocation densities (Skála 

and Jakeš 1999). 

Effects in the high-pressure regime are of fundamental 

importance, as calcite is known to decompose into solid 

CaO and gaseous CO 2 . The devolatilization of carbonates 

by impacts is considered to be important for the evolution 

of Earth’s atmosphere, climate, and life (Silver and 

Schultz 1980, Crutzen 1987). For example, the Chicxulub 

impact possibly perturbed the atmosphere with large 

amounts of CO 2 and SO x , changing its radiative balance 

and causing acid rain (Emiliani et al. 1981, Prinn and Fegley 

1987). This may have played an important role in the 

275


3 µm 1 µm 

a 

b 

Fig. 12. Shock effects in calcite: (a) bright-field TEM image of crossing deformation twins in calcite shocked to 85 GPa using a high-explosive setup; 

(b) dark-field TEM image of numerous dislocations in a laser-shocked calcite (see Langenhorst et al. 2002a). 

mass extinction scenario at the Cretaceous-Tertiary 

boundary (Alvarez et al. 1980). 

Numerous shock experiments have recently been conducted 

to address the question of shock-induced devolatilization 

(Martinez et al. 1995, Skála et al. 2001, 

Ivanov et al. 2002, Langenhorst et al. 2002a). The experiments 

indicate that strongly shocked calcite can melt at 

high pressure but degassing probably takes only place 

after decompression if the post-shock temperatures are 

sufficiently high. This result is strengthened by theoretical 

calculations of the phase diagram of calcite (Ivanov and 

Deutsch 2002) and the discovery of quenched carbonate 

melts in suevites of the Ries and Chicxulub craters (Graup 

1999, Jones et al. 2000). Furthermore, experiments indicate 

that back reactions between the decomposition products, 

CO 2 and CaO, are fast and efficient and therefore 

have to be taken into account in any quantification of impact-released 

CO 2 (Agrinier et al. 2001). To avoid the back 

reaction it is necessary to spatially separate the decomposition 

products, which is another complexity in a natural 

environment. 

Shock-melted calcite develops upon quenching a 

feathery texture (Jones et al. 2000). Under the optical microscope, 

the feathery texture appears as aggregates of radiating, 

elongated calcite crystals. At the TEM scale, 

quench crystals of calcite are usually devoid of lattice defects, 

indicating that they went through the liquid state 

(Langenhorst et al. 2002a). On the other hand, incipient 

decomposition in shocked calcite is manifested by the 

presence of numerous tiny dislocation loops, indicative for 

the mobilization of CO 2 (Langenhorst et al. 2002a). 

Estimation of shock conditions 

A prerequisite for deciphering the shock-metamorphic 

history of minerals and their host rocks is an assessment of 

their pressure-temperature-time paths. It is important to remember 

here that polymict impact breccias consist of rock 

and mineral fragments that have suffered different degrees 

of shock metamorphism. Therefore it is necessary to consider 

each fragment separately. 

The estimation of the shock duration is rather difficult, 

since no mineralogical speedometer is available. However, 

one can obtain a rough idea of the shock pulse if the size 

of the projectile is known from e.g. crater size. The shock 

pulse corresponds substantially to the time that the shock 

wave needs to propagate to the rear surface of the projectile 

and back to the point of impact. 

The estimation of temperatures is a difficult task as 

well, because it is influenced by a number of factors such 

as porosity. In the context of shock metamorphism, one 

can, in principle, distinguish three different temperatures: 

(1) pre-shock, (2) shock and (3) post-shock temperatures. 

The pre-shock temperature is not directly related to shock 

metamorphism but elevated pre-shock temperatures have 

the effect to lower the threshold pressure for certain shockmetamorphic 

effects (Huffman et al. 1989, Langenhorst et 

al. 1992). It can be elevated due to regional metamorphism 

at the time of impact and the temperature could be determined 

via classical petrologic concepts (mineral assemblages, 

element partitioning between coexisting minerals). 

It is also known that PDF orientations in quartz significantly 

change at elevated pre-shock temperature, when 

quartz is shocked in the β high-temperature structure 

(> 573°C; Langenhorst and Deutsch 1994, Grieve et al. 

1996). Quartz experimentally shocked in the β stability 

field contains PDFs parallel to planes of hexagonal pyramids, 

whereas shocked α-quartz exhibits PDFs that only 

belong to one rhombohedron (for further details see Langenhorst 

and Deutsch 1994). 

Shock and post-shock temperatures are directly related 

to the magnitude of shock compression. The shock temperature 

is the maximum temperature achieved during 

shock compression, whereas the post-shock temperature is 

the temperature prevailing directly after decompression. 

276


For compact materials, the pressure-dependence of shock 

and post-shock temperatures is fairly well known through 

calculations and pyrospectrometric measurements (e.g., 

Wackerle 1962, Raikes and Ahrens 1969, Martinez et al. 

1995, Holland and Ahrens 1997; Fig. 5). Therefore, temperatures 

can be estimated from these reference data, if the 

shock pressure has been determined by applying a shock 

barometer. 

It is more difficult to assess the shock and post-shock 

temperatures in porous rocks/minerals, because porosity 

has the general effect to significantly increase both temperatures. 

Shock compression leads however to compaction 

of rocks and loss of pore space, making it difficult 

to assess the initial porosity of rocks. Therefore, one can 

only apply simple arguments to estimate the temperatures 

from observations. For example, in case of mineral melting, 

the post-shock temperature has certainly exceeded the 

melting point of the mineral at ambient pressure. 

The estimation of shock pressure relies on calibration 

data obtained in well-controlled shock experiments. Shock 

experiments have provided quantitative information on (1) 

co(existence) of shock effects in certain pressure ranges, 

(2) changes in physical (bulk) properties (e.g. refractive 

index and density), (3) variations in PDF orientations (Fig. 

13), and (4) degree of mosaicism (Stöffler 1972, Stöffler et 

al. 1988, Hanns et al. 1978, Stöffler and Langenhorst 

1994, Grieve et al. 1996, Langenhorst and Deutsch 1998). 

To obtain a rough estimate of the shock pressure, it is 

often enough to use the petrographic microscope and to 

observe the shock effects in coexisting minerals. Data on 

the (co)existence of pressure-specific shock effects is 

available for most rock-forming minerals; compilations 

can be found in (Fig. 5): Stöffler (1972), Stöffler et al. 

(1988) and (1991), Bischoff and Stöffler (1992), Langenhorst 

and Deutsch (1998). 

Better constraints on shock pressures can be obtained 

if the refractive indices or densities of quartz or feldspars 

were measured, using a spindle stage and a density gradient 

column (Fig. 13; Medenbach 1985, Langenhorst 

1993). The change in these physical properties reflects the 

gradual conversion of quartz or feldspars into diaplectic 

glasses. This technique applies only over a small pressure 

range and requires that the glass is unaffected by postshock 

annealing or alteration. 

A more robust and widely used technique is to determine 

PDF orientations in quartz with an universal-stage 

(see appendix and Fig. 13, Stöffler and Langenhorst 1994, 

Grieve et al. 1996). Statistical data on PDF orientations in 

quartz have been provided, as a function of pressure, by 

various U-stage studies (Hörz 1968, Müller and Defourneaux 

1969, Robertson and Grieve 1977, Langenhorst 

and Deutsch 1994). These measurements reveal, for example, 

that, at pressures of about 20 GPa, PDFs are predominantly 

oriented parallel to {101 — 3} planes but change at 

higher pressure to {101 — 

2} orientations. A practical description 

on how to identify PDF orientations with an U- 

stage is given in the appendix of this paper. 

Mean refractive index 

1.55 

1.50 

1.45 

PDFs 

(1013) 

(1011) 

(1122) 

PDFs 

(1012) 

(1013) 

(1011) 

(1122) 

X-ray diffractometer studies on shocked rock-forming 

minerals reveal an increasing degree of mosaicism with increasing 

shock pressure as expressed by decreasing diffraction 

peak amplitude and pronounced line-broadening 

(Hörz and Quaide 1973, Hanns et al. 1978). Although this 

technique is capable of yielding relatively accurate pressure 

determinations, it has never been widely applied because 

it requires fresh, unaltered samples. 

Concluding remarks 

transformation regime 

PDFs 

(1012) 

(1013) 

20 25 30 35 

Shock pressure (GPa) 

diaplectic 

glass 

Fig. 13. Refractive index, density and combinations of PDF orientations 

of experimentally shocked quartz as function of pressure (from Langenhorst 

and Deutsch 1994). 

In the past decade, we have made considerable 

progress in understanding the nature and formation mechanisms 

of shock effects in minerals. These advances are, to 

a large extent, due to TEM studies of shocked minerals, 

providing a thorough characterization of the defect microstructures. 

Minerals show a unique response when subjected to 

strong shock waves. The effects range from deformation 

and transformation phenomena to decomposition, melting 

and vaporisation. Which shock effects are activated at 

what pressures and temperatures depends on the crystal 

structure and chemical composition of the mineral studied. 

Most shock effects have been reproduced in short laboratory 

experiments with various designs, although the 

shock pulses in experiments can be more than 6 orders of 

magnitude shorter than those prevailing in nature. Experimental 

limits are however reached if more time-consuming 

processes are simulated, such as reconstructive phase 

transformations to high-pressure silicates. On the other 

hand, the experimental simulation is very successful in reproducing 

deformation defects in shocked minerals. Furthermore, 

experiments led also to a better understanding of 

post-shock modifications of mineralogical shock effects in 

the natural environment. Finally, the shock signature of 

minerals is not only an unequivocal indicator for hypervelocity 

impact but can also be used to obtain quantitative 

constraints on the pressures and temperatures in natural 

impacts. 

2.6 

2.4 

2.2 

Density (g/cm 2 ) 

277


Acknowledgements. I am grateful for the financial support 

provided by the Deutsche Forschungsgemeinschaft, which funded part of 

this work (grants DE 401/15, HO 1446/3, LA 830/4). I am also indebted 

to many colleagues who provided samples or collaborated with me on the 

subject of shock metamorphism: A. Bischoff, M. Boustie, A. Deutsch, 

J.-C. Doukhan, H. Dypvik, U. Hornemann, B. Ivanov, V. L. Masaitis, 

J.-P. Poirier, G. Shafranovsky, and D. Stöffler. I also wish to thank 

F. Hörz, H. Leroux, and R. Skala for their constructive reviews and 

J. Hopf for technical assistance. 

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The following is a simple description on how to determine the crystallographic 

orientation of PDFs in shocked quartz grains, using an universal 

stage (von Federov 1896, Reinhard 1931, Nikitin 1936, Emmons 

1943, Phillips 1971). Although the use of universal stages has considerably 

declined in structural geology, it remains the only method, which 

provides statistical data on the orientation of PDFs. Other techniques 

such as TEM are also capable of measuring PDF orientations but TEM 

measurements are too time-consuming to obtain statistically meaningful 

data. 

There are four fundamental practical steps in the determination of 

PDF orientations in shocked quartz. One has to locate the spatial orientation 

of (1) the optic axis (parallel to c axis) and (2) the normals to the 

PDF planes in the grain studied. Once these directions are known from 

measurements, they are plotted and transformed in a stereographic Wulff 

net (3) and can then be indexed by comparison with the standard stereographic 

projection of quartz (4). 

(1) Since quartz is an uniaxial (trigonal) mineral it is only possible to locate 

the c axis optically. The a axes are in the plane perpendicular to 

c but their exact positions within this plane are not measurable. This 

causes a problem for indexing a PDF plane because the angle to the 

c axis can be measured but not the angle to the a axes. To circumvent 

this problem in part one needs at least two or more crossing PDF sets 

in a single quartz grain. To explain on how to locate the c axis of 

quartz, we will follow here the Reinhard notation of rotation axes for 

a 4-axis or 5-axis universal stage (Reinhard 1931, Fig. 14). Depending 

on the orientation of the quartz grain studied, the optic axis can 

be brought either parallel to the M (microscope) axis or parallel to 

the K (control) axis (E-W). Hence the first step is to find out whether 

the optic axis of the quartz grain is highly inclined to (polar position) 

or lies almost within the plane (equatorial position) of the thin section. 

Polar position: Rotate about the M axis to the diagonal position (45° 

to E-W) and then rotate about the H axis until the grain extincts under 

crossed Nicols. If it remains light, the grain is in the equatorial position 

APPENDIX: How to determine PDF orientations ? 

(see below). Rotate about the M axis to restore the 0° position, but keep 

H at its inclined position. The grain will become light again (unless the 

optic axis is already parallel to M). Now rotate about the N axis to 

achieve again the extinction position. Finally, repeat all the steps until the 

grain remains extinct when rotated about the M axis. 

Equatorial position: Rotate first about the N axis until the quartz 

grain extincts and its optic axis is in the E-W plane. This is the case if 

the grain becomes light by a further rotation about the K axis. Rotate 

then about the H axis until the extinction position of the grain is restored. 

The grain will again become light by a further rotation about the 

K axis in opposite direction (unless the optic axis is already parallel to 

E-W). Finally, repeat all the steps until the grain remains extinct when 

tilted about the K axis. The equatorial position is the more common situation 

in PDF measurements, because PDF poles mostly form a small 

angle to the c axis. 

(2) To determine the orientation of a PDF plane it is necessary to bring 

its normal parallel to E-W. This is achieved by rotating about the N 

axis until the trace of the PDF plane is parallel to N-S. Rotate then 

about the H axis until the trace becomes as sharp as possible. One 

can test whether the plane is exactly vertical by defocusing the PDF. 

It is vertical if the trace of the PDF does not move. Another way to 

test for the vertical position is to tilt about the K axis. If the PDF 

plane is exactly vertical, its trace will remain in N-S orientation. 

(3) The two important rotation angles to be plotted in the stereonet are the 

azimuth angle (i.e. rotation about N) and the ρ angle (i.e. rotation 

about H). When plotting the values for the optic axis, it has to be remembered 

whether it was tilted vertically or horizontally. Additionally, 

one needs to keep in mind the sense of tilting about H. Detailed 

descriptions on the plotting of U-stage data can be found in Bloss 

(1961) and Phillips (1971). Once the orientation of the optic axis and 

the normals to PDF planes are plotted, they need to be transformed 

into the standard stereographic projection with the c axis in the center 

(Fig. 15). The pole of the optic axis is transformed along the 

equatorial line of the Wulff net toward the center, and the normals to 

281


M 

A 

West 

K 

H 

K 

N 

K 

East 

1 (0001) 

2 {1013} 

3 {1012} 

4 {1011} 

5 {1010} 

6 {1122} 

7 {1121} 

8 {2131} 

9 {5161} 

10 {1120} 

a 3 

a 2 

left 

negative 

Fig. 14. Sketch of the rotation axes on a 5-axis universal stage. The notation 

of axes is according to Reinhard (1931): M = microscope axis, 

A = auxiliary axis, N = normal axis, H = horizontal axis, and K = control 

axis. 

a 1 

right 

positive 

Fig. 16. Standard stereographic projection of quartz with the c axis in the 

center. The circles have a 5° radius and mark the positions of the most 

abundant PDF orientations (modified after Stöffler and Langenhorst 

1994). Since quartz is trigonal, one can distinguish positive and negative 

(e.g., rhombohedra) as well as right and left (e.g. pyramids) forms. 

FN' 

ρ 

∆ 

FN' 

' – transformed positions 

OA – optic axis 

FN – face normal 

∆ – azimuth difference 

ρ – pole distance 

ϕ – azimuth angle 

ρ OA 

OA' 

FN 

ρ FN' 

ρ OA 

ϕ FN 

– ϕ OA 

OA 

a 3 

a 1 

(1012) 

(1013) 

(1011) 

(2112) 

a 2 

{1013} 

{1011} 

{0112} 

{1122} 

120° 

Fig. 17. Example for PDF orientations in quartz experimentally shocked 

to 25 GPa (from Langenhorst and Deutsch 1994). The right-hand 

diagram depicts schematically the full symmetry of coexisting forms. 

Fig. 15. An example for the transformation of a PDF pole (FN) and the 

optic axis (OA) into the standard stereographic projection of quartz with 

the optic axis (=c axis) in the center. 

the PDF planes are transformed along small circles by the same angle 

(see also von Engelhardt and Bertsch 1969). You can read now 

the angle between the PDF normals and the c axis. Alternatively, one 

can calculate this angle by the following equation: 

cos ρ FN’ = cos ρ FN cos ρ OA + sin ρ FN sin ρ OA cos (ϕ FN – ϕ OA) (5) 

In the literature, one will often find histograms in which this angle 

is plotted as function of the frequency of PDFs. However, as the a-axes 

cannot be located by polarizing microscopy, the angle to the c axis alone 

is insufficient for unequivocal indexing of PDF planes. This problem 

cannot be solved if the quartz grain of interest contains only one set of 

PDFs. However, if the grain contains at least two sets of PDFs, the next 

analytical step probably yields a reliable indexing result. 

(4) In this final step, the transformed stereoplot is compared to the standard 

stereographic projection of quartz, displaying the orientation of 

known PDF planes (Fig. 16). The latter are drawn as 5° circles, representing 

the estimated errors in the measurements. Table 1 shows 

that not only the pole distance varies for the crystallographic PDF 

planes but also the azimuth angles, e.g. rhombedral forms lie 30° off 

the pyramidal forms because they do not belong to the same crystallographic 

zone. This is the basis for indexing the PDF planes. In 

practice, the stereographic projection of measured PDF poles is rotated 

until all poles fall into the circles of the standard stereographic 

projection (Fig. 16). Usually, there should be only one indexing 

solution; if the measurements are not precise enough, some PDFs 

can remain unindexed. The stereographic projection yields not only 

the Miller indices of PDF planes but also provides information on 

the coexistence of positive and negative or right and left forms (Langenhorst 

and Deutsch 1994). Figure 17 shows an example for typical 

PDF combinations in quartz shocked at a pressure of 25 GPa. It 

demonstrates that the positive rhombohedra {101 — 1} and {101 — 3} are 

combined with the negative rhombohedron {01 — 12} and the right 

pyramid {112 — 2}. It is of course not possible to distinguish between 

a negative and a positive form but when indexing a stereogram one 

has to make an arbitrary choice for the first PDF pole (positive or 

negative) and can then consistently index the following poles. 

282

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