Fossil earthquakes: the formation and preservation of Pseudotachylytes

Lecture Notes in Earth Sciences Editors: S. Bhattacharji, Brooklyn H. J. Neugebauer, Bonn J. Reitner, Góttingen K. St¨...

Author: Aiming Lin

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Lecture Notes in Earth Sciences Editors: S. Bhattacharji, Brooklyn H. J. Neugebauer, Bonn J. Reitner, Góttingen K. Stüwe, Graz Founding Editors: G. M. Friedman, Brooklyn and Troy A. Seilacher, Tübingen and Yale

111

A. Lin

Fossil Earthquakes: The Formation and Preservation of Pseudotachylytes With 217 Figures 33 Tables

Aiming Lin Shizuoka University Graduate School of Science & Technology 836 Ohya Suruga-ku Shizuoka 422-8529 Japan

Library of Congress Control Number: 2007937505 “For all Lecture Notes in Earth Sciences published till now please see final pages of the book”

ISSN 0930-0317 ISBN 978-3-540-74235-7 Springer Berlin Heidelberg New York This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer. Violations are liable for prosecution under the German Copyright Law. Springer is a part of Springer Science+Business Media springer.com c Springer-Verlag Berlin Heidelberg 2008 The use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. Cover design: WMXDesign GmbH, Heidelberg Typesetting:by the authors and Integra using a Springer LATEX macro package Printed on acid-free paper

SPIN: 12054459

543210

Preface

Most books on earthquakes that are written by geophysicists focus on seismotectonics and analyses of earthquake waves recorded on seismographs in terms of seismic-source parameters such as seismic moment, focus location and depth, and rupture parameters. In contrast, traditional textbooks of tectonics and structural geology that are written by geologists are generally based on the principles of geology in terms of both their subject matter and the approach taken to studying the evolution of Earth. While Yeats et al. (1997) wrote a comprehensive textbook on the geology of earthquakes, including coverage of active global tectonics and paleoseismic studies, we have yet to see a book that focuses on earthquake-source materials that are produced or deformed by both seismic faulting and aseismic creep within seismogenic fault zones at different levels in the crust. The current book, Fossil Earthquakes: Formation and Preservation of Pseudotachylytes, addresses this shortcoming, focusing on the mechanisms and processes of formation of pseudotachylyte and related earthquake materials that form within natural fault zones and those that are generated artificially in high-velocity frictional experiments. The content of the book is largely based courses that I taught on Earthquake Geology and Structural Geology, beginning as lecture notes for undergraduate and graduate students of Earth Science at Shizuoka University, Japan. I hope that the book helps to bridge the gap between seismology and geology and that it encourages further studies of earthquake mechanisms and seismic faulting processes. The topics covered in this book encompass the principal results of field investigations, analyses of meso-scale and micro-scale textures and structures, laboratory experiments, chemical analyses, and conceptual fault models, leading to an analysis of the implications of fault-related pseudotachylyte and related earthquake materials in terms of our understanding of earthquakes themselves. The book is organized into twelve chapters. Chapter 1 is an introduction to pseudotachylyte and related fault rocks, while Chap. 2 presents the relevant terminology and reviews the historical controversy regarding the physical

VI

Preface

origin of pseudotachylyte. Chapter 3 is devoted to pseudotachylyte-related fault rocks, with a strong emphasis on the role of fabrics within fault rocks and the development of a conceptual fault-zone model. In the core of the book, Chaps. 4 to 7 deal with the tectonic environment, macro- to micro-scale structures, petrologic properties, and formation mechanisms of melt-origin pseudotachylyte, respectively, based on representative examples from the main faults worldwide from which pseudotachylyte has been reported, including the Outer Hebrides Thrust, England; the Woodroffe Thrust, Australia; the Fuyun Fault, China; and the Alpine Fault, New Zealand. The chemical composition of pseudotachylyte is considered in Chap. 8, and Chap. 9 explores mylonite- and granulite-associated pseudotachylyte that forms in deep-level fault shear zones within the semi-brittle to crystal-plastic regimes from the Woodroffe Thrust, Australia and Dahezhen Fault Shear Zone within an ultrahigh pressure complex in the Qinling-Dabie Shan collisional orogenic belt, China. Chapter 10 describes the meso- to micro-scale structures and petrologic properties of crushing-origin pseudotachylyte and related veinlet cataclastic rocks in the context of their occurrence as fossil earthquakes based on the representative examples from the Iida-Matsukawa Fault, the Nojima Fault, and the Itoigawa-Shizuoka Tectonic Line Active Fault System, Japan, as well as exploring their formation mechanisms. Chapter 11 details two representative examples of landslide-generated melt-origin pseudotachylyte: one from the Langtang Himalaya, Nepal, and another from Chiufener-Shan, Taiwan, with the latter being related to the 1999 Mw 7.6 Chi-Chi earthquake. Finally, Chap. 12 presents the principal results of high-velocity frictional melting experiments in terms of our understanding of pseudotachylyte generation. There are many organizations and individuals who helped to make this book possible. The bulk of the material presented in this book is based on the results of research projects undertaken by the author and supported financially by the Ministry of Education, Culture, Sports, Science, and Technology of Japan. I owe my personal development as a scientist to past and current associations with a great many people. I would particularly like to thank my two thesis supervisors at the graduate school of the University of Tokyo: T. Matsuda, who introduced me to fault rocks related to active faults, and T. Shimamoto, who advised me on the details of analyzing pseudotachylyte and related fault rocks, as well as high-velocity frictional melting experiments. Many colleagues assisted in the preparation of this book. I would like to thank K. Arita and H. Takagi for kindly providing samples and photographs of the Langtang Himalaya landslide-related pseudotachylyte, and I greatly appreciate assistance in the field provided by S. Ge (Fuyun Fault, China), Z. Sun (Qinling-Dabie Shan collisional orogenic belt, China), A. Camacho (Woodroffe Thrust, Australia), A. Stallard, (Alpine Fault, New Zealand), O. Fabbri (Outer Hebrides Thrust, Scotland and Saint-Barthélemy Massif, Pyrenees, France), and E. Ferre (Santa Rose mylonite shear zone, Southern

Preface

VII

California, USA). I am also grateful to A. Stallard for improving the English of the manuscript. I dedicate this book to my family, especially my wife Sujuan, who provided the comfortable environment in my personal life that helped to make this book a reality.

Contents

1

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1

2

Terminology and Origin of Pseudotachylyte . . . . . . . . . . . . . . . . 2.1 Terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2 Controversy Regarding the Physical Origin of Pseudotachylyte

5 5 8

3

Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2 Fault Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2.1 Classification of Fault Rocks . . . . . . . . . . . . . . . . . . . . . . . . 3.2.2 Mylonitic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2.3 Cataclastic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2.4 Formation of S-C Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3 Fault Zone Strength and Fault Model . . . . . . . . . . . . . . . . . . . . . . 3.3.1 Seismogenic Fault Zone Strength . . . . . . . . . . . . . . . . . . . . 3.3.2 Conceptual Fault Zone Model . . . . . . . . . . . . . . . . . . . . . . .

17 17 18 18 23 25 39 40 40 43

Tectonic Environment and Structure of Pseudotachylyte Veins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte 4.1.1 Tectonic Environment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.2 Field Occurrence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.3 Chilling-margin and Crack Textures . . . . . . . . . . . . . . . . . 4.2 Classification of Pseudotachylyte Veins . . . . . . . . . . . . . . . . . . . . . 4.2.1 Fault Veins and Injection Veins . . . . . . . . . . . . . . . . . . . . . 4.2.2 Pseudotachylyte Generation Zones . . . . . . . . . . . . . . . . . . . 4.3 Relation Between Fault Vein Thickness and Slip Amount . . . . .

47 47 47 48 55 60 60 64 70

4

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5

Contents

Pseudotachylyte Matrix . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2 Microstructural Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.1 Textural Classification of Pseudotachylyte Matrix . . . . . 5.2.2 Flow Structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.3 Vesicles and Amygdules . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3 Powder X-Ray Diffraction Analysis . . . . . . . . . . . . . . . . . . . . . . . . 5.3.1 X-Ray Diffraction Patterns for Pseudotachylyte . . . . . . . 5.3.2 Quantitative Analysis of Glass and the Crystalline Fraction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.3 Quantitative Analysis of Crystalline Material . . . . . . . . . 5.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.4.1 Properties of Glass and Glassy Matrix . . . . . . . . . . . . . . . 5.4.2 Effect of Frictional Melt on Fault Strength . . . . . . . . . . . 5.4.3 Estimation of the Formation Depth of Pseudotachylyte .

75 75 76 76 81 84 90 90 93 95 96 96 97 98

6

Microlites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 6.2 Texture and Morphology of Microlite . . . . . . . . . . . . . . . . . . . . . . 106 6.2.1 Texture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 6.2.2 Morphology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 6.3 Microlite Chemistry and Magnetic Properties . . . . . . . . . . . . . . . 118 6.3.1 Microlite Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 6.3.2 Magnetic Properties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128 6.4 Discussion of the Mechanism of Microlite Formation . . . . . . . . . 132

7

Fragments Within Pseudotachylyte Veins . . . . . . . . . . . . . . . . . . 139 7.1 Terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 7.2 Fragments that Resemble Conglomerate Clasts . . . . . . . . . . . . . . 139 7.3 Grain-size Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 7.3.1 Grain-size Distribution Within Melt-origin Pseudotachylyte . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 7.3.2 Grain-size Distribution: A Discussion . . . . . . . . . . . . . . . . 148 7.4 Fabrics of Fragments and Degree of Rounding . . . . . . . . . . . . . . . 151 7.4.1 Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 151 7.4.2 Degree of Rounding of Fragments . . . . . . . . . . . . . . . . . . . 151 7.5 Formation of Rounded Fragments: A Discussion . . . . . . . . . . . . . 155

8

Chemical Composition and Melting Processes of Pseudotachylyte . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 8.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 8.2 Bulk-Vein and Matrix Compositions . . . . . . . . . . . . . . . . . . . . . . . 160 8.2.1 Bulk Composition of Pseudotachylyte Veins . . . . . . . . . . 160 8.2.2 Chemical Composition of Pseudotachylyte Matrix . . . . . 162 8.2.3 Water Contents of Pseudotachylyte Veins . . . . . . . . . . . . . 168

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8.3 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 8.3.1 Melting Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 8.3.2 Melt Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171 8.3.3 Role of Water During Frictional Melting . . . . . . . . . . . . . . 173 9

Formation of Pseudotachylyte in the Brittle and Plastic Regimes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 177 9.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 177 9.2 Woodroffe Pseudotachylytes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 9.2.1 Tectonic Setting of the Woodroffe Thrust . . . . . . . . . . . . . 179 9.2.2 Field Occurrences of the Woodroffe Pseudotachylytes . . 181 9.2.3 Microstructures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 187 9.3 Dahezhen Pseudotachylytes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 197 9.3.1 Tectonic Setting of the Dahezhen Shear Zone . . . . . . . . . 197 9.3.2 Field Occurrence of the Dahezhen Pseudotachylytes . . . 198 9.3.3 Microscopy and Chemical Composition . . . . . . . . . . . . . . . 204 9.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 212 9.4.1 Formation Mechanisms of Large Volumes of Pseudotachylytes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 212 9.4.2 Conditions of Formation of the Dahezhen and Woodroffe M-Pt Veins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 216

10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 225 10.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 225 10.2 Occurrence of Crushing-Origin Pseudotachylyte and Cataclastic Veins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 226 10.2.1 Crushing-Origin Pseudotachylyte . . . . . . . . . . . . . . . . . . . . 226 10.2.2 Fault-Gouge Injection Veins . . . . . . . . . . . . . . . . . . . . . . . . 230 10.2.3 Layered Fault Gouge and Pseudotachylyte Veins . . . . . . 232 10.2.4 Crack-Fill Veins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 234 10.3 Petrologic Characteristics of Veinlet Cataclastic Rocks . . . . . . . 237 10.3.1 Microstructures of Veinlet Cataclastic Rocks . . . . . . . . . . 237 10.3.2 Powder X-ray Diffraction Analysis of Veinlet Material . . 244 10.3.3 Chemical Composition Data and Isotope Analyses . . . . . 250 10.3.4 Age Data for Crack-fill Veins . . . . . . . . . . . . . . . . . . . . . . . 252 10.4 Discussion on the Formation Mechanisms of Veinlet Cataclastic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 253 10.4.1 Formation Mechanism of Amorphous Material Within Veinlet Cataclastic Rocks . . . . . . . . . . . . . . . . . . . . 253 10.4.2 Coseismic Fluidization of Fine-grained Material Within Fault Zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 254 10.4.3 Repeated Events of Seismic Slip . . . . . . . . . . . . . . . . . . . . . 256 10.4.4 Repeated Coseismic Infiltration of Surface Water into Deep Fault Zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257

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11 Landslide-related Pseudotachylyte . . . . . . . . . . . . . . . . . . . . . . . . . 265 11.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 265 11.2 Occurrences of Landslides and Related Pseudotachylytes . . . . . 266 11.2.1 Langtang Himalaya Landslide and Related Pseudotachylyte . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 266 11.2.2 Chiufener-Shan Landslide and Related Pseudotachylyte 269 11.3 Petrographic Characteristics of Landslide-related Pseudotachylytes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 274 11.3.1 Petrography of the Langtang Himalaya Pseudotachylyte 274 11.3.2 Petrography of the Chiufener-Shan Pseudotachylyte . . . 277 11.3.3 Glass Contents of the Observed Pseudotachylytes . . . . . . 279 11.4 Discussion of the P-T Conditions during the Formation of Landslide-related Pseudotachylyte . . . . . . . . . . . . . . . . . . . . . . . . . 280 12 Experimentally Generated Pseudotachylyte . . . . . . . . . . . . . . . . 283 12.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 283 12.2 High-Velocity Frictional Experiments . . . . . . . . . . . . . . . . . . . . . . 284 12.2.1 Test Equipment and Experimental Conditions . . . . . . . . 284 12.2.2 Experiment Samples and Procedures . . . . . . . . . . . . . . . . . 290 12.2.3 High-Velocity Frictional Properties . . . . . . . . . . . . . . . . . . 292 12.3 Microstructures of Experimentally Generated Pseudotachylyte 293 12.3.1 Textures of the Fault Shear Plane . . . . . . . . . . . . . . . . . . . 293 12.3.2 Vein Geometry and Texture of Molten Material . . . . . . . 294 12.4 Powder X-ray Diffraction Analysis of Run Products . . . . . . . . . . 300 12.4.1 Diffraction Patterns of Run Products . . . . . . . . . . . . . . . . 300 12.4.2 Quantitative Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 301 12.5 Chemical Composition Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 304 12.5.1 Gabbro Samples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 304 12.5.2 Granite Samples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 307 12.5.3 Albitite–Quartz and Anorthosite–Anorthosite Pairs . . . . 308 12.6 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 315 12.6.1 Vein Geometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 315 12.6.2 Melting Textures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 315 12.6.3 Non-equilibrium Melting Processes . . . . . . . . . . . . . . . . . . 316 12.6.4 Melting Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 318 12.6.5 High-Velocity Slip Weakening . . . . . . . . . . . . . . . . . . . . . . . 319 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 321 Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 341

1 Introduction

It is well known that direct evidence of earthquakes within fault zones is limited to the occurrence of tectonic-generated pseudotachylyte. Pseudotachylyte is formed when frictional heat (e.g., McKenzie and Brune 1972; Sibson 1975; Spray 1987, 1995; Lin 1991, 1994a, b; Lin and Shimamoto 1998) and strong abrasion that are generated during rapid seismic faulting are sufficient to melt and/or crush rock within the fault zone and fluidize ultrafine-grained material (e.g., Lin 1996, 1997a; Kano et al. 2004). Since the 1970s, fault-related pseudotachylyte has become widely accepted as an indicator of high-velocity slip during earthquakes that occurred within seismogenic fault zones (e.g., Francis 1972; Sibson 1975, 1980a; Passchier 1982; Lin 1991, 1994a, b; Magloughlin 1992; McNulty 1995; Lin et al. 2003b, 2005b); consequently, pseudotachylyte can be thought of as a fossil earthquake. Although pseudotachylyte can also form from meteorite impact, this book focuses solely on tectonic pseudotachylytes which are related with earthquakes. When I was a graduate student in the 1980s, pseudotachylyte was a curious term that was not well known, even within the Earth science community in Japan. As my doctoral thesis involved a study of the origin of fault-related pseudotachylyte, puzzled colleagues often asked me “what is pseudotachylyte”. In the 1990s, however, pseudotachylyte became a familiar word, not just in the Japanese Earth science community, but also in the mass media. A Japanese science fiction movie was even filmed called Ultraman-daina, depicting a monster bird from Chinese mythology fought against Ultraman with the aid of the immense seismic energy of pseudotachylyte sourced from a large earthquake that occurred within the Chinese continent (Fig. 1.1). One of my thesis supervisors, Professor T. Shimamoto of the Institute of Earthquake Research, University of Tokyo, secured research funding from the Japanese Government to investigate the origin of pseudotachylyte, and I became involved in this project in 1989. The results of this project were published in a special issue of the journal Structural Geology (The Journal of Tectonic Research Group of Japan, 1994, Vol. 39) and

2

1 Introduction

Fig. 1.1. Cartoon image from the Japanese science fiction movie Ultraman-daina showing a monster bird from Chinese mythology fighting Ultraman (inside the shuttle plane) with the aid of the immense seismic energy of pseudotachylyte sourced from a large earthquake that occurred within the Chinese continent. Pt: Pseudotachylyte, W: seismic wave. Image courtesy Y. Lin

were widely introduced to the Seismological Society of Japan (Yoshioka 1994). This event marked the widespread acceptance of the term pseudotachylyte in Japan. Although many studies have investigated the nature and significance of pseudotachylyte over the past century, fault-related pseudotachylyte is still rarely found within exhumed fault zones throughout the world. It remains unclear as to whether this scarcity is merely apparent or whether seismic frictional melting within fault zones is inhibited by other mechanisms (Sibson 2002). It is also unclear as to whether fault-related pseudotachylyte can form from the passage of shock waves associated with hypervelocity impact, as with impact-originated pseudotachylyte (e.g., Francis 1972; Spray 1995). Frictional melt may well play an important role during coseismic slip as a lubricant on the fault plane (Lin et al. 2001a; Spray 2004; Di Toro et al. 2006).

1 Introduction

3

Current research on pseudotachylyte and related fault rocks will provide improved insight into the generation of earthquakes and the process of seismic rupture within fault zones located with the brittle and crystal-plastic regimes of the upper and lower crust.

2 Terminology and Origin of Pseudotachylyte

2.1 Terminology The term pseudotachylyte was first introduced by Shand (1916) to describe dark, aphanitic, glassy, and dike-like rocks that occur as veins and networks in the Parijs region of South Africa (Figs. 2.1 and 2.2). These rocks are located within the Vredefort Dome structure, which is one of the largest meteorite impact structures on Earth. Shand explained his choice of the term pseudotachylyte as follows: “. . . that these rocks have great similarity to tachylyte, also that such rocks have been mistaken for trap and tachylyte in Scotland and India as well as in South Africa, and for the further reason that no more suitable name is in existence.” The term pseudotachylyte was therefore originally used to describe a rock with a similar dark aphanitic appearance to a form of glassy basaltic rock known as tachylyte, although with different physical and chemical origins. It is clear that this original definition was intended to be used in the field rather than representing a rigorous petrologic name based on chemical composition or physical origin, as currently used in petrology, although Shand (1916) did favor a melting origin. In the Glossary of Geology (Jackson 1997), pseudotachylyte is defined as “a dense rock produced in the compression and shear associated with intense fault movements, involving extreme mylonitization and/or partial melting.” If we substitute cataclasis for mylonitization in the above definition, it would then be consistent with the current state of knowledge regarding pseudotachylyte, as mylonitization is an aseismic crystal-plastic mode of deformation that commonly overprints pseudotachylyte rather than a primary mechanism of pseudotachylyte generation (see Chap. 4 for further details). In fact, the pseudotachylyte veins reported to date generally show the distinct characteristics

6


Fig. 2.1. Photograph of pseudotachylyte-bearing zones (viewed in a vertical section) at Vredefort meteorite impact site, South Africa. There are at least three irregular horizontal pseudotachylyte-bearing zones (dark areas indicated by arrows) at intervals of 3–10 m; these zones are each made up of a network of numerous pseudotachylyte veins. The individual pseudotachylyte-bearing zones range in thickness from several tens of centimeters to several meters. The host rock is Archean granite. For scale, a person is standing in the lower center of the photograph

of both sedimentary rocks composed of numerous fragments or conglomeratic fragments set in a fine-grained matrix and intrusive igneous veins injected in the country rock in the same manner as dikes (Fig. 2.2). These types of rocks have been described using a number of different names since they were first documented toward the end of the 19th century, e.g., flinty crushrocks (Clough 1888; Clough et al. 1909), trap-shotten gneiss (Holland 1900), injection mylonite (Philpotts 1964), microlitic mylonite (Wallace 1976), and hyalomylonite (Scott and Drever 1953; Wallace 1976; Masch et al. 1985). These earlier studies attempted to distinguish the physical origin of pseudotachylyte from that of related tectonic rocks by using specific names; however, this approach led to confusion concerning nomenclature, and all of these earlier names have since been abandoned. Researchers who study fault-related rocks have had to deal with problems arising from the ambiguous definition of the term pseudotachylyte. Magloughlin and Spray (1992) advocated that the use of the term pseudotachylyte should be restricted to pseudotachylyte veins that originate from melting processes; however, other researchers have speculated that a gradation

2.1 Terminology

7

Fig. 2.2. Photograph of a network of pseudotachylyte veins (viewed in a vertical section) at Vredefort meteorite impact site, South Africa. The larger veins dipping at high angles are connected to numerous smaller veins to form a network within the Archean granitic host rocks. The hammer shown for scale is 35 cm in length

exists between melt-origin pseudotachylyte and pseudotachylyte-like cataclastic veinlets that form with little or no frictional melting (e.g., Philpotts 1964; Francis 1972; Wenk 1978; Lin 1989, 1996, 1997a, 2001; Shigetomi and Lin 1999; Wenk et al. 2000; Kano et al. 2004; Rowe et al. 2005). The proportion of host rock fragments to molten material within pseudotachylyte veins varies not only between different veins within the same fault zone but also between different parts of an individual vein (Lin 1997a). Clarification and classification of the physical origin of pseudotachylyte is complicated by its very fine-grained nature, the common presence of devitrified and recrystallized material and rock fragments, and the obscuring effects of subsequent deformation, alteration, and metamorphism. This is a major reason for the fact that the origin of pseudotachylyte had been disputed for over a century since it was first described toward the end of the 19th century. If the use of the term pseudotachylyte is to be restricted to those veins with a melting-related origin, it will then be necessary to formally define the term pseudotachylyte in terms of the proportion of melt (glassy matrix) within the vein. It is, however, difficult to quantify the melt content of pseudotachylyte veins because of the following two factors: i) large numbers of ultra-fine-grained fragments derived from the host rock are too small to be separated from the matrix, even with the aid of a microscope, and ii) devitrification commonly occurs during subsequent metamorphism and alteration. In fact, the term pseudotachylyte has been used to describe both melting- and

8


crushing-origin pseudotachylyte that has the general appearance of a dark aphanitic veinlet (e.g., Philpotts 1964; Irouschek and Huber 1982; Lin 1989, 1996, 1997a). Philpotts (1964) used the term pseudotachylyte to describe two different types of veinlets with a dark aphanitic appearance: one composed almost entirely of fine-grained fragments of wall-rock and another composed mainly of glassy melt material and fine-grained fragments of wall-rock. Given the difficulties involved in determining the physical origin of pseudotachylyte, Irouschek and Huber (1982) advocated the use of the term for all finegrained fault-related rocks of unknown origin. In a study of pseudotachylyte within granitic rocks along the active Iida–Matsukawa Fault, Central Japan (see Chap. 10 for details), Lin (1989, 1996, 1997a) used the term to describe dark aphanitic veinlets and vein networks that are similar in appearance to typical melt-origin pseudotachylyte but are in fact composed almost entirely of fine-grained fragments of the host rock. Wenk et al. (2000) used the term pseudotachylyte to describe “extremely fine-grained concordant or discordant veins that are injected into host rock and does not necessarily imply melting”; the authors applied the term to both melting- and crushing-origin pseudotachylyte found in the Santa Rose mylonite shear zone, Southern California, USA. For distinguishing the tectonic pseudotachylyte from Impact-generated pseudotchylyte, they also suggested to use a name such as ‘seismite’ to describe the rocks formed by cataclasis and melting in association with tectonic seismic activity. Two spellings are currently in use: pseudotachylyte and pseudotachylite. On the basis of the above review, and to avoid confusion in terms of its physical origin, in this book we use the term pseudotachylyte as a petrographic and field term to describe all dark-brown, aphanitic, and dike-like rocks of unknown origin that occur as veins and vein networks; this is consistent with the original definition proposed by Shand (1916) and applies to both meltingand crushing-origin pseudotachylytes. The term fault-related pseudotachylyte is defined as a dense, dark-brown, aphanitic, veinlet rock that formed from extreme cataclasis and/or partial melting within a fault zone.

2.2 Controversy Regarding the Physical Origin of Pseudotachylyte Most current researchers consider pseudotachylyte to be a typical product of frictional melting, but a majority consensus on its physical origin was only reached in the early 1990s after more than a century of debate. The major issue of controversy was whether pseudotachylyte was generated by melting along fault planes or formed as a consequence of tectonic events within fault zones. This is an important topic for solid-Earth science because any molten material formed by frictional heating in fault zones can be used as an indicator of paleoseismic events and to constrain earthquake source parameters such as dynamic frictional heating (McKenzie and Brune 1972), slip-weakening

2.2 Controversy Regarding the Physical Origin of Pseudotachylyte

9

distance (Hirose and Shimamoto 2005), shear stress resistance (Sibson 1975; O’Hara et al. 2006; Di Toro 2005), and fault-slip processes (Lin et al. 2003b, 2005b; Spray 2004; Di Toro et al. 2006). Shand (1916) described in detail the occurrences of intrusive veinlet pseudotachylyte that comprises distinctive conglomerate and breccia set in a dark, fine-grained matrix (Figs. 2.3 and 2.4). In analyzing the Vredefort pseudotachylyte, Shand found rounded and highly irregular embayed fragments, feldspar microlites and spherulites, and flow structures. In the absence of any obvious evidence of fault-related shearing, he concluded that “the pseudotachylyte originated from granite itself through melting caused not by shearing but by shocking, or alternatively, by gas flexing.” Figure 2.5 shows the typical mode of these injection veins of pseudotachylyte, for which there is no distinct displacement and no shearing structures in the granitic wall rock adjacent to the veins. Prior to Shand’s study, similar rocks had been described at the end of the 19th century and early 20th century at many localities along the Outer Hebrides Thrust Fault Zone in Scotland, associated with cataclastic flinty crush (e.g., Clough 1888, 1909). While working on rocks in India, it was Holland (1900) who first suggested that pseudotachylyte might form by the melting of rock via mechanically generated heat that arose from the confinement of dislocations to narrow bands. He described in details the

Fig. 2.3. Photograph of an individual pseudotachylyte-bearing zone (viewed in a horizontal section) at Vredefort meteorite impact site, South Africa. The hammer shown for scale is 35 cm in length

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Fig. 2.4. Photograph of a conglomerate-bearing network of pseudotachylyte veins (viewed in a vertical section) at Vredefort meteorite impact site, South Africa. Most of the larger boulders and smaller fragments of the host granite gneiss are rounded to subrounded and cemented by a dark matrix (pseudotachylyte), indicating a melting origin. The hammer shown for scale is 35 cm in length

petrographic characteristics and field occurrences of the rocks and elucidated that: The so-called “trap-shotten” bands coincided with lines of dislocation, and the black tongues and films which superficially resemble compact “trap” have the microscopical characters of mylonite which has been hardened-fritted and rarely half-fused-by the heat generated through the dislocation being confined to narrow bands, and thereby causing a higher local rise of temperature than would result from a general deformation of the rock-mass. During the early 20th century, many studies reported on pseudotachylyte observed along the great thrust faults developed within Lewisian gneiss in the Outer Hebrides, northwest Scotland (e.g., Clough 1909; Jehu and Craig 1923). Fieldwork in this area revealed that the thrusting was associated with movement along the Moine Thrust, thereby leading to the formation of pseudotachylyte. These early studies failed to comment on the origin of pseudotachylyte, with the exception of Holland’s (1900) suggestion. Waters and Campbell (1935), however, expressed doubt on the fusion origin of pseudotachylyte. They compared the textural features of mylonitic rocks from the San Andreas Fault, California, USA, with the Vredefort pseudotachylyte and concluded that both rocks were probably produced by extreme crushing rather than melting. This deduction was later supported by


11

Fig. 2.5. (a) Photograph of simple pseudotachylyte veins (indicated by white arrows) in the gate pillow-stone at Garden Park, Parys, South Africa. (b) Close-up photograph of the vein shown in (a). Note that there is no observable offset of the host rock across the pseudotachylyte vein in the center of the photograph (b). The horizontal ribs are the boring hole-walls for quarrying. The pen shown for scale is 15 cm long

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Willemse (1937), who re-examined the Vredefort pseudotachylyte using powder X-ray diffraction analysis and found that the rock was more typical of an extremely fine-grained crystalline powder than glass. Willemse clarified the crystalline nature of the aphanitic pseudotachylyte matrix, which up to then had been reported in many studies as glassy material or glass (such as volcanic glass) based on observations made using traditional optical microscopes. In a study of a thrust plane in the Himalayas, Scott and Drever (1953) described a truly glassy vein-like rock containing numerous vesicles and amygdule structures. The vein was gradational on both sides into the host granitic rock. This evidence represented the first conclusive proof that frictional melting can be generated upon a sliding plane, although later studies showed that the glassy pseudotachylyte vein described by Scott and Drever formed by landsliding rather than seismic faulting (Masch and Preuss 1977; Masch 1979; Masch et al. 1985; also see Chap. 11 for details). On the basis of a study of the Vredefort pseudotachylyte and other geological materials, Reynolds (1954) proposed a fluidized solid–gas system to explain the formation mechanism of pseudotachylyte veins and networks. Reynolds concluded that “the pseudotachylytes were composed of finely ground particles transported to their present locations as a suspension in a rapidly moving gas.” This process was used to explain the extreme mobility indicated by the commonly observed networks of intricate and narrow pseudotachylyte veins. The hypothesis of the fluidization and injection of fine-grained material is supported by the characteristic mode of veinlet cataclastic rocks within active fault zones in Japan, including fault gouge, fault breccia, and pseudotachylyte, which are almost entirely composed of fine-grained fragments and occur within massive rocks without distinct shear structures adjacent to the veins (e.g., Lin 1989, 1996, 1997a; Lin 1994; Shigetomi and Lin 1999; Kano et al. 2004; also see Chap. 10 for details). These studies demonstrated that the dark aphanitic veinlet mode of cataclastic fault rocks could form from the dynamic fluidization of fine-grained material within fault zones, even in the absence of melting (see Chap. 10 for details). In a study of pseudotachylyte veins within a Lewisian basement complex at Cairloch, Northwest Scotland, Park (1961) described glassy patches or veinlets containing spherulites and radial microlites, devitrification of the matrix, irregular embayed fragments, textures within veins, and recrystallized biotite, thereby concluding that “part at the vein material was melted at the time of intrusion”. Similar evidence of melting was also found by Philpotts (1964) in a study of pseudotachylyte in Quebec, Canada, in which vesicles and amygdules were observed. Philpotts also described crushing-related pseudotachylyte veins that formed by extreme “mylonitization” of the rock and associated injection of the finely pulverized material into fractures. On the basis of the structural


13

features described by Philpotts, it is apparent that he used the term mylonitization to describe brittle cataclasis or crushing rather than to describe crystal-plastic deformation, which is the definition of mylonitization that is currently used in structural geology based on our current understanding of fault rocks. This type of pseudotachylyte veins observed by Philpotts are similar to the crushing-origin pseudotachylyte described by Lin (1996, 1997a; also see Chap. 10 for details). On the basis of field and petrological studies of pseudotachylyte developed along the Outer Hebrides Thrust, Northwest Scotland, Sibson (1975) concluded that pseudotachylyte formed by seismic-related frictional melting upon a fault plane. This conclusion is consistent with earlier proposals of a melting origin for pseudotachylyte, such as the work of Francis (1972), which was based on a review of the extensive pseudotachylyte literature published prior to the 1970s. Two papers subsequently written by Sibson (1975, 1977) on pseudotachylyte and fault rocks have received widespread acclaim. It has also been shown theoretically that once slip within a fault zone exceeds several centimeters, frictional melt should be a relatively common phenomenon at known seismic slip rates, even for shear resistances as low as 100 bars (10 MPa) (e.g., McKenzie and Brune 1972; Richards 1976; Cardwebl et al. 1978). During the 1960s and 1970s, the frictional melting hypothesis for pseudotachylyte became widely accepted by the majority of workers in the Earth science community; however, on the basis of a TEM (transmission electron microscopy) study of the fine-grained matrix within pseudotachylyte, Wenk (1978) reached the same conclusion as Waters and Campbell (1935), in that pseudotachylyte may result from the ultra-comminution of host rocks. Wenk (1978) found only small amounts of glass and rare devitrification textures in the fine-grained matrix of pseudotachylyte veins that had previously been reported as being dominantly glass or glassy material on the basis of observations under an optical microscope. Wenk contended that the microstructure of pseudotachylyte veins displays many of the features of intensive brittle deformation, thereby indicating formation via shock deformation rather than melting and subsequent quenching to a glass accompanied by devitrification during cooling. He argued that the presence of minor pointed or pocked glass or glassy material does not provide conclusive evidence for the formation of pseudotachylyte via frictional melting. This hypothesis was receptively documented by Goode (1979) and Watts and Williams (1979). Subsequently, Wenk and Weiss (1982) and Weiss and Wenk (1983) claimed that previous studies “fail to confirm the presence of glass in any pseudotachylyte vein”. This new round of controversy concerning the physical origin of pseudotachylyte began with Wenk (1978) but was largely resolved by the end of the 1980s, as many studies demonstrated that the various textures found in natural pseudotachylyte indicate an origin associated with the melting of host rocks. Such features include the rounded and irregular embayed shapes of fragments (e.g., Gupta 1967; Maddock 1983; Lin 1991, 1994a, 1999b; Magloughlin 1992; Di Toro and Pennacchioni 2004), vesicles and amygdule structures (e.g.,

14


Maddock 1983; Lin 1991, 1994a), chilled margins (e.g., Lin 1994a), the occurrence of various microlite and spherulite morphologies that form only at high temperature and with rapid cooling or quenching of a melt (Maddock 1983; Macaudière et al. 1985; Toyoshima 1990; Lin 1991, 1994a, b), and flow structures (e.g., Lin 1994a, b). TEM analysis of pseudotachylyte veins from the South Mountains, Arizona, USA, revealed the presence of spot glass within the featureless regions of pseudotachylyte matrix that are characterized by amorphous rings in diffraction patterns (Goodwin et al. 1998). The presence of glass in pseudotachylyte veins unequivocally demonstrates a melt origin for pseudotachylyte, but the occurrence of minor spot glass is not necessarily indicative of a melt origin because spot glass or/and non-crystal material is also found in pseudotachylyte veins that formed from crushing (Ozawa and Takizawa 2007); such features may be involved in part in the formation of cataclastic veins, as argued by Wenk (1978). Despite these advances and the increasing number of examples of pseudotachylyte described in the literature, the various interpretations of structures and textures found within fault-related pseudotachylyte veins meant that their origin remained debated. Glass is sometimes produced by frictional heating during rock-drilling (e.g., Bowen and Aurousseau 1923; Kennedy and Spray 1992). Glass can also form from solid-state shocking during impact events, as is well documented for terrestrial and lunar impact sites (e.g., Chao 1968; Christie et al. 1973). Experimental investigations of shock metamorphism (e.g., Tomeoka et al. 1999) have demonstrated that molten layers and veins can result from impact at contact surfaces within wall rock that are oriented parallel to the shocking surface. It is clear that glass or glassy material within pseudotachylyte indicates the generation of melt during its formation, but the converse is not true: the absence of glass or glassy material is not a diagnostic test of the melting origin for pseudotachylyte because primary glass or glassy material can be devitrified during subsequent alteration and/or metamorphism. Field investigations and powder X-ray diffraction analyses of the Fuyun pseudotachylyte, China, revealed that up to 90 wt% of the veins consist of glass or glassy material (Lin 1991, 1994a). This finding was supported by the occurrence of associated melting-related textures such as vesicles and amygdules, rounded and irregularly embayed fragments, flow structures, and various shapes of microlites and spherulites that indicate an origin from primary melt (Lin 1991, 1994a, b; Lin et al. 2002; see Chap. 4 for details). Although many of the studies published in the decades prior to the work of Lin (1991, 1994a) and Lin et al. (2001a) reported the presence of glass or glassy material in pseudotachylyte veins (e.g., White 1974; Allen 1979; Magloughlin 1992), none had produced powder X-ray diffraction data to back up this claim (a point noted by Wenk 1978 and Wenk and Weiss 1982). The discovery of up to 90 wt% glassy material within pseudotachylyte veins left little doubt as to the frictional-melting origin of fault-related pseudotachylyte and demonstrated clearly that fault-related pseudotachylyte veins observed within the


15

active Fuyun Fault zone were generated by frictional melting on the fault plane during seismic slip (Lin 1991, 1994a, b). The results of high-velocity frictional melting experiments (Spray 1987, 1988, 1993; Lin 1991; Lin et al. 1992; Lin and Shimamoto 1994, 1998; see Chap. 12 for details) and observations made during drilling projects (Killick 1990; Kennedy and Spray 1992) also indicated that frictional melting occurs readily under conditions similar to those encountered during seismic faulting, even at depths of < 30 m (Lin, 1991; Lin and Shimamoto 1998). Melting-origin textures such as microlites, vesicles, and flow structures that are observed in natural pseudotachylyte have also been successfully reproduced in highvelocity frictional melting experiments (e.g., Spray 1987; Lin 1991). Recently, even Wenk, who originally advocated a crushing-origin for pseudotachylyte, reported a typical melting-origin pseudotachylyte within the Santa Rose mylonite shear zone in Southern California, USA (Wenk et al. 2000). The controversy surrounding the physical origin of pseudotachylyte continued for almost an entire century, and was finally brought to rest at the end of the 20th century. The studies of fault-related pseudotachylyte listed above demonstrate that fault rocks can melt via coseismic frictional melting on a slip plane. In another way, crushing-origin pseudotachylyte also occurs within fault zones. Such pseudotachylyte has a dark aphanitic appearance and occurs as simple and network veins within wall rock, as with typical melting-origin pseudotachylyte (Philpotts 1964; Lin 1996, 1997a, 2001; Shigetomi and Lin 1999; Wenk et al. 2000; Kano et al. 2004). In this case, the veins are mainly composed of fine-grained fragments of the wall rock, with little or no evidence of melting. The origin of these veins is ascribed to the fluidization and injection of fine-grained material, which contains little or no melt, in a gas–solid–fluid system during seismic faulting (Lin 1996, 1997a; see Chap. 10 for details). It is clear that there exists a gradation from melting-origin pseudotachylyte, which is mostly composed of glass or glass-derived material, to crushing-origin pseudotachylyte, which consists almost entirely of fine-grained fragments of the host rock, yet also occurs as dark and aphanitic veinlets. Injection veins of both melting- and crushing-origin pseudotachylyte can be interpreted to have formed from the fluidization of melt-fragments in a gas–solid–fluid system during seismic faulting (see discussion in Chap. 10).

3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

3.1 Introduction As most large intraplate earthquakes occur as slip on mature active faults, any investigation of the seismic faulting process requires an understanding of the nature of seismogenic fault zones. Modern seismic observation techniques have enabled seismologists to accurately determine important seismic source parameters such as seismic moment, focus location and depth, rupture length, and rupture parameters for earthquakes of varying magnitude (e.g., Kikuchi and Kanamori 1996; Kikuchi 2003). Despite these advances, the resolution of seismic methods is limited at short length scales because of complex nearsurface propagation and wave-attenuation effects; consequently, it is difficult to determine details of the rupture process and deformation features of seismogenic fault zones that are relatively small and occur relatively deep within the crust (Kanamori and Heaton 1999). An alternative source of information on seismic faulting is that directly recorded by fault rocks and fault-related rocks that form within fault zones themselves. Over the past two decades, much attention has focused on fault rocks that are exposed at the Earth’s surface and those recovered from drill cores that intersected seismic fault zones with recent histories of largemagnitude earthquakes. Drill core has been recovered from the Nojima Fault in Japan, which triggered the 1995 Kobe Mw 7.2 earthquake, the Chelungpu Fault in Taiwan, which triggered the 1999 Chi-Chi Mw 7.8 earthquake, and the San Andreas Fault in the USA, which has generated numerous large earthquakes over the past century. Core samples from seismic fault zones provide fresh samples that are free from the physical and chemical weathering that occurs close to the Earth’s surface; this enable us to study structures of fault rocks that formed within seismogenic fault zones at various depths soon after the structures were generated. The seismic history of active faults over recent geological time, generally during the Late Pleistocene and Holocene, can be understood with the help of trenching surveys; however, reconstructing the earlier seismic history and

18


rupture processes of long-lived active faults requires a different approach. As major continental fault zones are commonly rooted in the middle to lower crust (e.g., Ramsey 1980), seismic slip during large earthquakes generally propagates upward to the surface where coseismic surface rupturing occurs in association with distinct offset across the fault zone and downward to deep fault zone lower than hypocenters (e.g., Lin et al. 2003a, 2005b). Fault rocks commonly provide primary evidence of the faulting history and deformation process of seismic slip at all depths from the near-surface to deep levels in the crust. Fault rocks that form at relatively deep levels, even in the lower crust, are eventually uplifted by crustal movement and exhumed by erosion. If episodic fault movement occurs throughout this process of exhumation, the fault zone will contain a variety of fault rocks that formed under different conditions ranging from the brittle regime at shallow depths to the plastic flow regime at deeper levels of the crust (Scholz 2002). It is therefore possible to gain an insight into the formation processes of fault rocks that operated throughout the faulting history by studying the structures, textures, physical properties, and chemical compositions of fault-related rocks exposed at the surface or retrieved from deep drill cores that intersected fault zones. In this way, it may also be possible to constrain geophysical models of earthquake source faults, which to date are largely based on seismological data. In this chapter, we consider fault rocks that are closely related to pseudotachylyte and that formed within seismogenic fault zones at various crustal levels from shallow depths to deep regimes in the lower crust. The formation of these fault rocks is discussed and a related fault-zone model is presented.

3.2 Fault Rocks 3.2.1 Classification of Fault Rocks The terms fault rocks and fault-related rocks are commonly used to describe tectonic rocks that formed as a result of shear deformation within fault zones. The term fault rocks was first introduced by Sibson (1977) as a collective idiom for the distinctive rock types found in zones of shear dislocation in both the upper and lower crust whose textures are thought to arise at least in part from the shearing process. This term is currently used to describe all types of deformed rocks that form by brittle and/or crystal-plastic deformation mechanisms within fault-related shear zones. The term fault-related rocks is also used as a synonym of fault rocks (Wise et al. 1984; Snoke et al. 1998). In this book, the term fault rocks is used following Sibson’s original definition, i.e., for all fault-related rocks. Fault rocks are generally considered to arise from a concentration of strain within zones of shear, which are tabular or planar zones ranging in width from several millimeters to approximately 10 km. This zone of concentrated strain, or high-strain zone, is known as a shear zone. Shear zones may occur

3.2 Fault Rocks

19

in the brittle-dominated regime, comprising discrete fault planes marked by cataclastic rocks, or within the crystal-plastically dominated shearing deformation regime where mylonitic rocks occur. Fault rocks that form in a shear zone exhibit a variety of micro- to macro-structural characteristics that are determined by their parent rocks, the original deformation environment, and their exhumation path (Snoke et al. 1998). The textures and structures of fault rocks formed pass entirely throughout faulting periods, resulting in a broad suite of deformation processes being recorded by deformation structures and textures that formed at various depths within the fault zone. These textures and structures vary according to the physical and chemical conditions within which they formed, including temperature and pressure, which are largely related to the depth of faulting. Sibson (1977) proposed a classification of fault rocks based on their characteristic textures and structures (see Table 3.1); a modified version of this scheme is described below. With the exception of pseudotachylyte, there are general trends in the character of fault rocks with increasing depth: cataclastic rocks range from gouge to breccia and cataclasite with increasing depth, while mylonitic rocks range from protomylonite to mylonite and ultramylonite. Pseudotachylyte, which forms in association with seismic faulting at all levels of the crust, is described and discussed in detail in the following chapters. In Sibson’s classification, the presence or absence of a foliation is considered to be a crucial criterion in distinguishing cataclastic rocks that formed in the brittle-dominated frictional regime from mylonitic rocks that formed in the crystal-plastic dominated regime. Prior to the 1980s, cataclastic rocks were generally considered to be fault rocks with largely randomly oriented clasts that formed at shallow depths (e.g., Higgins 1971; Sibson 1977; Wise et al. 1984). In contrast, foliated fault rocks such as mylonites were considered to be characteristic of cohesive fault rocks that formed in deep fault-related shear zones dominated by crystalplastic deformation. During the past two decades, however, foliations similar to those observed in mylonitic rocks have also been widely recognized in both incohesive cataclastic rocks such as fault gouge (e.g., Chester et al. 1985; Chester and Logan 1987; Evans 1988; Kano and Sato 1988; Lin 1996, 1997a, b, 2001; Lin et al. 2005a) and cohesive cataclasite (e.g., Lin 1989, 1996, 1997b, 1999a; Kanaori et al. 1991; Tanaka 1992; Lin et al. 1998a, b, 2005a) formed as a consequence of cataclastic deformation within the brittle regime. Figure 3.1 shows typical examples of foliated cataclastic rocks, including fault gouge, breccia, and cataclasite, obtained from a drill core that penetrated the Chelungpu Fault Zone; this fault triggered the 1999 Mw 7.8 Chi-Chi (Taiwan) earthquake and an accompanying surface rupture over a length of approximately 100 km (Lin et al. 2001c, 2005a). The foliation in such rocks is generally observed in X–Z sections (X–Z plane of the finite strain ellipsoid), cut perpendicular to the main shear plane (Fig. 3.2). The foliations that form in cataclastic rocks are generally defined by the preferred orientation of rock fragments and the asymmetric shapes of fragments and aggregates of fine-grained

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models Table 3.1. Textural classification of fault rocks (Modified from Sibson 1977)

fragments, as shown in Fig. 3.1. The results of laboratory experiments also demonstrate the fact that foliations can develop within fault gouge at shallow crustal depths (e.g., Chester et al. 1985; Noda and Shimamoto 2005). Foliations are also found in pseudotachylytes that are associated with incohesive fault gouge and breccia (e.g., Lin 1999a, 2001; Shigetomi and Lin 1999; Kano et al. 2004, see Chap. 10 for details) and cohesive mylonitic rocks (e.g., Passchier 1982; Takagi et al. 2000; Lin et al. 2003b, 2005a; see Chap. 9 for

3.2 Fault Rocks

21

Fig. 3.1. Photographs of a polished section of drill core (obtained at depths from 39.4 to 40.0 m) from the Chelungpu Fault (Taiwan) showing the occurrence of foliated cataclastic rocks. Long white arrows indicate the shear sense across the fault zone. The right-hand side of b is continuous with the left-hand side of a (a continued core of A-B-C). The hanging wall (right side of a) is bounded by weakly consolidated silty mudstone, and the footwall (left side of b) extends into unconsolidated alluvial deposits. F: fault, S and C: S-C foliations of cataclastic rocks. (After Lin et al. 2005a)

details) that formed at various levels in the crust. It is therefore difficult to use the presence and/or absence of foliations as a decisive criterion in discriminating between mylonitic and cataclastic rocks. Despite this, there are essential differences in the types of foliations that develop in mylonitic and cataclastic rocks. Microstructurally, one of the most significant differences is the absence of dynamically recrystallized grains and crystal-plastic deformation in foliated cataclastic rocks: these are mainly characterized by the preferred orientation of clasts and cataclastic shear bands made up of fine-grained clasts that formed as a result of brittle deformation (Lin 1999a, 2001; see below in this chapter for further details). In terms of revising Sibson’s original classification, an explanation of the fundamental differences between cataclastic and mylonitic foliations is added to the scheme, and foliations are divided into cataclastic and plastic types (Table 3.1). In a second modification, foliated pseudotachylyte is added to the scheme, representing pseudotachylyte that formed from either cataclasis during the period of pseudotachylyte formation or subsequent aseismic crystalplastic deformation that accompanied the formation of mylonitic rocks. A reduction in grain size is another key criterion in Sibson’s classification. For a long time prior to the 1970s, cataclasis was considered to be the only

22


Fig. 3.2. Sketch of the kinematic coordinates used in this book to describe shear zones (a) and the geometric relationships among the main structural elements within a representative shear zone (b). S: primary foliation, R1 and R2 : secondary order Reidel faults, T: tension fracture, C: shear fracture parallel to the main fault direction. X: shear direction, Y: direction normal to X within the shear plane, Z: direction normal to the shear plane

deformation mechanism that resulted in grain-size reduction within cataclastic and mylonitic rocks (e.g., Higgins 1971). There remains a general consensus that cataclasis is the main deformation mechanism of grain-size reduction in cataclastic rocks; however, many studies have demonstrated that grain-size reduction within mylonitic rocks occurs via syntectonic recrystallization during plastic deformation rather than cataclasis (e.g., Bell and Etheridge 1973; White 1973; Sibson 1977). This finding is supported by experimental results (e.g., Carter et al. 1964; Christie et al. 1964; Tullis et al. 1973). Indeed, grainsize reduction via syntectonic recrystallization is a common feature of many mylonitic rocks deformed at relatively high stresses and low homogeneous temperatures, where subgrain rotation recrystallization dominates; however, this process may play only a minor role at high homogeneous temperatures where grain growth occurs in association with syntectonic grain-boundary migration recrystallization (Schmid and Handy 1991). The range of grain sizes in both

3.2 Fault Rocks

23

cataclastic and mylonitic rocks is also potentially affected by the texture and degree of homogeneity of the parent rock (e.g., Takagi 1985, 1986a, b). Accordingly, grain-size reduction is not suitable as a key criterion in classifying fault rocks; instead, with the aim of arriving at a suitable classification of fault rocks for a specific fault shear zone, it is more appropriate to analyze the degree of grain-size reduction by comparing the textures of the fault rocks with those of the host (parent) rocks. Previous studies considered pseudotachylyte to be a fault rock with primary cohesion of its host rock that arose from brittle deformation (e.g., Sibson 1977; Scholz 2002). In fact, the term cohesion, as used in the classification of fault rocks, signifies that the primary cohesion of the parent rock is maintained in the resulting fault rock. The degree of cohesion is used as a tentative indicator of the depth of formation of fault rocks, with a distinction being made between incohesive cataclastic rocks and cohesive cataclastic and mylonitic rocks. Incohesive fault breccia and fault gouge are generally considered to form at depths shallower than 4 km, whereas cohesive cataclastic and mylonitic rocks are thought to form within deeper fault zones below 4 km (Sibson 1977). The degree of cohesion of pseudotachylyte, however, is not related to that of the parent rock because pseudotachylyte forms by frictional melting and/or crushing of rocks, thereby destroying any primary cohesion of the parent rock, followed by cooling and/or consolidation via subsequent cementation following formation of the melt and/or fine-grained material. It is therefore apparent that interpretation of the degree of cohesion of pseudotachylyte differs from that of cataclastic and mylonitic rocks, as it depends mainly on its formation mechanism (i.e., frictional melting and/or crushing and subsequent cementation and consolidation) rather than its depth of formation (as in the case of cataclastic and mylonitic rocks). Accordingly, the cohesiveness of pseudotachylyte should be considered as an individual factor that is independent of the parent rock type and depth of burial, and is non-primary as shown in the revised scheme presented in Table 3.1. 3.2.2 Mylonitic Rocks Mylonitic rocks form from the dynamic recrystallization and crystal-plastic deformation of minerals within deep fault shear zones at >10–15 km within the crust, where crystal-plastic deformation is dominant. Such rocks are generally characterized by a distinctive lineation and foliation and occur in shear zones that range in width from several millimeters to several hundreds of meters, although the largest examples are up to 10 km in width. Although the definition of mylonite remains somewhat controversial, structural geologists recognize three features that are common to mylonitic rocks (Tullis et al. 1982; Snoke et al. 1998): i) evidence of grain-size reduction; ii) occurrence restricted to a tabular or planar shear zone; and iii) the presence of a distinct foliation with evidence of associated crystal-plastic deformation and dynamic recrystallization. Based on these features, Wise et al. (1984) defined mylonite as:

24


“Coherent rocks with at least microscopic foliation, with or without porphyroclasts, characterized by intense syntectonic crystal-plastic grain size reduction of the country rock to an average diameter less than 50 microns (0.5 mm) [Author’s note: this should be 0.05 mm] and invariably showing at least minor syntectonic recovery/recrystallization.” This definition emphasizes crystal-plastic deformation as a mechanism of reducing the grain size of the parent rock; however, it is known that grain-size reduction commonly results from the cataclasis of brittle minerals and synchronous crystal-plastic deformation and resultant dynamic recrystallization of weaker minerals (e.g., Simpson 1985; Nyman et al. 1992). Foliations within mylonitic rocks generally comprise two sets of planar structures: discrete and narrow shear bands defined by strongly deformed minerals (C and C’ or R1 surfaces) and a preferred orientation of porphyroclasts (S or P surfaces) (Fig. 3.2b; following the terminology of Berthé 1979). The shear bands are asymmetrically distributed around porphyroclasts (Figs. 3.3– 3.6). The geometry of such S-C fabrics is commonly used as a criterion in determining the sense of movement within fault shear zones (Figs. 3.3–3.6; e.g., Simpson and Schmid 1983). Based on the proportion of porphyroclasts within the rock mass, mylonitic rocks are generally subdivided into protomylonite (Figs. 3.3a and 3.4; 10– 50% matrix), mylonite (Figs. 3.3b and 3.5; 50–90% matrix),and ultramylonite

Fig. 3.3. Photographs of polished hand samples of granite-hosted protomylonite (a) and mylonite–ultramylonite (b) from the Futaba Fault and the Median Tectonic Line, Japan, respectively. White arrows indicate the sense of movement

3.2 Fault Rocks

25

Fig. 3.4. Photomicrographs of granitie-hosted protomylonite from the Futaba Fault Zone, Northern Japan. (a): plane polarized light, (b): crossed polarized light

(Figs. 3.3b and 3.6; > 90% matrix); these terms define a gradation in the intensity of deformation, with ultramylonite being the most strongly deformed (Table 3.1). 3.2.3 Cataclastic Rocks Cataclastic rocks comprise both cataclasite derived from parent rocks with primary cohesion and fault breccia and gouge derived from parent rocks without

26


Fig. 3.5. Photomicrographs of mylonite from the Tan-Lu Fault Zone, Central China. (a): plane polarized light, (b): crossed polarized light

primary cohesion. They represent tectonic breccias and fragments of various sizes that formed by crushing and fracturing in shallow fault zones at < 10– 15 km depth. Fault shear zones that contain cataclastic rocks vary in width from several millimeters to several kilometers, but are generally in the range of 1–50 m. Pseudotachylyte is a special type of cataclastic rock that forms by coseismic frictional melting and/or crushing associated with mechanical forces, as with other cataclastic rocks. Cohesive cataclasite generally forms

3.2 Fault Rocks

27

Fig. 3.6. Photomicrographs of ultramylonite from the Median Tectonic Line, Japan. (a): plane polarized light, (b): crossed polarized light

at depths of > 4 km, while incohesive cataclastic rock generally forms within shallow fault zones of < 4 km depth (Sibson 1975). Both cohesive cataclasite and incohesive fault breccia and gouge can be subdivided into two groups on the basis of textural criteria: non-foliated and foliated cataclastic rocks. These are described below.

28


Non-Foliated Cataclastic Rocks Non-foliated cataclastic rocks generally comprise angular to sub-angular fragments with numerous visible cracks and microcracks (Figs. 3.7 and 3.8). Under the microscope, non-foliated cataclasite shows a random fabric, with fragments ranging in diameter from several microns to several millimeters (Figs. 3.7b and 3.8b). The fine-grained matrix in non-foliated cataclastic rock is commonly gray, dark gray, brown, or dark brown in color and moderate hard.

Fig. 3.7. Photograph of a polished section (a) and photomicrograph (b) of nonfoliated cataclasite hosted in granite extracted within drill core from the Nojima Fault, Japan. (b): crossed polarized light

3.2 Fault Rocks

29

Fig. 3.8. Polished section (a) and photomicrograph (b) of grantite-hosted fault breccia extracted within drill core from the Nojima Fault, Japan. (b): plane-polarized light

Non-foliated fault breccia and gouge are incohesive and form the cores of fault shear zones, generally along the main fault plane where marked displacement occurs within the deformation zone of fault-related damage. Veinlets of fault breccia and gouge are also found in the wall rocks adjacent to fault zones where no clear displacement is recognized. Such veinlets show an injection mode of occurrence, as with injection veins of pseudotachylyte, and are therefore interpreted to form by rapid coseismic injection within fault zones (Lin 1996, 1997a; see Chap. 10 for details). The zone of fault breccia within

30


fault core zones is largely composed of angular to sub-angular randomlyoriented fragments that range in diameter from several millimeters to several tens of centimeters (Fig. 3.8). This zone is easily recognized and distinguished from cataclasite and fault gouge in field on the basis of the large fragments that are present within the zone. Non-foliated fault gouge consists of a fine-grained matrix with < 30% visible fragments, being incohesive and generally unconsolidated, and with no remnants of the primary structure and texture of the parent rock, whether at the meso- or micro-scale. The fault gouge zone usually occurs along the main fault plane, bounded by a zone of fault breccia and/or cataclasite (Fig. 3.9). The zone generally ranges in width from several millimeters to several tens of centimeters, although rare examples exceed one meter. The fault gouge zone generally has an unconsolidated mud-like or clay-like appearance in which fabrics are difficult to observe in outcrop; consequently, the rock is commonly described in the field as possessing a random fabric. It should be kept in mind that some fault breccia and gouge zones exposed at the surface are as equally consolidated and hard as adjacent host rocks. This increase in the cohesiveness of the fault rocks occurs progressively with the drying, cementation, and weathering of the fault zone, occasionally leading to the mistaken interpretation in the field of fault breccia and gouge as cohesive cataclasite. Another potential error is the misinterpretation of water-rich sedimentary rocks, particularly mudstone and sandstone bound with fault breccia and gouge within fault-fracture zones, as fault breccia and gouge zones. This error occurs because of the unconsolidated mud-like nature of the sedimentary rocks and the lack of visible structure or texture as a result of near-surface weathering and erosion associated with groundwater or surface water adjacent to or within the damage zone of the fault. If we carefully observe the structural and textural differences between the parent rock and fault gouge and use the definition of fault gouge provided in Table 3.1, in which the proportion of visible clasts in the rock is < 30%, it is possible to distinguish the difference between host rocks and fault gouge zones, which are generally less than several tens of centimeters in width. Problems associated with distinguishing between these different rock types may explain the fact that fault gouge zones are reported in the literature with widths of up to 2–5 m (e.g., Tanaka et al. 1996). When observing the textures of cataclastic rocks in the field, it is common for researchers in Japan to use a sickle to smooth the exposed walls of the trench or sections of unconsolidated deposits exposed within fault outcrops (Fig. 3.10). Weathering and weakening of the water-rich fault damage zone and core zone means that it is generally easy to smooth outcrops of cataclastic rocks, even those that formed from basement rocks (Figs. 3.9a and 3.10). The boundaries between the fault breccia zone and gouge and country rocks are generally sharp (Fig. 3.9), although some show a gradational change in texture from weakly or undeformed country rock to strongly fractured fault breccia and gouge zones.

3.2 Fault Rocks

31

Fig. 3.9. Photographs of the outcrop of the Gosukebashi Fault, Japan (a) and polished X–Z section of granite-hosted fault gouge and foliated cataclasite (b). The colored zone of fault gouge visible in (a) is ∼40 cm thick and is bound by foliated granite-derived cataclasite in the left side. The contact between the foliated cataclasite and fault gouge is generally sharp. The hammer shown in (a) for scale is 35 cm long

Foliated Cataclastic Rocks Foliated cataclastic rocks are generally characterized by a well-developed foliation oriented parallel to subparallel to the main fault plane. The foliation is defined by variations in color, the preferred orientation of clasts, and cataclastic

32


Fig. 3.10. Japanese-style twisted sickle (25 cm in length) that is commonly used to smooth the exposed walls of trench and sections of unconsolidated deposits exposed in outcrop. The photograph shows a smoothed outcrop section of granite-hosted cataclasite exposed within the granitic rocks in the Itoigawa-Tectonic Line Active Fault Zone at Tozawa site. Pt: Crushing-origin pseudotachylyte veins occurred within the fault zone (see Chap. 10 for details)

shear bands (visible cracks) (e.g., Lin 1999a, 2001; Figs. 3.1 and 3.9). Cohesive cataclasite and incohesive fault breccia and gouge commonly occur together in the same outcrop and even within a single hand specimen (Figs. 3.1 and 3.9). In quartzo-feldspathic cataclastic rocks, transparent and white-gray porphyroclasts of quartz and feldspar are predominantly oriented parallel or slightly oblique to the main fault plane (Fig. 3.9). S-C fabrics are defined by variations in color, the preferred orientation of fragments such as mica fish (S-surfaces), microshears (C-surfaces), and shear bands (C’-surfaces) developed parallel to R1 Riedel shears (Figs. 3.11 and 3.12). Foliated cataclasite that contains welldeveloped S-C fabric is termed S-C cataclasite (Lin 1999a). Transparent and light-gray fragments or aggregates of fine-grained clasts in host granitic rocks are commonly oriented predominantly parallel or slightly oblique to the microshears (C-surfaces). These aggregates of fine-grained clasts are generally asymmetric in shape and are commonly used as a criterion to deduce the sense of movement within the fault shear zone, as with comparable structures developed within mylonitic rocks (Figs. 3.11 and 3.12).

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33

Fig. 3.11. Sketches of S-C fabrics observed in X–Z sections of hand specimens of foliated cataclasite hosted in granite from the Nojima Fault, Japan. (After Lin 2001). c 2007, with kind permission from Elsevier Science Ltd

For cataclasite hosted in granitic rocks, foliations observed under the microscope are generally defined by the alignment of elongate minerals such as biotite and the preferred orientation of quartz and feldspar clasts and fine-grained mineral aggregates (Figs. 3.13 and 3.14). Fragments of quartz and feldspar show two general textural types: (1) well-oriented asymmetric fragments of a single grain or aggregates (S-surfaces), and (2) randomly oriented fragments of various sizes. The proportion of fine-grained matrix to fragments increases and is concentrated in microshears (C-surfaces) or shear bands (C’-surfaces). Unlike the brittle deformation that is commonly recorded by quartz and feldspar, most biotite crystals show evidence of crystal-plastic deformation and are elongate in the plane of the foliation, forming mica fish similar to those found in mylonites (Fig. 3.13). These mica fishes are usually linked to adjacent biotite grains by tails of mica fish; this defines the S-foliation and indicates the sense of shear upon the fault plane (Figs. 3.13 and 3.14).

34


Fig. 3.12. Photograph of polished X–Z section (a) and accompanying sketch (b) of a hand specimen of foliated cataclasite from the Gosukebashi Fault, Japan. The S-C fabrics are defined by asymmetric aggregates (S-surfaces), changes in color, c microshears (C-surfaces), and shear bands (C’-surfaces). (After Lin 1999a). 2007, with kind permission from Elsevier Science Ltd

The long axes of elongate biotite clasts are generally subparallel to their (001) planes. The deformation of biotite clasts varies with grain orientation, but the majority of grains tend to be aligned parallel to (001) and occur as aggregates, thereby defining the foliation. Cleavage fractures (001) generally form oblique to the bulk shear plane (Fig. 3.14a). The angle (φ) between the long axes of biotite clasts and (001) cleavage varies from 10 to 45◦ , and the ratio of length to width ranges from 3:1 to 10:1, with most grains around 5:2 (Fig. 3.14a; Kanaori et al. 1991; Lin 1997b).

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35

Fig. 3.13. Photomicrographs of microstructures within foliated cataclasite hosted in granite from the Gosukebashi Fault, Japan. Note that biotite grains occur as elongate mica fish, as commonly reported in mylonite. (a): plane-polarized light, c (b): crossed polarized light. (After Lin 1999a). 2007, with kind permission from Elsevier Science Ltd

Foliations such as those observed in cohesive foliated cataclasite also develop within incohesive fault breccia and gouge zones; this can be observed in outcrop (Fig. 3.9) and in drill cores (Figs. 3.1 and 3.15). The foliated fault breccia zone is composed of a fine-grained matrix that resembles fault gouge and angular to sub-rounded fragments that range in diameter from several millimeters to several tens of centimeters (Fig. 3.1). These fragments are aligned parallel or subparallel to shear bands made up of fine-grained clasts

36


Fig. 3.14. Schematic diagrams of broken and displaced cleavage-steps of biotite fragments (a) and resistant clasts of quartz and feldspar (b) observed in foliated cataclasite. Large arrows indicate sense of shear. Small arrows indicate antithetic offset along (001) cleavage planes in biotite clasts and micro-faults in quartz and c feldspar fragments. (After Lin 1997b). 2007, with kind permission from Elsevier Science Ltd

(Fig. 3.16). As shown in Fig. 3.1, ‘fish trails’ commonly develop at both ends of asymmetric aggregates of fragments within the foliated fault breccia zone, which generally contains very fine-grained clasts linked by microshears (Csurfaces). The ‘fish trails’ of mica fish and adjacent shadows of asymmetrical fragments or aggregates (Fig. 3.13), which are similar to pressure shadows developed in S-C mylonite, are generally composed of very fine-grained rigid clasts that also define the C-surfaces within the rock mass. Foliated fault-gouge zones are generally characterized by layers that are gray, dark-gray, and brown to reddish-brown in color; this is in contrast to the color of the parent rock (Figs. 3.1, 3.9, and 3.15). The variable colors of fault gouges probably reflect oxidation and the presence of alternating layers of mafic and felsic minerals and weathered material that has been affected by underground water that flowed through the fault zone at shallow depths. S-surfaces are generally characterized by aggregates of quartz and feldspar

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37

Fig. 3.15. Polished section (a) of drill core (b) (from a depth of 389.4 m) from the Nojima Fault, Japan, showing the occurrence of fault gouge. F: fault plane along which the primary slip occurred during the 1995 Kobe (Japan) Mw 7.2 earthquake. c (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

clasts, while C- and C’-surfaces are commonly defined by fine-grained angular to subangular material infilling microshears or shear bands (Figs. 3.1 and 3.17). These aggregates of clasts are asymmetrical and are flanked by pressure shadows similar to those developed in mylonite. The core part of each

Fig. 3.16. Photomicrograph of microstructures within foliated fault breccia hosted in granite from the Nojima Fault, Japan. Note that clasts are oriented parallel to the foliation. Plane polarized light

38


Fig. 3.17. Photomicrographs of microstructures within foliated fault gouge hosted in granite from the Nojima Fault, Japan. Note that clasts and aggregate of clasts are oriented parallel to the foliation. S, C, and C’: S-, C-, and C’-surfaces corresponding those of S, C, and R1 shear fractures shown in Fig. 3.2, respectively. Plane-polarized c light. (After Lin 2001). 2007, with kind permission from Elsevier Science Ltd

asymmetric aggregate consists of coarser clasts than those in the adjacent shadows, which usually house fine-grained material (Fig. 3.17). The asymmetry of textures within fault gouge is often used as a criterion for deducing the sense of movement within the fault shear zone (e.g., Lin 1997b, 1999a, 2001). Some fault gouge zones contain flow structures characterized by folded layers of contrasting color (Fig. 3.18).

3.2 Fault Rocks

39

Fig. 3.18. Photomicrograph of flow structures within foliated fault gouge hosted in granite from the Nojima Fault, Japan. Plane polarized light

3.2.4 Formation of S-C Fabrics The orientation of an S-C fabric reflects the geometry of the strain field in the shear zone, which commonly involves simple shear parallel to the shear zone; the fabric is therefore not diagnostic of a particular deformation mechanism (Scholz 2002). Foliations that develop during cataclastic flow have long been recognized in low-temperature experiments (e.g., Logan et al. 1979, 1981; Noda and Shimamoto 2005), naturally occurring low-temperature and lowpressure incohesive fault gouge zones (e.g., Chester et al. 1985; Kano and Sato 1988; Lin 1996, 1997b, 2001; Lin et al. 2005a), and cohesive cataclasites (e.g., Kanaori et al. 1991; Lin et al. 1998a, b, 2005a; Lin 1999a, 2001). Although cataclastic deformation at the scale of an individual particle primarily involves brittle fracture, cataclasis can produce ductile and macroscopically uniform flow of an aggregate as a whole (e.g., Griggs and Handin 1960). An S foliation, for example, may result from the preferred orientation of platy minerals such as biotite due to the mechanical rotation of rigid, inequant grains by fine-scaled cataclastic flow, especially in clay-rich gouges (e.g., Chester et al. 1985; Rutter et al. 1986). In naturally deformed rocks, the downward transition through the brittle–plastic transition commonly involves a gradual change to predominantly crystal-plastic deformation. It is clear that cataclastic processes are active during mylonitization, because although the matrix minerals deform mainly via crystal-plastic processes, other stronger minerals such as feldspar and garnet may undergo brittle deformation. This is especially commonly observed in mylonite hosted within granitic rocks (Figs. 3.4 and 3.5).

40


Conversely, crystal-plastic deformation processes are expected to be active during cataclasis. Although most rock-forming minerals (e.g., quartz and feldspar) in fault shear zones within continental crust deform via cataclastic processes at temperatures of 100 100 >300 10 10 16 3

[Chester et al. 1993] [Sibson et al. 1979] [Lin et al. 1998b] [Takagi 1985] [Kanaori et al. 1991] [Maruyama and Lin 2004] [Lin 2001; Lin et al. 2001c]

Length: total length of the active fault; W: width of the zone of cataclasitic rock that contains S-C cataclasite, as observed in the field; Displacement: total accumulated displacement along the fault; F: fault, SAFZ: San Andreas Fault zone, USA; Alpine F: Alpine Fault, New Zealand; Tan-Lu F: Tancheng–Lujiang Fault Zone, China; MTL: Median Tectonic Line, Japan; Atera F: Atera Fault, Japan; ATTL: Arima– Takatsuki Tectonic Line, Japan; RAF: Rokko–Awaji Fault, Japan.

3.3 Fault Zone Strength and Fault Model

τ = 0.5 + 0.6σn

(200 MPa < σn ≤ 2 GPa)

43

(3.2)

where μ is the frictional coefficient (mainly determined on the basis of frictional experiments on samples of massive rock) and σn is the normal stress perpendicular to the fault, which increases with depth. Equation (3.1) follows the Coulomb friction law, whereas (3.2) deviates slightly from the Coulomb friction law, as μ decreases more slowly with increasing σn than that for (3.1). These equations are well known as Byerlee’s Law. The above law is considered to be independent of lithology and is often used to estimate the strength of natural fault zones; however, experimental results also reveal that rocks containing amounts of montmorillonite, kaolinite, illite, chlorite, serpentinite, vermiculite, and holloysite have a much lower frictional coefficient than that determined in Byerlee’s Law (Byerlee 1978). Experiments involving fault gouge indicate that fine-grained gouge produced in zones of intense cataclasis typically displays a lower coefficient of friction than that of a clean-sliding gouge-free surface (Shimamoto and Logan 1981). Experiments involving S-C cataclasite reveal that S-C fabrics are the primary factor in terms of reducing fault strength and that the strength of S-C cataclasite is approximately half that of massive rocks used in traditional frictional experiments (Lin 2000). Accordingly, the value of frictional strength employed in Byerlee’s Law probably overestimates the strength of natural fault zones that contain weak cores of S-C cataclasite. As stated above, the bulk rheology and frictional strength of a seismogenic fault zone at the time of seismic activity is probably determined by mature active fault zones that contain a weak core of S-C cataclasite, including fault gouge with a low frictional coefficient. This suggests that Byerlee’s Law is not valid for mature active fault zones that contain S-C cataclastic rocks and that the strength and rheology of fault zones within the upper 5–10 km of the crust are largely controlled by the formation of S-C cataclasite. The universal occurrence of S-C cataclasite with a low friction coefficient at the depths at which large-magnitude earthquakes nucleate within seismogenic fault zones may explain the fact that mature large-scale active faults are rheologically weak. 3.3.2 Conceptual Fault Zone Model Mechanisms of earthquake generation have traditionally been studied based on a simple mechanical model of fault zones derived from friction experiments in which brittle fracturing is the predominant mechanism of deformation. This is despite the fact that abundant data on natural fault rocks clearly indicate that deformation mechanisms within fault shear zones are much more complex than those incorporated in fault models based on theoretical considerations and experimental data. Thus, integrated field and laboratory studies are essential in determining the frictional properties of faults with a view to developing a conceptual fault-zone model.

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The distribution depths, deformation textures, and structures of fault rocks in and around fault zones have been used in the past to construct conceptual models of fault zones, mostly for quartzo-feldspathic continental crust (e.g., Sibson 1977, 1982, 1983, 2002; Hanmer 1988; Scholz 1988a; Shimamoto 1988, 1989; Passchier and Trouw 1996; Lin 2000; Lin et al. 2005b). The first conceptual model of this type was developed by Sibson (1977) for a continental strike-slip fault zone (Fig. 3.19). In this model, Sibson divided a fault shear zone into two main deformation regimes with increasing depth within the crust, elastic-frictional (EF) and quasi-plastic (QP), and described the way in which the strength of the fault rocks is expected to vary with depth. The EF regime, which was later renamed the frictional regime (FR) (Sibson 1982), is composed of cataclastic rocks such as incohesive fault gouge, breccia, pseudotachylyte, and cohesive cataclasite, in which deformation via brittle cataclasis, influenced by confining pressure and seismic frictional sliding, is the main deformation mechanism. In this regime, fabric elements are generally randomly oriented, whereas the QP regime is composed of mylonitic rocks that form via crystal-plastic deformation under the dominant controls of temperature and aseismic creep. The textures of mylonitic rocks are mostly characterized by a lineation and a foliation that includes S-C fabrics, as described above. Sibson’s (1977) original model is termed a two-layer fault model (e.g., Shimamoto 1988); however, it must be remembered that the change in

c Fig. 3.19. Conceptual two-layer fault zone model. (After Sibson 1977). 2007, with kind permission from Geological Society of London

3.3 Fault Zone Strength and Fault Model

45

deformation mechanisms from brittle cataclasis in the frictional regime to crystal-plastic creep in the QP regime does not occur across a sharp boundary at a given depth; rather, the transition occurs gradationally across a wide range of depths in different parts of the crust. Experimental results demonstrated that there is a transition between frictional slip and plastic flow within a halite shear zone (Shimamoto 1986). Subsequently, Sibson (1983) replaced his initial two-layer fault model with a three-layer fault model by adding a transition regime that is a zone of overlap with the FR and QP regimes; this change was made on the basis of experimental and seismic data. The concept of this three-layer fault model is broadly accepted by both geophysicists and geologists (e.g., Scholz 1988a; Shimamoto 1988, 1989; Lin et al. 2005b), although certain aspects of the model require further attention. In revising Sibson’s three-layer model, Scholz (1988a) proposed a modified fault-zone model for quartzo-feldspathic continental crust by postulating a transition zone between the brittle and plastic regimes. This zone is located between the onset of quartz crystal plasticity at approximately 300◦C, which marks the transition between the brittle and semi-brittle regimes, and feldspar crystal plasticity at 450◦C, which marks the transition between the semi-brittle and plastic regimes. In his model, Scholz explains the observation that mylonite forms not only in the plastic regime but also in the intermediate regime. This proposal was supported by Shimamoto (1989), who, on the basis of previous experimental data and his own halite experiment data, proposed that mylonite forms over a wide depth range from the semibrittle and semi-ductile regimes to the crystal-plastic regime. In particular, Shimamoto’s (1989) model states that S-C mylonite forms in the semi-ductile regime. Both the revised fault models of Sibson (1983) and Scholz (1988) describe the downward propagation of seismic slip during large earthquakes to the depth that corresponds to a temperature of < 350◦ C, where pseudotachylyte is generated; however, recent studies of large volumes of coexisting cataclasite- and mylonite-ultramylonite-related pseudotachylyte veins along the Woodroffe Thrust, Australia, and along the Dahezhen shear zone within the ultra-high-pressure and high-temperature Dabieshan collisional zone, China, suggest that coseismic slip associated with large earthquakes nucleates near the base of the brittle-dominated seismogenic zone and propagates down through the brittle–plastic transition zone into the crystal-plastic deformation regime, even as far as the lower crust (Lin et al. 2003b, 2005b; see Chap. 9 for details). Lin (2000) proposed a simplified fault zone model for quartzo-feldspathic crust (Fig. 3.20) based on the experimental results of previous studies and the structural modes and mechanical properties of S-C cataclasite, mylonitic rocks, and mylonite-related pseudotachylyte. Field observations indicate that the thickness of a fault zone is greater in the deep regime of plastic flow deformation than that in the shallow brittle regime. It is presumed that S-C cataclasites form in the lower portion of the brittle regime and that

46


Fig. 3.20. Revised fault zone model for a major fault zone within continental crust. The shear strength of the fault zone is estimated to be approximately half that inferred using Byerlee’s Law. See text for details. Shear-strength Profile A: calculated using Byerlee’s Law. Shear-strength Profiles B and C: inferred for S-C cataclasite and S-C mylonite. (Modified from Lin 2000)

S-C mylonites mainly develop in the transitional zone between the regimes of brittle deformation and plastic flow. S-C foliations within mylonitic rocks begin to form at the onset of crystal plasticity, which is approximately 300◦C for quartz and >450◦C for feldspar (Simpson, 1985; Carter and Tsenn 1987). On this basis, the temperature of 300◦ C represents the upper boundary of the transition zone that marks the lower base of the seismogenic zone in which large-magnitude earthquakes nucleate (Fig. 3.20). Within this transition zone, quartz deforms via crystal-plastic shearing and dynamic recrystallization, whereas feldspar remains as porphyroclasts with discrete cracks. The upper limit of the plastic flow regime roughly coincides with the cessation of feldspar plasticity and the amphibolite–greenschist facies transition at 450◦C (Tullis and Yund 1985). The temperature of 450◦ C is a critical temperature above which the main rock-forming minerals within quartzo-feldspathic continental crust deform crystal-plastically within fault zones. Accordingly, it is possible that the lower boundary of the transition zone occurs at 450◦ C, which marks a depth limit of ∼15–20 km (for a geothermal gradient of 25–30◦C/km); this may represent the lower depth limit of aftershocks or micro-earthquakes within fault zones developed in quartzo-feldspathic continental crust. Seismic slip resulting from large earthquakes that nucleate in the lower portion of the brittle regime is able to propagate downward through the transition zone to the plastic flow regime where pseudotachylyte is generated and overprinted by subsequent mylonitization during aseismic crystal plastic deformation (Fig. 3.20; Lin et al. 2005b).

4 Tectonic Environment and Structure of Pseudotachylyte Veins

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte 4.1.1 Tectonic Environment Since it was firstly described at the end of 19th century, fault-related pseudotachylyte has been reported from all seven continents. In terms of tectonic environments, pseudotachylyte has been reported from numerous intracontinental fault zones (e.g., Sibson 1975; Swanson 1988; Magloughlin 1992; Lin et al. 1994a, b), collisional orogenic belts (e.g., Austrheim and Boundy 1994; Lin et al. 2003b), and subduction zones (e.g., Ikesawa et al. 2003; Austrheim and Anderson 2004; Rowe et al. 2005; Okamoto et al. 2006). It has long been known that plate boundaries at subduction and collision zones are characterized by large-magnitude earthquakes that contribute more than 95% of the total global seismic moment; however, more than 95% of pseudotachylyte occurrences described to date in the literature are located within intracontinental fault zones, with the remainder located in ancient and modern subduction zones and collisional orogenic belts that are partially exhumed and exposed at the Earth surface, as with intracontinental fault zones. The scarcity of pseudotachylyte in environments of tectonic subduction and collision is one of the major reasons that tectonic-related pseudotachylyte is considered to be an enigmatic and exceptional fault rock. Although most examples of intracontinental fault-related pseudotachylyte are found in ancient and inactive fault zones, some have been reported along active faults such as the Alpine Fault, New Zealand (Seward and Sibson 1985; Bossière 1991); the Fuyun Fault, Northwest China, which triggered a M 8 earthquake in 1931 (Lin 1994a, b; Lin and Ge 1994) and the Nojima Fault, Japan, which triggered a Mw 7.2 earthquake in 1995 (e.g., Shigetomi and Lin 1999; Lin 2001; Otsuki et al. 2003). Pseudotachylyte reported from ancient fault zones generally formed relatively deep in the crust, at > 4–5 km depth, and is commonly accompanied

48


by cataclasite and mylonitic rocks, without fault gouge, that have been exhumed and exposed at the surface. Fault-related pseudotachylyte also forms in the middle and lower crust under granulite facies conditions by the downward propagation of slip from the hypocenter, as indicated by Precambrian pseudotachylyte observed along the Woodroffe Thrust, Central Australia (Lin et al. 2005b, 2007; see Chap. 9 for details). In contrast, pseudotachylyte that forms in active fault zones is generally accompanied by unconsolidated fault breccia and gouge zones (e.g., Lin 1989, 1996; Lin 1994; Shigetomi and Lin 1999; Otsuki et al. 2003; Kano et al. 2004) that formed at relatively shallow depths of 3 km) is found along the Woodroffe Thrust, Central Australia (Camacho et al. 1995; Lin et al. 2005b; see Chap. 9 for details). Figure 4.3 shows a sketch of a typical outcrop of a pseudotachylyte-bearing fault zone from the Fuyun Fault, Northwest China (Lin 1994a; Lin and Ge 1994). Pseudotachylyte generally occurs as a single vein along a fault plane and network veins within adjacent country rocks. The network veins develop within fault-fracture zones and commonly occur over an interval of several meters to several tens of meters in sections oriented perpendicular to the general trend of the fault zone (Fig. 4.3). Single veins typically occur as a thin layer of several millimeters to 10 cm in thickness and locally as a mass of up to 30 cm in diameter upon a concave part of the fault plane, commonly in association with striations (Fig. 4.4). Network veins show evidence of highly irregular and branching intrusion patterns and may occur as isolated lenses (Fig. 4.5). Such isolated lenses and disconnected or connected vein networks commonly terminate sharply at a certain plane within the country rocks (Fig. 4.6). Figure 4.7 shows a typical mode of occurrence, originally traced at 1:1 scale from an outcrop along the Fuyun Fault, Northwest China, that consists

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte

51

Fig. 4.3. Sketch of pseudotachylyte-bearing zones from the Fuyun Fault, Northwest c China. (After Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd

of a complex network of injection veins within a fault-fracture zone. These structural features demonstrate that isolated lenses and complex networks of pseudotachylyte represent exotic material that was intruded into the country rock from a generation zone upon the source fault. It is generally difficult to recognize the source fault and generation zone of pseudotachylyte observed in the field because of strong weathering, post-pseudotachylyte deformation, and the disconnected nature of fault-fracture structures. Individual veins within networks vary in width from several millimeters to several tens of centimeters and may show variations in thickness related to irregularities in the generation zone upon the fault plane. Multiple generations of pseudotachylyte veins are commonly recognized within individual outcrops and even in hand specimen, in which younger veins cut and overprint older veins (Fig. 4.8). In some cases, the younger veins contain fragments of the older veins (Fig. 4.9). These overprinting relations indicate that seismic faulting events and the associated generation of pseudotachylyte occur repeatedly within fault zones; this confirms our current understanding that large earthquakes occur repeatedly along mature active faults.

52


Fig. 4.4. Pseudotachylyte occurring as a thin layer (a) and a lump (b) (indicted by small arrows) upon the generation fault surface of the Fuyun Fault. Horizontal striations and grooves (indicated by large black arrows) are visible upon the fault c plane. The hammer shown for scale is 35 cm long. (After Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd


53

Fig. 4.5. Network of pseudotachylyte veins (Pt) from the Woodroffe Thrust, Central Australia. Individual veins show relatively straight (a) and very irregular (b) geometric shape and vary in width from several millimeters to several centimeters (After Lin et al., 2005). (a) The pen shown for scale is 15 cm long. (b) The chisel c shown for scale is 16 cm long. 2007, with kind permission from Elsevier Science Ltd

54


Fig. 4.6. Isolated irregular pseudotachylyte (Pt) lenses (veins) within wall rock far from the thick main pseudotachylyte vein of the Woodroffe Thrust, Central Australia (a) and the Alpine Fault, New Zealand (b). The pseudotachylyte veins terminate sharply in the country rocks. (a) The coin shown for scale is 2.4 cm across. (b) The scale bar is 5 cm long


55

Fig. 4.7. Field sketch of a typical outcrop showing the irregular geometry of a network of pseudotachylyte veins from the Fuyun Fault, Northwest China. The sketch was originally made at 1:1 scale by tracing the features directly upon the outcrop (After Lin 1994b)

4.1.3 Chilling-margin and Crack Textures Pseudotachylyte veins are generally dark-brown to black in fresh outcrops, but may be gray, brownish-gray, green, or brown where weathered. The margins of weathered pseudotachylyte veins are commonly gray-brown to dark-brown or brown-green to pale–dark green in color, which contrasts with the dark color of the vein interior (e.g., Sinha-Roy and Ravindra Kumar 1985; Takagi et al. 2000; Shimada et al. 2001). The margins of fresh pseudotachylyte veins

56


Fig. 4.8. Multiple network veins from the Woodroffe Thrust, Central Australia. Note that younger network veins (Y-Pt) cut older veins (Old-Pt). The coin shown c for scale is 2.6 cm across. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

Fig. 4.9. Polished section of a hand sample showing the incorporation of early pseudotachylyte vein (Old-Pt) as fragments within younger veins (Y-Pt). (After Lin c et al. 2005b). 2007, with kind permission from Elsevier Science Ltd


57

are commonly dull relative to the vein center and are generally more compact and aphanitic in appearance (Fig. 4.10). When viewed in a polished section or under the microscope, the central zones of veins generally contain a higher proportion of visible fragments than the vein margins, which have a higher proportion of glassy matrix and small microlites (see Chap. 6 for details). These differences in color and texture between the margins and centers of pseudotachylyte veins may arise from rapid cooling, as with the chilled margins that develop within igneous dikes, and to some degree from subsequent metamorphism and weathering. Pseudotachylyte veins commonly contain cracks that resemble the cooling joints found within igneous rocks (Fig. 4.11; Shimamoto and Arai 1997). These cracks are generally oriented normal to the large vein wall (Fig. 4.11b) but partially oblique to the vein (Fig. 4.11a). In weathered outcrops, these cracks are gray to yellow-brown in color (Fig. 4.11b). These cracks are usually spaced at intervals of 10 m along the fault plane. Individual fault veins are typically several millimeters to several centimeters thick, but may locally exceed 20 cm in thickness. Injection veins branch off obliquely from the fault plane upon which the pseudotachylyte was generated and extend along minor tensional fractures in the wall rock that record no offset (Fig. 4.14). Injection veins are considered to be reservoir zones, providing space for melt and fine-grained fragments derived from elsewhere in the rock mass (Magloughlin and Spray 1992). In contrast to fault veins, injection veins generally form complicated networks and have an irregular and complex morphology (Figs. 4.5–4.8). They may also occur as simple shapes such as single veins, lenses, and isolated blobs within the country rock (Fig. 4.15). Some injection veins show a concordant relation with foliations within the country rock; these are termed concordant veins, whereas those that are oriented obliquely to the surrounding foliation are termed discordant veins (Fig. 4.13; Sibson 1975). Undeformed injection veins are easily recognized in foliated rocks such as gneiss and mylonite because of the geometric relation between the veins and the external foliation. Individual injection veins vary in thickness from less than a millimeter to several tens of centimeters and they can be traced up to several meters in length where exposure allows. Figure 4.14 shows thin (sub-millimeter to several millimeters thick), aphanitic, and compact injection vein networks of pseudotachylyte within the Woodroffe Thrust zone, Central Australia. The veins are composed of glassderived material and minor fine-grained fragments. The irregular nature of the

4.2 Classification of Pseudotachylyte Veins

Fig. 4.13. Classification of pseudotachylyte veins (After Sibson 1975)

61

62


Fig. 4.14. Network of thin pseudotachylyte veins developed within the Woodroffe Thrust zone, Central Australia. The coin shown for scale is 2.4 cm across

network of thin injection veins that extends for several meters from the generation plane demonstrates that the frictional melt and fine-grained fragments were rapidly injected along dilational fractures prior to cooling (Figs. 4.5 and 4.14). It is unlikely that the veins formed by a slow flow of melt because it is difficult to form a network of thin (< 10 mm thick) veins to an extent of several meters (as shown in Fig. 4.14) and in various directions via the slow flow of frictional melt that contains numerous fine-grained fragments of host rock. Such a hypothesis is also unlikely because of the high viscosity of siliceous melt and the effect of rapid cooling induced by cool wall rocks (generally 30◦) to the bounding shear planes, corresponding to the orientation of T1 and R’ shear fractures within a Riedel shear zone (Fig. 4.21). These types of tensional cracks that form within paired generation zones during seismic events provide a void space that acts as a melt reservoir. Fault veins generated upon a single generation zone are generally recognized on a distinctive fault shear plane marked by fault displacement structures such as striations and offset markers (Fig. 4.4). Locally, there may be a gradation between single and paired generation zones, and the two types may become joined along a fault zone (Fig. 4.19). In contrast to paired generation zones, single zones contain a higher proportion of melt and fine-grained matrix, with few fragments greater than 1 cm in size.

4.2 Classification of Pseudotachylyte Veins

69

Fig. 4.20. Strain partitioning and fracture orientations within idealized brittle fault structures that reflect the dominant role of layer anisotropy within fault shear zone. The different components consist of fracture sets associated with layer-parallel extension (X’, T2, X), layer-parallel shortening (P, P’, T3), and dextral layer-parallel simple shear (R, T1, R’, P*) that combine to form the potential fracture orientations expected within anisotropic rocks. (Modified from Swanson 1988). (b: After c Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

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Fig. 4.21. Pseudotachylyte veins developed within a generation zone along R, X, and Y (C) fractures along the Woodroffe Thrust, Central Australia. F: fault plane. The coin shown for scale is 2.4 cm across

Fig. 4.22. En echelon coseismic surface ruptures produced during the 2001 Mw 7.8 Kunlun earthquake, northern Tibet, China. Arrows indicate the sense of fault movement (Lin et al. 2002a, 2003a, 2004)

4.3 Relation Between Fault Vein Thickness and Slip Amount Estimates of apparent or true lateral displacement along pseudotachylyte-generating faults are sometimes possible when the equivalent marker can be recognized on both sides of a fault vein (Sibson 1975;

4.3 Relation Between Fault Vein Thickness and Slip Amount

71

Di Toro and Pennacchioni 2005). Sibson (1975) quantified the relation between vein thickness and fault displacement for pseudotachylyte veins along the Outer Hebrides Thrust, Scotland. He found an empirical linear relation between the thickness (a) of pseudotachylyte fault veins and displacement (d), where d = 436a2 (unit: cgs), after measuring values of d and a along microfaults (Fig. 4.23). In a study of the Gole Larghe Fault Zone, Italy, Di Toro and Pennacchioni (2005) reported that the thickness of pseudotachylyte fault veins increases with increasing displacement, although with a non-linear relation (Fig. 4.24). In most cases it is difficult to measure the net displacement because of difficulties involved in determining the slip direction (Sibson 1975). It is also impossible to accurately measure in the field the primary average

Fig. 4.23. Log-log plot of the displacement (d) and thickness (a) of pseudotachylyte layers along microfaults along the Outer Hebrides Thrust (both a and d are measured in cm) (After Sibson 1975)

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Fig. 4.24. Log-log plot of the displacement (d) and thickness (a) of pseudotachylyte layers (in meters) for faults along the Gole Larghe Fault Zone. (After Di Toro and c Pennacchioni 2005). 2007, with kind permission from Elsevier Science Ltd

thickness of melt generated during an individual seismic slip event. The three main reasons for this difficulty are as follows. i) Pseudotachylyte veins are generally heterogeneous in all dimensions and contain numerous fine-grained fragments of the host rock, ranging from ∼5 to >70 Vol% of the entire vein when observed under the microscope. Fragment-rich veins are still considered to have a melt-origin and have a similar dark, hard, and opaque appearance to that of glass-rich veins. There is also a complete gradation from cataclastic veins, which are composed mostly of fine-grained fragments of various sizes ranging from several nanometers to millimeters, have little or no glassy matrix, and which resemble melt-origin pseudotachylyte veins (Lin 1996, 1997a; see Chap. 10 for details), to melt-dominated veins that are composed mainly of primary glass or glassy material with a minor component of fine-grained fragments. ii) Pseudotachylyte fault veins commonly occur as ‘pinch and swell’ lenses and isolated irregular veins (Figs. 4.7, 4.14, and 4.15), and occur locally as aggregations upon concave sections of fault planes (Fig. 4.4). These complex geometric modes of pseudotachylyte veins make it difficult to accurately measure and estimate the average layer thickness of primary melt generated on the fault shear plane.

4.3 Relation Between Fault Vein Thickness and Slip Amount

73

iii) Most of the generated pseudotachylyte melt is generally lost from the generation zone by injection into microscale and/or mesoscale tension fractures whose sizes are difficult to accurately measure or estimate in any dimension in the field due to their complicated morphology (Figs. 4.5– 4.8). The amount of melt within injection veins and the total volume of injection material, including fragments within fault-fracture zones, are generally greater than the amounts that remain within the fault vein.

5 Pseudotachylyte Matrix

5.1 Introduction Microstructurally, the matrix within pseudotachylyte vein consists of fine-grained material such as glass or glassy material, devitrified matrix, microlites, and dispersed fragments, that are yellowish-brown to brown-dark in color when viewed under an optical microscope in plane polarized light. Shand (1916) noted early on that most pseudotachylyte veins are opaque when viewed under the microscope using a thin section of normal thickness (0.03 mm). This opacity makes it extremely difficult to determine the nature and properties of the fine-grained matrix as a means of understanding the formation of pseudotachylyte. To overcome this problem, ultra-thin sections ( 260.8bars)

(5.7)

Wp(H2 O) = (P + 5.663)/133.27(P < 260.8bar)

(5.8)

where Wp (H2 O) is the weight percent of water that is soluble in an andesitic melt under pressure (P). The model solubility function of CO2 employed here is derived from the empirical relation determined by Harris (1981): Wp(CO2 ) = 0.0005 + 0.059P (kb)

(5.9)

where Wp (CO2 ) is the weight percent of CO2 that is soluble in an andesitic melt.

100


Fig. 5.19. Mohr diagram showing the stress change during pseudotachylyte formation within a generation zone. (a) State of stress in a pseudotachylyte-generation zone at the time of melt generation. Lithostatic pressure (Plith ) is indicated by a stress circle. The differential stress is low, and a state of plane stress is assumed with lithostatic pressure equal to the mean stress. (b) Stress change caused by volume change of melt generated at high pressure. The difference between the lithostatic pressure (Plith ) and the effective lithostatic pressure (Ple ) approaches a value of – Plith + 1 , where 1 is the tensile strength. This is the maximum possible value of the liquid pressure. When fracturing occurs, liquid pressure is assumed to revert to lithostatic pressure; under this condition, the melt vesiculates. (After Maddock c et al. 1987). 2007, with kind permission from Elsevier Science Ltd

The amount of a given volatile component in the vapor W(v) is then given simply by W(v) = W∗ (H2 O) + W∗ (CO2 ) – [Wp(H2 O) + Wp(CO2 )]

(5.10)

where W*(H2 O) and W*(CO2 ) are the initial weight percents of H2 O and CO2 in solution, respectively.

5.4 Discussion

101

Fig. 5.20. Relationship between the volume percent of vesicles and pressure in andesitic melt obtained from Macphergon’ (1984) model. Initial CO2 concentrations are 0.1 wt% (a) and 0.2 wt% (b). Initial water contents are 0.2, 1.0, 1.5, 2.0, 2.5, 3.0, 3.5, and 4.0 wt%. Solid circles shown in (a) indicate the pressure inferred from the volume percent of vesicles and water contents for the Fuyun pseudotachylyte. (After Lin 1991)

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Table 5.1. Previously published reports of the typical occurrence of vesicles and amygdules within melt-origin pseudotachylyte veins Reference

Locality

Scott and Drever (1953) Langtang Himalaya

Bisschof (1962)

Philpotts (1964)

Beckholmen (1982)

Maddock et al. (1983)

Maddock et al. (1983)

Lin (1994a)

Description

Irregular/spherical to elliptical cavities up to 1–1.5 cm in length, sub-parallel and parallel to flow-banding and/or vein walls (see Fig. 11.4 in this book) Vredefort (South Africa) Irregular/elliptical cavities parallel to flow-banding, filled with drusy calcite + quartz + chlorite + chalcopyrite Canada Spherical/regular cavities up to 50 μm diameter. Occurs in microlitic pseudotachylyte; up to 20 modal vol%, filled with quartz + calcite + magnetite + pyrite; infilling attributed to vapor precipitate from pseudotachylyte melt; rims of feldspar porphyroclasts locally replaced by carbonate Sweden Irregular cavities up to 3 mm in diameter within spherulitic pseudotachylyte; filled with drusy quartz + calcite Scotland Spherical/irregular cavities up to 200 μm in diameter, filled with K-feldspar + sphene + epidote + quartz + carbonate Greenland Irregular/elliptical cavities up to 1 mm in length, filled with dolomite + barite + magnetite + celadonite + quartz + calcite, with zoning structures Fuyun, China Spherical/elliptical and irregular cavities up to 50 μm in diameter; occur within microlitic pseudotachylyte veins, filled with carbonate + quartz + magnetite

5.4 Discussion

103

Table 5.1. Continued Lin et al. (2001)

Taiwan

Shimada et al. (2001)

Japan

Spherical to elliptical/ irregular cavities up to 300 μm in diameter, parallel to subparallel to the vein walls, up to 20 modal% Spherical/elliptical and irregular cavities up to 70 μm in diameter, with long axes oriented oblique to the vein wall; filled with chlorite + quartz + hematite + pyrite + limonite, less than 10 modal%

The volumes (V) of H2 O and CO2 vapor per liter of andesitic melt were calculated independently using the ideal gas described by Macpherson (1984): V(vesicle percent) = 100[V(CO2 )+V(H2 O)(liters)]/[V(CO2 )+V(H2 O)(liters)] (5.11) Figure 5.20 shows the modified model curves for vesicles in glassy pseudotachylyte veins with an andesitic composition, assuming initial dissolved CO2 contents of 0.1 and 0.2 wt% and a range of H2 O contents. It is significant that the vesicle content of the pseudotachylyte that remained on the fault surface reflects the lithostatic pressure of vesiculation during pseudotachylyte formation. The volume percent of vesicles observed in the Fuyun glass pseudotachylyte, which was both calculated and measured directly from photomicrographs, is about 3% (Lin 1991, 1994a). Using the maximum water content as an initial water content, Lin (1991) obtained a minimum pressure estimate of 350 bars (35 MPa). Although the initial water content is generally unknown, it is assumed that vesicles form at a condition near that of solidification of the melt. Using the 3.5 wt% line as the upper limit of initial water content, a maximum pressure of 450 bar (45 MPa) is obtained (Fig. 5.20). The estimated pressure range of 350–450 bars obtained using the model that considers vesicle volume and pressure is consistent with that estimated on the basis of the relation between pressure and the solubility of water in andesitic melt. This result indicates that the modified Macpherson model is potentially useful in estimating the formation depth of pseudotachylyte melt of andesitic composition.

6 Microlites

6.1 Introduction Melt-origin pseudotachylyte generally contains numerous minute crystals known as microlites, which are also found in volcanic rocks. The presence of microlites of various shapes is considered to be a distinct characteristic of pseudotachylyte derived from a melt (Shand 1916; Park 1961; Philpotts 1964; Sibson 1975; Wallace 1976; Masch et al. 1985; Allen 1979; Maddock 1983; Macaudière et al. 1985; Magloughlin 1989, 1992; Toyoshima 1990; Lin 1991, 1994a, b; Techmer et al. 1992; Barker 2005). The nature and origin of microlites within pseudotachylyte have attracted widespread attention since they were first documented in the early 20th century. Although in the period prior to 1990 there was considerable controversy regarding the composition, structure, and origin of microlites in terms of the origin of fault-related pseudotachylyte, most researchers agreed that microlites represent primary crystals that formed from a melt generated by frictional melting of the rock during seismic faulting (e.g., Sibson 1975; Allen 1979; Maddock 1983; Spray 1988; Toyoshima 1990; Lin 1991, 1994a, b; Magloughlin and Spray 1992). Detailed analyses of natural pseudotachylytes reveal that most microlites are primary crystals of varying shape derived from rapid cooling or quenching of a primary melt produced by frictional fusion during the formation of pseudotachylyte within a seismic fault zone (Lin 1994b). High-velocity frictional experiments also verify that microlites can form under conditions of rapid cooling (15 μm in diameter. 6.2.2 Morphology Based on the morphological features described above, microlites can be divided into four main morphological groups on the basis of their degree of complexity: (1) simple, (2) skeletal, (3) dendritic, and (4) spherulitic (Lin 1994b). These different types are shown in Figs. 6.1 and 6.2 and listed in Table 6.1. Simple Group Microlites in this shape group are relatively simple in morphology, being either acicular, granular, trichitic, cross-shaped, lath-like, or spider-like in form (Figs. 6.1, 6.4, and 6.5). Acicular microlites are generally aligned parallel to subparallel to the vein margins and flow streaks within the matrix. Granular microlites are locally concentrated within microcrystalline pseudotachylyte veins in which the matrix is generally rich in magnetic minerals and appears opaque under the microscope. Spider-like microlites consist of two parts that can be thought of as the body (similar in appearance to acicular microlites) and the claws of a spider (Figs. 5.8 and 6.5). Similar trichitic and spider-like microlites are also observed in glassy volcanic rocks (Ross 1962). The cross-shaped microlites usually occur in association with spider-like microlites. Lath-like microlites are the most common form of microlites found within microlitic pseudotachylytes; they form local clusters that overgrow fragments of quartz and feldspar (Fig. 6.6).

6.2 Texture and Morphology of Microlite

107

Fig. 6.1. Sketches of microlite shapes of the simple, skeletal, and dendritic groups. (Modified from Lin 1991, 1994b)

108

6 Microlites

Fig. 6.2. Sketches of microlite shapes of the spherulite group (Modified from Lin 1991, 1994b)

6.2 Texture and Morphology of Microlite

109

Fig. 6.3. Sketch illustrating the textural variations in microlites from the margins to the center of a pseudotachylyte vein from the Fuyun Fault, China. Note that there is a gradation in the size and morphology of microlites from the margins to the center of the vein (After Lin 1991, 1994b)

Skeletal Group Skeletal microlites can be subdivided into four main forms on the basis of morphology: tabular-skeletal, dendritic-skeletal, box-skeletal and chain-skeletal (Figs. 6.1 and 6.7). These forms are very similar to those of quench crystals developed in submarine basalts, as reported by Bryan (1972). Tabular-skeletal microlites generally comprise stubby skeletal pillars that form a discontinuous chain. Dendritic-skeletal and chain-skeletal microlites are similar to dendritic microlites in form. Box-skeletal microlites occur as incomplete polygonal forms that are hollow within. Skeletal microlites generally occur within microlitictype pseudotachylyte veins and consist mainly of mafic minerals. Dendritic Group Microlites within the dendritic group occur as extremely complex and highly branched dendritic forms that can be subdivided into eight main shapes: fineand course-scoplitic, fine- and course-feathery, quartz-feathery, plumose, firlike, crotch, branching, and pine-like (Fig. 6.1). Scoplitic microlites comprise a number of trichites that are generally too small to be identified in terms of their mineral type, but they are known to be Fe-rich and appear opaque under the microscope. Feathery microlites comprise a central stalk (line) and crystal fibers that radiate from one end of the stalk. Quartz-feathery microlites consist of numerous fine quartz fibers that radiate from a core fragment of quartz or feldspar to form multiple feathers, generally in association with dendritic plagioclase microlites. Plumose microlites resemble the feathery varieties and consist of two parts: a plumose stalk of plagioclase and feathery parts of biotite intergrown with plagioclase that radiate from the stalk (Fig. 6.8). Twinning is usually visible within the plagioclase that makes up the plumose stalk. Firlike, pine-like, and branching microlites possess a general form that consists of two parts: a very straight trunk and numerous “thickets” that overgrow the trunk (Fig. 6.1). The trunk generally consists of plagioclase (An > 40), while the thickets consist of plagioclase and mafic minerals such as biotite and hornblende (Lin 1991, 1994b). This texture indicates that the microlites

Width (μm)

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