Università di Pisa
Dipartimento di Scienze della Terra
Scuola di Dottorato in Scienze di Base “Galileo Galile
i”
Programma in Scienze della Terra XXVII Ciclo
SSD GEO/07
S
HOCK
M
ETAMORPHISM
AND
I
MPACT
M
ELTING
AT
K
AMIL
C
RATER
, E
GYPT
Anno Accademico 2013-2014
Advisor
Prof. Massimo D’Orazio
Co-advisor
Dott. Luigi Folco
PhD Student
Agnese Fazio
Ricorda: “Quando stai per rinunciare, quando senti che la vita è stata troppo dura con te, ricordati chi sei. Ricorda il tuo sogno”. (Il Delfino - S. Bambarén)
T
ABLE OFC
ONTENTS ABSTRACT 7 RIASSUNTO 9 PREFACE 11 1. INTRODUCTION 131.1. IMPACT CRATERING AS A TERRESTRIAL GEOLOGICAL PROCESS 13
1.2. IMPACT CRATERING STAGES 17
1.3. SHOCK METAMORPHISM 21
1.3.1. Quartz 24
1.3.2. Deformation in other minerals 28
1.3.3. Selective and localized melting 29
1.4. IMPACT MELTING 31
1.5. SHOCK EFFECTS IN QUARTZ-BEARING ROCKS: CRYSTALLINE VS. SEDIMENTARY TARGETS 34
1.6. REFERENCES 37
2. SHOCK METAMORPHISM AND IMPACT MELTING IN SMALL IMPACT CRATERS ON EARTH: EVIDENCE FROM KAMIL CRATER, EGYPT 41
3. TARGET-PROJECTILE INTERACTION DURING IMPACT MELTING AT KAMIL CRATER,
EGYPT 89
4. MICROSCOPIC IMPACTOR DEBRIS IN THE SOIL AROUND KAMIL CRATER (EGYPT): INVENTORY, DISTRIBUTION, TOTAL MASS AND IMPLICATIONS FOR THE IMPACT
SCENARIO 131
5. CONCLUSIONS 161
6. FUTURE WORK 165
6.1. COMBINED MICRO-RAMAN AND TEM STUDY OF HIGH-PRESSURE PHASES FROM KAMIL CRATER
(EGYPT): IMPLICATIONS FOR THEIR FORMATION IN SMALL IMPACT CRATERS ON EARTH 165
6.2. LIQUID IMMISCIBILITY FEATURES IN IMPACT MELTS 165
6.3. REFERENCES 166
APPENDIX I. USE OF THE UNIVERSAL STAGE (U-STAGE) FOR INDEXING PLANAR
DEFORMATION FEATURES IN QUARTZ 169
APPENDIX II. THE EXTREMELY REDUCED SILICATE-BEARING IRON METEORITE NORTHWEST AFRICA 6583: IMPLICATIONS ON THE VARIETY OF THE IMPACT MELT
ROCKS OF THE IAB-COMPLEX PARENT BODY 175
APPENDIX III. OTHER ACTIVITIES 207
ACKNOWLEDGMENTS 209
A
BSTRACTShock effects in small terrestrial impact craters (diameter < 300 m) have been poorly studied because small craters are rare and often deeply eroded. Kamil is a young (< 5000 yr), small (45-m-in-diameter), and well preserved impact structure caused by the hypervelocity impact of the iron meteorite Gebel Kamil on sedimentary rocks in southwestern Egypt. Its pristine state of preservation makes Kamil a natural laboratory for the study of the cratering process of small impactors (about 1-m-in-diameter) on Earth, their consequences, and their impact on the terrestrial environment for hazard assessment.
This PhD Thesis deals with the definition of the shock metamorphism and impact melting in small terrestrial impact craters through a comprehensive mineralogical, petrographic, and geochemical study of shock-metamorphosed rocks and impact melts from Kamil. This study also allows us to constrain the impact cratering process related to the impact of meter-sized iron meteorites on Earth.
The results of this PhD Thesis highlight for the first time that a meter size iron body impacting on a sedimentary target can produce a wide range of shock features. These divide into two categories as a function of their abundance at the thin section scale: i) pervasive shock features (the most abundant), including fracturing, planar deformation features, and impact melt lapilli and bombs, and ii) localized shock features including high-pressure phases and localized impact melting in the form of intergranular melt, melt veins, and melt films in shatter cones. Pervasive shock features indicate the shock pressure suffered by rocks. The most shocked samples (impact melt lapilli and bombs) indicate that the shock pressure at the contact point between the projectile and the target was between 30 and 60 GPa. Based on the planar impact approximation model, this implies that the impact velocity of Gebel Kamil was at least 5 km s-1, for an impact angle of 45°. Localized shock features formed from the local enhancement of shock pressure and temperature at pores and/or at the heterogeneities of the target rocks. Thus, it is possible to find high-pressure phases and intergranular melting in sample that suffered low or moderate shock pressures.
In small meteorite impacts, the projectile may survive the impact through fragmentation. In addition, it may melt and interact with both shocked and melted target rocks. The interaction between target and projectile liquids is a process yet to be completely understood. Impact melt lapilli and bombs from Kamil are very fresh and their study can help constrain the target-projectile interaction. Two types of glasses constitute the impact melt
8 lapilli and bombs: a white glass and a dark glass. The white glass is inclusion-free, mostly
SiO2, and has negligible amounts of Ni and Co, suggesting derivation from the target rocks
with negligible interaction with the projectile liquid (<0.1 wt% of projectile contamination). The dark glass is made of a silicate glass with variable amounts of Al, Fe, and Ni. It also includes variously shocked and melted fragments from the target and projectile (Ni-Fe metal blebs). All this indicates an extensive interaction with the projectile liquid. The dark glass is thus a mixture of target and projectile (estimated projectile contamination 11-12 wt%) liquids. Based on the recently proposed models for the target-projectile interaction and for impact glass formation, we propose a model for the glass formation at Kamil. Between the contact and compression stage and the excavation stage, projectile and target liquids can chemically interact in a restricted zone. The projectile contamination affected only a shallow portion of the impacted target rocks. White glass formed out of this zone, escaping interaction with the projectile. During the excavation stage, due to a brief and chaotic time sequence and the high temperature, dark glass engulfed and coated white glass and target fragments and stuck on iron meteorite shrapnel fragments.
The microscopic impactor debris, systematically collected from the soil around Kamil, includes vesicular masses, spherules, and coatings of dark impact melt glass that is a mixture of impactor and target materials (Si, Fe, Al-rich glass), and Fe-Ni oxide spherules and mini shrapnel fragments. As a consequence of an oblique impact, this material formed a downrange ejecta curtain of microscopic impactor debris due SE-SW of the crater (extension ~300,000 m2, up to ~400 m from the crater), consistent with previous determination of the impactor trajectory. The Ni contents of the soil provided an estimate of the mass of the microscopic debris of the Gebel Kamil meteorite dispersed in the soil. This mass (<290 kg) is a small fraction of the total impactor mass (~10 t) in the form of macroscopic shrapnel. Kamil Crater was generated by a relative small impactor that is consistent with literature estimates of its pre-atmospheric mass (>20 t, likely 50-60 t).
R
IASSUNTOGli effetti di shock registrati in piccoli crateri di impatto terrestri (diametro < 300 m) sono stati poco studiati perché i piccoli crateri sono rari e spesso profondamente erosi. Kamil è una struttura di impatto giovane (<5000 anni), di piccole dimensioni (45 m di diametro) e ben preservata. È stata prodotta dall’impatto iperveloce della meteorite metallica Gebel Kamil su rocce sedimentarie dell’Egitto sudoccidentale. Il suo ottimo stato di preservazione permette di considerare Kamil un laboratorio naturale per studiare il processo di craterizzazione di piccoli impattori (circa 1 m di diametro) sulla Terra, le loro conseguenze e il loro impatto sull’ambiente terrestre per la valutazione del rischio.
La presente Tesi di Dottorato contribuisce alla definizione del metamorfismo da shock e della fusione da impatto in piccoli crateri terrestri attraverso uno studio comprensivo di tipo mineralogico, petrografico e geochimico delle impattiti e dei vetri da impatto di Kamil. Questo studio permette inoltre di ampliare le conoscenze sul processo di craterizzazione legato all’impatto di meteoriti metalliche di dimensioni metriche sulla Terra.
I risultati di questa Tesi di Dottorato evidenziano per la prima volta che un corpo metallico di dimensioni metriche può produrre una vasta gamma di strutture e associazioni mineralogiche da shock impattando su un target sedimentario. Queste sono state divise in due categorie, in funzione della loro abbondanza a scala della sezione sottile: i) effetti di shock pervasivi (più abbondanti), comprendenti fratturazione irregolare, piani di materiale amorfo orientati parallelamente agli indici cristallografici del quarzo o Planar Deformation Features (PDFs), lapilli e bombe di vetro da impatto; ii) effetti di shock localizzati (meno abbondanti) comprendenti fasi di alta pressione e fusione localizzata in forma di vetro intergranulare, vene e film di vetro su shatter cones (strutture coniche caratterizzate da strie disposte a coda di cavallo). Gli effetti di shock pervasivi indicano la pressione subita dalla roccia. I campioni più shockati (lapilli e bombe di vetro) indicano che la pressione al punto di contatto tra la meteorite e il proiettile era tra i 30 e 60 GPa. Sulla base del modello dell’approssimazione planare di impatto, la velocità minima di impatto della meteorite Gebel Kamil era ~5 km s-1, assumendo un angolo di impatto di 45°. Gli effetti di shock localizzati si sono formati come conseguenza di un aumento della pressione e temperatura di shock in corrispondenza dei pori e/o di eterogeneità delle rocce del target. Per questo è possibile trovare fasi di alta pressione e vetro intergranulare in campioni che hanno subito basse o moderate pressioni di shock.
10 In piccoli impatti di meteoriti, il proiettile può sopravvivere all’impatto,
frammentandosi e fondendo. Il suo fuso potrà interagire con rocce del target shockate e fuse. L’interazione tra i fusi del proiettile e del target è un processo ancora non pienamente compreso. I lapilli e le bombe di vetro da impatto di Kamil sono molto freschi e il loro studio può aiutare a vincolare il processo di interazione tra target e proiettile. Due tipi di vetri costituiscono i lapilli e le bombe di vetro da impatto: un vetro bianco e un vetro scuro. Il vetro bianco è privo di inclusioni, è principalmente costituito da SiO2 e ha contenuti trascurabili di
Ni e Co, suggerendo una derivazione dalle rocce target con un’interazione trascurabile con il proiettile (<0.1 wt% di contaminazione del proiettile). Il vetro scuro è costituito da un vetro silicatico con contenuti variabili di Al, Fe e Ni. Il vetro scuro è così una mistura dei liquidi del target e proiettile (contaminazione stimata del proiettile: 11-12 wt%). Il vetro scuro inoltre include frammenti fusi e variamente shockati del target e sferule metalliche di Ni e Fe. Tutto questo indica un’estensiva interazione con il proiettile. Sulla base dei recenti modelli sull’interazione tra target e proiettile e sulla formazione dei vetri da impatto, noi proponiamo un modello per la formazione dei vetri a Kamil. Tra lo stadio di contatto e compressione e lo stadio di escavazione, i fusi del proiettile e del target possono interagire chimicamente in una zona ristretta. La contaminazione del proiettile riguarda solo la porzione più superficiale delle rocce coinvolte dall’impatto. Il vetro bianco si forma fuori da questa zona, senza interagire con il liquido del proiettile. Durante lo stadio di escavazione, a causa della breve e caotica sequenza di eventi e delle alte temperature, il vetro scuro può inglobare e avvolgere il vetro chiaro e frammenti del target e può attaccarsi ai frammenti di meteorite.
Il detrito microscopico dell’impattore, sistematicamente campionato dal suolo circostante Kamil, comprende masse di vetro vescicolare, sferule, e pellicole di vetro scuro (vetro ricco di Si, Fe e Al, analogo al vetro scuro di lapilli e delle bombe) su frammenti di target e di vetro chiaro, sferule di ossidi di Fe e Ni, e piccoli frammenti di meteorite. Come conseguenza di un impatto obliquo questo materiale ha formato una coltre di ejecta verso SE-SW del cratere (estensione ~300,000 m2, fino a ~400 m dal cratere), questo è consistente con
gli studi precedenti sulla traiettoria di impatto. Il contenuto Ni del suolo fornisce una stima della massa del detrito microscopico del meteorite Gebel Kamil disperso nel suolo. Questa massa (<290 kg) è una piccola frazione della massa totale dell’impattore (~10 t) in forma di frammenti macroscopici. Il cratere Kamil è stato generato da un impattore relativamente piccolo, questo è consistente con le stime della sua massa pre-atmosferica (>20 t, probabilmente 50-60 t) riportate in letteratura.
P
REFACEThe aim of this PhD Thesis is the definition of the shock metamorphism and impact melting in small impact craters on Earth through the study of the impactites (rocks produced by hypervelocity impacts) found at Kamil Crater (Egypt). Kamil is a small impact crater (45-m-diameter) discovered in 2008. It was produced by the hypervelocity impact of the iron meteorite Gebel Kamil most likely < 5000 yr ago. This crater is very well preserved and is located in a very simple geological setting.
This PhD Thesis is organized into six chapters and three appendices. Chapter 1 is a brief digression on all the main aspects concerning impact cratering. In this chapter the answers to the following questions can be found: How many impact craters are there on Earth? Where are they located? How large are they? Why is impact cratering a unique process? How does an impact crater form? Which are the shock effects produced by impacts? Can the target rocks influence the shock effects? This chapter has the scope to support the understanding of the following three chapters. They are structured as scientific papers and they are accepted or submitted to international journals.
Chapter 2 is the paper entitled “Shock metamorphism and impact melting in small
impact craters on Earth: evidence from Kamil Crater, Egypt”, by Agnese Fazio, Luigi Folco,
Massimo D’Orazio, Maria Luce Frezzotti, and Carole Cordier. It is a compendium of all the shock effects that we found studying rocks from Kamil. The aim of this paper is to establish the impact velocity of the meteorite Gebel Kamil on the basis of shock effects. Moreover, this paper highlights that, due to the wide range and the freshness of the shock features, Kamil can be considered a natural laboratory for studying impact cratering and shock deformation processes in small impact structures. This paper was submitted to the international peer-reviewed journal Meteoritics & Planetary Science on 20th March 2014, accepted on 15th August 2014, and published on the December issue (vol. 49, pp. 2175-2200).
Chapter 3, is the paper entitled “Target-projectile interaction during impact melting
at Kamil Crater, Egypt”, by Agnese Fazio, Massimo D’Orazio, Carole Cordier, and Luigi
Folco. It is a detailed petrographic and geochemical study of the impact melt lapilli and bombs from Kamil. Impact melt masses are the most significant samples for the study of the process of the interaction between the target and the projectile liquids during the formation of Kamil, helping to constrain the general impact scenario. This paper will be submitted to an international peer-reviewed journal.
12
Chapter 4, is a paper entitled “Microscopic impactor debris in the soil around Kamil
Crater (Egypt): inventory, distribution, total mass and implications for the impact scenario”,
by Luigi Folco, Massimo D’Orazio, Agnese Fazio, Carole Cordier, Antonio Zeoli, Matthias van Ginneken, and Ahmed El-Barkooky. It is a report on the microscopic impactor debris collected around the crater. Beyond presenting an inventory of all microscopic particle types, this study allows to constrain the impactor mass at the time of the impact. This paper was submitted to the international peer-reviewed journal Meteoritics & Planetary Science on 18th
August 2014 and it was accepted on 11th December 2014.
The conclusions of this PhD Thesis are reported in Chapter 5. Finally, future activities and projects related to the present PhD topic are discussed in Chapter 6.
Appendix I provides detailed instructions for the use of the Universal Stage (U-Stage)
in indexing planar deformation features in shocked quartz. The U-Stage is considered an outdated technique by the most part of modern geologists. However, in the last years it is back to the top among planetary geologists as a low-cost and widespread technique for the indexing of planar deformation features in quartz.
Appendix II reports of the paper entitled “The extremely reduced silicate-bearing iron
meteorite Northwest Africa 6583: implications on the variety of the impact melt rocks of the IAB-complex parent body” by Agnese Fazio, Massimo D’Orazio, Luigi Folco, Jérôme
Gattacceca, Corinne Sonzogni). I started the study of this meteorite for my Master Thesis (Tesi di Laurea Magistrale), and I concluded it during the first months of my PhD. This paper was published in the international peer-reviewed journal Meteoritics & Planetary Science in the volume 48, number 12 of December 2013.
1. I
NTRODUCTION1.1. Impact cratering as a terrestrial geological process
Impact cratering is the most common geological process of the Solar System. On Solar System bodies, impact cratering plays an important role modeling their surfaces and shapes. Moreover, cosmic impacts can catastrophically destroy solar system bodies or form new ones (e.g., the Moon), change their surface geochemistry, generate abrupt climate changes and mass extinctions, and allow the formation of ore deposits, hydrocarbons reservoirs, and biological niches (Osinski 2008).
Impact craters are the main geological landforms of the solid bodies of the Solar System (Fig. 1) with the exception of the Earth. On Earth, impact structures are not preserved for long because they are continually obliterated by tectonic activity, erosion, burial, weathering, volcanic resurfacing, and vegetation. The first impact crater on Earth was recognized in 1905. It was the Barringer Crater (also known as Meteor; Fig. 2). However, it was only in the late 1960s that impact cratering was recognized as an important geological process not only of the Solar System but also of the Earth. Indeed, in those years, unique petrological and geochemical features produced by the passage of intense shock waves were recognized (see section 1.3. Shock metamorphism) and new high-resolution images of the Solar System bodies were obtained by satellites and modern telescopes (for review see French 1998; Osinski and Pierazzo 2013).
Fig. 1. The surface of Mercury. Credit: Nasa/Johns Hopkins University Applied Physics Laboratory/Carnegie Institution of Washington.
14
Fig. 2. Barringer Crater, Arizona, USA.
Nowadays, 184 impact structures have been recognized on Earth (Fig. 3;
http://www.passc.net/EarthImpactDatabase/index.html.). They range from 13.5 m (Carancas,
Peru) up to 300 km (Vredefort, South Africa). The surfaces of the Solar System solid bodies are characterized by a random distribution of impact craters. However, looking at the distribution of impact craters on Earth, the distribution is not random, but it reflects geological and social factors. This is due to three main reasons: (i) the age of the surfaces and their geological stability show that the best areas for the preservation of impact structures are continental shields and cratonic areas; (ii) the difficulties in searching for impact structure under the sea, due to the high exploration costs and the young age of oceanic crust; (iii) systematic search has been done only on specific continental areas (i.e., Canada, Russia, Ukraine, and Australia).
Fig. 3. Distribution of Earth’s craters. Small craters (< 300 m in diameter) have been highlighted by a larger circle. Modified after http://www.passc.net/EarthImpactDatabase/index.html.
The rate of the impacts follows an exponential trend with impactor size. Small bodies are more abundant and their collisions are more probable. Impacts producing small craters (< 300 m in diameter) on Earth occur on a decadal to a secular time scale; impacts producing craters > 200 km in diameter occur on a billion time scale (e.g., Bland and Artemieva 2006). However, Fig. 4 shows that craters between 0 and 1 km are a quarter of the impact craters between 1 and 10 km, indicating an incomplete record of small impact structures on Earth. In Fig. 3 the 17 small impact craters have been highlighted by a larger circle. Small craters are quickly eroded or buried in sediments and their features are easily lost (Fig. 5). This means that our knowledge about the formation mechanisms and the shock metamorphic features of craters of these proportions is not detailed (see section 1.3. Shock metamorphism). Moreover, this is an important limitation to the knowledge of the hazard that small impactors constitute to human populations.
Fig. 4. Histogram showing the number of impact craters on Earth for each diameter class.
0 10 20 30 40 50 60 70 80 90 100 Numer of Crater 0 - 1 km 1 - 10 km 11-100 km 101-1000 km
16
Fig. 5. Example of three terrestrial small craters (< 300 m in diameter). Source: http://www.passc.net/EarthImpactDatabase/index.html.
Impact cratering is a peculiar geological process. Its differences with respect to the other more familiar geological processes contributed to the initial skepticism for this process. These differences have been summarized in six points by French (1998):
(i) Rarity. In the whole human history, only two impact cratering events were observed by humans (Sikhote-Alin, Russia, in 1947, and Carancas, Peru, in 2007). In both cases, they are small impact craters (Sikhote-Alin formed a crater field and the largest crater is 26 m in diameter; Carancas formed one single crater 13 m in diameter). For this reason, we have only indirect evidence of the regional and global risk that a large meteorite impact can cause.
(ii) Immense energy. The energy of a large impact is not comparable to the energy release by earthquakes or volcanic eruptions. It has been calculated that the formation of the Barringer Crater released an energy of three orders of magnitudes higher than the Hiroshima atomic bomb, and that the impact responsible for the mass extinction of large reptiles (Chicxulub, Mexico, 180 km in diameter) was two orders of magnitude more energetic than the total annual energy release from Earth (heat flow, seismic, and volcanic eruption).
(iii) Instant effects. By contrast to the other geological processes, impact cratering is a virtually instantaneous process. It has been estimated that a crater of 1 km in diameter (hence similar to Barringer Crater) forms in a few seconds, while less than 10 minutes are necessary to form impact structures 200 km in diameter, like Sudbury (Canada) or Vredefort (South Africa). However, post-shock modification
(i.e., geological adjustments, mostly due to gravity) in these very large impact structures can continue for many years.
(iv) Concentrated energy release. Besides being immense and instantaneously released, the energy is also released virtually in a single point of Earth’s surfaces. Then, it is rapidly transferred into target rocks, atmosphere, and biosphere. (v) Extreme physical conditions. The shock waves propagate into the target rocks at
supersonic velocities of about 5–8 km/s. Rocks involved in this passage suffered pressures, temperatures and strain of several orders of magnitude higher than those reached by endogenic metamorphism (for details on shock metamorphism see section 1.3. Shock metamorphism).
(vi) Unique deformation effects. The effects produced by the propagation of shock waves into target rocks cause the formation of unique deformation effects. The occurrence of these unique features is nowadays the most important criterion for establishing the impact origin of a crater. For details on the shock metamorphic futures see section 1.3. Shock metamorphism.
1.2. Impact cratering stages
On Earth, impact craters form when a projectile is large enough (typically > 50 m for stony meteorites and > 20 m for iron meteorites) to pass through the atmosphere without a significant deceleration and disruption, thereby reaching the Earth’s surface at a velocity of > 11 km/s (hypervelocity). Hypervelocity impact starts to form as soon as an extraterrestrial object strikes the ground surface at its original cosmic velocity. Hypervelocity impact produces high-pressure shock waves, that radiate into the target at a velocity of 5–8 km/s. Hypervelocity impact craters are so characterized to produce unique shock metamorphic effects (Section 1.3.).
It is widely accepted that the mechanism of the formation of hypervelocity impact craters can be outlined into three main stages (Melosh 1989; French 1998; Osinski 2008; Osinski and Pierazzo 2013): (i) contact and compression; (ii) excavation; (iii) modification. For very large impact craters a fourth stage could be considered: hydrothermal and chemical alteration.
(i) Contact and compression. The first stage of impact cratering begins when the projectile hits the target (Figs. 6a and 6b). Models suggest that the projectile penetrates a solid target about one or two times its diameter. The impact pressure
18 at the contact point depends on the impact velocity. It can be estimated using the
planar impact approximation equations (Melosh 1989; Melosh 2013), when both the nature of the target and of the projectile are known. Typically, at the contact point shock pressure > 100 GPa.
After the contact with the ground, the kinetic energy of the impact is transferred to the target and to the projectile in form of shock waves. The shock waves travel in two opposite directions, as illustrated by the train model (Fig. 6c). Contrary to the normal compressional waves, shock waves are also characterized by the propagation of a material flow. The velocity of the shock wave and of the material flow are indicated by the symbols Us and Up in Fig. 6. Us is higher than Up. When the shock front of the shock wave travelling into the projectile reaches its free upper surface (Fig. 6d), the shock wave is reflected back initially into the projectile and immediately later into the target as a rarefaction wave (Fig. 6f). The rarefaction wave travels faster than the shock wave, because it travels in compressed material.
Fig. 6. Schematic representations showing the first two stages of impact cratering (contact and compression and excavation; sketch on the right; after Osinski 2008) and the formation and propagation of shock and rarefaction waves (sketch on the left; after Langenhorst and Deutsch (2012)). Us and Up indicate the velocities of the shock
wave and the velocity of the material flow, respectively.
In Fig. 7 a schematic representation of the propagation of the shock wave into the target is shown. From this image it is possible to note that the volume of rocks shocked between 1 and 5 GPa is greater than the volume of rocks shocked between 5 and >50 GPa.
Fig. 7. Initial shock-wave pressures and excavation flow lines around the impact point during the contact and compression stage. Image after French 1998.
The duration of the contact and compression stage is very short: 2 s for a 50-km-diameter projectile travelling object travelling at 25 km s-1, and less than 0.01 s for a 100-m-diameter object travelling at the same speed (French 1998).
(ii) Excavation. The passage from the contact and compression stage to the excavation stage is a continuum. The passage of the rarefaction stage decompresses the rocks causing the complete (only, for large impact events) melting and/or vaporization of the projectile and the acceleration of excavated target rocks in the opposite direction of the ground (Fig. 6e). During this stage the crater opens up, thanks to a complex interaction among shock and rarefaction waves. This interaction produces an “excavation flow-field” and generates a so-called “transient cavity” (Fig. 6f). The projectile plays no roles in the crater opening. The excavated material is ejected ballistically, following the trajectories showed in Fig. 7, to form the ejecta blanket.
Transient cavities with a maximum diameter of 2 km (for sedimentary rocks) and 4 km (for crystalline rocks) do not undergo further modification, more precisely they undergo minimum modification that stabilizes the structure (e.g., landslips). These craters are called simple impact craters (Fig. 8), in contrast to the complex impact craters (see point (iii) Modification). Simple impact craters are described by the following geometrical parameters: D = diameter, dA = apparent depth, tbr = ,
and dT = true depth (dA + tbr). The boundary between simple and complex impact
craters is not fixed, it can vary on Earth according to the lithology of the target material and on other solid planetary bodies because of the different gravitational
20 acceleration (e.g., on the Moon the transition between simple and complex crater
is between 15 and 27 km).
Fig. 8. Schematic representation of the final simple impact crater. Modified after Osinski 2008.
Although the excavation stage is longer than the contact and compression stage, it is still brief by geological standards. For example, a 1-km-diameter transient crater can be excavated in 6 s and a 200-km-diameter crater requires about 90 s (Melosh 1989).
(iii) Modification. The effects of the modification stage are governed by the size of the transient cavity and by the lithology and the features of the target rocks. Shock waves do not play a further part in crater development. For these reasons, the modification stage essentially concerns the complex impact craters. This stage is characterized by the gravitational collapse of the initial transient cavity. It determines an uplift of the crater floor and so the formation of a central peak (Fig. 9). Besides the central peak, a very large impact crater can show a peak ring or a multi-ring morphology. The modification stage does not have a defined end: processes relative to the modification of very large impact structures can involve widespread disturbances in the Earth’s crust and merge into endogenous geological processes. In some cases the modification stage can last for years.
Fig. 9. Schematic representations showing the modification stage for complex impact craters. Images after Osinski 2008.
In about 70 impact structures from 1.8 km (Lonar, India) to 250 km (Sudbury, Canada) in diameter impact-generated hydrothermal systems have been recognized. Due to their diffusion, it seems highly probable that any hypervelocity impact can generate a hydrothermal system, as long as sufficient H2O is present. For this reason, post-impact
hydrothermal activity could be considered as a fourth stage of the impact cratering process.
1.3. Shock metamorphism
Rocks involved in the impact cratering process can vaporize, melt, shock metamorphose, or deform. All rocks affected by one or more hypervelocity impact(s) resulting from collision(s) of planetary bodies are called "impactites" (Stöffler and Grieve, 2007). The classification of impactites follows the recommendation formulated by the IUGS Subcommission on the Systematic of Metamorphic Rocks. It is based on the degree of shock metamorphism, the occurrence of melting, and the geological setting of impactites. Fig. 10 shows the impactite classification suggested by the IUGS Subcommission on the Systematic of Metamorphic Rocks.
Fig. 10. Impactite classification according the IUGS Subcommission on the Systematic of Metamorphic Rocks (modified after Stöffler and Grieve 2007).
Shock metamorphism is a type of metamorphism of rocks and minerals caused by shock wave compression and decompression due to the hypervelocity impact of a solid body or due to the detonation of high-energy chemical or nuclear explosives. Contrary to endogenic metamorphism, shock metamorphism is characterized by pressure > 5 GPa and temperatures
22 up to 10000 °C (Fig. 11; Table 1). These enormous differences in the pressure and
temperature ranges, and the very short duration of the process cause unique irreversible changes to rocks and minerals.
Fig. 11 Pressure and temperature ranges of endogenic and shock metamorphism. Image after French 1998. Table 1. Differences between endogenic and shock metamorphism. Modified after French 1998.
Characteristic Endogenic metamorphism Shock metamorphism
Geological setting Typically to depths of 10-50 km Surface or near-surface regions of Earth’s crust
Pressures Typically <1-3 GPa
100-400 GPa near impact point; 10-60 GPa in large volumes of surrounding rocks
Temperatures Generally ≤ 1000°C
Up to 10000°C near impact point; typically 500-3000°C in
surrounding rocks
Strain rates 10-3-10-6 s-1 104-106 s-1
Time for completion of process From 105-107 yr Instantaneous (few seconds)
Reaction time Slow; minerals closely approach equilibrium
Rapid: abundant quenching and preservation of metastable minerals and glasses
The identification of shock effects is one of the main criteria for identifying impact structures. Other criteria for identifying a new impact structures are geophysical investigation and siderophile elements and isotopic anomalies, however their treatment is beyond the scope of this Thesis.
According to recent literature, the shock metamorphism does not include the impact melting (French 1998; Osinski and Pierazzo 2013). For this reason, in this dissertation shock metamorphism and impact melting features are presented in two different sections.
The most significant shock effect at the meso- to macro-scale is the occurrence of multiple sets of conical striated fractures surfaces. These structures are known as shatter
cones and are uniquely produced by hypervelocity impacts. Hence, their occurrence is a
strong clue for the identification of impact structures. They are best developed in fine-grained massive lithologies (such as limestone), but they can also occur in more coarse-grained lithologies, such as granite, in which, however, they are usually more poorly developed.
Fig. 12. Shatter cones in a very fine-grained carbonate form Steinheim (Germany). Nesting of multiple cones are visible. Image after French and Koeberl 2010.
Shatter cones usually occur in the central uplifts of complex impact structures and in some cases, isolated fragments/clasts of shatter cones have been found in impact breccias, within or outside the crater. Shatter cones are penetrative features; for this reason they are different to other striated features such as wind-abrasion structures. Although shatter cones are widely accepted as unequivocal proof of a meteorite impact crater, their formation mechanism is still debated (e.g., Sagy et al. 2002; Baratoux and Melosh 2003; Wieland et al. 2006). It is generally accepted that shatter cones form at relatively low shock pressures, generally between approximately 2 and 10 GPa, and exceptional up to 30 GPa. The shock pressure of shatter cones is estimated on the basis of the shock features recorded by minerals constituent the rock. Some occurrences of melt film among conical surfaces suggest that the formation of melt could play an important role in the formation mechanism of these structures (Gay 1978; Gibson and Spray 1998; Nicolaysen and Reimold 1999; Pittarello et al. 2011; Fazio et al. 2014).
24 Shock-induced changes in minerals could be classified in four typologies of effects: (i)
deformation, e.g., formation of dislocations, planar microstructures, mechanical twins, kink
bands, and mosaicism; (ii) phase transformation into high-pressure phases and diaplectic glass; (iii) decomposition into a solid residue and a gaseous phase; (iv) melting and
vaporization of entire mineral (Langenhorst 2002).
Shock-induced changes in minerals have been reported for various rock-forming minerals (Fig. 13). However, quartz is the mineral that shows the widest variety of shock effects over a wide range of pressures in a consecutive manner, because of its three-dimensional linked, corner-shared SiO4-tetrahedra crystal structures. Moreover, quartz is the
most common mineral on the Earth’s crust. The combination of these two factors determines that shock effects in quartz are widely studied and well pressure-calibrated (e.g., experiments on single quartz crystal by Langenhorst and Deutsch 1994 and Langenhorst 1994). In the next subsections the shock features occurring in quartz (1.3.1. Quartz) and other common minerals (1.3.2 Deformation in other minerals) from crystalline rocks are presented. The comparison between the shock metamorphic effects in crystalline (non-porous) rocks and sedimentary (porous) rocks is discussed in section 1.5. Shock effects in quartz-bearing rocks: crystalline vs. sedimentary target.
Fig. 13. Shock features in some of the most common rock-forming minerals. Source: http://www.lpi.usra.edu/exploration/training/resources/. Credit Kring D. A. (1989).
1.3.1. Quartz
The initial deformations occurring in quartz are the reduction of refractivity and
birefringence indices, and mosaicism. All these effects are not diagnostic features for the
of the occurrence of several sub-domains with slightly different optical axes. Mosaicism is produced by the distortion of the crystal lattice into small domains that are rotated by low angles against each other. Refractivity and birefringence indices decrease increasing the shock compression, drifting towards the amorphous values.
The most significant shock effect on quartz is the occurrence of planar microstructures. Planar microstructures are divided into planar fractures (PFs) and planar
deformation features (PDFs). Contrary to irregular fractures (ubiquitous in shocked quartz),
planar microstructures are crystallographically controlled: PFs and PDFs are oriented parallel to rational crystallographic planes. Each plane is indicated by its crystallographic index, following the four-digit notation of Miller-Bravais {hkil}.
Planar fractures are open fissures parallel to the rational crystallographic plane (0001) and
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. Planar fractures develop in sets of fractures with spacing of about 20 µm. Each fracture is generally wider than 3 µm. Planar fractures commonly control and/or limit the distribution of adjacent PDF sets; therefore PFs are formed earlier than PDFs. Planar fractures are supposed to form between 5 and 8 GPa. Contrary to PDFs, PFs are not considered as unequivocal evidence of shock, although they are rare and differently oriented in porous sedimentary rocks (see section 1.5).Planar deformation features are considered as one of the best criteria for the identification of new impact structures. Planar deformation features develop in sets parallel to the rational crystallographic plane as PFs, however they are generally thinner and narrower than PFs. Planar deformation features are less than 2 µm thick and between 2 and 10 µm spaced (Fig. 14). Up to ten differently oriented sets of PDFs per crystal have been counted. They form within a wide range of pressures: from 5-10 GPa to 35 GPa (Fig. 13). Planar deformation features are differently oriented according to their pressure of formation.
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may form between ~ 20 and ~ 25 GPa. Over ~ 25 GPa,€
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ones. The orientation of planes could be done by TEM or by universal stage analyses (Langenhorst 2002; Ferrière et al. 2009; Huber et al. 2011).26
Fig. 14. Planar deformation features (PDFs) in a quartz grain from Kamil crater sandstone.
In altered and geologically old impact structures, the original amorphous material in PDF planes is recrystallized back to quartz, and arrays of small (1-2 µm) fluid inclusions develop along the original planes. These PDFs are called decorated PDFs (Stöffler and Langenhorst 1994).
The most accepted model of formation of PDFs was proposed by Goltrant et al. (1992). This model is the result of TEM observations and of the estimated elastic properties of quartz. They observed that the magnitude of shear moduli changes as a function of pressure causing a separation of uncompressed and compressed regions in the crystal. Due to the different lattice dimensions, both crystal regions are incompatible with respect to their lattice parameters. The misfit is compensated by the formation of amorphous lamellae, namely PDFs, behind the shock front by solid-state amorphization. This mechanism operates up to 25 GPa (Fig. 15a). Above this pressure and up to 35 GPa, the shock temperature Th exceeds the
melting temperature of quartz Tm, causing the formation of melt and the subsequent
lengthening of PDFs, whose boundaries become wavy (Fig. 15b). Because the post-shock temperature Tr is below that of the Tm, diaplectic glass forms (glass formed by solid-state
transformation). Between 35 and 50 GPa the entire grain could be transformed into diaplectic glass (Fig. 15c). Above 50 GPa, also the Tr is higher than the Tm, determining the total
melting of the quartz grains, or rather the formation of lechatelierite (pure SiO2 glass; Fig.
15d). This model was proposed by Langenhorst (1994), and new evidence from Kamil impactites seems to confirm it (see Chapter 3 of this Thesis).
Fig. 15 Langenhorst model for the formation of PDFs, diaplectic glass and leshatelierite. Th= shock temperature, Tr= post-shock temperature, and Tm= melting temperature of quartz. Image from Langenhorst
1994.
During the passage of shock waves, quartz can transform into high-pressure phases that are coesite and stishovite. Both coesite and stishovite can also form during endogenic processes, but stishovite is generally very rare, hence their combined occurrence can be considered as a univocal feature of shock metamorphism. The formation pressures of these polymorphs during endogenic processes are about 2 GPa for coesite and around 7-8 GPa for stishovite (Heaney et al. 1994). Pressures of formation of 2-8 GPa are too low for shock metamorphism (few seconds) and are possible because rocks suffer these pressures for a very long time (million years) under static equilibrium conditions. In non-porous crystalline rocks coesite forms between 30 and 60 GPa, instead stishovite between 12 and 45 GPa. The formation pressures of these phases are very different if the rock is non-porous or porous (see Section 1.5). Moreover, the crystallization order is switched compared to the endogenic environment: stishovite generally forms at a lower pressure than coesite. Coesite is thought to form during the passage of the rarefaction wave, namely during the decompression stage (Stöffler and Langenhorst 1994). Both polymorphs were found within diaplectic glass, along
28 grain boundaries or in association with PDFs. High-pressure phases can be detected by X-ray
powder diffraction, µRaman or TEM.
Among the phase transformations, diaplectic glass can also be considered. Diaplectic glass is a glass that formed without melting but by solid-state transformation. Diaplectic glass typically forms from framework minerals, such as quartz and feldspar (diaplectic glass from feldspar is known as maskelynite). Diaplectic glass is characterized to be amorphous, to have a very low refractive index, to preserve the pre-shock shape, morphology and texture of the parent mineral, and being devoid of flow structures and vesicles. The minimum shock formation pressure of diaplectic glass is about 35 GPa for non-porous crystalline rocks. If rocks suffer a thermal annealing at temperatures above 1200°C, diaplectic glass is not preserved. It recrystallizes forming ballen α-cristobalite or ballen α-quartz (Ferrière and Osinski 2013 and reference therein).
Ballen α-cristobalite or ballen α-quartz, together called ballen silica, and toasted
quartz are not primary shock features. They form because of thermal annealing. The
identification of these post-shock thermal features is important for recognizing large, old and eroded impact structures. Ballen silica occurs as more or less spheroidal or rarely ovoid bodies of around a few micrometers up to 200 µm. These bodies can intersect or penetrate each other or abut each other (Ferrière and Osinski, 2013 and references therein). Toasted quartz is characterized to be orange-brown to grayish-reddish brown in color. The formation mechanism of toasted quartz is not clear. It probably results from hydrothermal or other post-shock modification or from the exsolution of water from glass, primarily along PDFs, during heat-driven recrystallization (Ferrière and Osinski, 2013 and reference therein).
1.3.2. Deformation in other minerals
Feldspar crystals, being tectosilicate as quartz, show a wide range of shock metamorphic effects. Similarly to quartz, they develop fracturing, PFs, PDFs, and diaplectic glass (called maskelynite). Planar deformation features have also been observed in olivine, zircon and tourmaline, pyroxene, amphibole, sillimanite, garnet, and apatite (Ferrière and Osinski, 2013 and reference therein). Shocked mica crystals commonly show kink bands, these features also develop by tectonic deformation, so they are not diagnostic criteria for establishing an impact origin of a structure.
Zircon under high shock pressure and temperature decomposes into baddeleyite (ZrO2) and a SiO2 phase. Zircon decomposition occurs at temperatures higher than 1750°C
into a high-pressure polymorph called reidite (about 20 GPa; Ferrière and Osinski, 2013 and reference therein).
Graphite transforms into diamond upon shock metamorphism. Impact diamonds are usually polycrystalline and rich in crystallographic defects. They commonly preserve some features of the precursor graphite. The origin of impact diamond is still debated. One of the widely accepted hypotheses, based on graphite-diamond textural relationship within shocked rocks from Ries Crater (Germany), suggests that diamonds are formed from graphite by shock-induced solid-state phase transformation during shock compression (El Goresy et al. 2001). Diamonds from Ries are supposed to have formed around 30-40 GPa. Shock formation pressures can vary a lot also for diamonds according to the porosity of the rock and the nature of graphite (degree of crystallinity; Ferrière and Osinski, 2013 and reference therein). Besides diamond, graphite can transform into lonsdaleite and other unknown C-phase(s). Lonsdaleite and the other C-phase(s) may represent intermediate phases before diamond formation. One of the best techniques for studying carbon phases is µRaman spectroscopy.
1.3.3. Selective and localized melting
Before whole-rock melting, melting starts in correspondence of rock heterogeneities as pores or boundaries between different mineral grains (interstitial impact melt glass). Selective melting could also occur in single mineral grains (mineral glass).
Interstitial impact melt glasses have compositions that are a mixture of adjacent
minerals. In meteorites these interstitial impact melt glasses are also known as melt pockets.
Mineral glasses have the same composition as the pre-existing minerals. Mineral glass is
different from diaplectic glass, because the latter forms by solid-state transformation and retains the shape of the mineral grain.
Localized melt can also be found in the form of melt veins and/or melt films. Melt veins are irregular veins of quenched melt produced by shock-induced localized melting in moderately to strongly shocked rocks (Stöffler and Grieve 2007).
One particular type of melt veins is represented by pseudotachylite veins. Pseudotachylite veins are specific melt veins produced by frictional melting (pseudotachylite is a genetic name). Frictional melting is controlled by the mechanical properties of a rock’s constituent minerals: mineral with the lowest fracture toughness and breakdown temperature are preferentially involved into frictional melting (Spray 2010; Fig. 16).
30
Fig. 16. Melting or breakdown temperature TMB versus fractures toughness KIC for selected minerals. Image
after Spray 2010.
Pseudotachylite veins vary from m-scale to µm-scale and they are characterized by a clear margin with the host rock, injection veins, fragment inclusions, and high-pressure mineral inclusions. Some authors classify pseudotachylites into two types. S-types (shock-type) pseudotachylite veins are typically thin (< 2 mm) and contain relict mineral grains and high-pressure polymorphs. They usually occur irregularly distributed within the innermost shock zone of the impact structures. E-types (endogenic-type) can be up to hundreds of meters wide, and occurs only in the periphery of the impact structures (Spray 2010; Fig. 17). Pseudotachylite veins can form networks of veins giving rise to “pseudotachylite breccias”.
Fig. 17. Vredefort (South Africa) is the type locality of pseudotachylite veins. The term pseudotachylite was introduced in literature in 1916 by Shand to identify and describe these veins in Vredefort impact structure.
Image after Spray 2010.
Other forms of localized melting in the form of veins and/or films were described at the surfaces of shatter cones from a few localities: Sudbury (Ontario, Canada) and Vredefort (South Africa; e.g., Gay 1978; Gibson and Spray 1998; Nicolaysen and Reimold 1999), Vista Alegre (Brazil; Pittarello et al. 2011), Kamil (Egypt; Fazio et al. 2014, Chapter 2 of this Thesis) and, possibly, from Santa Fe (New Mexico, USA; Fackelman et al. 2008).
1.4. Impact melting
Whole-rock melting is one of the most significant pieces of evidence of hypervelocity impact events (documented at approximately half of the known terrestrial impact structures). Impact melting is different from thermal melting (such as that produced by endogenic magmatic processes), because impact melting is a function of shock pressure, porosity and the compressibility of the target rock lithology and their constituent minerals. Impact melting occurs upon decompression from high shock pressure and temperature.
Impact melting can occur in a variety of settings: i) km-scale sheets and/or isolated bodies within the crater (impact melt rocks with variable amounts of clastic debris of different degree of shock metamorphism); ii) m- to cm-scale irregular and aerodynamically sculptured glassy particles, either within impact breccias inside the crater (e.g., suevite) or in nearby ejecta deposit around the crater in the form of impact melt lapilli and bombs. iii) tens of m- to cm-scale injection dikes and sills in the crater floor and walls; iv) distal cm- to µm-scale
32 discrete particles (spherules, tektites, and microtektites) distributed regionally to globally
(French 1998, Stöffler and Grieve 2007, and Osinski et al. 2013).
The main settings in which impact melt-bearing materials occur in a complex impact structure are shown in Fig. 18. The occurrence and the properties of impact melts provide a lot of information about the crater formation process.
Fig. 18. Main settings in which impact melt-bearing materials occur in a complex impact structure. Image after Osinski et al. 2013.
Besides the geological setting, impact melt-bearing rocks can be classified on the basis of the textural features of the groundmass or matrix and clast content. In Fig. 19 shows the classification scheme by Osinski et al. 2013. Clast-rich impact melts are also known as impact melt breccias. The prefix particulate is suggested if there is evidence that the groundmass remained molten during and after deposition.
Fig. 19 Classification scheme of impact melt-bearing impactites. Image after Osinski et al. 2013.
The impact melt process is essentially the same in sedimentary and crystalline rocks, although shock pressures necessary to melt those rocks are different (see next section 1.5). The mineralogical assemblage and the porosity of sedimentary rocks (e.g., the occurrence of phyllosilicate; Fig. 16) could facilitate the melting of sedimentary rocks respect to crystalline rocks. The consequence of these differences is that the products of melting in sedimentary and crystalline targets appear different. In crystalline targets, impact melts products have igneous structures. In large craters, they occur as large coherent melt bodies, with relatively homogenous compositions. In some cases, impact melt sheets can extend laterally many kilometers and can be several hundred meters thick. In sedimentary targets, impact melt products generally occur as isolated clasts within various impactite types and/or as clast-rich impact melt rocks. Impact melt products of sedimentary targets made of different lithologies result in unmixed and heterogeneous melts.
Most impact melt rocks have been produced as a consequence of large meteoritic impacts (projectile diameter larger than few meters). These rocks reflect the chemical composition of the target rocks; the contribution of the projectile is lower than 1 vol.%, because during large meteoritic impact the projectile vaporizes. For this reason, the nature of the projectile can possibly be detected only through the analysis of siderophile elements and of isotopic composition of osmium and chromium (e.g., for review Koeberl et al. 2012). During small impact events (final crater diameter < 1.5 km) impact melt rocks result both by the melting of the target and of the projectile; indeed the projectile survives the impact and partially melted impactor debris can be found in the proximity of the crater (e.g., Wabar (116
34 m in diameter), Kamil (45 m in diameter), Henbury (157 m in diameter), Aouelloul (390 m),
Barringer (1.2 km in diameter); Table 7 in Fazio et al. 2014 – Chapter 2 of this Thesis). Nineteen out to 27 small impact craters are formed by the impact of iron or stony-iron projectiles. The great geochemical difference between target rocks and iron projectiles allows us to study how the projectile interacts with the target, nevertheless this mixing process is not yet very well understood. In recent years new steps forward in this field have been made thanks to new experimental and numerical simulations (e.g., MEMIN experiments; Ebert et al. 2013 and 2014) and to new evidences from Wabar impactites (Hamann et al. 2013) and Kamil impactites (Chapter 3 of this Thesis).
1.5. Shock effects in quartz-bearing rocks: crystalline vs. sedimentary targets
Shock pressure calibration for single quartz crystal and for quartz-bearing crystalline was proposed for the first time by Stöffler (1971). This experimental calibration has not changed very much over the years, and it is still widely accepted (Table 2). Although quartz is the main component of several sedimentary rocks, shock effects in quartz-bearing sedimentary rocks are poorly calibrated for two main reasons: i) sedimentary porous rocks can display a very wide range in their highly mineralogical assemblage, porosity, type and amount of matrix, grain size, water contents, and fabric that make it difficult to compare rocks with a different lithology; ii) in naturally shocked sedimentary porous rocks, shock deformation effects typically attributable to very different shock pressures can coexist in crystalline rocks (Kieffer 1971; Kieffer et al. 1976; Grieve et al. 1996).
Table 2. Progressive shock metamorphism of quartz-feldspathic crystalline rocks. Table after Stöffler and Grieve 2007, modified after Stöffler 1971.
Modified after Stöffler 1971 Shock
stage Pressure (GPa)
Temperature
(°C) Shock produced phenomena
0 Fractured minerals
Ia ~ 10 ~ 100 Quartz and feldspar with planar deformation features
Ib ~ 20 ~ 170
Quartz and feldspar with planar deformation features and reduced refractive index stishovite and minor coesite
II ~ 35 ~ 300 Diaplectic SiO2 glass and feldspar glass; coesite and
traces of stishovite; cordierite glass
III ~ 45 ~ 900 Normal feldspar glass (vesiculated) and diaplectic SiO2
glass; coesite; cordierite glass
VI ~ 60 ~ 1500 Rock glasses or crystallized melt rocks (quenched from liquids) V ~ 80-100 > 2500 Rock glasses (melts condensed from silicate vapor)
Recently, new experiments and numerical simulations have been carried out to better understand how sedimentary porous rocks behave under the passage of a shock wave (e.g., Wünnemann et al. 2006; Schade and Wünnemann 2007; Wünnemann et al. 2008; Güldemeister et al. 2013; Kowitz et al. 2013). They demonstrate that shock pressure and temperature are enhanced in correspondence of pores and more in general in correspondence of heterogeneities of the rocks (i.e., grain boundary, pre-existing fractures, weakness planes). The shock pressure peak is recorded close to the pore, calculations by Güldemeister et al. (2013) show that here the shock pressure can be two to four times higher than elsewhere in the same sample. Hence, in rocks shocked at the nominal pressure of few GPa (2.5-17.5 GPa) it is possible to find shock features consistent with shock pressures in excess of 30 GPa (Kowitz et al. 2013) localized in correspondence of pores and eventually in correspondence of other heterogeneities.
The first shock pressure calibration for porous quartz-rich sandstone was proposed by Kieffer et al. (1976) for the Coconino Sandstone Formation. This classification is based on observations on progressively shocked Coconino Sandstone rocks from Barringer Crater. This classification was widely accepted, but it was never calibrated. Recently, Kowitz et al. (2013) calibrated and improved this classification scheme as shown in Table 3.
Table 3. Progressive shock metamorphism of porous sandstone (modified after Kieffer et al. 1976 and Kowitz et al. 2013).
Modified after Kieffer et al. 1976 Modified after Kowitz et al. 2013 Shock
stage Pressure (GPa)
Temperature
(°C) Shock produced phenomena
Nominal shock pressure (GPa)
Glass and high-pressure phases 0 ≤ 0.2-0-9 25 Undeformed sandstone
1a 0.2-0.9 – ~3 Deformed sandstone with remnant porosity 2.5 0 5 0.03 vol.% 7.5 0.4 vol.% 1b ~3 – 5.5
~250
Deformed sandstone compressed to zero porosity
10 2.2 vol.% 2 ~5.5 – ~13 ~350 Dense sandstone with 2-5 vol.% coesite, 3-10 vol.% glass, and 80-95
vol.% quartz
12.5 7.4 vol.%
3 ~13 – ~30 ~950
Dense sandstone with 18-32 vol.% coesite, traces of stishovite, 0-20 vol.%
glass, 15-45 vol.% quartz 15 24.5 vol.%
4 >~30 >1000
Dense sandstone with 10-30 vol.% coesite, 20-75 vol.% glass, and 15-45 vol.% quartz
17.5 80.6 vol.%
5
Vesicular (pumiceous) rock with 0-5 vol.% coesite, 80-100 vol.% glass
36 The main differences between the classification for quartz-bearing crystalline rocks
and for quartz-bearing sedimentary rocks relate to the absence of PDFs in quartz in sedimentary rocks and to the appearance of high-pressure silica polymorphs and glass at pressure > 5 GPa. More precisely, rare PDFs could be found in quartz of sedimentary rocks, for example in Coconino Sandstone from Barringer Crater, PDFs constitute about 5 vol.% (Robertson 1980). Moreover, when PDFs occur in sedimentary rocks, their poles have orientations dominated by high angles to c-axis (> 45°), like
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(e.g., in Oasis (Libya), Aorounga (Chad), Avak (Alaska, USA), Tookoonooka (Australia) impact structures; Robertson 1980; Grieve and Therriault 1995; Grieve et al. 1996).The lack of PDFs and their anomalous orientations in sedimentary targets have been attributed to the different impedance of the porous rocks and crystalline rocks: at shock pressure where the low angle to c-axis PDFs form (e.g. (0001) and
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), strain was taken up to close voids without PDFs formation (Robertson 1980). Planar deformation features can develop only when the porosity is zero. When sandstone rocks have a low porosity (e.g., silicified sandstone) they behave as crystalline rocks, hence PDFs are common and their poles are oriented at low angles to c-axis (Milton et al. 1972; Fazio et al. 2014).Coesite and stishovite can form starting from 5.5 GPa. High-pressure polymorphs are typically localized within so-called “symplectic regions”. These areas are roundish and opaque. They are made of a microscopic to sub-microscopic intergrowth of quartz, diaplectic glass, high-pressure polymorphs, and lechatelierite. The formation of these regions is the result of complex interactions during the collapse of the pore spaces (Kieffer et al. 1976). Diaplectic glass in porous sedimentary rocks also forms starting from 5.5 GPa.
In non-porous crystalline rocks, individual minerals start to melt around 50 GPa. Around 60 GPa the melting of the whole rocks can be considered concluded (Stöffler and Langenhorst 1994). In porous sedimentary rocks, individual quartz grains start to melt at around 20 GPa, the whole rock melting is concluded around 30-35 GPa (Kieffer et al. 1976). In order to understand how the shock pressure necessary to melt a rock changes with the rock porosity, Wünnemann et al. (2008) carried out numerical simulations for quartz and calcite (Fig. 20). As expected, the critical pressure for melting decreases with increasing porosity of the target rock. Calcite (dashed line) has a linear curve line, whereas quartz (solid line) shows a kink in the curve due to the solid-state phase transition. Gray lines represent the extrapolation of simulation results at porosity > 50% (Wünnemann et al. 2008).
Fig. 20 Critical pressure for melting vs. porosity for quartz-rich and calcite-rich rocks. Image after Wünnemann et al. (2008).
1.6. References
http://www.lpi.usra.edu/exploration/training/resources/ http://www.passc.net/EarthImpactDatabase/index.html
Baratoux D. and Melosh H. J. 2003. The formation of shatter cones by shock wave interference during impacting. Earth and Planetary Science Letters 216:43-54.
Bland P. A. and Artemieva N. A. 2006 The rate of small impact on Earth. Meteoritics &
Planetary Science 4:607–631.
Ebert M., Hecht L., Deutsch A. and Kenkmann T. 2013. Chemical modification of projectile residues and target material in a MEMIN cratering experiment. Meteoritics &
Planetary Science 48, 134–149.
Ebert M., Hecht L., Deutsch A., Kenkmann T., Wirth R. and Berndt J. 2014. Geochemical processes between steel projectiles and silica-rich targets in hypervelocity impact experiments. Geochimica et Cosmochimica Acta 133, 257–279.
El Goresy A., Gillet P., Chen M., Künstler F, Graup G., Stähle V. 2001. In situ discovery of shock-induced graphite-diamond phase transition in gneisses from the Ries Crater, Germany. American Mineralogy 86:611-621.
Fackelman S. P., Morrow J. R., Koeberl C., and McElvain T. H. 2008. Shatter cone and microscopic shock-alteration evidence for a post-Paleoproterozoic terrestrial impact structure near Santa Fe, New Mexico, USA. Earth and Planetary Science Letters 270:290-299.
Fazio A., Folco L., D’Orazio M., Cordier C., and Frezzotti M. L. 2014. Shock metamorphism and impact melting in small impact craters on Earth: Evidence from Kamil Crater,
38 Egypt. Meteoritics & Planetary Science DOI: 10.1111/maps.12385 (Chapter 2 of this
Thesis).
Ferrière L., Morrow J. R., Amgaa T., and Koeberl C. 2009. Systematic study of universal-stage measurements of planar deformation features in shocked quartz: Implications for statistical significance and representation of results. Meteoritics & Planetary Science 44:925-940.
French B. M. 1998. Traces of catastrophe: A handbook of shock- Metamorphic effects in terrestrial meteorite impact structures. Houston, Texas, USA. LPI Contribution No. 954, Lunar and Planetary Institute. 120 p.
French B. M. and Koeberl C. 2010. The convincing identification of terrestrial meteorite impact structures: What works, what doesn’t work, and why. Earth-Science Reviews 98:123-170.
Gay N. C. 1978. The composition of spherules and other features on shatter cone surfaces from the Vredefort structure, South Africa. Earth and Planetary Science Letters 41:3:372-380.
Gibson H. M. and Spray J. G. 1998. Shock-induced melting and vaporization of shatter cone surfaces: Evidence from the Sudbury impact structure. Meteoritics & Planetary
Science 33:2:329-336.
Goltrant O., Leroux H., Doukham J. C., and Cordier P. 1992. Formation mechanism of planar deformation features in naturally shocked quartz. Physics of the Earth and Planetary
Interiors 74:219-240.
Grieve R. A. F. and Therriault A. M. 1995. Planar deformation features in quartz: target effects (abstract). 26th Lunar and Planetary Science Conference 515-516.
Grieve R. A. F., Langenhorst F., and Stöffler D. 1996. Shock metamorphism of quartz in nature and experiment: II. Significance in geoscience. Meteoritics & Planetary
Science 31:6-35.
Güldemeister N., Wünnemann K., Durr N., and Hiermaier S. 2013. Propagation of impact-induced shock waves in porous sandstone using mesoscale modelling. Meteoritics &
Planetary Science 48:1:115-133.
Hamann C., Hecht L., Ebert M. and Wirth R. 2013. Chemical projectile–target interaction and liquid immiscibility in impact glass from the Wabar craters, Saudi Arabia. Geochiica
