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- Herausgeber: John Wiley & Sons
- Kategorie: Wissenschaft und neue Technologien
- Sprache: Englisch
Impact cratering is arguably the most ubiquitous geological process in the Solar System. It has played an important role in Earth’s history, shaping the geological landscape, affecting the evolution of life, and generating economic resources. However, it was only in the latter half of the 20th century that the importance of impact cratering as a geological process was recognized and only during the past couple of decades that the study of meteorite impact structures has moved into the mainstream. This book seeks to fill a critical gap in the literature by providing an overview text covering broad aspects of the impact cratering process and aimed at graduate students, professionals and researchers alike. It introduces readers to the threat and nature of impactors, the impact cratering process, the products, and the effects – both destructive and beneficial. A series of chapters on the various techniques used to study impact craters provide a foundation for anyone studying impact craters for the first time.
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Table of Contents
Cover
COMPANION WEBSITE
Title page
Copyright page
Dedication
Preface
List of contributors
ONE: Impact cratering: processes and products
1.1 Introduction
1.2 Formation of hypervelocity impact craters
1.3 Morphology and morphometry of impact craters
1.4 Impactites
1.5 Recognition of impact craters
1.6 Destructive effects of impact events
1.7 Beneficial effects of impact events
1.8 When a crater does not exist: other evidence for impact events
1.9 Concluding remarks
TWO: Population of impactors and the impact cratering rate in the inner Solar System
2.1 Introduction
2.2 Population of impactors in the inner Solar System
2.3 Impact frequency of NEOs with the Earth
2.4 Comparison with the impact record on terrestrial planets
2.5 Variability of the impact frequency during the last 3 Ga
2.6 The early cratering history of the Solar System
2.7 Conclusions
THREE: The contact and compression stage of impact cratering
3.1 Introduction
3.2 Maximum pressures during contact and compression
3.3 Jetting during contact and compression
3.4 The isobaric core
3.5 Oblique impact
3.6 The end of contact and compression
FOUR: Excavation and impact ejecta emplacement
4.1 Introduction
4.2 Excavation
4.3 Impact plume
4.4 Generation of continuous ejecta blankets
4.5 Rayed craters
4.6 Generation of multiple ejecta layers
4.7 Distal impact ejecta
4.8 Depth of excavation
FIVE: The modification stage of crater formation
5.1 Introduction
5.2 Morphology and morphometry of simple and complex impact craters
5.3 Kinematics of crater collapse
5.4 Subsurface structure of complex impact craters
5.5 Mechanics of cavity collapse: what makes the target so weak?
5.6 Effects of oblique impact incidences on cavity collapse
5.7 Effects of rheologically complex targets on cavity modification
SIX: Impact-induced hydrothermal activity
6.1 Introduction
6.2 Formation and development of the post-impact thermal field
6.3 Composition and evolution of the hydrothermal fluids and mineralization
6.4 Implications for extraterrestrial impacts and microbial life
SEVEN: Impactites: their characteristics and spatial distribution
7.1 Introduction
7.2 Autochthonous impactites
7.3 Parautochthonous impactites
7.4 Allochthonous impactites
7.5 Concluding remarks
EIGHT: Shock metamorphism
8.1 Introduction
8.2 Shock metamorphic features
8.3 Post-shock thermal features
8.4 Concluding remarks
NINE: Impact melting
9.1 Introduction
9.2 Why impact melting occurs
9.3 Terrestrial impact melt products
9.4 Planetary impact melt products
9.5 Impactor contamination
9.6 Concluding remarks
TEN: Environmental effects of impact events
10.1 Introduction
10.2 The impact hazard
10.3 The impact cratering process
10.4 Shock wave effects
10.5 Ejecta launch
10.6 Long-term atmospheric perturbation
10.7 The response of the Earth system to large impacts
10.8 Environmental impact effects favourable for life
10.9 Concluding remarks
ELEVEN: The geomicrobiology of impact structures
11.1 Introduction
11.2 Physical changes
11.3 Chemical changes
11.4 Impact events and weathering
11.5 Impoverishment or enrichment?
11.6 Astrobiological implications
11.7 Concluding remarks
TWELVE: Economic deposits at terrestrial impact structures
12.1 Introduction
12.2 Progenetic deposits
12.3 Syngenetic deposits
12.4 Epigenetic deposits
12.5 Hydrocarbon accumulations
12.6 Concluding remarks
THIRTEEN: Remote sensing of impact craters
13.1 Introduction
13.2 Background
13.3 Photogeology
13.4 Morphometry, altimetry, topography
13.5 Composition derived from remote sensing
13.6 Physical properties derived from remote sensing
13.7 General spectral enhancement and mapping techniques
13.8 Case studies
13.9 Concluding remarks
FOURTEEN: Geophysical studies of impact craters
14.1 Introduction
14.2 Geophysical signature of terrestrial impacts
14.3 The resolution of geophysical data
14.4 Modelling geophysical data
14.5 Case studies
FIFTEEN: Projectile identification in terrestrial impact structures and ejecta material
15.1 Introduction
15.2 Current situation: projectile identification at impact craters and ejecta layers
15.3 Methodology
15.4 Review of identified projectiles
15.5 Concluding remarks
SIXTEEN: The geochronology of impact craters
16.1 Introduction
16.2 Techniques used for dating terrestrial impact craters
16.3 Impact craters at the K–Pg boundary
16.4 Geochronology of impacts, flood basalts and mass extinctions
16.5 Using geochronology to identify clusters of impacts in the geological record
16.6 Concluding remarks
SEVENTEEN: Numerical modelling of impact processes
17.1 Introduction
17.2 Fundamentals of impact models
17.3 Material models
17.4 Validation, verification and benchmarking
17.5 Concluding remarks
EIGHTEEN: Comparison of simple impact craters: a case study of Meteor and Lonar Craters
18.1 Introduction
18.2 Meteor Crater, Arizona
18.3 Lonar Crater
18.4 Comparisons and planetary implications
18.5 Summary and concluding remarks
Acknowledgements
NINETEEN: Comparison of mid-size terrestrial complex impact structures: a case study
19.1 Introduction
19.2 Overview of craters
19.3 Comparisons and implications
19.4 Comparisons with lunar and Martian impact craters
19.5 Concluding remarks
TWENTY: Processes and products of impact cratering: glossary and definitions
20.1 Introduction
20.2 General definitions
20.3 Morphometric definitions and equations
20.4 Impactites
Index
COMPANION WEBSITE:
This book has a companion website:
www.wiley.com/go/osinski/impactcratering
with Figures and Tables from the book
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Library of Congress Cataloging-in-Publication Data
Impact cratering : processes and products / edited by Gordon R. Osinski and Elisabetta Pierazzo.
pages ; cm
Includes bibliographical references and index.
ISBN 978-1-4051-9829-5 (cloth)
1. Impact craters. 2. Cratering. I. Osinski, Gordon R., editor of compilation. II. Pierazzo, Elisabetta, editor of compilation.
QE612.I47 2013
551.3'97–dc23
2012018762
A catalogue record for this book is available from the British Library.
Wiley also publishes its books in a variety of electronic formats. Some content that appears in print may not be available in electronic books.
Main image: This Mars Global Surveyor image shows a cluster of impact craters in northwest Arabia Terra. NASA/JPL/Malin Space Science Systems.
Thumbnails. Left: The Haughton impact structure in the Canadian High Arctic. It is one of the best-exposed impact structures in the world and is estimated to be 39 million years old. Landsat 7 satellite image. NASA’s Goddard Space Flight Center/USGS.
Centre: Meteor crater near the Route 66, Arizona, US. ©iStockphoto.com/Helena Lovincic.
Right: The south pole of Saturn’s moon Enceladus. From afar, Enceladus exhibits a bizarre mixture of softened craters and complex, fractured terrains. NASA/JPL/Space Science Institute.
Cover Design by Steve Thompson
Dedication
This book is dedicated to my friend, colleague, and co-editor, Elisabetta Pierazzo, known to everyone as Betty.
Elisabetta (Betty) Pierazzo
1963–2011
Betty was recognized internationally as an expert on impact cratering. She was a Senior Scientist at the Planetary Science Institute in Tucson, Arizona. Betty attended graduate school at the Department of Planetary Sciences at the University of Arizona, receiving her Ph.D. in 1997 and was awarded the Gerard P. Kuiper Memorial Award for this research. After a few years as a Research Associate at the University of Arizona, she joined the Planetary Science Institute as a Research Scientist in 2002. It is not an exaggeration to say that Betty played an integral role in the successful expansion of the Planetary Science Institute and making it the renowned research organization that it is today.
Betty was an exceptional scientist and she made numerous important research contributions throughout her career. Amongst her many achievements, she generated the first robust models of melt production during impact events and of oblique impact events. In more recent years Betty increasingly focused on furthering our understanding of the environmental effects of impact events. Another area of research that Betty played a pioneering role in was in modeling the astrobiological consequences of impact events. This ranged from investigating the delivery of organics to planets to the generation of hydrothermal systems by impacts that may have been favorable sites for life on Mars. Betty’s community-minded spirit also led her to lead an international effort to benchmark and validate the various different numerical codes used in the impact cratering community.
This sense of bringing together the community is also reflected in Betty’s leadership in developing the Bridging the Gap series of conferences. I remember attending the first of these conferences in 2003 as a graduate student and being incredibly impressed by the level of community participation and the time dedicated to discussion and the highlighting of major research problems facing the community. The Lunar and Planetary Institute Contribution (#1162) that arose from this conference stands as a testament to Betty’s leadership.
In addition to her research, Betty was passionate about education and public outreach. One of her earliest projects was the Explorers Guide to Impact Craters program. Betty hired me to help develop this program during my time as a Research Associate at the University of Arizona. One of my fondest memories of Betty was the trip we took together to Meteor Crater with Frank Chuang, where we did the filming for the virtual tour of this crater. Since that time, this program has grown, reached thousands of school children, and served as a model for other programs at the Planetary Science Institute.
Finally, any dedication to Betty would be incomplete without mentioning Betty’s lust for life. Whether it was interacting with colleagues on field trips to impact craters around the world, on the pitch playing soccer, or during the many social gatherings at her house in Tucson, Betty always had a smile on her face that invoked the same sense of enthusiasm and joy in everyone she met. Her approach to life made her an ideal mentor for all budding scientists, but especially for women. She proved by her actions that there were many opportunities available to all hard working scientists. She prioritized the happiness of herself and her family, and by doing so ensured her success in work and life, leading the way for others to do the same. The impact community has lost an invaluable member; we all miss her.
Gordon Osinski
Preface
Despite first being described on the Moon by Galileo Galilei in the seventeenth century, it was not until the late nineteenth century that the American geologist Grove Gilbert proposed an impact origin for these lunar ‘craters’. This is long after the birth of modern geology. It was not until 1906 that the first impact crater on Earth – Meteor or Barringer Crater – was recognized by Daniel Barringer in northern Arizona. By the 1930s, several small craters on Earth were suspected as being of impact origin, based on their association with fragments of meteorites. Robert Dietz established the first reliable geological criterion for the identification of impact structures, in the absence of meteorites, in 1947, with the recognition of shatter cones. The increased recognition of impact sites resulting from this discovery, together with the impetus provided by the Apollo landings on the Moon, led to a more complete understanding of the formation of impact craters in the 1960s thanks to the pioneering work of a small group of scientists, including Ralph Baldwin, Carlyle Beals, Robert Dietz, Bevan French, Eugene Shoemaker and others. An improved understanding of the impact cratering process continued during the 1970s and 1980s, with the recognition of several shock metamorphic criteria and dozens more terrestrial impact sites. Finally, in the 1990s, two events finally resulted in impact cratering entering the geological mainstream. Namely, the spectacular impact of 21 fragments of the comet Shoemaker–Levy 9 into Jupiter over 6 days in July 1994 and the discovery of the approximately 200 km diameter Chicxulub impact structure, Mexico, with its link to the Cretaceous–Palaeogene mass extinction event. As has been argued by others, impact cratering can be considered geology’s latest revolution.
The initial idea for this book dates back to my time as a graduate student, in the late 1990s. As a student new to the field of impact cratering, I was struck by the general absence of books on this topic. The few that were in circulation were mostly either conference volumes on particular sub-topics or were out of print (e.g. Impact Cratering: A Geologic Process, published by H. Jay Melosh in 1989). I also became aware of the general lack of appreciation in the scientific community as to the importance of impact cratering as a geological process, particularly on the Earth. Three events then occurred in quick succession that led to the development of this book. First, in 2007, I co-organized the second Bridging the Gap conference in Montreal with Betty Pierazzo and Robbie Herrick. In January 2008, shortly after moving to Western University, I developed a graduate course on Impact Cratering: Processes and Products for the first time. The structure of this book is modelled on this course. Then, in late 2008, I was contacted by Ian Francis from Wiley–Blackwell following the publication of a general article entitled ‘Meteorite impact structures: the good and the bad’ in Geology Today. Ian proposed the concept of a book on impact cratering. With some hesitation – as I was only in my second year of being a professor and not even in a tenure track position – I agreed. Soon afterwards I contacted Betty Pierazzo with the idea of continuing the spirit of the Bridging the Gap theme of bringing together scientists from various disciplines who study the impact cratering process, its products and effects. Betty gladly agreed and this book is testament to Betty’s knowledge, integrity and professionalism.
Sadly, Betty passed away in May 2011, after a battle with a rare form of cancer. A dedication to Betty can be found on the preceding pages. This book would not have been possible without Betty. I am greatly indebted to Richard Grieve for stepping in to assist with editing during Betty’s illness and subsequent passing. I would like to thank Richard and all the authors and co-authors of the various chapters in this book for their contributions and taking time out of their busy schedules to make their contributions happen. In addition to various authors and co-authors, who provided peer reviews of other chapters, I would like to thank Claire Belcher, Veronica Bray, Michael Dence, J. Wright Horton, Fred Jourdan, David King, Lucy Thompson, Mark Pilkington, Uwe Reimold, Ralf-Thomas Schmidt, John Spray, Axel Wittman and Michael Zanetti for providing thoughtful and constructive reviews of various chapters. Lastly, I would like to thank the publishing team at Wiley–Blackwell for their unwavering support and patience throughout this process.
Dr Gordon R. OsinskiWestern UniversityLondon, Ontario
List of Contributors
Natalia Artemieva Institute for Dynamics of Geospheres, Russian Academy of Science, Leninsky pr. 38, blg.1, Moscow 119334, Russia.
Anna Chanou Departments of Earth Sciences/Physics and Astronomy, Western University, 1151 Richmond Street, London, ON, N6A 5B7, Canada.
Philippe Claeys Earth System Science, Vrije Universiteit Brussel, Pleinlaan 2, BE-1050 Brussels, Belgium.
Charles S. Cockell School of Physics and Astronomy, University of Edinburgh, Edinburgh, EH9 3JZ, UK. Email: [email protected].
Gareth S. Collins IARC, Department of Earth Science and Engineering, Imperial College London, London, SW7 2AZ, UK. Email: [email protected].
Ludovic Ferrière Natural History Museum, Burgring 7, A-1010 Vienna, Austria. Departments of Earth Sciences/Physics and Astronomy, University of Western Ontario, 1151 Richmond Street, London, ON, N6A 5B7, Canada. Email: [email protected].
Steven Goderis Earth System Science, Vrije Universiteit Brussel, Pleinlaan 2, BE-1050 Brussels, Belgium. Department of Analytical Chemistry, Universiteit Gent, Krijgslaan 281 – S12, BE-9000 Ghent, Belgium. Email: [email protected].
Richard A. F. Grieve Earth Sciences Sector, Natural Resources Canada, Ottawa, Ontario, K1A 0E8, Canada, and Departments of Earth Sciences/Physics and Astronomy, Western University, 1151 Richmond Street, London, ON, N6A 5B7, Canada. Email: [email protected].
Justin J. Hagerty United States Geological Survey, Astrogeology Science Center, 2255 N. Gemini Drive, Flagstaff, AZ 86001, USA. Email: [email protected].
Simon P. Kelley Centre for Earth, Planetary, Space and Astronomical Research, Department of Earth and Environmental Sciences, Open University, Milton Keynes, MK7 6AA, UK. Email: [email protected].
Thomas Kenkmann Institut für Geowissenschaften – Geologie, Albert-Ludwigs-Universität Freiburg, Albertstrasse 23-B, D-79104 Freiburg, Germany. Email: [email protected].
Kalle Kirsimäe Department of Geology, University of Tartu, Ravila 14a, 50411 Tartu, Estonia. Email: [email protected].
Cassandra Marion Departments of Earth Sciences/Physics and Astronomy, Western University, 1151 Richmond Street, London, ON, N6A 5B7, Canada.
H. Jay Melosh Earth, Atmospheric and Planetary Sciences Department, 550 Stadium Mall Drive, Purdue University, West Lafayette, IN 47907, USA. Email: [email protected].
Patrick Michel University of Nice–Sophia Antipolis, CNRS, Côte d’Azur Observatory, BP 4229, 06304 Nice Cedex 4, France. Email: [email protected].
Saumitra Misra School of Geological Sciences, University of Kwazulu-Natal, Private Bag X54001, Durban 4000, South Africa. Email: [email protected].
Alessandro Morbidelli University of Nice–Sophia Antipolis, CNRS, Côte d’Azur Observatory, BP 4229, 06304 Nice Cedex 4, France.
Joanna Morgan Department of Earth Science and Engineering, Imperial College London, South Kensington Campus, London, SW7 2AZ, UK. Email: [email protected].
Horton E. Newsom Institute of Meteoritics and Department of Earth and Planetary Sciences MSC03-2050, University of New Mexico, Albuquerque, NM 87131, USA. Email: [email protected].
Gordon R. Osinski Departments of Earth Sciences/Physics and Astronomy, Western University, 1151 Richmond Street, London, ON, N6A 5B7, Canada. Email: [email protected].
François Paquay Department of Geology and Geophysics, University of Hawaii at Manoa, Honolulu, HI, USA.
Elisabetta Pierazzo (Deceased) Formerly of Planetary Science Institute, 1700 E. Fort Lowell Road, Suite 106, Tucson, AZ 85719, USA.
Michael S. Ramsey Department of Geology and Planetary Science, University of Pittsburgh, 4107 O’Hara Street, SRCC, Pittsburgh, PA 15260-3332, USA.
Mario Rebolledo-Vieyra Centro de Investigación Científica de Yucatán, A.C., CICY, Calle 43 No 130, 97200 Mérida, Mexico.
Sarah C. Sherlock Centre for Earth, Planetary, Space and Astronomical Research, Department of Earth and Environmental Sciences, Open University, Milton Keynes, MK7 6AA, UK.
Ann M. Therriault Earth Sciences Sector, Natural Resources Canada, Ottawa, Ontario, K1A 0E8, Canada.
Livio L. Tornabene Departments of Earth Sciences/Physics and Astronomy, Western University, 1151 Richmond Street, London, ON, N6A 5B7, Canada.
Mary A. Voytek US Geological Survey, MS 430, 12201 Sunrise Valley Drive, Reston, VA 20192, USA.
Shawn P. Wright Institute of Meteoritics and Department of Earth and Planetary Sciences MSC03 2050, University of New Mexico, Albuquerque, NM 87131, USA. Email: [email protected].
Kai Wünnemann Museum für Naturkunde, Leibniz-Institut an der Humboldt-Universität Berlin, Invalidenstrasse 43, 10115 Berlin, Germany.
ONE
Impact Cratering: Processes and Products
Gordon R. Osinski* and Elisabetta Pierazzo†
*Departments of Earth Sciences/Physics and Astronomy, Western University, 1151 Richmond Street, London, ON, N6A 5B7, Canada
†Planetary Science Institute, 1700 E. Fort Lowell Road, Suite 106, Tucson, AZ 85719, USA
1.1 Introduction
Over the past couple of decades, it is has become widely recognized that impact cratering is a ubiquitous geological process that affects all planetary objects with a solid surface. Indeed, meteorite impact structures are one of the most common geological landforms on all the rocky terrestrial planets, except Earth, and many of the rocky and icy moons of Saturn and Jupiter. A unique result of the impact cratering process is that material from depth is brought to the surface in the form of ejecta deposits and central uplifts. Impact craters, therefore, provide unique windows into the subsurface on planetary bodies where drilling more than a few metres is not a viable scenario for the foreseeable future. On many planetary bodies where planetary-scale regoliths can develop through micrometeorite bombardment, aeolian or cryogenic processes, the crater walls of fresh impact craters also provide unique sites where in situ outcrops can be found. It should not be surprising, therefore, that impact craters have been, and remain, high-priority targets for planetary exploration missions to the Moon, Mars and elsewhere.
The impact record on Earth remains invaluable for our understanding of impact processes, for it is the only source of ground-truth data on the three-dimensional structural and lithological character of impact craters. However, Earth suffers from active erosion, volcanic resurfacing and tectonic activity, which continually erase impact structures from the rock record. Despite this, 181 confirmed impact structures have been documented to date, with several more ‘new’ impact sites being recognized each year (Earth Impact Database, 2012). Although we lack ground truth, apart from a few lunar and Martian sites visited by human and robotic explorers, the results of planetary exploration missions continue to provide a wealth of new high-resolution data about the surface expression of impact craters. The driving paradigm is that impact cratering is governed by physics and the fundamental processes are the same regardless of the planetary target (Melosh, 1989). However, variations in planetary conditions permit the investigation of how different properties lead to slightly different end results. The Moon represents an end-member case with respect to the terrestrial planets. Low planetary gravity and lack of atmosphere result in cratering efficiency, for a given impact, that is higher than on the other terrestrial planets (Stöffler et al., 2006). The relatively simple target geology combined with the lack of post-impact modification by aqueous and aeolian processes makes the Moon an ideal natural laboratory for studying crater morphology and morphometry. Mercury is similar to the Moon, except for a higher impact velocity, and new data from the MESSENGER spacecraft (see Solomon et al. (2011) and references therein) are providing a wealth of new information on the mercurian impact cratering record (Strom et al., 2008). Venus is almost the antithesis of the Moon and Mercury. The relatively high planetary gravity and thick atmosphere reduce cratering efficiency for a given impact relative to these bodies (Schultz, 1993). Hotter surface and subsurface temperatures affect numerous aspects of the cratering process on Venus, the most spectacular outcome of which is the production of vast impact melt flows (Grieve and Cintala, 1995). The final terrestrial planet, Mars, has a thinner atmosphere, a more complex geology, including the presence of volatiles, and more endogenic geological processes to modify craters (Carr, 2006). It is more Earth-like in this respect, which comes with the associated complications, but its impact cratering record is vastly better preserved and exposed than on Earth (Strom et al., 1992).
Notwithstanding the prior discussion of the ubiquity of impact craters throughout the Solar System, it is important to recognize that, despite being first observed on the Moon by Galileo Galilei in 1609, it was not until the 1960s and 1970s that the importance of impact cratering as a geological process began to be recognized. In 1893, the American geologist Grove Gilbert proposed an impact origin for these lunar craters, but it was not until the 1900s that the first impact crater was recognized on Earth: Meteor or Barringer Crater in Arizona (Barringer, 1905; Fig. 1.1a). In the decades that followed, there remained little awareness in the geological community of the importance of impact cratering and there was a general view that impact events were not important for Earth evolution. Indeed, even G. Gilbert, himself, initially disputed the impact origin of Barringer Crater. It was not until the recognition of shock metamorphic criteria (French and Short, 1968; see Chapter 8), which resulted in the increased recognition of terrestrial impact sites, together with the impetus provided by the Apollo landings on the Moon in the 1960s and 1970s, that increasing awareness and a more complete appreciation of the formation of impact craters commenced.
Figure 1.1 Simple impact craters. (a) Panoramic image of the 1.2 km diameter Meteor or Barringer Crater, Arizona. (b) Schematic cross-section through a simple terrestrial impact crater. Fresh examples display an overturned flap of near-surface target rocks overlain by ejecta. The bowl-shaped cavity is partially filled with allochthonous unshocked and shocked target material. (c) A 2 m per pixel true-colour image of Barringer Crater taken by WorldView2 (north is up).
Image courtesy of Livio L. Tornabene and John Grant.
(d) Portion of Lunar Reconaissance Orbiter Camera (LROC) image M122129845 of the 2.2 km diameter Linné Crater on the Moon (NASA/GSFC/Arizona State University).
Discussion of the importance of meteorite impacts for Earth evolution finally entered the geological mainstream in 1980, with evidence for a major impact as the cause of the mass extinction event at the Cretaceous–Paleogene (K–Pg) boundary 65 Ma ago (Alvarez et al., 1980). The actual impact site, the approximately 180 km diameter Chicxulub crater, was subsequently identified in 1991, buried beneath approximately 1 km of sediments in the Yucatan Peninsula, Mexico (Hildebrand et al., 1991). The spectacular impact of comet Shoemaker–Levy 9 into Jupiter in July 1994 reminded us that impact cratering is a process that continues to the present day. The result is that it is now apparent that meteorite impact events have played an important role throughout Earth’s history, shaping the geological landscape, affecting the evolution of life and producing economic benefits. As summarized in Chapter 2, the evolutions of the terrestrial planets and the Earth’s moon have been strongly affected by changes in the population of impactors and in the impact cratering rate through time in the inner Solar System.
To summarize, our understanding of the impact cratering process has come a long way in the past century, but several fundamental aspects of the processes and products of crater formation remain poorly understood. One of the major reasons for this is that, unlike many other geological processes, there have been no historical examples of hypervelocity impact events (French, 1998). This is, of course, fortunate, as impacts release energies far in excess of even the most devastating endogenous geological events (Fig. 1.2). Our understanding is also hindered by the major differences between impact events and other geological processes, including (1) the extreme physical conditions (Fig. 1.3), (2) the concentrated nature of the energy release at a single point on the Earth’s surface, (3) the virtually instantaneous nature of the impact process (e.g. seconds to minutes) and (4) the strain rates involved (∼104–106 s−1 for impacts versus 10−3–10−6 s−1 for endogenous tectonic and metamorphic processes) (French, 1998). Impact events, therefore, are unlike any other geological process, and the goal of this chapter, and this book, is to provide a modern up-to-date synthesis as to our current understanding of the processes and products of impact cratering.
Figure 1.2 A comparison of the energy released during impact events with endogenous geological processes and man-made explosions. Note that only the frequency of impact events is shown. The vertical axis represents the frequency of impact events expressed as the estimated interval in years for a particular size of event. For example, an impact event of the size that formed Barringer Crater is expected once every 1900 years.
Data from French (1998).
Figure 1.3 Pressure–temperature (P–T) plot showing comparative conditions for shock metamorphism and ‘normal’ crustal metamorphism. Note that the pressure axis is logarithmic. The approximate P–T conditions needed to produce specific shock effects are indicated by vertical dashed lines below the exponential curve that encompasses the field of shock metamorphism.
Modified from French (1998).
1.2 Formation of Hypervelocity Impact Craters
The formation of hypervelocity1 impact craters has been divided, somewhat arbitrarily, into three main stages (Gault et al., 1968): (1) contact and compression; (2) excavation; and (3) modification (Fig. 1.4). These are described below. A further stage of ‘hydrothermal and chemical alteration’ has also sometimes been included as a separate, final stage in the cratering process (Kieffer and Simonds, 1980), and is also described below.
Figure 1.4 Series of schematic cross-sections depicting the three main stages in the formation of impact craters. This multi-stage model accounts for melt emplacement in both simple (left panel) and complex craters (right panel). For the modification stage section, the arrows represent different time steps, labelled ‘a’ to ‘c’. Initially, the gravitational collapse of crater walls and central uplift (a) results in generally inwards movement of material. Later, melt and clasts flow off the central uplift (b). Then, there is continued movement of melt and clasts outwards once crater wall collapse has largely ceased (c).
Modified from Osinski et al. (2011).
1.2.1 Contact and Compression
The first stage of an impact event begins when the projectile, be it an asteroid or comet, contacts the surface of the target (Fig. 1.4) – see Chapter 3 for details. Modelling of the impact process suggests that the projectile penetrates no more than one to two times its diameter (Kieffer and Simonds, 1980; O’Keefe and Ahrens, 1982). The pressures at the point of impact are typically several thousand times the Earth’s normal atmospheric pressure (i.e. >100 GPa) (Shoemaker, 1960). The intense kinetic energy of the projectile is transferred into the target in the form of shock waves that occur at the boundary between the compressed and uncompressed target material (Melosh, 1989). These shock waves, travelling faster than the speed of sound, propagate both into the target sequence and back into the projectile itself. When this reflected shock wave reaches the ‘free’ upper surface of the projectile, it is reflected back into the projectile as a rarefaction or tensional wave (Ahrens and O’Keefe, 1972). The passage of this rarefaction wave through the projectile causes it to unload from high shock pressures, resulting in the complete melting and/or vaporization of the projectile itself (Gault et al., 1968; Melosh, 1989). The increase in internal energy accompanying shock compression and subsequent rarefaction also results in the shock metamorphism (see Chapter 8), melting (see Chapter 9) and/or vaporization of a volume of target material close to the point of impact (Ahrens and O’Keefe, 1972; Grieve et al., 1977). The point at which the projectile is completely unloaded is generally taken as the end of the contact and compression stage (Melosh, 1989; Chapter 3).
1.2.2 Excavation Stage
The transition from the initial contact and compression stage into the excavation stage is a continuum. It is during this stage that the actual impact crater is opened up by complex interactions between the expanding shock wave and the original ground surface (Melosh, 1989) – see Chapter 4 for details. The projectile itself plays no role in the excavation of the crater, having been unloaded, melted and/or vaporized during the initial contact and compression stage.
During the excavation stage, the roughly hemispherical shock wave propagates out into the target sequence (Fig. 1.4). The centre of this hemisphere will be at some depth in the target sequence (essentially the depth of penetration of the projectile). The passage of the shock wave causes the target material to be set in motion, with an initial outward radial trajectory. At the same time, shock waves that initially travelled upwards intersect the ground surface and generate rarefaction waves that propagate back downwards into the target sequence (Melosh, 1989). In the near-surface region an ‘interference zone’ is formed in which the maximum recorded pressure is reduced due to interference between the rarefaction and shock waves (Melosh, 1989).
The combination of the outward-directed shock waves and the downward-directed rarefaction waves produces an ‘excavation flow-field’ and generates a so-called ‘transient cavity’ (Fig. 1.4 and Fig. 1.5) (Dence, 1968; Grieve and Cintala, 1981). The different trajectories of material in different regions of the excavation flow field result in the partitioning of the transient cavity into an upper ‘excavated zone’ and a lower ‘displaced zone’ (Fig. 1.5). Material in the upper zone is ejected ballistically beyond the transient cavity rim to form the continuous ejecta blanket (Oberbeck, 1975) – see Chapter 4. Experiments and theoretical considerations of the excavation flow suggest that the excavated zone comprises material only from the upper one-third to one-half the depth of the transient cavity (Stöffler et al., 1975). It is clear that the excavation flow lines transect the hemispherical pressure contours, so that ejecta will contain material from a range of different shock levels, including shock-melted target lithologies. In simple craters (see Section 1.3.1), the final crater rim approximates the transient cavity rim (Fig. 1.1). In complex craters (see Section 1.3.2), however, the transient cavity rim is typically destroyed during the modification stage, such that ejecta deposits occur in the crater rim region interior to the final crater rim (Fig. 1.6).
Figure 1.5 Theoretical cross-section through a transient cavity showing the locations of impact metamorphosed target lithologies. Excavation flow lines (dashed lines) open up the crater and result in excavation of material from the upper one-third to one-half the depth of the transient cavity.
Modified after Grieve (1987) and Melosh (1989).
Figure 1.6 Complex impact craters. (a) Landsat 7 image of the 23 km (apparent) diameter Haughton impact structure, Devon Island, Canada. (b) Portion of Apollo 17 metric image AS17-M-2923 showing the 27 km diameter Euler Crater on the Moon. Note the well-developed central peak. (c) Thermal Emission Imaging System (THEMIS) visible mosaic of the 29 km diameter Tooting Crater on Mars (NASA). Note the well-developed central peak and layered ejecta blanket. Scale bars for (a) to (c) are 10 km. (d) Schematic cross-section showing the principal features of a complex impact crater. Note the structurally complicated rim, a down-faulted annular trough and a structurally uplifted central area (SU). (e) Schematic cross-section showing an eroded version of the fresh complex crater in (d). Note that, in this case, only the apparent crater diameter can typically be defined.
Ejecta deposits represent one of the most distinctive features of impact craters on planetary bodies (other than Earth), where they tend to be preserved. It is notable that the continuous ejecta deposits vary considerably in terms of morphology on different planetary bodies. For example, on Mars, many ejecta blankets have a fluidized appearance that has been ascribed due to the effect of volatiles in the subsurface (Barlow, 2005). Indeed, the volatile content and cohesiveness of the uppermost target rocks will significantly affect the runout distance of the ballistically emplaced continuous ejecta blanket, with impact angle also influencing the overall geometry of the deposits (e.g. the production of the characteristic butterfly pattern seen in very oblique impacts) (Osinski et al., 2011). In terms of the depth of excavation de, few craters on Earth preserve ejecta deposits and/or have the distinct pre-impact stratigraphy necessary for determining depth of materials. Based on stratigraphic considerations, de at Barringer Crater is greater than 0.08D (Shoemaker, 1963), where D is the final rim diameter. For complex craters and basins, the depth and diameter must be referred back to the ‘unmodified’ transient cavity to reliably estimate the depth of excavation (Melosh, 1989). The maximum de of material in the ballistic ejecta deposits of the Haughton and Ries structures, the only terrestrial complex structures where reliable data are available, yield identical values of 0.035Da, where Da is the apparent crater diameter (Osinski et al., 2011). If the initial final rim diameter D is used, which is the parameter measured in planetary craters, a value 0.05D is obtained for Haughton.
Based on experiments, it was generally assumed that material in the displaced zone remains within the transient cavity (Stöffler et al., 1975); however, observations from impact craters on all the terrestrial planets suggest that some of the melt-rich material from this displaced zone is transported outside the transient cavity rim during a second episode of ejecta emplacement (Osinski et al., 2011). This emplacement of more melt-rich, ground-hugging flows – the ‘surface melt flow’ phase – occurs during the terminal stages of crater excavation and the modification stage of crater formation (see Chapter 4). Ejecta deposited during the surface melt flow stage are influenced by several factors, most importantly planetary gravity, surface temperature and the physical properties of the target rocks. Topography and angle of impact play important roles in determining the final distribution of surface melt flow ejecta deposits, with respect to the source crater (Osinski et al., 2011). A critical consideration is that the upper layer of ejecta reflects the composition and depth of the displaced zone of the transient cavity (Fig. 1.4). At Haughton, this value is a minimum of 0.08Da or 0.12D.
A portion of the melt and rock debris that originates beneath the point of impact remains in the transient cavity (Grieve et al., 1977). This material is also deflected upwards and outwards parallel to the base of the cavity, but must travel further and possesses less energy, so that ejection is not possible. This material forms the crater-fill impactites within impact craters (see Chapter 7 for an overview of impactites). Eventually, a point is reached at which the motions associated with the passages of the shock and rarefaction waves can no longer excavate or displace target rock and melt (French, 1998). At the end of the excavation stage, a mixture of melt and rock debris forms a lining to the transient cavity.
1.2.3 Modification Stage
The effects of the modification stage are governed by the size of the transient cavity and the properties of the target rock lithologies (Melosh and Ivanov, 1999) – see Chapter 5 for an overview. For crater diameters less than 2–4 km on Earth, the transient cavity undergoes only minor modification, resulting in the formation of a simple bowl-shaped crater (Fig. 1.1). However, above a certain size threshold the transient cavity is unstable and undergoes modification by gravitational forces, producing a so-called complex impact crater (Fig. 1.4 and Fig. 1.6; Dence, 1965) – see Chapter 5. Uplift of the transient crater floor occurs, leading to the development of a central uplift (Fig. 1.4 and Fig. 1.6). This results in an inward and upward movement of material within the transient cavity. Subsequently, the initially steep walls of the transient crater collapse under gravitational forces (Fig. 1.4). This induces an inward and downward movement of large (∼100 m to kilometre-scale) fault-bounded blocks. The diameter at which the transition occurs from simple to complex craters on Earth occurs at approximately 2 km for craters developed in sedimentary targets and approximately 4 km for those in crystalline lithologies. This transition diameter is dependent on the strength of the gravitational field of the parent body and increases with decreasing acceleration of gravity (Melosh, 1989). Thus, the transition from simple to complex craters occurs at approximately 5–10 km on Mars and at approximately 15–27 km on the Moon (Pike, 1980).
It is generally considered that the modification stage commences after the crater has been fully excavated (Melosh and Ivanov, 1999). However, numerical models suggest that the maximum depth of the transient cavity is attained before the maximum diameter is reached (e.g. Kenkmann and Ivanov, 2000). Thus, uplift of the crater floor may commence before the maximum diameter has been reached. As French (1998) notes, the modification stage has no clearly marked end. Processes that are intimately related to complex crater formation, such as the uplift of the crater floor and collapse of the walls (Chapter 5), merge into more familiar endogenous geological processes, such as mass movement, erosion and so on.
1.2.4 Post-Impact Hydrothermal Activity
Evidence for impact-generated hydrothermal systems has been recognized at over 70 impact craters on Earth (Naumov, 2005; Osinski et al., 2012), from the approximately 1.8 km diameter Lonar Lake structure, India (Hagerty and Newsom, 2003), to the approximately 250 km diameter Sudbury structure, Canada (Ames and Farrow, 2007). Based on these data, it seems highly probable that any hypervelocity impact capable of forming a complex crater will generate a hydrothermal system, as long as sufficient H2O is present (see Chapter 6 for an overview). Thus, the recognition of impact-associated hydrothermal deposits is important in understanding the evolution of impact craters through time. There are three main potential sources of heat for creating impact-generated hydrothermal systems (Osinski et al., 2005a): (a) impact melt rocks and impact melt-bearing breccias; (b) physically elevated geothermal gradients in central uplifts; and (c) thermal energy deposited in central uplifts due to the passage of the shock wave. Interaction of these hot rocks with groundwater and surface water can lead to the development of a hydrothermal system. The circulation of hydrothermal fluids through impact craters can lead to substantial alteration and mineralization. It has been shown that there are six main locations within and around an impact crater where impact-generated hydrothermal deposits can form (Fig. 1.7): (1) crater-fill impact melt rocks and melt-bearing breccias; (2) interior of central uplifts; (3) outer margin of central uplifts; (4) impact ejecta deposits; (5) crater rim region; and (6) post-impact crater lake sediments.
Figure 1.7 Distribution of hydrothermal deposits within and around a typical complex impact crater.
Modified from Osinski et al. (2012).
1.3 Morphology and Morphometry of Impact Craters
1.3.1 Simple Craters
Impact craters are subdivided into two main groups based on morphology: simple and complex. Simple craters comprise a bowl-shaped depression (Fig. 1.1). When fresh, they possess an uplifted rim and are filled with an allochthonous breccia lens that comprises largely unshocked target material, possibly mixed with impact melt-bearing lithologies (Fig. 1.1b; Shoemaker, 1960). The overall low shock level of material in the breccia lens suggests that it formed due to slumping of the transient cavity walls, and is not ‘fallback’ material (Grieve and Cintala, 1981). Simple craters typically have depth-to-diameter ratios of approximately 1 : 5 to 1 : 7 (Melosh, 1989). It is important, however, to make the distinction between the true and apparent crater (Fig. 1.1b). Morphometric data from eight simple impact structures (i.e. Barringer, Brent, Lonar, West Hawk, Aouelloul, Tenoumer, Mauritania and Wolfe Creek) define the empirical relationships: da = D1.06 and dt = 0.28D1.02, where da is the depth of the apparent crater, dt is the depth of the true crater and D is the rim diameter of the structure (Grieve and Pilkington, 1996). As diameter increases, so-called ‘transitional craters’ form. Such craters have not been recognized on Earth, but on the Moon and Mars, where they are abundant, spacecraft observations show that, while they lack a central peak, they possess some of the other characteristics of complex craters (see below), such as a shallower profile and terraced crater rim. As such, they are neither simple nor complex and the exact mechanism(s) responsible for their appearance remain poorly understood.
1.3.2 Complex Craters
Observations of lunar craters first revealed that, as diameter increases yet further, a topographic high forms in the centre of a transitional crater, signifying the progression to a so-called complex impact crater. Such craters generally have a structurally complicated rim, a down-faulted annular trough and an uplifted central area (Fig. 1.6). These features form as a result of gravitational adjustments of the initial crater during the modification stage of impact crater formation (Chapter 5). Owing to these late-stage adjustments, complex impact craters are shallower than simple craters, with depth-to-diameter ratios of approximately 1 : 10 to 1 : 20 (Melosh, 1989). The so-called annular trough in complex craters is filled with a variety of impact-generated lithologies (impactites) that will be introduced in Section 1.4.
A unique result of complex crater formation is that material from depth is brought to the surface. As noted above, for many impact sites, these ‘central uplifts’ provide the only samples of the deep subsurface. This is particularly important on other planetary bodies, but even on Earth they provide vital clues as to the structure of the crust. For example, the central uplift of the approximately 250 km diameter Vredefort impact structure, South Africa, provides a unique profile down to the lower crust (Tredoux et al., 1999). On the Moon and other planets, where post-impact modification of craters is generally minimal, there is a progression with increasing crater size from central peak, central-peak basin (i.e. a fragmentary ring of peaks surrounding a central peak), to peak-ring basins (i.e. a well-developed ring of peaks but no central peak) (Fig. 1.8; Stöffler et al., 2006). On Earth, erosion has modified the surface morphology of all impact craters and it is, therefore, typically not possible to ascertain the original morphology. As such, the term central uplift is preferred. Related to this is the fact that a number of relatively young (i.e. only slightly eroded) terrestrial complex structures (e.g. Haughton (Fig. 1.6a), Canada; Ries, Germany; Zhamanshin, Kazakhstan) lack an emergent central peak (Grieve and Therriault, 2004). These structures are in mixed targets of sediments overlying crystalline basement, with the lack of a peak most likely due to target strength effects. This highlights the problems with making direct comparisons between impact craters on Earth and those on other planetary bodies.
Figure 1.8 Series of images of lunar craters depicting the change in crater morphology with increasing crater size. (a) The 27 km diameter Euler Crater possesses a well-developed central peak. Portion of Apollo 17 metric image AS17-M-2923 (NASA). (b) The 165 km diameter Compton Crater is one of the rare class of central-peak basin craters on the Moon. Clementine mosaic from USGS Map-A-Planet. (c) Clementine mosaic of the 320 km diameter Schrödinger impact crater, which displays a peak ring basin morphology (NASA). (d) The approximately 950 km diameter Orientale Basin is the youngest multi-ring basin on the Moon (NASA/GSFC/Arizona State University).
Based on observations of 24 impact craters developed in sedimentary rocks on Earth, the structural uplift of the target rocks in the centre of the crater (Fig. 1.6d) was determined to be 0.086D1.03, where D is the crater ‘diameter’ (Grieve and Pilkington, 1996). According to this estimate, a good working hypothesis is that the observed structural uplift is approximately one-tenth of the rim diameter at terrestrial complex impact structures. It is important to note that no data exist on the amount of structural uplift in craters developed in crystalline targets for the obvious reason that stratigraphic markers, upon which this calculation relies upon, are lacking. Despite its widespread application, there is also currently no data to support the hypothesis that this formula for structural uplift holds for craters on other planetary bodies, at least in its current form.
A key descriptor for complex craters is ‘diameter’. As noted above, defining the size, or diameter, of a crater is critical for estimating stratigraphic uplift, in addition to energy scaling and numerical modelling of the cratering process. Unfortunately, there is considerable confusion about crater sizes within the literature. This arises largely from the fact that most craters on Earth are eroded to some degree, whereas most craters on other planetary bodies are relatively well preserved. For a discussion of what crater diameter represents, the reader is referred to Turtle et al. (2005) and the nomenclature recommended here comes from this synthesis paper. In short, the rim (or final crater) diameter is defined as the diameter of the topographic rim that rises above the surface for simple craters, or above the outermost slump block not concealed by ejecta for complex craters (Fig. 1.6d). This is relatively easy to measure on most planetary bodies, where the topographic rim is usually preserved due to low rates of erosion (e.g. Fig. 1.6b,c). On Earth, however, such pristine craters are rare and the rim region is typically eroded away (e.g. Fig. 1.6a). The apparent crater diameter, in contrast, is defined as the diameter of the outermost ring of (semi-) continuous concentric normal faults (Fig. 1.6e). For the majority of impact structures on Earth this will be the only measurable diameter. It is not always clear how the apparent diameter is related to the rim diameter, although one would expect the rim diameter to be smaller than the apparent crater diameter. This is consistent with observations at the Haughton impact structure, where an apparent crater diameter of 23 km and a rim diameter of 16 km have been reported (Osinski and Spray, 2005).
Returning to the previous discussion on stratigraphic uplift and its application to planets other than Earth, D in Grieve and Pilkington’s (1996) formula actually is predominantly based on apparent crater diameter estimates (R. A. F. Grieve, personal communication, 2012) and not rim diameter estimates, further complicating the discussion about its application to other planets.
1.3.3 Multi-Ring Basins
The largest impact ‘craters’ in the Solar System are typically surrounded by one or more concentric scarps or fractures and are known as multi-ring basins (Fig. 1.8d). Multi-ring basins are best studied on the Moon and Callisto, where a large number exist, although these structures remain the least understood crater morphology. There are two basic morphological types (e.g. Melosh and McKinnon, 1978). The first type, as exemplified by the Orientale basin on the Moon, exhibits a few to several inward-facing scarps with gentle outward slopes. The second type exhibit tens to hundreds of closely spaced rings consisting of a graben or outward-facing scarps surrounding a central, flat basin (e.g. Valhalla, Callisto). An important observation is that very few multi-ring basins have been documented on Ganymede, despite the obvious similarities with Callisto, and there is no clear evidence for their existence on Mercury, Mars or Venus (Melosh, 1989). In this respect, it is critical to understand that just because an impact crater is very ‘large’ (e.g. Hellas, Mars; South Pole-Aitken, Moon), this does not necessarily mean that a structure is a multi-ring basin; to be categorized as such, multiple rings must be clearly observable. It is also important to note that the rings that define multi-ring craters are distinct from the peak rings described in the previous section. In particular, it is thought that rings characteristic of multi-ring craters form outside the final crater. Several mechanisms have been proposed to account for the formation of multi-ring basins, but no agreement exists in the literature to date – see Melosh (1989) for a discussion. Melosh (1989) preferred the so-called ring tectonic theory, where the thickness of the lithosphere plays a dominant role in determining whether or not a ring forms. More recently, a nested melt-cavity model has been proposed to account for transition from complex craters to multi-ringed basins on the Moon (Head, 2010).
Complications arise, as external rings have been documented around much smaller impact structures, such as the proposed (but not confirmed) approximately 20 km diameter Silverpit structure in the North Sea (Fig. 1.9; Stewart and Allen, 2002). Numerical modelling suggests this morphology formed due to an impact into a layer of brittle chalk overlying weak shales (Stewart and Allen, 2002; Collins et al., 2003). Whether multi-ring basins exist on the Earth also remains a topic of debate. Of the three largest structures on Earth (Chicxulub, Sudbury and Vredefort), Chicxulub is the best-preserved large terrestrial impact structure, due to burial. As such, however, the definition of its morphological elements depends on the interpretation of geophysical data. It has an interior topographic ‘peak-ring’, a terraced rim area and exterior ring faults and, therefore, appears to correspond to the definition of a multi-ring basin, as on the Moon (Grieve et al., 2008).
Figure 1.9 Perspective view of the top chalk surface at the Silverpit structure, North Sea, UK, a suspected meteorite impact structure. The central crater is 2.4 km wide and is surrounded by a series of concentric faults, which extend to a radial distance of approximately 10 km from the crater centre. False colours indicate depth (yellow: shallow; purple: deep).
Image courtesy of Phil Allen and Simon Stewart.
1.4 Impactites
In terms of the products of meteorite impact events, the above considerations of the impact cratering process reveal that pressures and temperatures that can vaporize, melt, shock metamorphose2 and/or deform a substantial volume of the target sequence can be generated. The transport and mixing of impact-metamorphosed3 rocks and minerals during the excavation and formation of impact craters produces a wide variety of distinctive impactites that can be found within and around impact craters (see Fig. 1.10; ‘rock affected by impact metamorphism’) (Stöffler and Grieve 2007) – see Chapter 7. Much of our knowledge of impactites comes from impact craters on Earth and, to a lesser extent the Moon, where large numbers and volumes of samples from known locations are available for study.
Figure 1.10 Field images of impactites. (a) Oblique aerial view of the approximately 80 m high cliffs of impact melt rock at the Mistastin impact structure, Labrador, Canada.
Photograph courtesy of Derek Wilton.
(b) Close-up view of massive fine-grained (aphanitic) impact melt rock from the Discovery Hill locality, Mistastin impact structure. Camera case for scale. (c) Coarse-grained granophyre impact melt rock from the Sudbury Igneous Complex, Canada. Rock hammer for scale. (d) Impact melt-bearing breccia from the Mistastin impact structure. Note the fine-grained groundmass and macroscopic flow-textured silicate glass bodies (large black fragments). Steep Creek locality. Marker/pen for scale. (e) Polymict lithic impact breccias from the Wengenhousen quarry, Ries impact structure, Germany. Rock hammer for scale. (f) Carbonate melt-bearing clast-rich impact melt rocks from the Haughton impact structure. Penknife for scale. This lithology was originally interpreted as a clastic or fragmental breccia (Redeker and Stöffler, 1988), but was subsequently shown to be an impact melt-bearing impactite (Osinski and Spray, 2001; Osinski et al., 2005b).
The transient compression, decompression and heating of the target rocks lead to shock metamorphic effects (see Chapter 8 for an overview), which record pressures, temperatures and strain rates well beyond those produced in terrestrial regional or contact metamorphism (Fig. 1.8 and Fig. 1.9). Given the highly transient nature of shock metamorphic processes, disequilibrium and metastable equilibrium are the norm. The only megascopic shock features are shatter cones, which are distinctive, striated and horse-tailed conical fractures ranging in size from millimetres to tens of metres (Fig. 1.11a). The most-documented shock metamorphic feature is the occurrence of so-called planar deformation features, particularly in quartz (Fig. 1.11b), although they do occur in other minerals (e.g. feldspar and zircon). When fresh, the planar deformation features are parallel planes of glass, with specific crystallographic orientations as a function of shock pressures of approximately 10–35 GPa. At higher pressures, the shock wave can destroy the internal crystallographic order of feldspars and quartz and convert them to solid-state glasses, which still have the original crystal shapes. These are ‘diaplectic’ glasses (Fig. 1.11c,d), with the required pressures being 30–45 GPa for plagioclase feldspar (also known as maskelynite) and 35–50 GPa for quartz. The extremely rapid compression and then decompression also produces metastable polymorphs, including coesite and stishovite from quartz and diamond and lonsdaleite from graphite (Chapter 8).
Figure 1.11 Shock metamorphic effects in rocks and minerals. (a) Shatter cones in limestone from the Haughton impact structure. Penknife for scale. (b) Planar deformation features in quartz.
Image courtesy of L. Ferriëre.
(c) and (d) Plane- and cross-polarized light photomicrographs, respectively, of diaplectic quartz glass from the Haughton impact structure, Canada. Note the original grain shape of the sandstone quartz grains has been preserved, which is diagnostic for diaplectic glass, but not whole rock glasses.
1.4.1 Classification of Impactites
As part of the IUGS Subcommission on the Systematics of Metamorphic Rocks, a study group formulated a series of recommendations for the classification of impactites (Stöffler and Grieve, 2007). This group suggested that impactites from a single impact should be classified into three major groups irrespective of their geological setting: (1) shocked rocks, which are non-brecciated, melt-free rocks displaying unequivocal effects of shock metamorphism; (2) impact melt rocks (Fig. 1.10a–c), which can be further subclassified according to their clast content (i.e. clast-free, -poor or -rich) and/or degree of crystallinity (i.e. glassy, hypocrystalline or holocrystalline); (3) impact breccias (Fig. 1.10d,e), which can be further classified according to the degree of mixing of various target lithologies and their content of melt particles (e.g. lithic breccias and ‘suevites’).
It is apparent from the literature that substantial problems exist with the current IUGS nomenclature of impactites, particularly those including impact melt products (see Chapters 7 and 9 for detailed discussions). This is due to several reasons, including the erosional degradation of many impact structures on Earth such that outcrops of impact melt-bearing lithologies preserving their entire original context are relatively rare (Grieve et al., 1977). Other complicating factors are introduced due to inconsistent nomenclature and unqualified use of terms (such as ‘suevite’ – Fig. 1.10d) for several types of impactites with somewhat different genesis; for example, impactites with glass contents ranging up to approximately 90 vol.% have been termed suevites at the Popigai impact structure (Masaitis, 1999). It is also important to note that the framework for the IUGS classification scheme was developed in the 1990s and remained little changed up to its publication in 2007, despite several major discoveries and advancements in our understanding of impactites. In particular, in recent years, the effect(s) of target lithology on various aspects of the impact cratering process, in particular the generation and emplacement of impactites, has emerged as a major research topic (Osinski et al., 2008a).
1.4.2 Impact Melt-Bearing Impactites
The production of impact melt rocks and glasses is a diagnostic feature of hypervelocity impact, and their presence, distribution and characteristics have provided valuable information on the cratering process (Dence et al., 1977; Grieve et al., 1977; Grieve and Cintala, 1992) – see Chapter 9. Within complex impact structures formed entirely in crystalline targets, coherent impact melt rocks or ‘sheets’ are formed. These rocks can display classic igneous structures (e.g. columnar jointing) and textures (Fig. 1.10a–c). Impact craters formed in ‘mixed’ targets (e.g. crystalline basement overlain by sedimentary rocks) display a wide range of impact-generated lithologies, the majority of which were typically classed as ‘suevites’ (Fig. 1.10d; Stöffler et al., 1977; Masaitis, 1999); the original definition of a suevite is a polymict impact breccia with a clastic matrix/groundmass containing fragments and shards of impact glass and shocked mineral and lithic clasts (Stöffler et al., 1977). Minor bodies of coherent impact melt rocks are also sometimes observed, often as lenses and irregular bodies within larger bodies of suevite (e.g. Masaitis, 1999). In impact structures formed in predominantly sedimentary targets, impact melt rocks were not generally recognized, with the resultant crater-fill deposits historically referred to as clastic, fragmental or sedimentary breccias (Masaitis et al., 1980; Redeker and Stöffler 1988; Masaitis 1999). These observations led to the conclusion that no, or only minor, impact melt volumes are apparently present in impact structures formed in predominantly sedimentary targets. However, more recent work suggests that impact melting is more common in sedimentary targets than has, hitherto, been believed and that impact melt rocks are produced (Fig. 1.10f; Osinski et al., 2008b). These observations are generally consistent with numerical modelling studies (Wünnemann et al., 2008), which also suggests that the volume of melt produced by impacts into dry porous sedimentary rocks should be greater than that produced by impacts in a crystalline target. Thus, it seems that the basic products are genetically equivalent regardless of target lithology, but they just appear different. That is, it is the textural, chemical and physical properties of the products that vary (Osinski et al. 2008a,b); for example, compare Fig. 1.10f with Fig. 1.10b.
1.5 Recognition of Impact Craters
Several criteria may be used to recognize hypervelocity impact structures, including the presence of a crater form and/or unusual rocks, such as breccias, melt rocks and pseudotachylite; however, on their own, these indicators do not provide definitive evidence for a meteorite impact structure. Geophysics can also provide clues (see Chapter 14), and a geophysical anomaly is often the first indicator of the existence of buried structures. The most common geophysical anomaly is a localized low in the regional gravity field, due to lowering of rock density from brecciation and fracturing (Pilkington and Grieve, 1992). Larger complex impact structures tend to have a central, relative gravity high, which can extend out to approximately half the diameter of the structure. In terms of magnetics, the most common expression is a magnetic low, with the disruption of any regional trends in the magnetic field. This is due to an overall lowering of magnetic susceptibility and the randomizing of pre-impact lithologic trends in the target rocks (Pilkington and Grieve, 1992). Seismic velocities are reduced at impact structures, due to fracturing, and reflection seismic images are extremely useful in characterizing buried structures in sedimentary targets. There is, however, no geophysical anomaly that can provide definitive evidence for a meteorite impact structure.
The general consensus within the impact community is that unequivocal evidence for hypervelocity impact takes the form of shock metamorphic indicators (French and Koeberl, 2010), either megascopic (e.g. shatter cones) or microscopic (e.g. planar deformation features, diaplectic glass) (Fig. 1.11
