Fossil earthquakes: the formation and preservation of Pseudotachylytes

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

Author: Aiming Lin

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Lecture Notes in Earth Sciences Editors: S. Bhattacharji, Brooklyn H. J. Neugebauer, Bonn J. Reitner, G´ottingen K. St¨uwe, Graz Founding Editors: G. M. Friedman, Brooklyn and Troy A. Seilacher, T¨ubingen 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-Barth´elemy 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

2 Terminology and Origin of Pseudotachylyte

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

2 Terminology and Origin of Pseudotachylyte

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

10

2 Terminology and Origin of Pseudotachylyte

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

2.2 Controversy Regarding the Physical Origin of Pseudotachylyte

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

12

2 Terminology and Origin of Pseudotachylyte

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

2.2 Controversy Regarding the Physical Origin of Pseudotachylyte

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

2 Terminology and Origin of Pseudotachylyte

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; Macaudi`ere 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

2.2 Controversy Regarding the Physical Origin of Pseudotachylyte

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

3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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

20

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

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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:

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

“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 Berth´e 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

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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.

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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).

3.2 Fault Rocks

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).

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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).

3.2 Fault Rocks

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

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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).

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3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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 <300oC, some minerals within cataclasite (e.g., biotite) may deform via crystal-plastic processes at temperatures of 150– 300◦C. For example, S-foliations that develop in granite-hosted cataclasite are mainly deﬁned by the preferred orientation of rigid fragments or aggregates and biotite ﬁsh, cleavage-steps, and dragged biotite fragments (Lin 1999a; Figs. 3.13 and 3.14). S-foliations within granite-hosted S-C cataclasite, which are deﬁned by crystal-plastically deformed biotite, mainly develop via a combination of brittle and crystal-plastic deformation mechanisms. Microstructurally, one of the most signiﬁcant diﬀerences between mylonitic rocks and foliated cataclastic rocks is the absence of dynamically recrystallized grains in foliated cataclasite, as described above. C-foliations mainly comprise aggregates of very ﬁne-grained clasts and microshears, as well as narrow shear bands. The progressive nature of the deformation is indicated by the fact that visible cracks deﬁne the foliation at the mesoscale and microscale and that ﬁne-grained clasts are concentrated and aligned along microshears and shear bands. S-surfaces in mylonitic rocks commonly form prior to C-surfaces (Platt and Vissers 1980; Lister and Snoke 1984). The results of halite experiments demonstrate that both S-C mylonitization and brittle faulting occur in the semi-ductile regime and that the S-surfaces form prior to the development of R1 shear surfaces (corresponding to C’-surfaces): the S-surfaces are deformed by R1 shear (e.g., Shimamoto and Logan 1986; Hiraga and Shimamoto 1987; Shimamoto 1989). S-surfaces that are deﬁned by crystal-plastically deformed biotite are also commonly boundinaged along C-surfaces (or C’-surfaces) in natural foliated cataclasites, as described in previous sections. This also demonstrates that S-surfaces in foliated granitic cataclasite form prior to Cand C’-surfaces.

3.3 Fault Zone Strength and Fault Model 3.3.1 Seismogenic Fault Zone Strength Based on the coexistence of dominantly plastically deformed biotite and cataclastically fractured quartz and feldspar, it has been suggested that foliated granitic cataclasite forms at temperatures of 150–300◦C, corresponding to depths of 5–10 km assuming a geothermal gradient of 30˚C/km (Lin 1997b, 1999a). We should bear in mind that the geothermal gradient within continental crust varies at diﬀerent locations; for example, the average geothermal gradient beneath Japan is about 30◦ C/km, but it is about 25◦ C/km in western parts of the USA close to the San Andreas Fault, and less than 20◦ C/km in Northern China. Accordingly, the deformation temperature of fault rocks at a deﬁned fault depth depends mainly on the local geothermal gradient.

3.3 Fault Zone Strength and Fault Model

41

Seismic data reveal that intracontinental earthquakes tend to occur at depths of <10–15 km (Sibson 1983). Larger earthquakes appear to nucleate close to the base of the seismogenic zone that is deﬁned by background seismicity. In this case, seismic rupture can propagate some distance downward into deeper parts of the fault zone as well as upward to shallow depths and ultimately the surface (Sibson 1983; Scholz 1988; Lin et al. 2003b, 2005b). This suggests that an intracontinental fault zone that is active at a depth of 5–10 km (at a temperature of <300◦C) is a potential nucleation zone for large earthquakes. Accordingly, it is possible that the rheology of a seismogenic fault zone within continental crust is principally controlled by the deformation mechanisms of S-C cataclasite at the nucleus depths of large-magnitude earthquakes (Lin 1999a). The solid-state rheology of rocks depends primarily on the relative proportions of weak and strong rock-forming minerals, as well as the physical conditions of deformation (e.g., Handy 1990; Schmid and Handy 1991). In nonfoliated cataclasite within quartzo-feldspathic rocks, relatively strong quartz and feldspar clasts are randomly distributed and generally in contact with one another; consequently, the shear resistance may be expected to depend on the deformation mechanism of the strong mineral clasts rather than that of the ﬁne-grained matrix or weak biotite clasts. In S-C cataclasites, however, ﬁnegrained matrix minerals are concentrated along microshears and shear bands that form interconnected weak zones (C- and C’-surfaces). Deformed clasts of weak minerals such as biotite are interconnected to form S-foliations, whereas quartz and feldspar clasts within the ﬁne-grained matrix of S-C cataclasite are not in direct contact (Figs. 3.13 and 3.14). Therefore, the shear strength of S-C cataclasite is expected to depend on the deformation mechanisms of weak minerals such as biotite and the ﬁne-grained matrix, including clay minerals. Laboratory and ﬁeld observations suggest that mature active fault zones such as the San Andreas Fault zone are weak relative to the surrounding country rock (e.g., Chester and Logan 1986; Zoback et al. 1987; Sleep and Blanpied 1992). There are two general explanations for this weak-fault hypothesis: i) fault zones contain a core of weak material that has a low frictional coeﬃcient (e.g., Sleep and Blanpied 1992; Chester et al. 1993), and ii) high pore-ﬂuid pressures are localized within fault zones, thus reducing the eﬀective normal stress (e.g., Sleep and Blanpied 1992; Byerlee 1993). Scholz (1992) argued against the latter case on the grounds that brittle material within fault zones undergoes hydrofracturing when the pore pressure exceeds the minimum compressive stress in the absence of drainage. A consequence of the dilatancy/ﬂuid diﬀusion mechanism for large shallow earthquakes is that considerable volumes of ﬂuid are rapidly redistributed in the crust following seismic events; this is manifest as outpourings(Sibson 1975) and falling groundwater levels around fault traces following moderate to large-magnitude earthquakes (Lin et al. 2003c). These features are commonly observed across wide areas centered upon epicenters, both before and after moderate- to largemagnitude earthquakes. This observation does not support the hypothesis that

42

3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

high pore-ﬂuid pressures can be maintained within fault zones and thereby trigger large earthquakes. In terms of the ﬁrst of the two explanations listed above, a lack of data means that it is unclear as to whether weak core materials are universally developed within mature fault zones; however, Table 3.2 shows that weak cores consisting of cataclastic rocks are universally developed in well-studied mature active faults that record large displacements. These cores have widths of 30–300 m and are marked by S-C foliations. Cataclasitic rocks form at depths of <10–15 km and are exhumed by uplift and erosion. Foliations that form during cataclastic deformation have long been recognized in low-temperature experiments(Logan et al. 1985) and in naturally occurring low-temperature and low-pressure fault gougeand cataclasite (Lin 1999a, 2001). S-foliations that form in S-C cataclasite are mainly deﬁned by the preferred orientation of rigid fragments or aggregates and fragments of deformed weak minerals such as biotite, whereas C-surfaces are mainly deﬁned by aggregates of very ﬁne-grained clasts, microshears, and shear bands. These observations indicate that S-C fabrics form during progressive deformation within fault zones; consequently, we can expect that weak cataclastic cores within fault zones (that contain well-developed S-C foliations) occur universally in the seismogenic regime within mature and long-lived active fault zones that record large displacements. At a normal stress of less than 2 GPa, corresponding to a lithostatic depth of <70 km in continental crust, the value of frictional strength (τ) is mainly based on experimental results, as follows: τ = 0.85σn

(σn ≤ 200 MPa)

(3.1)

Table 3.2. Width of cataclasitic fault zones within a number of typical mature active faults exposed at the surface (Data from Lin 2000) Fault

Length (km)

Width (m)

Displacement Reference (km)

SAFZ Alpine F Tan-Lu F MTL Atera F ATTL RAF

2000 400 >2000 >200 66 60 70

30–100 >100 50–200 ∼300 300 >300 50

>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.

44

3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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

3 Pseudotachylyte-Related Fault Rocks and Conceptual Fault Models

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; Bossi`ere 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

4 Tectonic Environment and Structure of Pseudotachylyte Veins

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 <4–5 km. An example of typical melt-origin pseudotachylyte can be found along the active Fuyun Fault in China, which triggered a M 8 earthquake in 1931 and produced a coseismic surface rupture zone that extended for over 170 km (Figs. 4.1 and 4.2). Veins of pseudotachylyte observed within fault zones located close to the 1931 surface rupture along the Fuyun Fault typically contain up to 90 wt% glass material (Fig. 4.2), which is estimated to have formed at a depth of <1.5 km (Lin 1991, 1994a; see Sect. 5.4.3 for details). The youngest melt-origin pseudotachylyte reported in the literature is that found along the Alpine Fault, New Zealand, a major active strike-slip fault; this pseudotachylyte formed within the past 1.1 Ma (Seward and Sibson 1985). The youngest crushing-origin pseudotachylyte veins are those observed within a thrust zone along the Itoigawa–Shizuoka Tectonic Line, an active fault system in Central Japan; these veins are estimated to have formed in association with seismic faulting over the past 1550–2350 years, at shallow depths of several tens of meters (Kano et al. 2004; see Chap. 12 for details). The above studies demonstrate that it is possible for fault-related pseudotachylyte to form at shallow depths, close to the surface, even during recent large-magnitude earthquakes. The results of high-velocity frictional experiments also demonstrate that pseudotachylyte can form at pressures as low as 1 MPa, corresponding to a depth within the crust of just 30 m (Lin 1991; Lin and Shimamoto 1998; see Chap. 12 for details).

4.1.2 Field Occurrence Fault-related pseudotachylyte forms by frictional melting and/or crushing during seismic sliding within a fault shear zone. The frictional melt and/or crushed ﬁne-grained material occurs as planar veins along the generation plane and is also injected into the surrounding country rock along coseismic fractures as individual and/or network veins. Field investigations reveal that pseudotachylyte veins are generally exposed within fault-fracture zones that record distinct oﬀset of markers; the pseudotachylyte may occur as simple or complicated networks of veins in outcrop at scales ranging from several meters to several kilometers (e.g., Sibson 1975; Lin 1994a; Allen 2005; Lin et al. 2005b). Individual pseudotachylyte-bearing fault zones observed in outcrop vary in

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte

49

Fig. 4.1. Geological map of the Fuyun region, Northwest China. (Modiﬁed from Lin 1991, 1994a and the Regional Surveying Team of the Xingjiang Geological Bureau 1978)

50

4 Tectonic Environment and Structure of Pseudotachylyte Veins

Fig. 4.2. Photograph showing the fault topography of the Fuyun Fault at Akesanyi, where an outcrop of pseudotachylyte occurs close to the coseismic surface rupture (indicated by small arrows) produced by the 1931 Fuyun M 8.0 earthquake. For scale, a person is standing in the center of the photograph (indicated by an open circle) (After Lin 1991, 1998)

thickness from several centimeters to >3 km, but are generally less than 10 m thick; such zones can generally be traced for several kilometers along the strike of the fault (e.g., Allen 2005; Lin et al. 2005b). The widest pseudotachylytebearing fault zone (>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

4 Tectonic Environment and Structure of Pseudotachylyte Veins

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

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte

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

4 Tectonic Environment and Structure of Pseudotachylyte Veins

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

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte

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

4 Tectonic Environment and Structure of Pseudotachylyte Veins

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

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte

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 <1 to 10 mm and are mostly restricted to the vein itself (Fig. 4.11b), but some cut throughout veins into the wall rocks (Fig. 4.12a). Locally, the cracks are inﬁlled by the fractured pseudotachylyte

Fig. 4.10. Chilled margins of pseudotachylyte veins (Pt). The margins are more compact, aphanitic (a), and glassier (b) than the vein center. (a): Hand specimen from the Woodroﬀe Thrust, Central Australia. The coin is 2.4 cm across. (b): Hand specimen from the Fuyun Fault, Northwest China. The pen shown for scale is 15 cm long. Fo: mylonitic foliation of the country rock

58

4 Tectonic Environment and Structure of Pseudotachylyte Veins

Fig. 4.11. Cracks (indicated by white arrows) that resemble cooling joints developed within an isolated blob (a) and a vein of pseudotachylyte (Pt) (b) from the Woodroﬀe Thrust, Central Australia. Note that the cracks are restricted within the pseudotachylyte vein and terminate sharply the boundary (indicated by arrows and white dotted line) between the vein and wall rock (b). The coin shown in (a) for scale is 2.4 cm across

4.1 Tectonic Environment and Field Occurrence of Pseudotachylyte

59

Fig. 4.12. Photograph (a) showing cracks (indicated by red arrows) cutting perpendicularly throughout a vein of pseudotachylyte (Pt) into the wall rocks and SEMBSE image (b) showing that a crack is inﬁlled by the fractured pseudotachylyte fragments from the Woodroﬀe Thrust, Central Australia

fragments along the cracks (Fig. 4.12b). These textural features indicate that these cracks are probably not caused by rapid cooling as those formed within igneous rocks but formed by ongoing tectonic stress due to the strength difference between the vein and wall rocks.

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4 Tectonic Environment and Structure of Pseudotachylyte Veins

4.2 Classiﬁcation of Pseudotachylyte Veins 4.2.1 Fault Veins and Injection Veins Sibson (1975) proposed a basic classiﬁcation of pseudotachylyte veins based on ﬁeld occurrence and vein geometry (Fig. 4.13). In this scheme, the source vein generated upon the fault plane is termed a fault vein, while individual veins and network veins within country rocks are termed injection veins. It is important to correctly distinguish these two types of pseudotachylyte veins in the ﬁeld in order to gain an understanding of seismic rupture structures (e.g., Grocott 1981; Swanson 1988, 1992), the kinematic evolution of the pseudotachylyte generation zone (Magloughlin and Spray 1992), estimate the amount of melt generated during an individual seismic event, and infer earthquake parameters (e.g., Sibson 1975; Di Toro et al. 2005). Fault veins generally have a simple veinlet morphology, occurring as a planar vein or a thin ﬁlm along the fault plane upon which the frictional melt and ﬁne-grained material were generated (Fig. 4.4a). Some fault veins are discontinuous, occurring with an irregular ‘pinch and swell’ morphology with local concentrations upon concave parts of the fault surface. This presumably arises due to surface irregularities upon the fault plane (Fig. 4.4b). Fault veins are generally discrete and variable in thickness; they can commonly be traced for >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

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4 Tectonic Environment and Structure of Pseudotachylyte Veins

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 <300◦C). The intrusion mode of the veins is considered to involve rapid injection under thermal pressurization of the melt associated with expansion, which in turn is related to frictional heating within the pseudotachylyte generation zone and syn-dilational fractures formed by seismic rupture. Once generated, the fragment-bearing frictional melt moves from the generation plane to the void spaces within dilational fractures that formed at the same time as the pseudotachylyte. Such injection veins are the most common reservoir of melt and/or ﬁne-grained material in the cooler country rocks (Fig. 4.15b). In some instances, dilational fractures contain a zone of cataclasite that is cut by pseudotachylyte veins. It is commonly observed that minor fractures continue from the termination of the pseudotachylyte vein. As some veins are cut by shear oﬀsets, it may be diﬃcult to determine if the fractures that they occupy are those that propagated ahead of the intruding vein, or if the veins prompted the renewed dilation of older fractures, or if the fractures in fact represent the generation surface of the pseudotachylyte.

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63

Fig. 4.15. Isolated pseudotachylyte blobs that acted as melt reservoirs within a fault-fracture zone. (a) Example from the Woodroﬀe Thrust, Central Australia. (b) Example from the Outer Hebrides Thrust, Scotland. The coin shown for scale is 2.4 cm across

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4.2.2 Pseudotachylyte Generation Zones Fault veins are referred to as a pseudotachylyte generation zone (Grocott 1981; Magloughlin and Spray 1992), which may be a narrow fault shear zone that is restricted to a fault plane or a wide fault shear zone bound by two fault planes (Fig. 4.16. Distinctive and possibly characteristic fault structures associated with the generation of pseudotachylyte veins have been reported from many fault zones (e.g., Grocott 1981; Swanson 1988, 1992; Magloughlin 1989; Lin 1994a, b; Lund and Austrheim 2003). There is a close geometric relation between fault veins and generation fault-fracture zones. Field relations indicate that pseudotachylyte generation zones are generally fault-bounded by parallel slip surfaces that serve as the dominant displacement structures (Fig. 4.16).

Fig. 4.16. Fault structures associated with the generation of pseudotachylyte. (a) Single generation zone within which a fault vein occurs along the generation fault plane and injection veins intrude the wall rock away from the fault plane. (b) Paired generation zone with two distinct shear fault planes along which frictional melt is generated, including ﬁne-grained material. (c) En echelon linked duplex with a typical extensional geometry. (d) Sidewall ripouts with leading and trailing structural assemblages (Modiﬁed from Swanson 1988)

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65

Two basic types of generation zones can be recognized on the basis of geometric features: a wide shear zone bounded by paired slip planes that serve as the dominant displacement structures, here termed a paired generation zone (Figs. 4.16b and 4.17), and a narrow fault shear zone of < 5 cm width that occurs along a single distinct fault shear plane, here termed a single generation zone where a fault vein is developed (Figs. 4.16a and 4.18). Paired generation zones were ﬁrst observed along the Outer Hebrides Thrust, Northwest Scotland (Sibson 1975), and distinctive structures have been described in detail from a pseudotachylyte-bearing fault zone in the Ikertoq Belt, West Greenland (Grocott 1981). Similar structures within generation zones, deﬁned by paired shear planes, have also been described from strike-slip fault zones in West Greenland (Swanson 1988, 1992, 2005); these are termed slab duplex structures and sidewall ripouts (Fig. 4.16c and d). Detailed ﬁeld mapping reveals that slab duplex structures develop between overlapping fault-parallel slip surfaces, which are themselves characterized by extensional and/or contractional fault geometries. Contractional duplexes tend to thicken during displacement via internal imbrication, whereas extensional duplexes that contain listric faults may thin catastrophically, leading to fragmentation within the pseudotachylyte matrix (Swanson 1992). Sidewall ripouts consist of asymmetric plano-convex lenses and slab geometries that act as reservoirs for frictional melt and/or ﬁne-grained material (Figs. 4.16d and 4.19). Ripout lenses are interpreted to represent an asperity moving through the adjoining wall rock along the leading contractional ramp until the cassation of slip or until it is sheared oﬀ during continued displacement (Swanson 1992). Translational plowing of this type causes a distortion of the fault surface at the leading ramp and an extensional reservoir at the trailing ramp for pseudotachylyte produced along the dominant slip surface. Similar melt-dominant zones with a plano-convex geometry and a high degree of assimilation are also described from a fault zone in the Nason terrane, Washington, USA (Magloughlin 1989) and the Woodroﬀe Thrust, Central Australia (Fig. 4.19; Lin et al. 2005b). The structural conﬁguration of paired generation zones can be interpreted in terms of structural features within Riedel shear zones (Fig. 4.20; Swanson 1988). Pseudotachylyte fault veins are generated not only on the main shear plane but also along conjugate fractures (R, R’, T, Y) upon which coseismic oﬀset occurs during earthquakes (Fig. 4.20). Internal fracture systems consisting of orthogonal dilatant veins and conjugate shear fractures indicate fault-parallel extension associated with the injection of pseudotachylyte (Fig. 4.21; e.g., Park 1961; Sibson 1975; Swanson 1988, 1992). Most frictional melt and ﬁne-grained fragments within fault veins that developed within paired generation zones are expected to have formed upon the major paired shear planes. Pseudotachylyte veins that develop along conjugate Riedel shears (R, R’) and tensional fractures (T1, X’) within a paired generation zone are generally oriented oblique or perpendicular to the main

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Fig. 4.17. Paired generation zone along the Outer Hebrides Thrust, Scotland (a) and Woodroﬀe Thrust, Central Australia (b). Broken lines indicate the fault planes where frictional melt and ﬁne-grained fragments were generated. Pt: pseudotachylyte vein. The marker pen shown in (a) and hammer shown in (b) for scale are c 14 cm and 35 cm long, respectively. (b: After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

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67

Fig. 4.18. Single generation zones and injection veins of pseudotachylyte developed within (a) the Sanda Rose mylonite shear zone, Southern California, USA and (b) the Alpine Fault, New Zealand. The marker pen shown in (a) for scale is 14 cm long. The scale bar shown in (b) is 5 cm long

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Fig. 4.19. Sidewall ripout vein structures of pseudotachylyte generation zone developed within the Woodroﬀe Thrust zone, Central Australia. The coin shown for scale is 2.4 cm across

generation surface, and the vein width is generally small relative to that generated upon the main generation zone. Paired generation zones vary in width from several centimeters to several meters and locally extend up to several tens of meters. They generally contain 10–50% pseudotachylyte matrix, including very ﬁne-grained fragments; locally, this ﬁgure is as high as 90%. The structural geometry of generation zones is often observed in coseismic surface rupture zones that show an en echelon array of paired shear planes on both sides of the seismic shear zone (Fig. 4.22; Lin et al. 2002a, 2003a, 2004). Dilational cracks within coseismic shear zones are generally oblique (angle of >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.

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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.

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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 (<15 μm thick) are routinely used to observe microstructures within pseudotachylyte veins under the microscope. The matrix of pseudotachylyte is generally heterogeneous, varying from aphanitic ultra-ﬁne-grained sub-micron scale fragments to crystalline, microlitic, and glass (or glassy) material. Observations of microstructures within pseudotachylyte and analyses of the chemical composition of the pseudotachylyte matrix are generally undertaken using an electron probe microanalyzer (EPMA), scanning electron microscope-back scattered electron-energy dispersive X-ray (SEM-EDX), transmission electron microscopy (TEM), and powder-X-ray diﬀraction (XRD). To select suitable samples for electron microprobe analysis, which enables the observation of distinctive melting textures, it is generally necessary to examine a large number of thin sections under the optical microscope. This is because of the degree of heterogeneity within the veins and the alteration and metamorphism of pseudotachylyte that occurs subsequent to its formation. This chapter focuses on microstructures observed within the ﬁne-grained glassy matrix of pseudotachylyte veins and quantitative analyses of the matrix in the context of its formation.

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5.2 Microstructural Characteristics 5.2.1 Textural Classiﬁcation of Pseudotachylyte Matrix Microstructurally, pseudotachylyte veins occur as both simple veins and complex networks, as also observed in the ﬁeld and in hand specimen (Fig. 5.1). When viewed under the microscope, it is common to observe overprinting relations between diﬀerent pseudotachylyte veins, with younger veins injected into older ones; the contacts between veins of diﬀerent ages are generally sharp but are locally irregular in form (Fig. 5.2). Younger veins tend to terminate

Fig. 5.1. Photomicrographs of microstructures within injection pseudotachylyte veins from the Fuyun Fault, Northwest China. (a): Crossed polarized light, (b): plane polarized light (a: After Lin 1998)

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77

Fig. 5.2. Photomicrograph of microstructures within cryptocrystalline-type matrix (Type-II) and overprinting vein within pseudotachylyte from the Fuyun Fault, Northwest China. The photomicrograph shows a new pseudotachylyte vein (Y-Pt) injected into an old vein (Old-Pt). Plane polarized light (Modiﬁed from Lin 1998)

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sharply within older veins. These overprinting relations also indicate that pseudotachylyte-generating events occur repeatedly within individual fault zones, as also observed in the ﬁeld and in hand specimen (see Figs. 4.8 and 4.9). The ﬁne-grained matrix within pseudotachylyte veins generally appears yellowish-brown to brown-dark in color and typically contains numerous fragments of various shapes ranging from angular to rounded, as well as irregular embayed forms of diﬀerent sizes ranging from the nanometer- to centimeterscale. Based on microscope observations, Shand (1916) was the ﬁrst to describe in detail the microstructures of ﬁne-grained pseudotachylyte matrix (in this case pseudotachylyte from the Vredefort impact site, South Africa). Shand divided the matrix into three textural types: opaque, crystallime and microlite. Similar petrographic and microstructural features have been observed in numerous fault-related pseudotachylyte veins reported worldwide over the past century. Based on the microstructural features of typical melt-origin pseudotachylyte veins found along the active Fuyun Fault, Northwest China (see Chap. 4), which contain all of the diﬀerent types of matrices described in previous studies, Lin (1994a, b) subdivided pseudotachylyte matrix into ﬁve types: glass (Type-I), cryptocrystalline (Type-II), microcrystalline (Type-III), microlitic (Type-IV), and mixed (Type-V). These diﬀerent matrix types are described in detail in the following paragraphs. Glass and microcrystalline matrices form individual veins, commonly occupying the marginal zones of microcrystalline, microlitic and mixed types of veins. Type-I matrix (Fig. 5.3) is sub-isotropic under the microscope and variable in color from transparent to gray, yellowish brown, and dark brown. Flow structures are commonly observed and are generally deﬁned by layers of contrasting color, including yellowish-brown, brown, and dark brown, as well as oriented fragments (Figs. 5.3 and 5.4a). Locally, the matrix is transparent or translucent and exhibits the optical character of glass found in volcanic rocks. Transparent and translucent glass matrix is typically observed around irregular and rounded feldspar and quartz fragments that survived fusion (Fig. 5.5). Glass matrix can also occur as isolated spots and pockets or slight streaks oriented parallel to ﬂow streaks. Such transparent and translucent glass matrix consists of pure silica or chemical component of feldspar in terms of chemical composition (Lin 1994a; see Chap. 8 for details). Locally, the glass matrix may be black in color and vitreous, with a refractive index of approximately 1.56 and a speciﬁc gravity of 2.76–2.73, as identiﬁed from glassy material in the Quebec pseudotachylyte, Canada (Philpotts 1964). Powder X-ray diﬀraction patterns reveal that Type-I matrix in the Fuyun pseudotachylyte is entirely glass or glassy material; this makes up about 90 wt% of the entire vein (Lin 1994a; see Sect. 5.3 for details). This is the only glass type of pseudotachylyte vein to have been identiﬁed by X-ray diﬀraction patterns, although there are many qualitative descriptions of glass or glassy matrix based on observations undertaken using an optical microscope. In hand specimen, glass-type

5.2 Microstructural Characteristics

79

Fig. 5.3. Photomicrographs of microstructures within glass-type pseudotachylyte matrix (Type-I) from the Fuyun Fault, Northwest China. Note that the fragments, all of which are quartz crystals (Qz), are generally rounded and irregularly embayed. Gl: Glass matrix. (b) Enlargement of an area of glass-type matrix shown in (a). Plane polarized light (a: After Lin 1998)

pseudotachylyte veins have a distinct vitreous luster that is similar to that of fresh obsidian. Type-II matrix (Fig. 5.2) is composed of ﬁne-grained material that is too small to identify using an optical microscope; however, powder X-ray diﬀraction patterns reveal that this type of matrix consists mainly of ﬁne-grained crystals with little or no glassy (or non-crystal) material (see Sect. 5.3 for

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Fig. 5.4. Photomicrograph of microstructures within glass-type pseudotachylyte matrix (Type-I) from the Fuyun Fault, Northwest China. The ﬂow structures are deﬁned by glass matrix layers of contrasting color and oriented quartz fragments. Plane polarized light

details). The matrix generally appears brown to dark brown in color under plane polarized light. Veins of this type do not possess a vitreous luster, even on freshly exposed sections. Type-III matrix (Fig. 5.6) is composed mainly of very ﬁne-grained crystals with small microlites that can be recognized under the microscope. The microlitic crystals, however, are generally so small that the mineral type cannot be identiﬁed. The matrix varies in color from gray to yellowish brown and dark brown, and locally appears transparent, translucent, or opaque. Powder X-ray diﬀraction patterns reveal that the matrix is mainly crystalline, with little glassy (or non-crystal) material (see Sect. 5.3 for details). Type-IV matrix (Fig. 5.7) consists mainly of microlites that range in size from several nanometers to several hundreds of microns, with most being several microns to several tens of microns in size. The microlites possess a variety of shapes that resemble those of microlites found in volcanic rocks, and they contain many diﬀerent minerals (see Sect. 5.3 and Chap. 6 for details); however, only plagioclase can be identiﬁed (on the basis of twinning structure), as most of the other minerals are too small to be identiﬁed under the microscope. Chemical compositions determined via electron microprobe reveal that the microlites are mainly composed of biotite, hornblende, feldspar, quartz, pyroxene, garnet, and other maﬁc minerals (see Chap. 6 for details).

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81

Fig. 5.5. Photomicrographs of transparent glass (Tg) within a silica halo surrounding an irregular quartz fragment (Qz) in a glass pseudotachylyte vein from the Fuyun Fault, Northwest China. Gl: Glass matrix. (a): Plane polarized light, (b): crossed polarized light (After Lin 1991, 1998)

Type-V matrix (Fig. 5.8) is optically non-isotropic and is composed of a mixture of glass or glassy material and microlites, being intermediate in texture between Type-I and Type-III matrix (Fig. 5.8). Powder X-ray diﬀraction data reveal a mixed character that is intermediate between that of glass- and microcrystalline-type matrix. 5.2.2 Flow Structures Flow structures developed in pseudotachylyte veins are generally characterized by thin alternating layers of contrasting color (Figs. 5.9 and 5.10) and streaks that locally curve around fragments (Figs. 5.3 and 5.4). These features are apparent in both polished specimen (Fig. 5.10a) and thin sections observed under the microscope (Figs. 5.3, 5.4, 5.6, and 5.8). Colored ﬂow streaks within pseudotachylyte veins are thought to form by the viscous ﬂow of material with non-uniform chemical composition (Lin 1991, 1994a, b) when in a molten state (Philpotts 1964; Wallace 1976; Toyoshima 1990; Lin 1991; Berlenbach and Roering 1992; Theunissen et al. 2002). Within injection veins, fragments are

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Fig. 5.6. Photomicrograph of microstructures within microcrystalline-type pseudotachylyte matrix (Pt, Type-III) from the Fuyun Fault, Northwest China. The rounded quartz fragment (Qz) that survived melting is surrounded by ﬂow structures of silica matrix that resemble the strain shadows commonly observed in mylonitic rocks. Bo: Boundary between the pseudotachylyte vein and host mylonite. c Plane polarized light. (After Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd

generally re-oriented parallel or subparallel to ﬂow streaks (Figs. 5.3 and 5.4). Flow trails similar to those observed in running water moving around boulders within a river channel are developed at both ends of oriented fragments; these structures indicate the ﬂow direction during the formation of pseudotachylyte (Fig. 5.3). Flow fabrics have also been recognized from the anisotropy of low-ﬁeld magnetic susceptibility (AMS) data for Proterozoic pseudotachylyte veins (Craddock and Magloughlin 2005). The geometry of ﬂow structures within pseudotachylyte veins commonly resembles that of similar-type folds (ﬂow folds; O’Hara 2001), in which the ﬂow streaks are curved in the central part of the vein and are parallel to the vein–wall rock contact in the vein margin (Figs. 5.9 and 5.10). Sheath-fold-like structures within pseudotachylyte veins are explained by the injection of silicate melt (which formed within the shear zone itself) due to the formation of zones of contrasting pressure (Fig. 5.10a; Berlenbach and Roering 1992). The distances between adjacent ﬂow streaks increase toward the center of the vein (Figs. 5.9 and 5.10); this is attributed to a gradient in the ﬂow velocity of the melt liquid from the margin to the center of the vein

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83

Fig. 5.7. Photomicrograph of microstructures within microlitic-type pseudotachylyte matrix (Type-IV) from the Fuyun Fault, Northwest China. Spherulitic microlites are well developed upon a rounded quartz fragment (Qz). Plane polarized light

(Lin 1991). The curved ﬂow streaks indicate a ﬂow direction of frictional melt formed during the pseudotachylyte formation (Fig. 5.10). Reynolds (1954) suggested that the transportation of ﬁne-grained fragments within a gas–solid system leads to turbulent ﬂow structures within certain geological media. The formation of ﬂow structures in melt-origin pseudotachylyte veins that resemble similar-type folds can be explained as a consequence of a wall eﬀect, termed the Bagnold eﬀect or Magnus eﬀect (Komar 1972; Barri`ere 1976) that occurs during the injection of a frictional melt (Lin 1991; Fig. 5.11). In considering the Bagnold eﬀect, the ﬂow velocity of a Newtonian ﬂuid ﬂowing within the longitudinal cross-section of a channel

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Fig. 5.8. Photomicrograph of microstructures within mixed-type pseudotachylyte matrix (Type-V) from the Fuyun Fault, Northwest China. Spider-like microlites are developed within a glassy matrix (Gl). Note that the trichites (Tr, claw-like parts of spider-like microlites) are oblique or perpendicular to ﬂow lamellae within the glassy matrix. Sr: Sanidine microlite. Plane polarized light

can be calculated from the following equation (Barri`ere 1976): U = Umax (2RY − Y2 )/R2

(5.1)

where U is the velocity of a Newtonian suspension within a lamellar ﬂow at a distance Y from the margin, Umax is the velocity in the center of channel, and 2R is the width of the channel (Fig. 5.11). The ﬂow velocity at the center of the channel, Umax , is generally greater than that in the marginal zones. The relation between U and Umax in (5.1) explains the observation that the distance between adjacent ﬂow streaks increases toward the center of the vein (e.g., Figs. 5.9 and 5.10). The gradient in ﬂow velocity is the main factor leading to the formation of the U-shaped ﬂow streaks. The direction in which the U-shaped ﬂow streaks point represents the ﬂow (injection) direction of the melt liquid (and supported ﬁne-grained material) during the formation of pseudotachylyte veins. 5.2.3 Vesicles and Amygdules Vesicles and amygdules are commonly observed in melt-origin pseudotachylytes and are generally considered to form via the extrusion of gas from a melt. The presence of vesicles and amygdules is one of the most convincing lines of evidence in support of the former molten state of the veins (e.g., Philpotts 1964; Maddock et al. 1987; Lin 1991, 1994a, b; Magloughlin 2005). Distinct vesicles and bubbles within a frictional melt were ﬁrst reported by

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85

Fig. 5.9. Photomicrograph of ﬂow structures with the geometry of similar-type folds within pseudotachylyte (Pt) from the Fuyun Fault, Northwest China. The ﬂow streaks marked by yellowish-brown and dark-brownish layers are oriented parallel to the margins of the pseudotachylyte vein but curve away from this orientation in the center of the vein. Bo: Boundary between the pseudotachylyte vein and wall mylonite. Fo: Mylonitic foliation. Plane polarized light (After Lin 1998)

Scott and Drever (1953) from the Langtang pseudotachylyte. In this case, the vesicles and bubbles are generally suﬃciently large in size that they can be observed on polished sections of a hand specimen (Fig. 5.12; see Chap. 11 for details). The concentration of vesicles and amygdules within pseudotachylyte veins commonly increases from minor amounts within marginal zones to as much as 20 Vol% within the center of the vein (Philpotts 1964), although values of 3–5 Vol% are most common (Fig. 5.13; Maddock et al. 1987; Lin 1991, 1994a). This gradient in the concentration of vesicles and amygdules demonstrates

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Fig. 5.10. Photograph (a) and photomicrograph (b) of ﬂow structures with the geometry of Sheath-fold-like structures from the central part of a pseudotachylyte vein from the Outer Hebrides Thrust, Scotland (a) and similar-type folds from the Fuyun Fault, Northwest China (b). (a) Polished section of pseudotachylyte hand sample. Arrows indicate the ﬂow direction of frictional melt formed during the c pseudotachylyte formation. (b) Plane polarized light. (b: After Lin 1994a). 2007, with kind permission from Elsevier Science Ltd

unequivocally that they formed primarily within melts of contrasting cooling rates. Vesicles and amygdules generally vary in shape with increasing size from almost perfectly circular to elliptical; this occurs over a size range of 20–200 μm in terms of the maximum dimension (Maddock et al. 1987). Elliptical vesicles have aspect ratios as high as 10 and are generally aligned parallel or subparallel to ﬂow streaks. Locally, they overlap each other to form an imperfect circular shape when viewed under the microscope. The contrast in the chemical compositions of the matrix and amygdules or vesicles means

5.2 Microstructural Characteristics

87

Fig. 5.11. Schematic cross-section of the velocity proﬁle of a Newtonian suspension within a cylinder or channel bounded by two parallel planes. The velocity in the center (Umax ) is higher than that in the margins. U: The velocity of a Newtonian suspension. 2R: The width of the channel (After Lin 1991)

Fig. 5.12. Photomicrograph of elliptical vesicles (Ve) observed on the polished surface of a hand specimen from the Langtang pseudotachylyte, Nepal. The vesicles are elongate parallel to ﬂow streaks. Sample courtesy K. Arita

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5 Pseudotachylyte Matrix

Fig. 5.13. Photomicrograph of vesicles and amygdules (Am) within a Fuyun pseudotachylyte vein. Plane polarized light (After Lin 1991)

that their contacts are generally sharp when viewed on SEM-BSE images (Fig. 5.14). Vesicles are often partly or fully ﬁlled with ﬁne-grained material and are usually associated with hydrous alteration; they are termed amygdules that often possess a zoned structure (Fig. 5.14). Amygdules within pseudotachylyte veins are generally ﬁlled by carbonate, quartz, or euhedral grains of maﬁc minerals such as magnetite and pyrite (e.g., Philpotts 1964; Maddock et al. 1987; Lin 1991, 1994a). The amygdules in some pseudotachylyte veins are inﬁlled with the assemblage of quartz-episode-chlorite-carbonate (Maddock et al. 1987), indicating that the partial recrystallization of minerals that ﬁlled the amygdules and fractures may occur at shallow depths under conditions of low temperature and pressure. The common occurrence of carbonate suggests that inﬁlling of the cavities takes place under conditions of moderate partial pressure of CO2 ; this favors the breakdown of Ca-silicates, yielding calcite and an aluminous phase (commonly chlorite) (Maddock et al. 1987). The occurrence of aligned elliptical vesicles suggests shear ﬂattening of the hot and viscous melt, probably arising from a sudden reduction in pressure

5.2 Microstructural Characteristics

89

Fig. 5.14. SEM-BSE image showing the structure of amygdules within the Fuyun pseudotachylyte. The circular patterns within the amygdule, deﬁned in the ﬁgure by variations in color, formed by the in-ﬁlling of carbonate- (Ca) and iron-rich material (Fe). The vesicle (Ve) in the lower left of the image is only partially ﬁlled (After Lin 1991, 1998)

that caused an ebullition of water vapor from solution and the partial collapse of vesicles during ﬂow of the frictional melt liquid (Maddock et al. 1987; Lin 1991, 1994a; Bjørnerud and Magloughlin 2004; Magloughlin 2005). As vesicles are generally uncommon within pseudotachylyte veins, it has been suggested that their presence indicates that gas played a secondary rather than primary role in the formation of the pseudotachylyte (Francis 1972). The absence of vesicles indicates that cooling was not suﬃciently rapid to prevent the vesicles from escaping, although it might be inferred from the absence of vesicles in some thin pseudotachylyte veins that a more rapid congelation tended to inhibit their escape (Scott and Drever 1953). Such escape structures of bubbles are found in relatively thick pseudotachylyte veins at Wenatchee Ridge, Washington, USA, where they are generally characterized by a central elliptical quartz monocrystal or polycrystal surrounded by a light-colored halo (Magloughlin 2007).

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5 Pseudotachylyte Matrix

5.3 Powder X-Ray Diﬀraction Analysis 5.3.1 X-Ray Diﬀraction Patterns for Pseudotachylyte The opaque and dark aphanitic matrix of pseudotachylyte is commonly described as glass or glassy material because when viewed under the optical microscope it shows the general optical features of glassy volcanic rocks; however, powder X-ray diﬀraction analyses reveal that most of the pseudotachylyte veins that are considered to have a melt-origin contain little or no glass or glassy/non-crystalline amorphous material (e.g., Willemse 1937; Lin 1991). Figure 5.15 shows spectra for typical melt-origin pseudotachylytes from the Outer Hebrides Thrust (Northwest Scotland) and the Woodroﬀe Thrust (Central Australia). These pseudotachylytes contain many distinct textures that are indicative of a melt origin, including the shapes of microlites, rounded, and irregularly shaped fragments, and ﬂow structures (e.g., Sibson 1975; Camacho et al. 1995; Lin et al. 2005b). The crystal peaks in these spectra are very sharp, but there is no broad band in the range of low 2θ values as seen in the Fuyun glass pseudotachylyte vein (Fig. 5.16). Some of the crystal peaks in powder X-ray diﬀraction patterns may represent the devitriﬁcation of glass or glassy material during alteration and metamorphism over the period since the primary glass material formed during the formation of the pseudotachylyte.

Fig. 5.15. Powder X-ray diﬀraction spectra of pseudotachylyte veins from the Woodroﬀe Thrust, Central Australia (a) and the Outer Hebrides Thrust, Scotland (b). Note that neither spectra show the characteristic diﬀraction patterns of noncrystalline material. The scale used for the vertical axis is the same as that used for (a–g). Qz: Quartz crystal; Pl: plagioclase crystal

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91

Fig. 5.16. Powder X-ray diﬀraction spectra of Fuyun pseudotachylyte veins and a sample of volcanic glass. (a) Volcanic glass (obsidian) from the Wadatoge volcanic area, Central Japan. (b–f) Types-I, -V, -III, -IV, and -II pseudotachylyte veins from the Fuyun Fault, Northwest China. The scale used for the vertical axis is the same as that used for (a–g). Qz: Quartz crystal; Pl: plagioclase crystal; Mi: Mica. Type-I (b) and Type-V veins show a distinct broad band in the range of low 2θ values as c that of obsidian (a). (After Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd

Convincing evidence of fresh glass and a glassy matrix within fault-related pseudotachylytes was ﬁrst reported by Lin (1991, 1994a) in samples obtained from the active Fuyun Fault, China; this observation was conﬁrmed by X-ray diﬀraction patterns. Figure 5.16 shows the powder X-ray diﬀraction spectra of ﬁve types (Type-I–V) of pseudotachylyte veins from the Fuyun Fault. The spectra show a distinctive broad band ranging from 2θ values of 12 to 42◦ , similar to that observed for volcanic glass (obsidian) as shown in Fig. 5.16a. The integrated intensity of glass or non-crystalline material (the area of the broad band) in glass pseudotachylyte (Type-I, Fig. 5.16b) is the largest, and that of the mixed-type pseudotachylyte (Type-V, Fig. 5.16c) is intermediate between that of Type-I and Types-II, -III, and -IV pseudotachylytes. This means that the amount of remnant glass or non-crystalline material is greatest in the Type-I pseudotachylyte, followed by the Type-V pseudotachylyte. This result is consistent with those derived from qualitative observations made using an optical microscope, as described above.

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5 Pseudotachylyte Matrix

Fig. 5.17. Powder X-ray diﬀraction spectra of calibration samples. (a) Measurement of intensity of glass fraction (Ig) and quartz crystal (Iq). (b–g) Standard samples containing 0, 50, 80, 90, 95, and 100 wt% glass material fused from host granitic rock within the Fuyun Fault, Northwest China. The scale used for the vertical axis is the same as that used for (a–g). Mi: mica; Qz: quartz; Pl: plagioclase; Iq: intensity of quartz; Ig: integrated intensity of glass material; If: intensity of feldspar. (After c Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd

The crystalline peaks for Type-IV pseudotachylyte consist mainly of those for quartz, feldspar minerals, and some mica minerals (Fig. 5.16e). The X-ray diﬀraction spectra for this type of pseudotachylyte reveal that the amount of intensity scattered by mica is weaker than that of the granitic country rock (Fig. 5.17b). The spectrum for Type-I pseudotachylyte (Fig. 5.16b) reveals only quartz crystal peaks, without peaks for feldspar, mica and other minerals; this is also consistent with observations made under the microscope and using an electron microprobe. The absence of feldspar and mica in the glass pseudotachylyte vein (Fig. 5.6b) indicates that they were preferentially fused during frictional melting due to their low melting temperature relative to quartz (Lin 1991, 1994a). It is possible that some or all of the micas and feldspars within Type-II and Type-IV pseudotachylyte are primary microlites formed by crystallization from frictional melt during the formation of the pseudotachylyte.

5.3 Powder X-Ray Diﬀraction Analysis

93

In contrast, the pseudotachylytes from the Outer Hebrides Thrust and Woodroﬀe Thrust show X-ray diﬀraction patterns that are typical of crystalline material, with no indication of non-crystalline or glassy material (Fig. 5.15). The diﬀraction crystal peaks in these spectra reveal that the crystalline material consists mainly of quartz and feldspar. TEM analyses conﬁrm the absence of a massive amount of glassy and non-crystalline material within the pseudotachylyte from the Outer Hebrides Thrust (Macaudi`ere et al. 1985; Wenk, 1978). 5.3.2 Quantitative Analysis of Glass and the Crystalline Fraction When undertaking a quantitative analysis of the glassy matrix within pseudotachylyte veins, pure glass samples are commonly used for calibration. In this regard, it is best to use artiﬁcially generated glass fused from the host rock as a calibration material, as this will have the same chemical composition as the glass-type pseudotachylyte. In quantitatively analyzing the Fuyun pseudotachylytes, Lin (1991, 1994a) used six standard samples made up of 0, 5, 10, 20, 50, and 100 wt% of host granitic rock mixed with artiﬁcially generated glass powder. The X-ray diﬀraction spectra of the standard samples showed a distinct broad band over 2θ values of 12 to 42◦ (Fig. 5.17c–g), which is very similar to that for obsidian (Fig. 5.17a). The integrated intensity (Ig) of the glass fraction diﬀraction (broad band over low values of 2θ) increases with increasing amount of glass (Fig. 5.17a). The scattering intensity of crystal peaks for a given 2θ position decreases with increasing amount of glass. The peaks for quartz and feldspar can be recognized even if the amount of crystalline granitederived material is as little as 5 wt% (Fig. 5.17f), but peaks for micas can only be recognized when the amount of crystalline material (standard host rock sample) is greater than 10 wt% (Fig. 5.17e). This means that the main rockforming minerals (quartz, feldspar, and mica) can be identiﬁed from the X-ray diﬀraction spectra for granite-hosted pseudotachylyte veins as long as 10 wt% of the vein consists of fragments of the original crystalline granitic rock. It is well known that i) the total integrated intensity is a constant that is independent of the relative proportions of crystalline and non-crystalline material; ii) the integrated intensity of a crystalline material increases with increasing amount of crystalline material; and iii) the integrated intensity of a glass matrix increases with increasing amount of glass having a chemical homogeneity (Klug and Alexander 1954). Accordingly, calibration curves can be obtained by measuring the integrated intensity scattered by non-crystalline or crystalline material in the spectra of the calibration substances; quantitative analyses of crystalline and non-crystalline material can then be carried out for the test samples. The integrated intensities of glass fractions in standard samples are measured using the triangle approximation method (Fig. 5.17a). It is clear that the integrated intensity (Ig) of glass increases linearly with increasing amount of glass (Fig. 5.18):

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5 Pseudotachylyte Matrix

Fig. 5.18. Relations between (a) the intensity and weight percentage of glass matrix and (b) the intensity and weight percentage of crystalline quartz in calibration substances and pseudotachylyte. The data points (b–f) correspond to the spectra c (b–f) shown in Fig. 5.16. (After Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd

Ig = K Wg

(5.2)

where K is a coeﬃcient that reﬂects the composition of the rock and Wg is the weight percent of glass in the measured material. The integrated intensity (Ig) of the glass fraction in pseudotachylyte is also measured using the triangle approximation method and is plotted on

5.3 Powder X-Ray Diﬀraction Analysis

95

a calibration curve (Fig. 5.18a). The amount of glass read directly from the calibration curve and that calculated using (5.2) vary from several weight percent to as much as 89 wt%. Fragments that produce sharp crystalline peaks make up just 11 wt% of glass pseudotachylyte veins; this ﬁnding is consistent with observations made using the optical microscope (see Chap. 7 for details). It is possible that the crystalline material within Types-II to -V Fuyun pseudotachylyte veins consist of the relicts of fragments that originated from the granitic country rocks, microlites that formed from the friction melt, and clay minerals that formed during subsequent metamorphism (Fig. 5.16). 5.3.3 Quantitative Analysis of Crystalline Material The diﬀraction spectra shown in Fig. 5.17 demonstrate that the scattered intensities of quartz, feldspar, and mica peaks for a given 2θ position increase with increasing amount of crystalline material. To quantitatively determine the relation between the scattered intensity and the amount of crystalline material, Lin (1991, 1994a) measured the intensities of the quartz peaks in the calibration samples for a given 2θ position (see Fig. 5.17a). The results are plotted in a diagram that shows the relation between intensity and weight percent in the standard samples (Fig. 5.18b). The intensity of quartz peaks (Iq) for a given 2θ position also increases linearly with increasing amount of crystalline material: Iq = k Wq (5.3) where k is a coeﬃcient that reﬂects the composition of the rock and Wq is the weight percent of quartz. The total amount of quartz within the country rock is taken as 100 wt%. This calibration curve can then be used to quantitatively determine the amounts of the most common minerals in the country rock (quartz, feldspar, and mica) within the pseudotachylyte veins. Figure 5.18b shows the intensities of quartz peaks for a given 2θ angle (26–27◦) for Types-I to -V pseudotachylytes. It is apparent from the ﬁgure that the Type-I Fuyun pseudotachylyte contains 34 wt% quartz, while the types –II to -IV pseudotachylyte contain 28–80 wt% quartz (where the amount of quartz within the granitic country rock represents 100 wt%). This means that 66 wt% of the quartz grains were melted within the glass pseudotachylyte during its formation. As stated above, the volume percentage of fragments that survived in the analyzed pseudotachylyte is about 10 vol%, corresponding to 10 wt% (based on an average density of 2.70 g/cm3 for quartz and a granitic country rock) in the glass pseudotachylyte. This is consistent with the results derived from X-ray diﬀraction data. The glass fraction in the Types-II to -V Fuyun pseudotachylyte varies from several weight percent to several tens of weight percent. This variation can be explained by the fact that crystallization occurred in the original melt and that some of the melt was quenched to form glass during the formation of the pseudotachylyte.

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5.4 Discussion 5.4.1 Properties of Glass and Glassy Matrix One of the most important questions related to the physical origin of pseudotachylyte is whether fault-generated pseudotachylyte, including that found in fractures and upon the fault plane, originated as a melt associated with frictional heating on the shear plane or/and as an assemblage of very ﬁnegrained fragments, as small as nanometers in size, that formed by crushing of the wall rock. Most of the studies carried out during the 1970s and 1980s on the origin of pseudotachylyte focused on the physical origin of the ﬁnegrained matrix (e.g., Wenk 1978; Allen 1979; Wenk and Weiss 1982; Weiss and Wenk 1983; Maddock 1983; Macaudi`ere et al. 1985; Maddock et al. 1987; Killick and Roering 1998). As described above, the powder X-ray diﬀraction patterns for the Fuyun pseudotachylytes clearly demonstrate the characteristic features of the diﬀraction patterns of glass or non-crystalline material. These observed patterns are very similar to those for volcanic glass, landslide-generated molten material (see Chap. 11 for details), and artiﬁcially generated glass fused from the granitic country rock that hosts these veins (see Chap. 12 for details). The fact that these diﬀraction patterns resemble those of glass or non-crystalline material does not completely discount the possibility that the matrix of the Fuyun pseudotachylyte consists mostly of nanometer-sized fragments of the host granitic rocks; however, the three following lines of evidence suggest that this is not the case. i) If the matrix is considered to be mainly composed of ﬁnely crushed fragments of the country rock, as proposed by Wenk (1978), then the powder X-ray diﬀraction spectra for the type-I pseudotachylyte vein should show the characteristic crystal peaks of micas and feldspars that are evident in X-ray spectra for the host granitic rock; however, with the exception of the quartz peaks shown in Fig. 16b, the peaks of the other rock-forming minerals (e.g., feldspar) in the host rock are not recognized in the spectrum for the glass-type pseudotachylyte. Accordingly, it is impossible that the non-crystalline X-ray diﬀraction patterns for the matrix represent the mineral assemblage that would be present in ﬁnely crushed fragments of the country granitic rocks. ii) The results of quantitative analyses reveal that the matrices in the ﬁve types of Fuyun pseudotachylytes record a gradation from almost completely crystalline to almost completely glass, with the glass content of the pseudotachylytes varying from just several weight percent to 89 wt%. The type-I pseudotachylyte consists of just 11% crushed fragments, which are generally larger than 2 μm in size; 89% of the matrix shows an X-ray diﬀraction pattern that is typical of glass or non-crystalline material. If the matrix was considered to consist of ultra-small fragments

5.4 Discussion

97

of crushing origin, then close to 90% of the fragments in the pseudotachylyte must occur at the submicron scale. In general, ﬁne-grained fragments are generated simultaneously with coarser fragments (Shimamoto and Nagahama 1992), producing a gradation in grain size similar to that seen in cataclastic rocks (e.g., Lin 1996, 1997a). Thus, it is considered unlikely that 90% of the matrix in the glass-type pseudotachylyte vein occurs as crush-origin ultra-ﬁne-grained fragments of submicron to nanometer size that cannot be detected by powder X-ray diﬀraction analysis. iii) A melt origin is also supported by the coexistence of ﬂow structures, vesicles or amygdules, high-temperature microlites that show a wide variety of crystal forms indicative of quenching and rapid cooling (Lin 1994b), highly irregular embayed and rounded fragments in the pseudotachylyte veins (Lin 1999a, b) that indicate a pre-melt state during pseudotachylyte formation, and the fact that the Fuyun pseudotachylyte is completely impervious to light under crossed polars and shows the optical features of glass material similar to that found in volcanic rocks. Therefore, it is concluded that the X-ray glass-diﬀraction patterns for the Fuyun pseudotachylyte represent a glass matrix and/or non-crystalline material rather than ﬁne-grained fragments that arose from crushing (Lin 1991, 1994a). The presence of glass pseudotachylyte veins demonstrates beyond doubt that fault-related pseudotachylyte can be generated by frictional melting during seismic faulting within seismogenic zones. 5.4.2 Eﬀect of Frictional Melt on Fault Strength The shear resistance (τ) of a dry fault zone is generally described as: τ

=

μσn

(5.4)

where μ is the coeﬃcient of friction, which is generally determined from the results of frictional experiments involving massive rock samples, and σn is the stress perpendicular to the fault plane, which increases with depth in the crust. In considering the presence of ﬂuid (H2 O) in the fault zone, the shear strength (τf ) is as follows: τf = μ(σn − Pf ) (5.5) where Pf is the pore ﬂuid pressure. The shear resistance (τf ) of a fault zone in the presence of ﬂuid is less than or equal to the shear strength τ of a dry fault zone. One of the major mechanisms that acts to increase the pore ﬂuid pressure within a fault zone is frictional heating during an earthquake (e.g., McKenzie and Brune 1972; Sibson 1975; Rice 2006). High-velocity frictional experiments show that the shear resistance in the presence of a fault-zone ﬂuid changes as the result of gasiﬁcation of the ﬂuid during frictional heating associated with seismic faulting (O’Hara et al. 2006). The change in shear resistance (Δτ) is described as follows:

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5 Pseudotachylyte Matrix

Δτ = μ(σn − ΔPg )

(5.6)

where ΔPg is the change in gas pressure associated with the gasiﬁcation of ﬂuid due to pore collapse. The high water contents and numerous vesicles and amygdules observed in melt-origin pseudotachylyte indicate that devolatilization and thermal pressurization occurred once frictional melt was generated during pseudotachylyte formation; consequently, the change in shear resistance (Δτ) might be expected to follow the pattern described in (5.6). A rapid reduction in shear stress (τf ) occurs with the emission of gas and the formation of large amounts of melt; the shear stress became stable once the emission of gas had ceased (O’Hara et al. 2006). High-velocity frictional experiments on frictional melting also indicate that the frictional melt that forms during earthquakes may act as either a coseismic fault lubrication or a viscous brake (Di Toro et al. 2006). These experiment results demonstrate that the shear strength of a fault zone will largely decrease once suﬃcient frictional melt forms in a fault zone and that the frictional melt plays an important role in increasing the gas pressure (ΔPg ) and as a lubricant upon the slip plane during seismic faulting. 5.4.3 Estimation of the Formation Depth of Pseudotachylyte The presence of vesicles and amygdules in pseudotachylyte indicates the waterrich character of the frictional melt, extensive hydration of fault rocks in the fault zone prior to melting, and a shallow depth of pseudotachylyte formation (Lin 1991, 1994a; Magloughlin 2005). Two methods are commonly used to estimate the formation depth of frictional melt on the basis of vesicle and amygdule structures. One method makes use of the relation between the water content of a melt and lithostatic pressure (formation depth) at the time of formation of the vesicles (e.g., Toyoshima 1990; Lin 1991, 1994a), while a second method involves estimating the lithostatic pressure from the volume percent of vesicles and amygdules preserved in the pseudotachylyte (Maddock et al. 1987; Lin 1991, 1994a). The solubility function of water determined in experiments of andesite melting has previously been used to estimate the formation depth of pseudotachylyte (Toyoshima 1990). The problem with using this method is the difﬁculty involved in determining the water content of the frictional melt at the time of pseudotachylyte formation. Although Toyoshima (1990) estimated the water content for the Hidaka pseudotachylytes, Japan, from the total amount of oxides in the bulk composition, as determined by microprobe analysis, the heterogeneous ﬁne-grained matrix makes it diﬃcult to accurately determine water content via microprobe analysis. To overcome this limitation, Lin (1991, 1994a) inferred the formation depth of the Fuyun pseudotachylyte by directly measuring the water content. The Fuyun glass pseudotachylytes contain 2.3–2.65 wt% water (H2 O+ ), which is

5.4 Discussion

99

considered to be the initial water within the frictional melt during pseudotachylyte formation. The highest measured content of 2.65 wt% is used to estimate the maximum depth of pseudotachylyte formation. Although the glass matrices are chemically heterogeneous, most are andesitic in composition (see Chap. 7 for details; Lin 1991, 1994a). Using the relation between the weight percent of water and pressure for andesitic glass (Hamilton et al. 1964), a maximum lithostatic pressure of ∼400 bars is obtained for the Fuyun glass pseudotachylytes, corresponding to a depth of 1.5 km. Maddock et al. (1987) attempted to estimate the formation depth of the Greenland pseudotachylyte on the basis of the volume percent of vesicles and amygdules. They proposed that as a large proportion of the stress drop occurs prior to the generation of melt, the magnitude of diﬀerential stress decreases and a state of plane stress is assumed, with lithostatic pressure (Plith ) equal to the mean stress (Fig. 5.19). When melt is generated, liquid pressure (Pliq ) can exceed lithostatic pressure (Plith ) due to volume change associated with frictional heating. This proposal has been conﬁrmed by high-velocity friction experiments on bituminous coal gouge, in which the high pore pressure caused by thermal pressurization and devolatilization of gouge ﬂuids contained within the fault shear zone exceeded the normal pressure (O’Hara et al. 2006) that corresponded to lithostatic pressure (Plith ). Structural relations observed between pseudotachylyte veins and fault-fractures indicate that frictional melt may continue to be generated on fault surfaces despite the expected reduction in friction. If the diﬀerential stress is small, Pliq may approach a value of –Plith + σ1 , whereσ1 is the tensile strength (Maddock et al. 1987). When extension failure occurs, the value of Pliq decreases within injection veins. The pressure at which vesicles form can be estimated using a model based on the percentage of vesicles within submarine pillow basalt (Macpherson, 1984). Using the relation between the solubility of water and pressure during pseudotachylyte formation, Lin (1991) modiﬁed Macpherson’s model by assuming H2 O and CO2 to be the major vapor components in the andesitic frictional melt. The solubility function of water was obtained by taking an approximately linear function from the plot provided by Hamilton et al. (1964). The relations obtained in this way are as follows: Wp(H2 O) = (P + 360.12)/310.95(P > 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.

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5 Pseudotachylyte Matrix

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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5 Pseudotachylyte Matrix

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; Macaudi`ere 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 (<1 minute) from a frictional melt (Spray 1987). In this chapter, we describe the morphology and chemical composition of microlites found within melt-origin Fuyun pseudotachylyte (Lin 1991, 1994b), which occurs along the Fuyun Fault, China, and that contains microlites of various shapes and mineralogy that are representative of the range of microlites described in previous reports. The formation mechanism of such microlites is also considered.

106

6 Microlites

6.2 Texture and Morphology of Microlite 6.2.1 Texture Microlites within pseudotachylyte generally occur in a wide variety of morphologies from simple forms such as acicular to complex forms such as dendritic and spherulitic (Figs. 6.1 and 6.2). It is also common to observe a gradation in the size and morphology of microlites from the margins to the centers of pseudotachylyte veins (Fig. 6.3). In the margins of veins of Fuyun pseudotachylyte, where the matrix is mostly glass or glassy material that appears opaque when viewed in plane polarized light under the optical microscope, microlites are generally smaller than 6–7 μm in size, occurring as single acicular and lath-shaped microlites. They are usually aligned parallel to subparallel to ﬂow streaks within the matrix and parallel or subparallel to the vein margin. These textural features indicate that the microlites formed during the ﬂow of a melt liquid. Toward the center of vein, the forms of the microlites become more complex, progressing from simple to skeletal, dendritic, and spherulitic shapes of up to >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)

<6 <6 <3 <6 <15 <6

<10 <8 <6 <6

<2 <2 <8 <8 <8 <8 <7 <8

Microlite habit

Simple shape group Acicular Granular Trichite Cross-shaped Lath-like Spider-like

Skeletal. group Tabular Dentritic Box Chain

Dentritic group Scoptic Quartz-feathery Feathery Plumose Fir-like Crotch Branching Pine-like <10 <50 <50 <50 <100 <100 <100 <80

<100 <150 <10 <10

<25 <10 <30 <20 <150 <30

Length (μm)

pl, bi il, opa il, opa bi, il, opa

opa qz pl, bi, opa pl, bi, am, px pl, hb, bi pl, hb, bi gr, bi, px pl, bi, hb, opa

hb, hb, hb, hb,

sa, pl il, sp, ma opa al, pl al, pl, bi, ma sa, pl

Composition

mx mi mc, mc, mc, mc, mc, mc,

mi, mi, mi, mi, mi, mi,

mx mx mx mx mx mx

cr, mc, mi mc, mi cr, mc, mi cr, mc, mx

mx, mc mx mx mx, mi mi, mx mx

Vein type

Table 6.1. Microlite morphologies and mineral types in the Fuyun pseudotachylyte veins

margin center all parts center center center center center

center center center center

margin margin margin margin all parts margin

Location within vein

110 6 Microlites

<6 <6 <5 <5 <5 <15 <15 <10 <10 <10 <10

<10 <130 <130 <130 <130 <150 <200 <150 <150 <200 <200

al, pl pl pl pl, bi pl, bi pl pl pl, bi, am, px, opa pl, bi, am, px, gr, opa gr, pl, bi, am, px, pl, gr, bi, am, px, opa

mi, mx mx mx mx mc, mi, mc, mi, mc, mi, mc, mi, mc, mi, mc, mi, mc, mi,

mx mx mx mx mx mx mx

margin center center center center center center center center center center

Mineral abbreviations: sa: sanidine, pl: plagioclase, il: ilmenite, sp: spinel, opa: opaque mineral, bi: biotite, al: alkali feldspar, hb: hornblende, qz: quartz, gr: grossular garnet, px: pyroxene, ma: magnetite. Vein types, mc: microcrystalline (Type-II), mi: microlite (Type-III), cr: cryptocrystalline (Type-IV), mx: mixed-type (Type-V) (Modiﬁed from Lin 1991, 1994b)

Sperulitic group Globular Fan Bow-tie Spherulitic Circular Sheaf Sheaf aggregate Overgrowth Stellate Dentritic compound

6.2 Texture and Morphology of Microlite 111

112

6 Microlites

Fig. 6.4. Photomicrograph showing variations in microlite sizes and shapes from the margin to the center of a pseudotachylyte vein from the Fuyun Fault, China. Plane polarized light. Bo: boundary, Mi: microlite, Qz: quartz fragment derived from host rock (After Lin 1991, 1998)

formed over at least three distinct stages. The plagioclase trunk formed ﬁrst, followed by the growth of biotite and hornblende thickets over the trunk; ﬁnally, branched thickets overgrew the earlier thickets. Spherulitic Group Spherulitic microlites are commonly observed and exhibit a wide variety of forms that have been reported from many diﬀerent localities (e.g., Maddock 1983; Lin 1994b). Such spherulites generally consist of aggregates of ﬁbrous crystals of one or more mineral species radiating from a nucleus (Figs. 6.2 and 6.9–6.11). The basic morphologies of the spherulitic microlite group include globular, fan, bow-tie, spherulitic, circular, sheaf, sheaf aggregate, overgrowth, stellate, dendritic-spherulite, and compound spherulitic (Fig. 6.2). This group of spherulites is very similar to those found in glassy volcanic rocks, as described by Lofgren (1971a). Globular forms are gradational to circular forms. It appears that globular microlites form from the ongoing crystallization of spiculate microlites. The fan, bow-tie, and spherulitic microlites consist of a dense mass of very ﬁne intergrown needles of both anorthoclase and Ferich minerals that radiate from a common nucleus. Fan spherulites consist of a single cone of needles, while bow-tie spherulites consist of two cones joined at their apices; this form can be gradational to circular spherulites as the opposing fans enlarge (Fig. 6.9).

6.2 Texture and Morphology of Microlite

113

Fig. 6.5. Photomicrographs showing the textures of microlites with simple. (a) Acicular microlites aligned parallel to ﬂow lamellae within the glassy matrix. (b) Trichitic and spider-like microlites are oriented parallel to ﬂow lamellae within the glassy matrix. Sr: spider-like sanidine microlite; Gl: glass; Tr: trichite; Qz: quartz fragment. Plane polarized light (After Lin 1991)

114

6 Microlites

Fig. 6.6. Photomicrograph showing the textures of acicular and lath-like microlite clusters. Plane polarized light

Fig. 6.7. SEM-BSE image showing the textures of tabular and skeletal hornblende microlites

6.2 Texture and Morphology of Microlite

115

Fig. 6.8. Photomicrograph showing the textures of dendritic microlites. In the branching microlites, Photomicrograph showing the textures of dendritic microlites. In the branching microlites, biotite and hornblende microlites (black) overgrow plagioclase. Pl: plagioclase microlite. Plane polarized light

The spheruical spherulites are generally composed of i) a central rounded core made up of fragments of quartz or/and feldspar (Fig. 6.10) and ii), several concentric circular zones of very ﬁne-grained crystal zones intergrown with very ﬁne crystal ﬁbers of maﬁc minerals such as biotite and hornblende that radiate from a common nucleus or central core (Fig. 6.2). Sheaf spherulites consist of two parts: a central bar and plumose ﬁbers or trichites that extend from the bar as plumes (Fig. 6.2). This texture is also commonly observed in glassy volcanic rocks (Lofgren 1971a). The sheaf aggregate spherulites consist of numerous sheaf spherulites that are generally gradational in form to sheaf aggregate spherulites (Fig. 6.11). Overgrowth and stellate spherulites consist of two parts: a central core of fragments of quartz and/or feldspar, usually with a rounded shape, and crystal ﬁbers that radiate from the core fragment (Fig. 6.2). The faces of the fragments act as sites for the nucleation of microlitic overgrowths of feldspar, biotite, hornblende, and other maﬁc minerals (Fig. 6.12). The dendritic-spherulite forms consist of spherical aggregates of dendrites that exhibit feathery, plumose, ﬁr-like, branching, and pine-like forms that radiate from a common central point of a rounded quartz or feldspar fragment (Fig. 6.2). Dendritic-spherulitic forms are generally irregular in outline and

116

6 Microlites

Fig. 6.9. Photomicrograph showing the textures of fan and bow-tie spherulites. Mi: microlite, Qz: quartz fragment. Plane polarized light

Fig. 6.10. Photomicrograph showing the textures of circular spherulites that contain zoning structures. Qz: quartz fragment. Plane polarized light

6.2 Texture and Morphology of Microlite

117

Fig. 6.11. Photomicrographs showing the textures of sheaf-shaped (a) and sheavesaggregate (c–d) plagioclase microlites. Plane polarized light (After Lin 1991, 1994b)

118

6 Microlites

Fig. 6.12. Photomicrographs showing the textures of overgrowth-shaped spherulitic microlites. Acicular biotite (black areas) and plagioclase microlites have overgrown quartz (Qz) and feldspar fragments. Mi: microlite. Plane polarized light (After Lin 1991, 1998)

have an open spherulitic form similar to that of overgrowth spherulites (Fig. 6.13). Compound spherulites are mixed aggregates of acicular and trichitic microlites and spherical, overgrowth, and stellate spherulites (Fig. 6.2). Trichitic microlites generally overgrow the spherulites to form a trichitic microlite zone as an outer zone of the compound spherulite (Figs. 6.14 and 6.15). Acicular microlites show partial intergrowths with fragments of both quartz and feldspar (Fig. 6.15).

6.3 Microlite Chemistry and Magnetic Properties 6.3.1 Microlite Chemistry The chemical compositions and mineral types of microlites are generally measured using EPMA, SEM-EDX, TEM, and powder X-ray diﬀraction (XRD) methods. Most microlites are so small that measurements of their compositions are contaminated by the surrounding matrix; however, the mineral types

6.3 Microlite Chemistry and Magnetic Properties

119

Fig. 6.13. SEM-BSE image showing the textures of dendritic-spherulitic microlites (After 1991)

of some small microlites can be identiﬁed by comparing and calculating the compositions of the microlite and its surrounding matrix. Microlites consist mainly of feldspar, biotite, hornblende, pyroxene, garnet, and other maﬁc minerals such as ilmenite, spinel, and magnetic minerals that can be identiﬁed on the basis of their chemical compositions and recognized to some degree under the optical microscope. Feldspar Feldspar microlites within the Fuyun pseudotachylyte occur as both plagioclase or alkali feldspar (Fig. 6.16; Tables 6.2 and 6.3) and possess various shapes ranging from simple to dendritic, spherical, and spherulite. Plagioclase microlites have a large compositional range from oligoclase to bytownite (An24−70 ) (Fig. 6.16), although plagioclase fragments derived from host granitic rocks occupy a limited range of oligoclase compositions (An21−29 ). Microlites of alkali feldspar also possess a wide compositional range, from Or17 to Or93 , representing anorthoclase, sanidine, and probably orthoclase and microcline (Lin 1991, 1994b). Anorthoclase microlites have a high Fe content (up to 4.34 wt% FeO), although this probably represents a degree of contamination by intergrown biotite. Fragments of alkali feldspar derived from the host

120

6 Microlites

Fig. 6.14. Photomicrographs showing the textures of compound spherulitic microlites. (a) Crystal ﬁbers of feldspar radiate from quartz fragments (Qz); these ﬁbers are then overgrown by trichitic crystals ﬁbers. (b) Acicular biotite and plagioclase microlites radially overgrew the central quartz fragment. An outer zone of trichitic crystal ﬁbers overgrew the zone of biotite (black patches) and feldspar microlite. Plane polarized light (After Lin 1991)

6.3 Microlite Chemistry and Magnetic Properties

121

Fig. 6.15. SEM-BSE images showing the textures of compound spherulite (a) and dendritic spherulite (b). Qz: Quartz. (b: After Lin 1994b)

granitic rocks all occupy a small range of microcline compositions (Or88−96 ) (Fig. 6.16); these can also be recognized optically under the microscope. There are diﬀerences in the oxide compositions of feldspar microlites and typical feldspar crystals found in volcanic rocks. These diﬀerences are also observed in biotite, hornblende, pyroxene, and grossular garnet microlites that are generally enough large in size (> several microns) and can be measured

122

6 Microlites

Fig. 6.16. Or-Ab-An ternary diagram showing the compositional relations of feldspar microlites and feldspar fragments (After Lin 1991, 1994b)

using an electronic microprobe analyzer. It is possible that these diﬀerences reﬂect contamination by ﬁne-grained inclusions within the microlites that formed during rapid non-equilibrium crystallization within a frictional melt. The presence of anorthoclase, sanidine, and An-rich plagioclase microlites, which are not present in the host granitic rocks within the Fuyun Fault zone, indicates that these microlites formed under higher temperature conditions than those of the host rocks. Biotite Biotite microlites usually occur in association with hornblende and plagioclase microlites and generally show acicular to lath-like and spherulite forms which developed within microlitic and microcrystalline pseudotachylyte veins. The chemical compositions of biotite microlites diﬀer to those of biotite fragments derived from the host granitic rocks (Table 6.4). One of the main diﬀerences is that most of the biotite microlites have a higher Ti content than that of biotite fragments derived from the host granite-mylonite rock (Fig. 6.17a). The biotite microlites having a higher Ti content than those of biotite crystals derived from the host rock are also reported from the Alpine (New Zealand) pseudotachylyte (Bossi`ere 1991).

62.59 0.65 19.68 0.00 3.82 0.16 1.25 3.21 5.83 2.69 99.74

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Total

61.95 0.68 19.70 0.05 4.34 0.00 1.87 2.87 6.10 3.04 100.50

Anor2

64.27 0.55 19.74 0.05 2.51 0.20 0.64 2.99 6.44 2.62 99.93

Anor3

63.23 0.67 21.39 0.00 1.92 0.07 0.42 3.13 6.17 3.61 100.52

Anor4

16.35 62.96 20.69

16.85 65.55 17.60

16.83 60.07 23.10

0.00 25.19 74.81

2.972 0.002 1.022 0.000 0.000 0.001 0.000 0.000 0.232 0.689

65.09 0.07 18.99 0.00 0.00 0.03 0.00 0.00 2.62 11.83 98.83

Alsp1

7.61 56.44 35.95

3.009 0.004 1.006 0.000 0.000 0.001 0.000 0.006 0.438 0.334

66.73 0.12 19.49 0.00 0.00 0.04 0.00 0.12 5.01 5.80 97.31

Alsp2

0.81 55.67 43.52

3.015 0.006 1.027 0.000 0.000 0.006 0.001 0.005 0.417 0.326

66.70 0.16 19.35 0.00 0.00 0.15 0.01 0.23 4.76 5.65 97.01

Alsp3

Alkali feldspar microlites

0.05 16.70 83.25

2.866 0.033 1.027 0.001 0.020 0.001 0.012 0.049 0.154 0.719

62.55 0.94 18.89 0.02 1.55 0.03 0.53 0.98 1.72 12.21 99.41

Alsp4

4.16 11.37 84.47

2.854 0.009 1.072 0.007 0.060 0.006 0.015 0.045 0.123 0.916

60.54 0.26 19.29 0.19 1.53 0.15 0.22 0.89 1.35 15.22 99.65

Alsp5

4.30 13.03 82.67

2.887 0.004 1.045 0.007 0.033 0.009 0.010 0.046 0.140 0.892

60.85 0.19 18.70 0.19 0.93 0.22 0.14 0.91 1.53 14.74 98.34

Alsp6

1.24 2.91 95.85

2.930 0.007 0.971 0.003 0.040 0.007 0.022 0.014 0.033 1.086

62.83 0.21 17.67 0.03 1.03 0.17 0.32 0.28 0.37 18.25 101.20

Frag1

0.00 7.63 92.37

2.955 0.006 1.015 0.002 0.003 0.000 0.000 0.000 0.072 0.872

64.43 0.16 18.72 0.04 0.07 0.00 0.00 0.00 0.81 14.85 99.08

Frag2

0.00 8.80 91.20

2.954 0.003 1.029 0.000 0.000 0.002 0.000 0.000 0.084 0.870

64.13 0.24 18.95 0.00 0.00 0.04 0.00 0.00 0.94 14.80 99.40

Frag3

Alkali feldspar fragments

0.00 11.18 88.82

2.968 0.004 1.019 0.000 0.000 0.002 0.000 0.000 0.105 0.834

64.34 0.12 18.75 0.00 0.00 0.05 0.00 0.00 1.17 14.16 98.60

Frag4

FeO∗ : total Fe as FeO. Anor1–Anor4: anorthoclase microlites, Alsp1–Alsp6: alkali feldspar microlites, Frag1–Frag4: alkali feldspar fragments derived from host rock. (After Lin 1991, 1994b)

An:Ab:Or ratios An 18.90 Ab 62.20 Or 18.90

Number of ions on the basis of eight oxygens Si 2.822 2.790 2.863 2.835 Ti 0.022 0.023 0.018 0.022 Al 1.046 1.046 1.036 1.096 Cr 0.000 0.002 0.002 0.000 Fe 0.144 0.167 0.094 0.070 Mn 0.006 0.000 0.008 0.003 Mg 0.084 0.113 0.042 0.028 Ca 0.155 0.138 0.143 0.146 Na 0.510 0.533 0.556 0.520 K 0.155 0.448 0.149 0.200

Anor1

wt%

Anorthoclase microlites

Table 6.2. Chemical compositions of alkali feldspar microlites and fragments analyzed by EDX

6.3 Microlite Chemistry and Magnetic Properties 123

51.88 0.43 29.43 0.06 0.99 0.00 0.16 11.67 3.92 0.45 99.00

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Total

50.42 0.25 31.01 0.05 0.98 0.08 0.18 13.01 3.30 0.37 99.65

Lab2

50.46 0.32 30.10 0.06 1.10 0.00 0.07 12.25 3.23 0.42 98.27

Lab3 50.99 0.44 30.28 0.10 0.91 0.01 0.16 12.44 3.66 0.38 99.37

Lab4

66.97 30.79 2.24

66.15 30.97 2.88

63.75 33.95 2.30

27.56 70.53 1.91

2.699 0.002 1.290 0.001 0.016 0.001 0.003 0.259 0.665 0.018

60.79 0.06 24.65 0.03 0.44 0.04 0.05 5.46 7.72 0.31 99.54

Olig1

27.11 71.41 1.48

2.703 0.002 1.293 0.000 0.021 0.002 0.004 0.261 0.673 0.008

60.48 0.05 24.47 0.00 0.55 0.06 0.07 5.43 7.74 0.15 99.00

Olig2

24.19 73.13 2.68

2.734 0.006 0.267 0.001 0.014 0.000 0.000 0.226 0.682 0.025

61.33 0.18 24.10 0.03 0.38 0.00 0.00 4.72 7.89 0.44 99.07

Olig3

69.73 26.68 3.59

2.315 0.014 1.617 0.004 0.041 0.024 0.013 0.680 0.294 0.026

50.38 0.40 29.02 0.11 1.04 0.20 0.18 13.43 2.84 0.58 97.18

Bct1

32.78 65.13 4.09

2.722 0.014 1.509 0.002 0.066 0.009 0.035 0.287 0.571 0.018

60.99 0.40 22.36 0.06 1.78 0.08 0.53 6.01 6.60 0.32 99.11

And1

49.41 47.16 3.43

2.482 0.003 1.489 0.000 0.058 0.000 0.030 0.461 0.440 0.032

55.01 0.08 27.99 0.00 1.53 0.00 0.45 9.54 5.02 0.56 100.18

And2

22.02 75.64 2.34

2.762 0.006 1.233 0.001 0.017 0.000 0.000 0.207 0.711 0.022

62.36 0.18 23.62 0.04 0.45 0.00 0.00 4.35 8.27 0.38 99.83

Frag5

21.60 77.73 1.67

2.838 0.004 1.146 0.002 0.012 0.000 0.000 0.194 0.698 0.006

64.69 0.13 22.17 0.07 0.32 0.01 0.00 4.12 8.20 0.11 99.82

Frag6

25.94 72.55 1.51

2.713 0.010 1.276 0.003 0.027 0.003 0.000 0.241 0.674 0.014

61.14 0.29 24.40 0.07 0.73 0.10 0.00 5.08 7.83 0.25 99.89

Frag7

Plagioclase fragments

29.20 68.21 2.59

2.688 0.006 1.306 0.002 0.017 0.002 0.005 0.271 0.633 0.024

61.97 0.19 24.57 0.06 0.48 0.05 0.08 5.82 7.53 0.44 101.19

Frag8

FeO∗ : total Fe as FeO. Lab1–Lab4: labradorite microlites, Olig1–Olig3: oligoclase microlites, Bet1: betonite microlite, And1–And2: andesine microlites, Frag5–Frag8: plagioclase fragments derived from host rock. (After Lin 1991, 1994b).

An:Ab:Or radios An 60.07 Ab 37.16 Or 2.77

Number of ions on the basis of eight oxygens Si 2.384 2.305 2.327 2.332 Ti 0.015 0.009 0.011 0.015 Al 1.594 1.671 1.644 1.632 Cr 0.002 0.002 0.002 0.004 Fe 0.034 0.034 0.043 0.035 Mn 0.000 0.003 0.000 0.001 Mg 0.006 0.013 0.005 0.011 Ca 0.565 0.673 0.622 0.609 Na 0.349 0.293 0.290 0.325 K 0.026 0.021 0.025 0.022

Lab1

wt%

Plagioclase microlites

Table 6.3. Chemical compositions of plagioclase microlites and fragments

124 6 Microlites

94.87

–

35.53 5.48 15.67 0.04 13.57 0.15 11.04 0.36 0.68 7.35

Bio1

36.67 2.25 17.43 0.00 22.74 0.51 10.40 0.32 0.60 7.91 0.31 99.24

Bio2

37.20 2.27 16.28 0.00 17.73 0.23 12.43 0.34 0.34 8.55 0.34 95.71

Bio3

38.76 6.90 15.93 0.04 11.62 0.17 14.13 0.76 1.80 7.51 0.16 97.65

Bio4 37.52 4.28 17.89 0.14 15.30 0.20 11.38 0.23 0.76 8.79 0.01 96.46

Bio5 36.70 4.44 17.78 0.09 14.40 0.24 13.22 0.16 0.86 8.45 0.16 96.54

Bio6 37.32 4.15 17.39 0.00 13.65 0.08 13.93 0.10 0.87 8.60 0.10 96.17

Bio7 37.43 7.27 16.66 0.00 11.35 0.26 14.11 0.23 1.21 7.60 0.06 96.17

Bio8 36.99 2.82 17.65 0.02 17.65 0.24 12.95 0.18 0.58 8.57 0.12 96.35

Bio9 36.48 2.42 17.33 0.08 21.01 0.32 8.01 0.00 0.00 9.45 0.18 95.28

Bio10 37.00 2.32 18.08 0.00 21.19 0.50 8.44 0.00 0.05 9.35 0.27 97.20

Bio11

Biotite fragments

Number of ions on the basis of twenty-two oxygens Si 5.352 5.401 5.555 5.515 5.512 5.367 5.453 5.403 4.133 5.854 5.545 Ti 0.624 0.250 0.251 0.738 0.473 0.488 0.456 0.790 0.237 0.278 0.262 Al 2.797 3.028 2.866 2.671 3.098 3.065 2.996 1.834 2.137 3.126 3.193 Cr 0.005 0.000 0.000 0.004 0.013 0.011 0.000 0.000 0.003 0.009 0.000 Fe 2.353 2.802 2.214 1.382 1.880 1.765 1.668 1.370 1.650 2.688 2.653 Mn 0.019 0.063 0.030 0.020 0.025 0.029 0.010 0.032 0.022 0.042 0.064 Mg 2.492 2.305 2.767 2.975 2.491 2.882 3.033 3.036 2.156 1.829 1.884 Ca 0.059 0.050 0.055 0.115 0.036 0.024 0.015 0.036 0.022 0.000 0.000 Na 0.200 0.173 0.096 0.497 0.216 0.244 0.245 0.340 0.127 0.000 0.015 K 1.424 1.488 1.628 1.362 0.648 1.578 1.603 1.400 1.221 1.846 1.787 P – 0.039 0.041 0.019 0.001 0.022 0.012 0.006 0.011 0.024 0.035 XMg 0.51 0.45 0.55 0.68 0.59 0.62 0.61 0.68 0.56 0.40 0.42 ∗ FeO : total Fe as FeO. Bio1–Bio9: biotite microlites, Bio10–Bio14: biotite fragments derived from host rock. (After Lin 1991, 1994b)

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 Total

wt%

Biotite microlites

Table 6.4. Chemical compositions of biotite microlites and fragments

5.531 0.278 3.135 0.000 2.682 0.045 1.856 0.003 0.064 1.802 0.043 0.41

36.97 2.49 17.78 0.00 21.44 0.36 8.32 0.02 0.22 9.58 0.31 97.26

Bio13

–: undetermined

5.543 0.253 3.215 0.000 2.612 0.001 1.848 0.001 0.053 1.847 0.021 0.41

37.13 2.25 18.28 0.00 20.93 0.45 8.30 0.00 0.18 9.70 0.17 97.39

Bio12

5.573 0.272 3.177 0.000 2.664 0.038 1.802 0.000 0.000 1.754 0.027 0.40

36.94 2.40 17.89 0.00 21.12 0.30 8.01 0.00 0.00 9.82 0.21 96.67

Bio14

6.3 Microlite Chemistry and Magnetic Properties 125

126

6 Microlites

Fig. 6.17. Ti–Al diagrams showing chemical variations and compositional relations between biotite and hornblende microlites and equivalent fragments of host rock from microlitic pseudotachylyte veins along the Fuyun Fault, China. (a) Ti–Al plot of biotite microlites and host-rock fragments. (b) Ti–Al plot of hornblende microlites and host-rock fragments. Note that the biotite and hornblende microlites have higher Ti-contents than that of fragment derived from the host grantic rock (After Lin 1991, 1994b)

Hornblende Hornblende microlites commonly show skeletal and spherulitic forms. Although measurements were performed carefully under the same conditions as those for feldspars, the totals for hornblendes from the Fuyun pseudotachylyte are lower than those for common standard hornblende (Table 6.5). Low totals were also obtained for pyroxene and garnet microlites. The low totals reﬂect the high water contents within the Fuyun pseudotachylytes, which are about 2 wt% higher than those in the host rocks (Lin 1994a). As with the biotite microlites, hornblende microlites have higher Ti contents than that of hornblende fragments derived from the host granitic rock (Fig. 6.17b). The presence of Ti-rich hornblende and biotite microlites indicates that they formed under higher temperature conditions than those of the host granite-mylonite rocks.

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127

Table 6.5. Chemical composition of hornblende microlites and fragments Hornblende microlites wt% SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 Total

Horn1 42.17 2.88 10.74 0.00 14.50 0.44 2.26 9.17 2.40 0.71 0.88 96.19

Fragments

Horn2

Horn3

Horn4

Horn5

Horn6

Horn7

Horn8

37.85 6.14 3.00 0.08 13.38 0.93 10.32 9.46 1.88 1.70 0.47 95.67

36.30 7.02 11.10 0.13 17.70 0.73 9.48 9.45 2.03 1.30 0.36 95.62

44.17 1.62 13.27 0.00 15.93 0.34 11.33 6.25 2.79 0.64 0.24 96.58

44.24 1.25 14.65 0.00 14.65 0.34 11.98 6.81 2.33 0.61 0.47 97.33

45.18 0.88 13.47 0.00 16.74 0.30 11.27 9.01 1.41 0.88 0.36 99.50

46.16 0.65 11.37 0.01 14.90 0.26 11.21 8.48 2.45 1.11 0.28 96.68

42.83 1.08 14.92 0.15 15.68 0.36 12.34 3.74 2.12 3.17 0.29 96.68

Number of ions on the basis of twenty-three oxygens Si 6.340 5.830 5.705 6.546 6.475 6.787 6.905 Ti 0.326 0.711 0.830 0.181 0.138 0.100 0.085 Al 1.902 2.361 2.056 2.317 2.478 1.804 1.942 Cr 0.000 0.000 0.017 0.000 0.000 0.000 0.001 Fe 1.824 1.723 2.326 1.974 1.973 2.103 1.865 Mn 0.056 0.121 0.097 0.042 0.042 0.038 0.033 Mg 0.748 2.368 2.221 2.503 2.614 2.524 2.477 Ca 1.477 1.562 1.592 0.993 1.068 1.450 1.353 Na 0.710 0.560 0.619 0.803 0.661 0.411 0.431 K 0.135 0.334 0.261 0.121 0.114 0.168 0.192 P 0.112 0.062 0.048 0.030 0.059 0.046 0.023 XMg 0.600 0.580 0.490 0.550 0.590 0.541 0.570 FeO∗ : total Fe as FeO. Horn1–Horn5: hornblende microlites, Horn6–Horn8: blende fragments derived from host rock. (After Lin 1991, 1994b).

6.352 0.123 2.657 0.015 2.006 0.034 2.817 0.609 0.609 0.635 0.037 0.580 horn-

Pyroxene Pyroxene microlites within the Fuyun pseudotachylytes have low Ca contents and extraordinarily high Al contents; compositionally, they are clinohypersthene and magnesian pigeonite (Fig. 6.18; Table 6.6). Al-rich pyroxene microlites have also been reported from the Hidaka pseudotachylytes (Toyoshima 1990), while Al-rich clinopyroxene microlites have been reported from the Musgrave pseudotachylyte veins (Wenk and Weiss 1982) and within a supercooled, artiﬁcially generated frictional melt (Spray 1988). Such Alrich pyroxene microlites are suggested to have formed from non-equilibrium crystallization within a melt (Toyoshima 1990); consequently, the presence of these microlites within pseudotachylytes indicates an origin of quenching of a melt under non-equilibrium chemical conditions (Lin 1991, 1994b).

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6 Microlites

Fig. 6.18. Compositional variations among pyroxene microlites within pseudotachylyte veins from the Fuyun Fault, China (Modiﬁed from Lin 1991)

Garnet Garnet microlites are grossular in composition and Ca-poor (Fig. 6.19; Table 6.7). The chemical compositions of garnet microlites diﬀer to those of minerals from other petrogenetic settings such as Al-rich and Ca-poor pyroxene, Fe-rich feldspar, and Ca-poor grossular garnet. This may reﬂect the fact that the microlites formed under conditions of non-equilibrium crystallization during the rapid cooling of melt during pseudotachylyte formation. Garnet microlites generally coexist with pyroxene microlites and possess dendriticspherulitic forms. Other Maﬁc Minerals Other maﬁc microlitic minerals such as ilmenite, spinel, and magnetite are rarely observed in pseudotachylyte veins; where present, they generally coexist with biotite and hornblende microlites. Ilmenite microlites are generally of dendritic-skeletal form, although some are granular in shape. Although these maﬁc microlites are too small to determine their mineralogy, it can be inferred from their chemical compositions and that of the matrix (Table 6.8). 6.3.2 Magnetic Properties Studies of melt-origin pseudotachylyte reveal that most natural pseudotachylyte and some black fault-gouge veins are more highly magnetic than their protoliths. This occurs because of the presence of ﬁne-grained ferromagnetic minerals within the veins (e.g., Francis and Sibson 1973; Camacho et al. 1995; Nakamura et al. 2002; Ferre et al. 2005). Such veins exhibit high natural remnant magnetization (NRM) carried by single domain to pseudo-single domain

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129

Table 6.6. Chemical compositions of pyroxene microlites (Py1–Py6) Pyroxene microlites wt%

Pyr1

Pyr2

Pyr3

Pyr4

Pyr5

Pyr6

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 Total

45.50 1.11 14.80 0.14 15.00 0.19 20.17 1.03 0.72 0.31 0.32 99.29

47.26 1.08 14.50 0.04 14.73 0.33 16.41 1.73 1.10 0.23 0.31 97.58

45.92 1.07 14.18 0.08 14.59 0.28 17.83 1.68 1.68 0.39 0.20 97.01

46.22 1.03 13.36 0.00 15.52 0.21 19.41 1.31 0.71 0.37 0.33 98.52

46.62 0.92 13.20 0.16 15.15 0.08 20.63 0.92 0.42 0.16 0.26 98.46

45.75 1.09 13.68 0.01 16.18 0.30 19.58 1.50 0.38 0.02 0.32 98.80

Number of ions on the basis of six oxygen Si 1.660 1.742 1.714 Ti 0.031 0.030 0.030 Al 0.636 0.631 0.624 Cr 0.004 0.002 0.001 Fe 0.458 0.504 0.505 Mn 0.006 0.011 0.008 Mg 1.097 0.902 0.991 Ca 0.040 0.069 0.067 Na 0.051 0.079 0.059 K 0.015 0.011 0.019 P 0.010 0.009 0.007 Ca:Fe:Mg ratios Ce 2.51 Fe 28.69 Mg 68.80 FeO∗ : total Fe as FeO. (After Lin 1991, 1994b)

4.81 31.87 63.32

4.43 30.03 65.55

1.704 0.029 0.585 0.000 0.478 0.008 1.067 0.052 0.053 0.017 0.010 3.24 29.96 66.80

1.708 0.025 0.570 0.003 0.464 0.003 1.127 0.036 0.030 0.008 0.008 2.21 28.53 69.26

1.684 0.030 59487 0.000 0.498 0.009 1.074 0.509 0.027 0.001 0.010 3.62 30.53 65.85

(PSD) magnetite (e.g., Piper and Poppleton 1988; Enomoto and Zheng 2001; Ferre et al. 2005; Nakamura et al. 2005). Figure 6.20 shows an example of the magnetic properties of melt-origin pseudotachylyte from the Santa Rose mylonite shear zone, Southern California (Ferre et al. 2005), whose petrologic characteristics and ﬁeld structures are reported in detail by Wenk et al. (2000). The magnetic susceptibility of the pseudotachylyte vein is much higher than that of the wall rock (Fig. 6.21). The Santa Rose pseudotachylyte displays distinct hysteresis properties that are consistent with the grain size of pseudo-single domains. Such magnetic properties are also found in artiﬁcially generated molten materials (Nakamura

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6 Microlites

Fig. 6.19. Mn-Mg-Fe ternary diagram showing compositional variations among grossular microlites within the pseudotachylyte veins from the Fuyun Fault, China (Modiﬁed from Lin 1991)

et al. 2002) and black fault gouges (Enomoto and Zheng 2001; termed crushorigin pseudotachylyte in Chap. 9) from the Nojima Fault (Japan), as shown in Fig. 6.22. Many pseudotachylyte veins that possess stable remnant magnetism show distinctive magnetic properties, but their formation process and mechanism of magnetization remain poorly understood and continue to be debated. Some natural pseudotachylyte veins contain higher proportions of maﬁc microlitic minerals such as biotite and ilmenite (and locally magnetite) than that of the wall rocks (Warr and van der Pluijm 2005). Alteration and metamorphism of pseudotachylyte veins may also lead to the formation of magnetic minerals such as hematite (e.g., O’Hara and Sharp 2001). The results of high-velocity frictional melting experiments reveal that microlites that are mainly composed of maﬁc minerals such as hornblende and magnetite preferentially form from frictional melt during seismic faulting over a short period of <1 minute (Spray 1988; Nakamura et al. 2002). These concentrations of ferromagnesian microlite minerals within melt-origin pseudotachylyte are considered to have formed by the preferential fractional crystallization of maﬁc minerals within a friction melt during formation of the pseudotachylyte (Warr and van der Pluijm 2005). Based on the results of

6.3 Microlite Chemistry and Magnetic Properties

131

Table 6.7. Chemical compositions of garnet microlites (Gar1–Gar5) Garnets (Grossular) In spherulitic microlites wt%

Gar1

Gar2

Gar3

Gar4

Gar5

49.38 0.49 18.78 0.09 10.17 0.00 0.61 17.19 1.14 0.11 0.57 98.68

39.13 0.04 24.76 0.08 10.14 0.20 0.05 22.69 0.10 0.02 0.21 97.07

39.09 0.06 25.83 0.12 9.80 0.22 0.06 22.99 0.12 0.00 0.16 98.43

38.69 0.11 38.69 0.07 10.37 0.21 0.18 21.99 0.00 0.27 0.20 97.73

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 Total

38.21 0.02 25.44 0.11 9.61 0.24 0.18 22.66 0.13 0.00 0.14 96.78

Number Si Ti Al Cr Fe Mn Mg Ca Na K P

of ions on the basis of twelve oxygens. 2.952 3.633 2.966 2.967 0.002 0.027 0.002 0.003 2.316 1.629 2.270 2.312 0.006 0.006 0.004 0.008 0.620 0.626 0.660 0.623 0.016 0.000 0.014 0.014 0.020 0.066 0.006 0.008 1.876 1.355 1.892 1.871 0.024 0.162 0.016 0.017 0.000 0.011 0.002 0.000 0.008 0.036 0.014 0.011

Mn:Fe:Mg ratios Mn 2.46 0.00 Fe 94.51 90.46 Mg 3.05 9.54 ∗ FeO : total Fe as FeO. (After Lin 1991, 1994b)

2.06 97.06 0.84

2.17 96.59 1.24

2.990 0.006 2.258 0.005 0.671 0.014 0.021 1.821 0.000 0.023 0.014 2.13 94.90 2.97

high-velocity frictional melting experiments and studies of natural fault gouges and pseudotachylyte, it is inferred that the ﬁne-grained magnetite minerals within the veins formed as a new magnetic carrier in association with extensive oxidation of maﬁc silicates during rapid cooling from a frictional melt; accordingly, pseudotachylytes hold great potential in terms of paleo-intensity reconstructions of the magnetic ﬁeld during seismic faulting or impact events (Nakamura et al. 2002). Ferre et al. (2005) proposed an alternative magnetization mechanism in which the high magnetic remanence in pseudotachylyte veins is related to coseismic electrical currents. It is possible that the maﬁc

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6 Microlites

Table 6.8. Chemical compositions of ilmenite (Ilm1–Ilm2), spinel (Spi1) and magnetic (Mag1) mineral microlites Ilmenite wt% SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 NiO Total

Ilm1

Ilm2

Spinel Spi1

Magnetic Mineral (Mag1)

2.58 47.33 3.18 0.02 36.37 0.13 4.35 0.13 0.34 0.26 0.11 0.00 94.88

4.10 55.63 3.28 0.00 27.20 1.27 1.79 0.31 0.53 0.13 0.10 0.01 94.37

11.33 0.31 46.18 0.07 27.83 0.55 7.46 0.75 1.09 0.61 0.00 0.00 96.18

19.39 5.56 9.03 0.15 50.43 0.96 3.65 0.50 0.71 3.40 0.20 0.28 94.26

Atomic ratios (Ilmenite:6; Spinel and magnetic mineral:24) Si 0.122 0.147 1.715 4.024 Ti 2.049 2.005 0.040 0.868 Al 0.178 0.185 9.203 2.208 Cr 0.001 0.000 0.010 0.024 Fe 1.045 1.089 3.934 8.752 Mn 0.005 0.051 0.078 0.168 Mg 0.309 0.128 1.879 0.282 Ca 0.009 0.016 0.135 0.112 Na 0.032 0.049 0.356 0.288 K 0.016 0.008 0.131 0.225 P 0.008 0.008 0.000 0.026 Ni 0.002 0.000 0.000 0.038 Total 3.776 3.686 17.680 17.400 ∗ FeO : total Fe as FeO. (After Lin 1991, 1994b)

microlite minerals act as a recorder of the past or current magnetic ﬁeld under which the pseudotachylyte vein formed or was metamorphosed.

6.4 Discussion of the Mechanism of Microlite Formation Given the above, this section considers the origins and formation mechanisms of the various microlite morphologies and mineral species. For the Fuyun pseudotachylytes, a melting origin for the microlites rather than a metamorphic origin is indicated by the following factors.

6.4 Discussion of the Mechanism of Microlite Formation

133

Fig. 6.20. Hysteresis properties of pseudotachylyte veins from the Santa Rose mylonite shear zone, Southern California. (A) Hysteresis plot reproduced from Day et al. (1977). Note that most of the data display hysteresis properties that are consistent with the grain size of the pseudo-single domain. (B) Hysteresis loop of pseudotachylyte specimens showing the lack of a goose-neck shape and a saturation acquisition consistent with a PDF magnetic population. (Modiﬁed from Ferre c et al. 2005). 2007, with kind permission from Elsevier Science Ltd

i) The high-temperature microlite minerals such as sanidine, anorthoclase, An-rich plagioclase (An40−70 ), pyroxene, grossular garnet, Ti-rich biotite, and hornblende are not derived from the host granitic rocks and clearly formed at high temperatures.

134

6 Microlites

Fig. 6.21. Spatial variations in magnetic susceptibility across a pseudotachylyte vein from the Santa Rose mylonite shear zone, Southern California. The map was produced by interpolation over a grid of 170 measurements taken on a smooth planar surface using a Bartington MS probe. (A) Rock slab with an overlying 10 mm grid. Note the injection vein at the upper left. (B) Magnetic susceptibility map (10−3 SI). (C) Proﬁle of magnetic susceptibility along the section X–X’. Note the increase in susceptibility across the vein. (After Ferre et al. 2005). Color ﬁgure courtesy c E. Ferre. 2007, with kind permission from Elsevier Science Ltd

6.4 Discussion of the Mechanism of Microlite Formation

135

Fig. 6.22. Three-dimensional hysteresis plots reproduced from Day et al. (1977) on the basal plane and low-ﬁeld ferromagnetic susceptibility along the vertical axis for artiﬁcially generated pseudotachylytes and natural fault gouges with black and c blue-gray veins. (After Nakamura et al. 2002). 2007, with kind permission from Elsevier Science Ltd

ii) From the margins to the centers of the microcrystalline and microlitic pseudotachylyte veins, the sizes of microlites increase and the shapes vary from simple acicular shapes to complex dendritic-spherulitic forms. iii) Within glassy pseudotachylyte veins, the microlites consist almost entirely of sanidine, anorthoclase, and An-rich plagioclase (An > 40–70). Individual crystals and crystal ﬁbers within glassy veins are generally smaller than those within microcrystalline and microlitic pseudotachylyte veins. iv) Ti-rich biotite and hornblende crystals overgrow plagioclase microlites (An > 40–70) and show complex dendritic-spherulitic textures that are absent in the host granitic mylonitic rocks. v) The fact that spider-like microlites are stretched parallel to subparallel to streak-like ﬂow structures in the glassy pseudotachylyte veins indicates that these microlites formed during the ﬂow of a melt during pseudotachylyte formation. In metamorphic rocks, high-temperature minerals such as sanidine and anorthoclase indicate metamorphic conditions of the sanidinite facies. The anorthite component in plagioclase generally increases with increasing temperature in metamorphic rocks (Sen 1959); this is also observed in crystal-growth experiments (Lofgren 1974). Plagioclase fragments derived from the host granite-mylonite rocks generally makes up less than 30% An component (see Fig. 6.16), but most plagioclase microlites have a higher An component, varying from An30 to An70 . This indicates that the plagioclase microlites formed at much higher temperatures than those of the plagioclase crystals within the host granite-mylonite rocks. Analyses of microstructures within fault rocks in the Fuyun Fault zone and their deformational history suggest that the host granite-mylonite formed at

136

6 Microlites

metamorphic conditions of the epidote-amphibolite facies, corresponding to depths of < 10–15 km (Lin and Fan 1984). The mylonites were overprinted by cataclastic rocks that formed along the fault zone with the injection of carbonate, quartz, and chlorite veins, which are in turn cut by pseudotachylyte veins (Lin 1991, 1994a). This indicates that the depth of formation of the Fuyun pseudotachylytes was less than 10 km. Assuming a continental geothermal gradient of 30◦ C/km, the temperature of the host granite-mylonite rocks immediately prior to injection of the pseudotachylyte veins is estimated to have been < 300◦C. It is inferred that the Fuyun pseudotachylyte formed at a depth of ∼1.5 km as described in Chap. 5. Accordingly, the high-temperature minerals within the pseudotachylyte clearly formed from a primary melt during the formation of the pseudotachylyte rather than during overprinting metamorphism and alteration. The observed crystal forms, such as skeletal, dendritic, and spherulitic, appear to be a function of the cooling rate or the degree of supercooling of the melt liquid within which the crystal grew; this has been conﬁrmed by direct experiments involving feldspars, pyroxene, and olivine (Lofgren 1980). Crystal-growth experiments reveal that the morphology of crystals grown from an An-Ab-H2 O system varies with the degree of undercooling (Lofgren 1974, 1980; Doherty 1980). With more pronounced undercooling, individual crystal ﬁbers become progressively ﬁner until the morphology changes from polyhedral to dendritic and ﬁnally spherulitic. This explains the changes in microlite morphology and size observed from the margins to the centers of veins and the diﬀerences observed between glassy and microcrystalline veins within the Fuyun pseudotachylytes. The fact that microlites within glass in the margins of veins are small in size and simple in shape indicates that they formed primarily via supercooling or the quenching of a melt. The complex dendritic and coarse spherulitic textures found in the central parts of veins represent the ﬁnal products due to diﬀerent cooling stages of the melt. The overgrowth of Ti-rich biotite and hornblende on An-rich plagioclase clearly indicates multi-stage development of the microlites; this is also indicated by the textural variations observed between the outer edge and center of individual sheaf plagioclase microlites. Such textures are commonly developed in glassy volcanic rocks (e.g., Lofgren 1971b). The body bars of spider-like microlites found in areas of glassy matrix are largely oriented parallel to ﬂow structures within the pseudotachylyte veins (Figs. 5.8 and 6.5). The fact that trichites that grew on the edges of the body bars are bent, in combination with the occurrence of streak-like ﬂow structures in the veins (Fig. 5.8), shows clearly that these textures formed within a ﬂowing melt. Some of the trichites that grew at the edges of the body bars, however, are oblique or even perpendicular to the body bars and matrix ﬂow structures (Fig. 6.5b). The above textures indicate that the microlites formed over a three-stage process. In the ﬁrst stage, the body bars grew and were reoriented parallel to the ﬂow structures within the ﬂowing melt. During the second stage, the body

6.4 Discussion of the Mechanism of Microlite Formation

137

bars acted as nucleation sites for crystal growth. The trichites that grew on the side faces of the body bars were reoriented, together with the body bars themselves, parallel to ﬂow lamellae during ongoing ﬂow of the melt. In the third and ﬁnal stage, randomly oriented trichites grew on existing body bars once ﬂow had ceased. Similar ﬂow textures have been described from glassy volcanic rocks (Ross 1962). Metamorphic processes could not have formed these ﬂow textures at the shallow depths at which the glass was preserved. Therefore, it is assumed that the ﬂow textures of the spider-like microlites are primary products derived from the ﬂow of a melt during pseudotachylyte formation (Lin 1994b). Although experimental results demonstrate that ﬁne spherulitic textures can form from the devitriﬁcation of volcanic glass (Lofgren 1971b), this origin can be rejected in the case of the Fuyun pseudotachylytes. The textures of the microlites and the presence of high-temperature microlite minerals indicate that these microlites formed via the quenching or supercooling of a melt. The presence of pure silica glass indicates that the temperature of the melt was at least 1450◦C (Lin 1994a; see Sect. 8.3.2 for details). Microlites observed within the Fuyun pseudotachylytes have a varied mineralogy but are mainly composed of An-rich plagioclase (An20−70 ) and alkali feldspar (including sanidine and anorthoclase), as well as lesser clinopyroxene, grossular garnet, biotite, hornblende, and magnetic minerals. Biotite and hornblende microlites have variable Ti contents that are generally higher than those of biotite and hornblende fragments derived from the host rocks. The fact that these microlites possess a wide variety of morphologies and textures, when combined with the presence of sanidine, anorthoclase, An-rich plagioclase, pyroxene, Ti-rich biotite, and hornblende, indicates that they are primary crystals that formed during the rapid cooling or quenching of a melt during pseudotachylyte formation.

7 Fragments Within Pseudotachylyte Veins

7.1 Terminology A number of diﬀerent terms have been used for the fragments found in pseudotachylyte veins, including fragments (Shand 1916; Lin 1994a, b, 1999b), xenoclasts (Wallace 1976), lithic porphyroclasts (Maddock 1983), xenoliths (Magloughlin 1989), lithic clasts or lithic fragments (Magloughlin and Spray 1992), lithic fragments and lithoclasts (O’Hara 1992), and clasts (Shimamoto and Nagahama 1992; Karson et al. 1998; Tsutsumi 1999; Ray 1999, 2004). The term fragment is deﬁned as a piece of rock that has been detached or broken from a pre-existing mass (Jackson 1997); this reﬂects the mechanical origin of the fragments that occur in all kinds of rocks, including fault rocks. The term porphyroclast is generally used to describe a relatively large crystal within a metamorphic rock that originated by metamorphic recrystallization rather than a mechanical process. The other terms listed above are generally used for sedimentary rocks (e.g., lithoclast and lithic clast ) and for volcanic rocks (e.g., xenolith and xenoclast ), with their usage restricted to clasts that formed by mechanical or chemical disintegration of a larger rock mass during sedimentary and volcanic processes rather than by fault movement. Faultrelated pseudotachylytes generally contain a substantial volume of angular to rounded fragments of the wall rock that survived abrasion and/or fusion during seismic faulting events within the fault zone; therefore, in this book the term fragment is used according to the above deﬁnition to avoid any confusion that might arise in the use of alternative terms that are commonly used in the description of metamorphic, sedimentary, and volcanic rocks and their weathered and altered products.

7.2 Fragments that Resemble Conglomerate Clasts Rock fragments within pseudotachylyte veins generally vary from nanometerto centimeter-sized fragments within an individual vein to several tens of centimeters in size where found within generation zones. Fragments may be

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angular, subangular, irregularly embayed, or rounded. A number of pseudotachylyte-bearing fault zones are characterized by abundant rounded pebble- and cobble-sized fragments that are randomly scattered throughout the ﬁne-grained matrix of the pseudotachylyte vein (Fig. 7.1). Such faultrelated pseudotachylyte zones are similar in texture to sedimentary conglomerates and are comparable in terms of fragment texture, size, and rounding to those observed in pseudotachylyte veins associated with meteorite impact (Fig. 7.2). Sibson (1975) described pebble and cobble fragments within pseudotachylyte veins as quasi-conglomerate and proposed that they form in situ within the jog area between en echelon fault veins that originated during the ﬁnal stages of deformation during a single slip increment (Fig. 4.12). Pseudotachylyte zones that contain quasi-conglomerate were ﬁrst reported by Shand (1916) from the Vredefort impact crater, South Africa (Figs. 2.2 and 7.2). Such rocks contain a gradation in fragment shape from angular to subangular, irregularly embayed, and rounded. For both breccia and quasi-conglomerate that occur within pseudotachylyte zones, small fragments generally show signs of rotation, with the amount

Fig. 7.1. Photograph of paired pseudotachylyte-generation zones along the Outer Hebrides Thrust, Scotland. Note the numerous rounded pebble- to cobble-sized fragments of contrasting lithology to the wall rock. The coin shown for scale is 2.4 cm across

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141

Fig. 7.2. 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 pseudotachylyte matrix, thereby indicating a melt origin. The hammer shown for scale is 35 cm long

of rotation increasing with increasing ratio of matrix to fragments (Fig. 7.3). A complete gradation can be seen in the ﬁeld between ladder networks of echelon fault veins linked by arrays of tensional injection veins, breccias, and quasi-conglomerates (Sibson 1975). In some pseudotachylyte veins, fragments are concentrated in the central part of the vein. This can be explained by the wall eﬀect on a Newtonian melt ﬂuid during pseudotachylyte formation (Macaudi`ere et al. 1985). Some conglomerate-bearing pseudotachylyte zones contain rock fragments of contrasting lithology to the adjacent wall rocks (Figs. 7.1 and 7.3). This indicates that the fragments were transported from a generation zone located far from the current location of the vein. The presence of foreign pebble- and cobble-sized fragments indicates a melting origin for the pseudotachylyte, involving rapid injection under high pressure into coseismically dilated fractures within massive basement rocks followed by cooling of the melt liquid. Microscopically, pseudotachylyte veins are universally characterized by fragments of various shapes and sizes within an ultra-ﬁne-grained or glassy matrix (Fig. 7.4). For melt-origin pseudotachylytes, fragments are generally sub-rounded to rounded, surrounded by zoned spherulitic microlites and ﬂow structures (Figs. 5.3 and 7.5).

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7 Fragments Within Pseudotachylyte Veins

Fig. 7.3. Pseudotachylyte-bearing zone along the Outer Hebrides Thrust, Scotland (a) and the Alpine Fault, New Zealand (b). Note the numerous and small-sized and rounded and irregular pebble-sized fragments that show evidence of rotation. The foliations within the diﬀerent fragments have contrasting orientations and are oblique to the foliation within the host schist in both (a) and (b). The coin shown in (a) and scale bar shown in (b) for scale are 2.4 cm across and 5 cm long, respectively

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Fig. 7.4. Photomicrographs showing the various shapes of fragments within meltorigin pseudotachylyte veins along Outer Hebrides Thrust, Scotland (a) and the Langtang Himalaya landslide zone (b). Note that fragments make up approximately 70% of the vein shown in (a) (a: After Lin 1991)

7.3 Grain-size Analysis 7.3.1 Grain-size Distribution Within Melt-origin Pseudotachylyte The grain-size reduction of fragments is a ubiquitous phenomenon within fault rocks within both brittle and ductile fault shear zones. Many studies have analyzed the grain-size distribution of fragments within pseudotachylyte veins

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7 Fragments Within Pseudotachylyte Veins

Fig. 7.5. Photomicrographs showing the rounded shapes of fragments within the melt-origin Fuyun (a) and Woodroﬀe (b) pseudotachylytes. Note that zoned spherulitic microlites surround the rounded and irregular fragments. Plane polarized light

with the aim of understanding the formation mechanisms of pseudotachylyte veins and associated mechanical processes (Schwarzman et al. 1983; Lin 1991, 1992, 1996, 1997; Shimamoto and Nagahama 1992; Lin 1994; Tsutsumi 1999; Ray 1999, 2004; Kawamoto 2004). Analyses of the grain sizes of fragments are generally based on microscope observations and to a lesser degree SEM-EDX images.

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In their study of the Outer Hebrides pseudotachylyte, Scotland, Okamoto and Kitamura (1990, 1996) were the ﬁrst to document a linear relationship between grain size and the cumulative number of fragments. Shimamoto and Nagahama (1992) developed a unique method of measuring the grain size of fragments within the Woodroﬀe pseudotachylyte, Central Australia, using optical photomicrographs. As the proportion of the ﬁne-grained matrix was found to be incompatible with the grain-size distribution of the fragments, they suggested that the pseudotachylytes formed by melting rather than ultracomminution. Shimamoto and Nagahama (1992) found that for certain ranges of grain size, the sizes of fragments show a fractal distribution on a grain size–frequency distribution diagram. The authors proposed the following relationships among grain size, fractal dimension D, and the cumulative number of fragments:

or

N = N r−D ,

(7.1)

D = (log N − logN )/logr,

(7.2)

where N is the cumulative number of fragments with sizes greater than r and N is a constant that depends on the number of measurements. This represents a basic equation for the fractal analysis of fragments within pseudotachylyte, as modiﬁed from the equation deﬁned by Mandelbrot (1983) for self-similar fragmented isotropic parallelpipeds. Fractal fragmentation yields a linear distribution of grain size–frequency when plotted in log–log coordinates. The value of D for cataclastic rocks is generally obtained by determining the slope of a least-squares ﬁtted straight line on a grain size–frequency diagram. It should be noted that fractal fragmentation theory is based on the assumption that the fragmented mass is isotropic in all dimensions. In contrast to above, Ray (1999, 2004) found that grain-size distributions for the Rajashan pseudotachylytes, West India, depart slightly from a power law relationship; instead, they follow a modiﬁed power-law distribution (Fig. 7.6). Log–log plots of fragment size vs. cumulative frequency yield gently curving trends that reﬂect a decrease in the slope of the trend for ﬁner grain sizes (Fig. 7.6). This trend is similar to those reported for experimentally generated pseudotachylyte (Fig. 7.7; Tsutsumi 1999). Ray (2004) concluded that this pattern is a characteristic feature of all fault-generated pseudotachylytes that form via the transformation of cataclastically deformed rocks by the pervasion of frictional melt. Based on grain-size distributions, Ray (2004) inferred that fragments and ﬁnely crushed matrix material are the products of grain-size reduction via cataclasis, which generates a fractal set of fragments that obey a power-law size distribution (Ray 2004). Tsutsumi (1999) measured the grain sizes of relict fragments within experimentally generated pseudotachylyte and found that the grain-size distribution is fractal, as with natural pseudotachylytes (Fig. 7.7). He found that

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Fig. 7.6. Cumulative frequency plots of fragment sizes within the Rajashan pseudotachylyte, Western India. (a–h): Fragment sizes as measured in diﬀerent parts of the pseudotachylyte vein. The numbers preceded by SKR are the sample numbers. c (After Ray 1999). 2007, with kind permission from Elsevier Science Ltd

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Fig. 7.7. Cumulative frequency plots of fragment sizes within experimentally generated pseudotachylyte. N is the cumulative number of fragments whose mean diameter (r, open circles) or major axis (A, open triangles) is greater than A. (a) Cumulative frequency plots of fragments within experimentally derived pseudotachylyte vein that remained within the generation (slip) zone. (b) Cumulative frequency plots of feldspar, magnetic, and all fragments within experimentally derived molten material that fell from the generation zone during the experiment. (c) Cumulative frequency of fragments within experimentally derived pseudotachylytes plotted again 1+r/r’, on the basis of the form: N = N’ (1+r/r’)−D , where N’ and r’ are constant, D is a powder-law exponent (Nagahama et al. 1992). HFR (040, 069, 070): c Run number. (After Tsutsumi 1999). 2007, with kind permission from Elsevier Science Ltd

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the proportion of ultra-ﬁne-grained matrix is less than 0.5% of the measured area and concluded that the majority of the ﬁne-grained matrix formed by melting rather than crushing. 7.3.2 Grain-size Distribution: A Discussion Theoretical and applied studies of the size distribution of fragments within pseudotachylyte are mainly based on an empirical power-law relationship, as shown in (7.1) and (7.2). This law is also widely used in studies of brittle fault rocks such as natural and experimental fault gouges and cataclastic rocks (e.g., Sammis et al. 1986, 1987; Turcotte 1986a, b; Biegel and Sammis 1989; Marone and Scholz 1989; Zhao et al. 1990). The empirical power-law relationship applied to the grain-size distribution of fragments within pseudotachylyte is mainly derived on the basis of the following two premises proposed by Shimamoto and Nagahama (1992): i) that the shapes of fragments do not change signiﬁcantly with varying size, and ii) that the same type of power law relationship holds for ultra-ﬁne submicroscopic fragments. Most studies of the grain-size distribution of fragments within pseudotachylyte report that the power law relationship evident in the grain-size distribution can be used as a quantiﬁed index of frictional melting in understanding the formation of the fragments (e.g., Shimamoto and Nagahama 1992; Tsutsumi 1999; Ray 1999, 2004); however, there are two basic questions that remain unanswered: i) does the grain size of fragments always follow the power law relationship as shown in (7.1) and (7.2) for a grain size–frequency diagram, and ii) if the answer to the preceding question is yes, what is the geological and physical signiﬁcance of the fractal dimension D? Let us now consider the ﬁrst question. Although from an empirical basis it is conventionally considered that the grain-size distribution of brittle fault rocks, including fault gouges, is fractal, a number of studies have yielded results that suggest otherwise (e.g., Lin 1994; Lin 1996, 1997a; Wilson et al. 2005). Wilson et al. (2005) demonstrated that plots of the grain number–size distribution for 250 samples of fault gouge obtained from the San Andreas Fault, USA, and the fault rupture zone that triggered a recent earthquake within a South African gold mine do not ﬁt a systematic fractal distribution; this is because diﬀerent slopes (corresponding to diﬀerent fractal dimensions) are attained with progressive disaggregation. This means that fragments of diﬀerent size ranges within fault gouges may have contrasting grain-size distributions. Lin (1994) and Lin (1996, 1997a) measured the grain sizes of cataclastic rocks including fault gouge, breccia, cataclasite, and crushing-origin pseudotachylyte from the Iida–Matsukawa Fault, an active fault in Central Japan, and found similar grain-size distributions in the diﬀerent rock types; however, only some of the grain sizes in the range from 10 to 100 μm showed a fractal pattern (Fig. 7.8; Lin 1996). These grain-size distributions are similar to those reported from experimentally generated pseudotachylyte (Fig. 7.7a, b).

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Fig. 7.8. Cumulative frequency plot of fragment sizes within cataclastic rocks including crushing-origin pseudotachylyte, fault gouge, foliated cataclasite, and nonfoliated cataclasite from the Iida–Matsukawa Fault zone, central Japan. (After c Lin 1996). 2007, with kind permission from Elsevier Science Ltd

Comparison of the grain-size distributions of natural pseudotachylyte and fault gouge reveals that the grain-size distributions within pseudotachylyte do not conform to a uniform power-law size distribution with the same slope for all grain sizes (Figs. 7.6, 7.7a, b); instead, diﬀerent grain sizes show curved or linear trends with varying slopes. It appears that the grain-size distribution within cataclasitic fault rocks including pseudotachylytes described in previous studies is largely dependent on the sizes and numbers of fragments. In both fault gouges (Fig. 7.8) and natural and experimentally derived pseudotachylytes (Figs. 7.6 and 7.7), the number of fragments decreases sharply for fragments smaller than several microns. This sharp reduction may reﬂect the limit of resolution of photomicrographs and SEM-EDX images. In addition,

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in the above studies the number counts of fragments > 100 μm in size are generally less than 10, which may be too small a number to be statistically signiﬁcant. In summary, it is clear that the grain-size distribution of fragments does not always obey a power law relationship and that diﬀerent types of power laws may govern the grain-size distribution of diﬀerent size classes, especially for ultra-ﬁne fragments. The fractal dimension D shown in (7.1) and 7.2) has been calculated theoretically to be ≤2 in terms of area-frequency (two-dimensional measurements) and ≤3 in terms of volume-frequency (three-dimensional measurements) (Zhao 1990). The theoretical power law relationship assigned to the grain-size distribution of fragments within pseudotachylyte is based on the proposal of Shimamoto and Nagahama (1992) that host rocks within seismic fault zones are crushed prior to melting; this generates a fragment grain-size distribution that follows a power-law relationship (Ray 2004). Based on experimental studies, Sammis et al. (1986) attributed a decrease in particle size with increasing conﬁning pressure to a suppression of the extension of microfractures and subsequent reduction in the spacing of axial microfractures. Their gouge particles obeyed both log-normal and power-law size distributions. Sammis et al. (1987) measured particle sizes from unaltered natural fault gouge from the Lopez Fault Zone, within the San Gabriel Fault Zone, USA, and found that the particle size distribution for the size class from 10 μm to 1 cm (in diameter) is self-similar, with a fractal dimension of D = 2.6 ± 0.1 in volume-frequency (three-dimensional measurements). Based on the fact that for all size classes the fractal dimension characterizes a distribution that minimizes the number of nearest neighbors of the same size, Sammis et al. (1987) proposed a comminution mechanism in which the probability of particle fracture is determined solely in terms of the relative size of the nearest neighbors to the grain of interest rather than its size or mineralogy. In contrast, Wilson et al. (2005) reported that for two-dimensional measurements the grain-size distributions of fault gouges from the San Andreas Fault Zone do not follow a single fractal distribution but vary from D = 1 to D = 6. Experimental results for deformed granite reveal that the grain-size distribution of cracked fragments and mature gouge have diﬀerent values of D on a log-log plot; for grain sizes of <2 μm, D is ∼1 for both cracked fragments and gouge, whereas for grain sizes of >2μm D is 1.56–1.68 for cracked fragments and 2.0–2.26 for mature fault gouge in a two dimension measurement (Heilbronner and Keulen 2006). The lower D values for grain sizes of <2 μm may reﬂect the fact that some fragments were too small to be measured. The value of D tends to increase with increasing deformation of the granite from cracked fragments to mature gouge. This result indicates that the dimension D is inﬂuenced by both grain size and intensity of deformation. In a study of cataclastic rocks such as cataclasite, breccia, fault gouge, and pseudotachylyte from the active Iida–Matsukawa Fault, Central Japan, Lin (1996) found that the grain-size distributions (for two-dimensional

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151

measurements) only showed self-similar patterns for the size range of 10–100 μm; the fractal dimension D varied between 1 and 5 (Fig. 7.8). The grain sizes of relict fragments within experimentally derived pseudotachylyte also show a wide range of D values, between 2.4 and 6.0, depending on the mineral type of the fragments (Fig. 7.7c; Tsutsumi 1999). On the basis of the above review, it appears that the geological signiﬁcance of the fractal dimension D for fragments within pseudotachylyte veins and related fault rocks remains unclear and requires further study.

7.4 Fabrics of Fragments and Degree of Rounding 7.4.1 Fabrics Many studies have documented that fragments set in ﬁne-grained and/or glassy matrices within melt-origin pseudotachylyte hosted in continental quartzo-feldspathic rocks are mostly composed of quartz and to a lesser degree feldspars (e.g., Sibson 1975; Lin 1991, 1994a). Fragments of maﬁc minerals are generally rare within such pseudotachylyte hosted within quarto-feldspathic rocks . Fragments of biotite are occasionally observed close to the margins of such pseudotachylyte veins, but only on SEM-EDX images. The volume percentages of fragments within pseudotachylyte veins vary from 10 to 70% (Figs. 7.4 and 7.5; e.g., Lin 1991, 1996). On the basis of the total area of fragments measured under the microscope, it is estimated that the volume percent of fragments within Types-I, -III, and -V Fuyun pseudotachylytes (see Sect. 5.1 for the deﬁnition of these types of pseudotachylytes) is 10, 14, and 16% respectively (Lin 1991, 1994a), with up to 70% recorded in a sample of Outer Hebrides pseudotachylyte (Fig. 7.4a). Quartz fragments within glassy-type pseudotachylyte veins commonly exhibit transparent or translucent rims (Figs. 5.4 and 5.5), while those within microlitic veins are commonly zoned (Fig. 7.5). The transparent rims are isotropic and have the optical characteristics of glass within the matrix, similar to that of the glass of the thin section (mainly silica with minor feldspar) when viewed under the microscope (Lin 1994a). The zoned rims are generally composed of ﬁne microlites of feldspar and biotite (see Chap. 8 for details). Fragments of quartz and feldspar are also commonly surrounded by iron-rich opaque material that occurs as spindle shapes that resemble pressure shadows that are observed in mylonitic rocks and commonly aligned parallel to ﬂow structures within the adjacent matrix. The rims are variable in thickness, ranging from narrow submicron-scale rims that represent just a small fraction of the total size of the grain to thick rims that make up the bulk of the grain (Fig. 7.5). 7.4.2 Degree of Rounding of Fragments In sedimentology, the shapes of grains and fragments are used to infer transport processes and the depositional environment (e.g., Wadell 1932;

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Krumbein 1941; Research Group of Sedimentary Rocks of Japan 1983), while in structural geology the shapes of grains and fragments are used to interpret deformation processes and the formation mechanisms of fault rocks (e.g., Lin 1991, 1999b). Fragments within pseudotachylyte veins form synchronously with pseudotachylyte formation via crushing, melting, and partial assimilation during transportation, prior to ﬁnal injection and cooling. Such fragments possess a variety of sizes and shapes, including rounded, embayed, and irregular outlines (Figs. 7.2–7.5). Large (cobble-sized) fragments of up to several tens of centimeters across are also found in large-scale fault-generated pseudotachylyte zones (Figs. 7.1 and 7.3) and impact-related zones (Figs. 7.2–7.4). Most of the fragments within melt-origin pseudotachylyte are more rounded and possess greater numbers of irregular embayments (Figs. 7.4 and 7.5) than those within cataclasitic rocks (Fig. 7.9). Such rounded and embayed fragments set within a ﬁne-grained matrix are generally considered to be indicative of a melting origin rather than crushing (e.g., Shand 1916; Philpotts 1964; Sibson 1975; Wallace 1976; Allen 1979; Maddock 1983; Lin 1991, 1994a, b, 1996, 1999b; Magloughlin 1992). The degree of rounding of fragments derived from brittle fault rocks can be used as a quantitative index in understanding the formation process and mechanism of pseudotachylyte formation (Lin 1999b). Thus, a high degree of rounding is considered to indicate wear or melt of the fragment; the degree of wear or melt depends upon the fragment’s size and hardness of fragments.

Fig. 7.9. Photomicrograph of the microstructure of non-foliated cataclasite from the Iida–Matsukawa Fault, Central Japan. Note that all fragments possess angular to sub-angular shapes. Crossed polarized light

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153

Wadell (1932) deﬁned roundness (Rd) observed in a single plane (twodimensional measurement) as Rd =

n

(ri/R)/n

(7.3)

i=1

where ri is the radius of curvature of a corner of the grain, R is the radius of the maximum inscribed circle in the plane of measurement, and n is the number of corners apparent in the observed plane (Fig. 7.10). In this scheme, a perfectly round object has a roundness of 1.0. The measured or calculated roundness of a fragment can also be checked with reference to a roundness chart (Fig. 7.11). Lin (1999) measured the roundness of fragments from the following three natural pseudotachylytes and related cataclasitic rocks: the well-known meltorigin Outer Hebrides pseudotachylyte (e.g., Sibson 1975), the Fuyun pseudotachylyte (Lin 1994a, b), the Musgrave pseudotachylyte (Camacho et al. 1995; Lin et al. 2005b), the Osumi pseudotachylyte (Fabbri et al. 2000), crushingorigin pseudotachylyte (Lin 1996), and cataclastic rocks such as cataclasite, fault breccia, and fault gouge from the Nojima Fault and the Iida-Matsukawa Fault, Japan (Lin 1996, 1997a; Lin et al. 1998a, b). As the host rocks of all of the analyzed samples are granitic rocks dominated by quartz and feldspar, measurements of roundness were performed on quartz and feldspar fragments to avoid the aﬀect of the varying hardness and melting-point temperatures of other minerals. Measurements of roundness are generally undertaken from

Fig. 7.10. Sketch of the method employed to determine the roundness (Rd) of fragments, where ri is the radius of curvature of a corner and R is the radius of the maximum inscribed circle in the plane of measurement. (After Lin 1991, 1999b). c 2007, with kind permission from Elsevier Science Ltd

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Fig. 7.11. Chart for which to estimate the roundness (Rd) of fragments. Rd varies from 0.1 to 1.0. (Modiﬁed from Research Group of Sedimentary Rocks of Japan 1983; Lin 1999b)

photomicrographs taken under plane polarized light and SEM images of thin sections (Fig. 7.12). The degree of precision involved in measuring the radii of corner curvatures is largely determined by fragment size. To avoid measurement errors related to grain size, measurements were performed over a wide area of the sample for all fragments larger than 10 μm in diameter. To

Fig. 7.12. SEM image showing the rounded shape of fragments within the Fuyun pseudotachylyte. Spherulitic microlites are developed around the fragments (After Lin 1994b)

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155

minimize bias in the measurements, the roundness of fragments was calculated using (7.3) and with reference to a rounding chart (Fig. 7.11). The obtained roundness values vary from 0.0 to 1.0 (Fig. 7.13). Between 35 and 90% of fragments within melt-origin pseudotachylyte have a roundness of >0.4, whereas all of the fragments within crushing-origin pseudotachylyte and cataclastic rocks have a roundness of <0.4; 80–95% of fragments in such rocks have a roundness of <0.3 (Fig. 7.13). It is clear that there is a marked diﬀerence in the degree of rounding between fragments in melt-origin pseudotachylyte and those in cataclastic rocks such as crushing-origin pseudotachylyte and fault gouge.

7.5 Formation of Rounded Fragments: A Discussion The shape fabrics of fragments have been reported from many natural cataclastic rocks (e.g., Anderson et al. 1983; Mitra 1984; Chester and Logan 1986; Sammis et al. 1987; Evans 1988; Chester et al. 1993; Evans and Chester 1995; Lin 1992, 1994, 1999). Based on a comparison of fragment shapes derived from cataclastic rocks (plate IIa in Anderson et al. 1983; Fig. 5.5d in Chester and Logan 1986; Figs. 4.4–8.8 in Sammis et al. 1987; Fig. 1.1a in Evans 1988) and crushed grains generated in triaxial compression experiments (Figs. 5.5 and 6.6 in Menendez et al. 1996) using the roundness chart shown in Fig. 7.11, it is estimated that the majority of fragments with a size of >10 μm reported in these previous studies have a roundness of <0.4, although this approach provides only an approximate roundness. This ﬁnding suggests that a fragment roundness of >0.4 is unlikely to form by fracturing or chipping during cataclastic deformation. The surface textures of fragments within fault gouges are generally aﬀected by corrosion and dissolution within groundwater (e.g., Kanaori 1982). Based on the surface smoothness of quartz grains from 250 fault-gouge samples obtained from 14 major faults in Japan, Kanaori (1982) categorized quartz grains into four groups (I to IV) that reﬂect progressive corrosion and dissolution within groundwater during periods of fault movement from the Late Pleistocene to the Miocene; this transformation timing was estimated on the basis of geological evidence. Similar surface shapes upon quartz grains have been reported from fault gouges obtained from several major active faults in China (Yang 1986; Yang et al. 1985a, b, 1994a, b). Comparison of the shapes of these grains with those in the roundness chart shown in Fig. 7.11 reveals that the grains have a roundness of <0.3. Thus, it is unlikely that fragments with a roundness of >0.4 achieved such a degree of rounding via corrosion and dissolution within groundwater percolating through fault gouges developed along active fault zones. Between 35 and 90% of quartz and feldspar fragments within intensivelystudied melt-origin pseudotachylyte have a roundness of >0.4 (Fig. 7.13), whereas all fragments that originate via cataclasis have a roundness of <0.4.

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Fig. 7.13. Histograms of the degree of rounding of fragments from various sources. (a–d) Melt-origin pseudotachylyte (Pt). (e–h) Cataclastic rocks. Sources of the samples are as follows: (a) Fuyun pseudotachylyte, (b) Outer Hebrides pseudotachylyte, Scotland, (c) Woodroﬀe Thrust, Central Australia, (d) Osumi Fault Shear Zone, Southern Japan, (e) Iida–Matsukawa cataclasite, Japan, (f) Iida–Matsukawa Fault breccia, Japan, (g) Iida–Matsukawa crushing-origin pseudotachylyte, Japan, (f) Nojima Fault gouge, Central Japan (Modiﬁed from Lin 1999b)

7.5 Formation of Rounded Fragments: A Discussion

157

This pattern strongly suggests that fragments with a roundness of >0.4 formed via melting rather than fracturing, abrasion, or chipping. In conclusion, the origin of highly rounded fragments (roundness >0.4) found within pseudotachylytes are best explained in terms of frictional melting. The favored explanation for the generation of frictional melt within fault zone rocks is that friction-related heat generated at a seismic slip rate of >∼1 m per second on the fault plane accumulates because of the low thermal conductivity of rocks, thus leading to melting. Frictional melting may aﬀect the size and shape of fragments, or it may lead to the melting of a uniform thickness of the fragment rim, thereby reducing its size and increasing the degree of rounding (Ray 2004). Sibson (1975) suggested that fragments within pseudotachylytes are rounded by the marginal meltdown of angular fragments within a melt liquid phase. On the basis of simple conduction theory, he explained that the isothermal surface of angular fragments, while in general being subparallel to the cooling surface or contacts, would tend to be “rounded oﬀ” at sharp angular projections and corners (Fig. 7.14). The large surface area of fragment corners in contact with a melt liquid, relative to that of the fragment edges and sides, means that the corners absorb more heat energy than the sides during melting; consequently, the corners of fragments are melted more rapidly than the straight edges and sides of the fragment, leading to the corners being rounded oﬀ.

Fig. 7.14. Sketch of the process of thermally induced rounding of angular rock fragments by marginal meltdown. Dashed line A represents the transient isotherm, solid line B within the fragment indicates the outline of rounded fragment (Modiﬁed from Sibson 1975)

158

7 Fragments Within Pseudotachylyte Veins

In summary, the above review indicates that fragments within pseudotachylyte veins with a roundness of >0.4 are rounded via frictional melting rather than fracturing or chipping; the degree of rounding can be used as a special index of disintegration associated with attrition and melting when evaluating the formation process and mechanism of pseudotachylytes (Lin 1991, 1999).

8 Chemical Composition and Melting Processes of Pseudotachylyte

8.1 Introduction Obtaining an understanding of the properties of frictional melt is one of the major goals in studying the mechanisms and processes of coseismic frictional melting that occurs within a seismogenic fault zone. One approach in this regard is to study the chemical composition of melt-origin pseudotachylyte. Many previous studies have investigated the bulk-vein and matrix compositions of natural and experimentally derived pseudotachylyte veins and related host rocks. In early studies, workers sought to compare the bulk chemical compositions of pseudotachylyte veins with that of their host rocks using conventional X-ray ﬂuorescence (XRF) methods. It was Shand (1916) who ﬁrst noted that the bulk composition of pseudotachylyte is similar to that of the host rock; this observation has since been conﬁrmed in numerous studies (e.g., Philpotts 1964; Sibson 1975; Lin 1991, 1994a; Bossi`ere 1991; Maddock 1992; Camacho et al. 1995; Wenk et al. 2000). However, prior to the 1970s it proved diﬃcult to measure the chemical composition of ﬁne-grained pseudotachylyte matrix using conventional analytical methods because it was impossible to measure the composition of the glassy matrix independently of fragments and crystals that also occurred within the veins. This problem has been addressed since the 1970s, however, as use of the electron microprobe enables the measurement of the chemical composition of heterogeneous matrix, even at the nanometer scale, while avoiding contamination from ﬁne-grained fragments and crystals (e.g., Allen 1979; Spray 1988; Toyoshima, 1990; Bossi`ere, 1991; Lin 1991, 1994a; Maddock 1992; Warr and van der Pluijm 2005). Thus, it is now possible to quantitatively access the properties of frictional melt, including its degree of heterogeneity, as well as diﬀerences in the chemical compositions of the ﬁne-grained matrix and the pseudotachylyte vein as a whole. This chapter reviews previous studies of the chemical composition of pseudotachylyte veins and related host rocks and discusses the properties of frictional melt and melting processes that occur during the formation of pseudotachylyte.

160

8 Chemical Composition and Melting Processes of Pseudotachylyte

8.2 Bulk-Vein and Matrix Compositions 8.2.1 Bulk Composition of Pseudotachylyte Veins The bulk chemical compositions of natural pseudotachylyte veins and associated host rocks are generally determined via XRF analysis. Previous studies report that pseudotachylyte is generally chemically heterogeneous, both among and within veins (e.g., Ermanovics et al. 1972; Toyoshima 1990; Lin 1991,1994a). Pseudotachylyte veins vary in composition from silica-poor basic types to silica-rich acidic compositions; this variation is largely controlled by the host rock type (e.g., Lin 1991; Maddock 1992; Magloughlin 1992). The average bulk compositions of typical melt-origin pseudotachylyte veins reported in literature vary from basic to acidic, with SiO2 contents of 40 to 75 wt%, similar to that of the host rocks (Tables 8.1–8.3; e.g., Philpotts 1964; Lin 1991, 1994a). The ﬁve types of pseudotachylyte veins recognized within the Fuyun Fault zone (see Chap. 5 for details) are also heterogeneous in terms of bulk compositions, with SiO2 contents ranging from 59 to 69 wt% (Table 8.1b; Fig. 8.1). The concentrations of SiO2 , Na2 O, FeO, MgO, CaO, Na2 O, and K2 O vary widely, but no systematic diﬀerences exist in the concentrations of these oxides in pseudotachylyte veins and host rocks; however, the host granitic rocks are enriched in SiO2 relative to the veins (Table 8.1; Fig. 8.2). The similarity in bulk composition between the pseudotachylyte veins, including fragments of host rock within the veins, and the host rocks themselves is generally interpreted to indicate that pseudotachylyte is derived from the host Table 8.1a. Bulk compositions of the host rocks (a: H1–H7) where the pseudotachylytes (b: P1–P7) injected analyzed by XRF. The water contents (H2 O+ and H2 O− ) were determined by independently ignition loss wt%

H1

H2

H3

H4

H5

H6

H7

SiOe2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 H2 O + H2 O −

68.12 0.42 15.15 3.65 0.09 1.24 3.60 3.67 2.10 0.13 0.54 0.38

72.77 0.25 13.98 2.10 0.05 0.49 2.00 3.57 3.88 0.07 0.51 0.34

71.32 0.32 14.10 2.95 0.07 0.63 2.28 3.77 3.63 0.10 0.83 −

69.75 0.48 14.28 3.73 0.06 0.67 2.31 3.61 4.01 0.14 0.52 0.43

74.13 0.21 13.83 1.81 0.03 0.47 0.95 3.73 3.80 0.12 0.94 −

68.00 0.50 15.83 4.87 0.10 2.43 0.99 3.10 3.14 0.14 0.93 −

67.15 0.50 15.83 3.95 0.08 1.94 2.42 4.48 2.66 0.11 0.89 −

0.92 0.84 0.83 0.95 0.94 0.93 Total H2 O FeO∗ : total Fe as FeO, -: unmeasured (Data from Lin 1991 1994a)

0.89

8.2 Bulk-Vein and Matrix Compositions

161

Table 8.1b. wt%

P1

P2

P3

P4

P5

P6

P7

SiOe2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 H 2 O+ H 2 O−

68.52 0.46 14.55 3.60 0.08 0.64 2.28 3.51 3.96 0.14 2.34 0.26

67.82 0.47 15.53 4.42 0.09 1.28 3.24 3.11 2.93 0.13 2.65 0.55

67.57 0.48 15.31 4.10 0.08 1.20 3.19 3.77 2.26 0.14 2.05 0.81

62.86 0.70 16.84 5.33 0.07 2.47 3.01 4.40 2.69 0.22 2.10 0.81

59.77 0.54 17.74 5.04 0.06 4.52 2.31 4.64 2.51 0.12 2.11 0.83

63.53 0.71 16.50 6.72 0.10 3.39 1.03 1.40 4.07 0.17 3.24 −

63.90 0.59 16.66 6.20 0.13 2.78 0.75 2.85 3.48 0.13 3.15 −

Total

100.34

3.21

100.15

2.91

2.94

3.24

3.15

rock in which it occurs (e.g., Philpotts 1964; Sibson 1975; Lin 1991, 1994a; Maddock 1992; Techmer et al. 1992; Wenk et al. 2000). Table 8.2. Chemical composition of Quebec (Canada) pseudotachylytes (P1–P4) and their host rocks (H1–H2) Location 1

Location 2

Location 3

wt%

H1

P1

H2

P3

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 05 CO2 H2 O+ H2 O−

49.45 2.95 14.28 5.14 7.71 0.17 5.39 8.62 3.57 1.52 0.64 n.d. 0.32 0.04

51.58 2.85 15.98 5.03 6.83 0.12 4.36 6.20 1.10 1.91 0.91 0.53 1.92 0.18

46.10 3.75 15.36 7.28 8.00 0.18 2.69 7.65 3.28 1.95 0.54 1.77 1.38 0.00

50.22 2.60 14.96 6.42 5.89 0.11 2.33 5.77 4.34 2.69 0.56 3.42 0.99 0.06

51.02 2.41 14.62 5.82 5.34 0.13 2.17 5.62 3.88 2.66 0.84 4.37 1.53 0.06

58.33 1.70 16.18 2.69 4.95 0.09 2.10 4.20 2.61 4.24 0.65 0.36 1.69 0.04

Total

99.80

99.53

99.93

100.36

100.47

99.88

(Data from Philpotts 1964)

P2

P4

162

8 Chemical Composition and Melting Processes of Pseudotachylyte

Table 8.3. XRF (samples 1 and 2) and electron microprobe analyses (samples 3 and 4) in wt% of felsic granulite and pseudotachylyte veins from the Woodroﬀe Thrust Sample

1

2

3

4

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O

73.91 0.31 13.48 1.44 0.06 0.65 1.54 3.21 4.87

72.20 0.34 12.36 3.19 0.07 1.41 2.31 2.57 4.47

63.03 0.77 18.05 5.11 0.24 2.05 2.96 4.77 2.92

62.34 0.66 18.00 5.39 0.22 2.12 3.00 5.00 2.66

Total

99.47

98.92

99.89

99.38

∗

FeO : total Fe as FeO (Data from Camacho et al. 1995) 1– Average felsic gneiss (4 samples, host rocks). 2– Pseudotachylyte vein. 3– Representative analysis of pseudotachlyte matrix free of clasts. This fraction contains crystals that appear to have crystallized from a melt. Area measured 50 μm2 . 4– As for analysis 3. Area measured 12.5 μm2 .

8.2.2 Chemical Composition of Pseudotachylyte Matrix While the bulk chemical compositions of pseudotachylyte veins are generally similar to those of their host rocks, many studies have documented a signiﬁcant

Fig. 8.1. SiO2 –Al2 O3 –All other components diagram showing the relation between the glass matrices within Fuyun pseudotachylytes and their granitic host rocks. c (After Lin 1991, 1994a). 2007, with kind permission from Elsevier Science Ltd

8.2 Bulk-Vein and Matrix Compositions

163

Fig. 8.2. Concentrations of oxides within Fuyun pseudotachylytes and granitic host rocks (After Lin 1991)

diﬀerence in composition between the ﬁne-grained matrix within pseudotachylyte veins and that of associated host rocks (e.g., Toyoshima 1990; Lin 1991, 1994a; Wenk et al. 2000). The chemical composition of the pseudotachylyte matrix generally depends upon the bulk composition of the host rock. Accordingly, in accessing the properties of melt and melting processes, the chemical composition of the matrix is generally compared with that of the host rock. As most of the pseudotachylytes reported in the literature are derived from granitic rocks, this section focuses on granite-hosted Fuyun

164

8 Chemical Composition and Melting Processes of Pseudotachylyte

Table 8.4. Chemical composition of glass matrix in glass pseudotachylyte (Type-I) vein (Gla1–Gla8) analyzed by EDX Glass matrix in glass pseudotachylyte wt%

Gla1

Gla2

Gla3

Gla4

Gla5

Gla6

Gla7

Gla8

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 05 NiO

63.90 0.59 17.44 0.07 3.35 0.14 0.85 2.81 3.04 4.13 0.59 0.11

63.50 0.64 17.26 0.00 3.43 0.10 0.92 3.01 3.02 3.83 0.38 0.01

64.94 0.60 18.03 0.02 3.40 0.13 0.86 2.93 3.16 3.99 0.55 0.00

64.41 0.58 16.92 0.00 3.49 0.09 0.68 2.68 2.97 4.07 0.44 0.02

65.73 0.59 16.41 0.01 3.31 0.08 0.75 2.48 2.84 4.30 0.62 0.06

94.41 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.09 0.00 0.73 0.04

92.68 0.07 2.24 0.03 0.81 0.00 0.03 0.70 0.36 0.00 0.16 0.00

88.52 0.30 6.13 0.00 0.63 0.08 0.08 0.31 2.00 1.52 0.11 0.00

Total

97.02

96.08

98.95

96.36

97.18

98.14

97.08

99.73

Number Si Ti Al Cr Fe Mn Mg Ca Na K P Ni

of ions on the basis of twenty-four 8.827 8.848 8.801 8.940 0.061 0.067 0.061 0.061 2.840 2.834 2.879 2.768 0.008 0.000 0.002 0.000 0.387 0.399 0.419 0.405 0.017 0.011 0.015 0.011 0.174 0.190 0.173 0.140 0.416 0.449 0.426 0.398 0.814 0.815 0.831 0.799 0.728 0.681 0.690 0.721 0.069 0.045 0.063 0.052 0.012 0.001 0.000 0.003

oxygens 9.018 11.90 0.061 0.000 2.654 0.000 0.001 0.000 0.380 0.000 0.009 0.000 0.154 0.000 0.364 0.022 0.756 0.000 0.753 0.000 0.072 0.075 0.007 0.003

11.60 0.006 0.331 0.003 0.331 0.000 0.005 0.093 0.088 0.000 0.017 0.000

11.04 0.028 0.000 0.000 0.070 0.008 0.000 0.041 0.485 0.243 0.012 0.000

Total

14.354

14.230

12.243

12.831

14.341

14.359

14.297

12.000

FeO∗ : total Fe as FeO (Data from Lin 1991, 1994a)

pseudotachylyte in considering chemical heterogeneity within the matrix of melt-origin pseudotachylyte. As stated above, the matrix of granite-hosted pseudotachylyte is SiO2 poor relative to the bulk chemical composition of the host rock. This is the case for glass matrices in Types-I (Table 8.4) and Type-V (Table 8.5) Fuyun pseudotachylyte veins (Fig. 8.3), except for local areas of matrix that are pure SiO2 . The average SiO2 component in the glass matrix is 60–64 wt%, which is 5–10 wt% less than that of the host granitic rocks (Tables 8.4 and 8.5). Similar trends have been reported by Philpotts (1964), Sibson (1975), and Camacho et al. (1995). For example, Table 8.3 shows that the SiO2 component of ﬁne-grained matrix within the melt-origin Woodroﬀe pseudotachylyte is

8.2 Bulk-Vein and Matrix Compositions

165

Table 8.5. Chemical composition of glass matrix in Type-V pseudotachylyte vein (Gla9–Gla16) analyzed by EDX Glass matrix within type-V vein wt%

Gla9

Gla10

Gla11

Gla12

Gla13

Gla14

Gla15

Gla16

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 05 NiO

61.77 0.62 18.92 0.07 4.32 0.11 1.55 4.32 2.22 3.33 0.35 0.92

61.59 0.55 18.91 0.04 4.28 0.05 1.56 4.28 2.29 3.32 0.31 0.17

61.16 0.68 18.76 0.00 3.67 0.00 1.56 3.67 2.03 3.62 0.66 0.03

60.09 0.57 18.00 0.02 4.62 0.06 2.14 4.62 1.81 3.16 0.376 0.02

60.84 0.42 19.25 0.00 5.24 0.12 1.71 5.24 1.80 2.84 0.36 0.00

64.75 0.36 13.26 0.01 7.47 0.08 1.08 2.17 2.85 3.34 0.18 0.04

57.06 0.50 15.63 0.00 8.55 0.17 2.73 2.43 3.43 4.38 0.20 0.00

59.05 0.28 9.65 0.00 16.88 0.17 2.29 2.21 1.02 2.13 0.37 0.09

Total

97.64

97.42

96.90

97.57

98.79

98.14

95.08

94.84

Number Si Ti Al Cr Fe Mn Mg Ca Na K P Ni

of ions on the basis of twenty-four 8.532 8.531 8.499 8.417 0.064 0.057 0.072 0.060 3.080 3.087 3.072 2.972 0.002 0.004 0.000 0.002 0.512 0.504 0.549 0.769 0.013 0.005 0.000 0.007 0.318 0.323 0.321 0.446 0.639 0.635 0.546 0.684 0.590 0.615 0.549 0.491 0.587 0.586 0.642 0.560 0.041 0.037 0.078 0.044 0.002 0.019 0.003 0.020

oxygens 8.514 0.044 3.010 0.000 0.587 0.014 0.357 0.786 0.488 0.508 0.044 0.000

Total

14.387

14.369

14.411

14.496

14.342

9.190 0.038 2.218 0.001 0.886 0.010 0.228 0.329 0.784 0.504 0.022 0.005

8.400 0.055 2.713 0.000 1.503 0.021 0.600 0.979 0.979 0.824 0.024 0.000

8.865 0.031 1.708 0.000 2.120 0.021 0.669 0.355 0.297 0.407 0.047 0.011

14.315

15.053

14.531

FeO∗ : total Fe as FeO (Data from Lin 1991, 1994a)

about 10 wt% less than that of the host rock, a felsic granitic gneiss (Camacho et al. 1995). The chemical composition of glass matrix represents the primary composition of the frictional melt that cooled rapidly following formation of the pseudotachylyte (Lin 1991, 1994a). The commonly observed depletion in SiO2 within the glass matrix of pseudotachylyte veins indicates that the proportion of quartz crystals that survived melting and remained as fragments was greater than that of other rock-forming minerals; consequently, the glass matrix is SiO2 -poor relative to the host rock (Lin 1991, 1994a). Locally, the matrix within Fuyun pseudotachylyte is silica-rich or even composed of pure silica glass (Table 8.4, Gla6–Gla8) that is translucent to transparent; such

166

8 Chemical Composition and Melting Processes of Pseudotachylyte

Fig. 8.3. Concentrations of oxides within the glass matrix of Type-I and Type-V pseudotachylytes and their granitic host rocks (After Lin 1991)

zones occur mainly as ring-shaped bodies around fragments or as isolated patches within glass pseudotachylyte veins (Fig. 5.5). Within the Fuyun pseudotachylytes, the Al2 O3 content of the glass matrix is generally 3–5 wt% higher than that in the host granite rock, although the matrix contains the same range of feldspars as that found in the host rock. There is a minor increasing trend in CaO and K2 O concentrations from the host rocks to pseudotachylyte veins to areas of glass matrix. Locally,

8.2 Bulk-Vein and Matrix Compositions

167

the composition of the matrix matches that of alkaline and plagioclase feldspar (Tables 8.4 and 8.5). FeO and MgO concentrations are generally higher in the glass matrix than in the host rocks (Fig. 8.3); this trend has also been reported from volcanic rock-hosted pseudotachylyte along a Tertiary normal fault in East Greenland (Curewite and Karson 1999). As stated in Chap. 7, the fragments that survived melting to remain as relics within glass pseudotachylyte veins are almost entirely made up of quartz, even though feldspar and biotite are prominent rock-forming minerals within the host granitic rocks. The fact that the glass matrix is depleted in SiO2 and enriched in Al2 O3 , CaO, K2 O, FeO, and MgO indicates that the melt formed by the preferential melting of maﬁc minerals such as biotite and feldspar rather than the complete melting of the host rock (Lin 1994a). Flow layers and streaks observed within the glass matrix, as deﬁned by colored layers, mainly reﬂect variations in the chemical composition (i.e. FeO and MgO) of maﬁc components. The concentrations of FeO and MgO within the veins increase progressively from 0.5 to 15 wt% from transparent to translucent zones to opaque layers and streaks. With the exception of glass-type pseudotachylyte, most pseudotachylytes are mainly composed of ﬁne-grained crystalline matrix and microlites. In biotite-rich Type-IV veins, the ﬁne-grained matrix proximal to microlites is similar in composition to biotite (Table 8.6, Column M7–M8). Likewise, the ﬁne-grained matrix that surrounds plagioclase microlites is similar in composition to the plagioclase itself (Table 8.6, Columns M1–M5). This indicates that the constituent elements of biotite and feldspar diﬀused through the matrix toward sites at which biotite and feldspar microlites grew. High concentrations of mineral-forming elements are also observed in zonal structures developed around fragments (Fig. 8.4). Spherical spherulites that surround rounded quartz and feldspar fragments consist of three main zones. Zones a and c consist of ﬁne-grained crystals with plagioclase-type compositions (An < 35). Zone b consists of crystal ﬁbers that are too small (< 3 μm) to be accurately identiﬁed under the microscope, but they are inferred to be biotite on the basis of chemical composition and SEM-BSE images (Lin 1994b). Both Zone a, which is in contact with the fragment, and Zone c are similar in composition to plagioclase (Table 8.6, column M6), while Zone b has a composition similar to that of biotite (Table 8.6, column M8). Similar zoned structures, termed corona, are commonly developed around quartz xenocrysts in basalt or andesite (Sato 1975). On the basis of a diﬀusion model, Sato (1975) explained the formation of such coronas by the preferred crystallization of certain mineral species from a melt. It is therefore possible that the coronas observed within glassy or ﬁne-grained matrix in pseudotachylyte formed via a similar mechanism.

168

8 Chemical Composition and Melting Processes of Pseudotachylyte

Table 8.6. Chemical composition of ﬁne-grained matrix in Type-IV pseudotachylyte vein (M1–M8) analyzed by EDX Fine-grained matrix within type-IV vein wt%

M1

M2

M3

M4

M5

M6

M7

M8

SiO2 TiO2 Al2 O3 Cr2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 05 NiO

63.58 0.73 21.38 0.77 0.05 0.00 2.48 9.24 0.12 0.00 0.34 0.07

63.38 0.80 21.02 1.17 0.00 0.00 2.34 9.32 0.00 0.00 0.50 0.00

64.18 0.87 21.54 0.67 0.00 0.01 2.47 3.54 0.07 0.00 0.52 0.13

62.61 1.05 21.73 0.85 0.12 0.00 2.56 9.17 0.24 0.04 0.05 0.05

63.81 0.81 21.04 0.62 0.00 0.00 2.31 9.34 0.09 0.00 0.44 0.00

63.70 0.79 21.11 0.90 0.00 0.18 2.46 8.97 0.25 0.05 0.60 0.00

39.73 1.69 15.34 15.55 0.19 12.24 1.20 0.07 4.46 0.00 0.12 0.12

41.81 1.08 14.48 15.78 0.26 12.44 3.74 1.02 3.27 0.13 0.29 0.14

Total

98.76

98.53

100.00

98.84

98.46

99.01

91.34

94.80

Number Si Ti Al Cr Fe Mn Mg Ca Na K P Ni

of ions on the basis of twenty-four 8.514 8.508 8.487 8.406 0.075 0.081 0.087 0.105 3.375 3.327 3.357 3.438 0.087 0.132 0.075 0.096 0.006 0.000 0.000 0.012 0.000 0.000 0.003 0.000 0.357 0.336 0.348 0.369 2.400 2.427 2.445 2.367 0.021 0.000 0.012 0.039 0.000 0.000 0.000 0.006 0.039 0.057 0.057 0.048 0.009 0.000 0.012 0.006

oxygens 8.547 0.081 3.321 0.069 0.000 0.000 0.330 2.439 0.015 0.000 0.051 0.000

Total

14.888

14.853

14.868

14.883

14.892

8.499 0.078 3.321 0.099 0.000 0.036 0.351 2.322 0.042 0.006 0.001 0.000

6.545 0.209 2.979 2.143 0.025 0.004 0.212 0.224 0.938 0.000 0.016 0.015

6.628 0.128 2.772 2.093 0.035 2.940 0.636 0.313 0.662 0.016 0.038 0.018

14.835

16.310

16.279

∗

FeO : total Fe as FeO (Data from Lin 1991, 1994a)

8.2.3 Water Contents of Pseudotachylyte Veins Analytic results reveal that pseudotachylyte veins contain more water than their host rocks (e.g., Toyoshima 1990; Lin 1991, 1994a; Techmar et al. 1992), although some studies report similar water contents for veins and their hosts (Camacho et al. 1995). The water contents of Fuyun pseudotachylyte veins and host rocks are shown in Table 8.1 and Fig. 8.5. Here, crystalline water (H2 O+ ) and fractured water (H2 0− ) are measured independently on the basis of ignition loss. The crystalline water (H2 O+ ) contents in Type-I to -V veins show minor variations between 2.0 and 2.65 wt%. These values are 1.5–2.0 wt%

8.3 Discussion

169

Fig. 8.4. SEM-BSE image of zoning structures within a circular spherulite developed within pseudotachylyte from the Fuyun Fault, China (After Lin 1991, 1994b)

higher than that of the host granitic rock, which contains about 0.5 wt% (H2 O+ ) (Fig. 8.5). As stated above, the glass matrix that contains vesicles and amygdules is not metamorphosed; therefore, it can be inferred that the crystalline water contents of Type-I and Type-V Fuyun pseudotachylyte veins represent the initial water contents that entered the melt as hydroxyl ions associated with quenching of the melt. These high crystalline-water contents, in combination with the presence of vesicles and amygdules, indicate that melting occurred under water-saturated conditions. In a study of the Ivrea Verbano Fault Zone in the Italian Alps, Techmar et al. (1992) reported water contents in a pseudotachylyte vein that were 2–3 wt% higher than those in the host rock. These studies demonstrate that some of the water within the veins was previously housed in fractures within the host rock along the fault zone.

8.3 Discussion 8.3.1 Melting Processes As stated above, given that the bulk chemical composition of pseudotachylyte veins is commonly similar to that of the host rocks, it was inferred that

170

8 Chemical Composition and Melting Processes of Pseudotachylyte

Fig. 8.5. Water contents of diﬀerent pseudotachylytes from the Fuyun Fault, China, and their granitic host rocks (Data from Lin 1991, 1994a)

frictional fusion involves complete rather than selective melting of the host rock (e.g., Philpotts 1964; Ermanovics et al. 1972; Masch et al. 1985). Comparable chemical compositions have also been found for molten material and its amphibole host rock during a frictional melting experiment; this ﬁnding was also interpreted to indicate the complete melting of the host rock (Spray 1988). As it is impossible to separate all of the fragments of host rock from a ﬁnegrained matrix formed from a primary melt, it is unsurprising that the bulk composition of pseudotachylyte is similar to that of the host rock. Accordingly, this similarity in bulk composition does not provide conclusive proof of the complete melting of the host rock; in fact, it is commonly reported that the chemical composition of pseudotachylyte matrix hosted within granitic rock is heterogeneous and has a lower SiO2 content than that of the host rock (e.g., Shand 1916; Sibson 1975; Ermanovics et al. 1972; Lin 1991, 1994a; Camacho et al. 1995). In their study of the Langtang pseudotachylyte, Scott and Drever (1953) were the ﬁrst to report that biotite crystals within host granitic rocks are absent from molten material derived from the host granitic rock. The authors explained this in terms of the selective melting of biotite, which yielded the ferrous iron component of the frictional melt. This suggestion is also supported by Bossi`ere (1991) in a study of pseudotachylyte along the Alpine Fault, New Zealand.

8.3 Discussion

171

On the basis of microstructural observations and the composition of the glassy matrix, previous studies explain the compositional heterogeneity of pseudotachylyte in terms of the melting process. Wallace (1976) suggested that melt composition is mainly controlled by the preferential melting of those minerals with melting points below the minimum temperature achieved by frictional heating during faulting. On the basis of fusion microstructures and the chemical compositions of fused micas from the Harry Creek pseudotachylytes of Central Australia, Allen (1979) concluded that the critical factors controlling pseudotachylyte composition are the composition of the host rock and the cooling rate of the frictional melt. The glassy pseudotachylytes found along the Fuyun Fault, China, contain 5–10 wt% less SiO2 than the granitic host rock (Lin 1991, 1994a). As stated in Chap. 5, powder X-ray diﬀraction data and petrologic features indicate that 11 wt% of quartz fragments survived melting along the Fuyun Fault, whereas other rock-forming minerals within the granitic host rock such as mica and feldspar are absent from the glass pseudotachylyte. This observation may explain why the matrix has a lower SiO2 content than the host rock. The high Al2 O3 , CaO, K2 O, FeO, and MgO contents can be explained in terms of the depletion in SiO2 , as this would lead to an increase in the ratios of the other components to SiO2 . The above ﬁndings indicate that minerals with lower melting temperatures such as mica and albite were melted entirely during pseudotachylyte formation, whereas those with high melting temperatures such as quartz survived as fragments (Lin 1991, 1994a). Accordingly, it is concluded that the melt-origin pseudotachylytes formed mainly by the preferential melting of minerals with low melting temperatures rather than the complete melting or partial melting of the host rocks (Lin 1991, 1994a). Such a chemical melting process has also been demonstrated in high-speed frictional melting experiments (see Chap. 12 for details). 8.3.2 Melt Temperature The temperature of frictional melt during pseudotachylyte formation is generally estimated using mineral geothermometers applied to microlites (e.g., Maddock 1983; Toyoshima 1990; Lin 1991, 1994b) or the chemical composition of the matrix (e.g., Wallace 1976; Lin 1991, 1994a). Previously estimated temperatures of frictional melts range from 750 to >1400◦C (Table 8.7). The presence of a pure silicate glass matrix in the Fuyun pseudotachylyte indicates that quartz crystals were melted without any chemical reaction or contamination by other minerals. The occurrence of numerous vesicles and amygdules within melt-origin pseudotachylyte indicates that melting occurred in an excess ﬂuid phase. Melting experiments for quartz reveal that its melting point at atmospheric pressure is 1723◦C; the equivalent ﬁgure at a pressure of >200 MPa in the system SiO2 –H2 O is ∼1100◦C (Fig. 8.6; Kennedy et al. 1962).

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8 Chemical Composition and Melting Processes of Pseudotachylyte

Table 8.7. Previous estimates of the depth and temperature of formation of meltorigin pseudotachylyte Location

Formation depth (km)

Temperature (◦ C)

Reference

Outer Hebrides Thrust Woodroﬀe Thrust, Australia Alpine Fault, New Zealand Central Otago, New Zealand Ikecaqu, Greenland Hiddaka Metamorphic Zone, Japan Fuyun Fault Zone, China (Modiﬁed from Lin and

4–5 <5

1100

2.2 2–7 2–6

750 900–1100

Sibson (1975) Allen (1979) Camacho et al. (1995) Seward (1985) Wallace (1976) Barker (2005)

1.6 4

1100

Maddock et al. (1987) Toyoshima (1990)

1.5

>1400

Lin (1991, 1994a, b)

1200

Shimamoto 1998)

Based on the above results (Fig. 8.6), a minimum temperature of 1450◦C is estimated for melt generated by frictional heating during the formation of the Fuyun pseudotachylyte. This estimate is based on the presence of pure SiO2 glass at a depth of 1.5 km (corresponding to a pressure of ∼400 bars), as described in Chap. 5. Melt that forms from frictional heating has a transient existence, with the cooling half-life of an injection vein with a width of 1 cm being in the order of 40 s (Sibson 1975). Considerable overstepping of the melting temperature is therefore necessary to melt quartz grains, meaning that the temperature of 1450◦ C represents a minimum estimate. The melting temperature can also be directly measured using a thermocouple and radiation thermometer set within a simulated fault shear zone during high-velocity frictional experiments (Tsutsumi 1994; Tsutsumi and Shimamoto 1997a, b). The results obtained in this way reveal that the average temperature of frictional melting of a gabbro sample is 1100–1150◦C; this is the highest temperature that can be measured by the thermometer used in the experiment. Accordingly, the increase in temperature that occurs during frictional melting within a fault zone is likely to be higher than this ﬁgure. This experimental result is consistent with estimates based on natural pseudotachylytes, as listed in Table 8.7. Previous estimates of melt temperatures within pseudotachylytes are based on the assumption that the frictional melts were generated under conditions of chemical equilibrium; however, high-velocity frictional melting experiments reveal that this is not the case (Lin 1991; Lin 1994; Lin and Shimamoto 1998; see Chap. 12 for details). This means that frictional melting only occurs at the melting temperature higher than that of each individual rock-forming mineral.

8.3 Discussion

173

Fig. 8.6. Melting pressure of SiO2 in equilibrium with H2 O, or projection of the univariant equilibrium curve for a SiO2 –H2 O system on a P–T plane (modiﬁed from Kennedy et al., 1962). Solid circle indicates the inferred formation P-T conditions of the Fuyun pseudotachylyte (P = ∼400 bars, T = ∼ 1450◦ C) (Data from Lin 1991, 1994a)

The frictional melt temperatures reported to date are therefore likely to be underestimates. 8.3.3 Role of Water During Frictional Melting During seismic faulting events, water plays an important role in the genesis of pseudotachylyte by inﬂuencing frictional resistance (e.g., Francis 1972; Ermanovics et al. 1972; Sibson 1973, 1975; Lachenbruch 1980; Lin 1991, 1994a, b; Maddock 1992). It is generally suggested that the formation of pseudotachylyte is favored in crystalline rocks that lack an intergranular ﬂuid, as ﬂuids may cause hydraulic weakening and reduced eﬀective normal stress upon a fault plane due to ﬂuid pressure (Francis 1972; Sibson 1975); however, it is also argued that the presence of ﬂuids may favor melting because of the eﬀect that ﬂuid has on lowering the melting point of minerals and generating lowviscosity melt (Ermanovics et al. 1972; Allen 1979; Lin 1991; Maddock 1992).

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8 Chemical Composition and Melting Processes of Pseudotachylyte

Fig. 8.7. Slip displacement–stress–sample shortening data obtained from highvelocity frictional experiments on coal gouges. Four stages (I–IV) can be recognized during the experiment. The stages are characterized by major stable shortening (I), minor erratic shortening (II), major erratic shortening (III), and, stable shortening (IV). A sharp reduction in shear stress occurred during Stage III, synchronous with an increase in sample shortening; this was caused by the thermal pressurization of gouge ﬂuid during at the time when a large amount of gas was emitted. The sharp reduction in both normal stress and shear stress is attributed to pore collapse causc ing gas pressurization.(After O’Hara et al. 2006) 2007, with kind permission from Elsevier Science Ltd

Indeed, experimental results reveal that artiﬁcially generated friction melts can be produced in a water-rich environment (Killick 1990; Kennedy and Spray 1992). It is possible that maﬁc minerals act as a source of water that facilitates the generation of friction melt (e.g., Bossi`ere 1991; Maddock 1992). The release of water from hydrous minerals such as biotite may also aﬀect the course of melting, as described in the following points (Maddock 1992). i) The release of water from a hydrous mineral species during dehydration acts to lower the melting points of other minerals. This in turn enhances the frictional melting of other rock-forming minerals during seismic faulting. ii) Released water acts to reduce the viscosity of the melt, thereby assisting heat transfer by convection and advection. This may also assist the ﬂow of the melt and injection into fractures within the fault zone. iii) The incorporation into the melt of network-modifying cations such as Fe2+ , Mg+ , Ca2+ , Na+ , and K+ will also act to reduce the viscosity of

8.3 Discussion

175

the melt. The occurrence of such a process is supported by the Fe- and Mg-rich composition of the glass matrix within pseudotachylyte. iv) The products of dehydration reactions are typically extremely ﬁne-grained, making them susceptible to melting. Sibson (1973, 1980b) suggested that the general scarcity of pseudotachylyte arises from the fact that most seismic faulting takes place at an approximately constant water volume around the fault plane. Frictional heating tends to expand pore ﬂuid and lead to a transient increase in pressure, which in turn leads to a sharp reduction in both the eﬀective normal stress and dynamic friction on the fault surface (Sibson 1980b; Lachenbruch 1980). The assumption regarding the change of pore ﬂuid pressure is based on the hypothesis that water is held at the fault plane and does not migrate rapidly from this site; however, experiments involving the high-velocity frictional melting of gabbro and granitic rocks reveal that melt formed by frictional heating on the sliding plane is accompanied by the formation of numerous small thermal cracks on both sides of sliding plane; frictional melt is then injected into these cracks (Lin 1991; Lin and Shimamoto 1998; also see Chap. 12). High-velocity frictional experiments on coal gouges also reveal a sharp reduction in the eﬀective normal stress, shear stress, and dynamic friction on the fault surface; this is attributed to pore collapse and associated gas pressurization in the experiment (Fig. 8.6; O’Hara et al. 2006). These experimental results suggest that the change in pore ﬂuid pressure within a fault zone, which leads to a reduction in the eﬀective normal stress and shear stress, is mainly caused by frictional heating associated with the thermal pressurization of pore ﬂuids; these conditions occur following suﬃcient slip, during the ﬁnal stages of seismic faulting. Melts that form via frictional fusion have a transient existence, with the cooling half-life of injection veins of 1 cm in width being in the order of 40 (Sibson 1975). The results of high-velocity frictional melting experiments undertaken at room temperature reveal that melts are generated within several seconds of the initiation of sliding (Lin 1991; Lin and Shimamoto 1998). Under such conditions, the presence of water is critical to the generation of melt, as it acts to reduce the temperature of fusion, lowers the viscosity of the melt, and acts as a catalyst during fusion (Allen 1979). Analyses of water content also reveal the presence of abundant water during pseudotachylyte formation (e.g., Toyoshima 1990; Lin 1991, 1994a; Magloughlin, 1992; Techmar et al., 1992). For the Fuyun glassy pseudotachylyte (Table 8.1b) and pseudotachylyte from the Italian Alps (Techmar et al. 1992), the crystal water content is approximately 2 wt% greater than that of the host rocks. The presence of vesicles and amygdules also indicates that the frictional melt was water-saturated during pseudotachylyte formation. It is impossible that all of the water within pseudotachylytes is derived from hydrous minerals within the host rock, as the total water content of the granitic rocks that host the Fuyun pseudotachylytes is <1 wt%, less than that of the pseudotachylyte

176

8 Chemical Composition and Melting Processes of Pseudotachylyte

veins (Table 8.1). The preexistence of water within the generation zones of pseudotachylyte is also conﬁrmed by Magloughlin’s (1992) study of pseudotachylyte hosted within pelitic sediments in the Nason terrain, Washington State, USA, in which the author noted the presence of sulﬁde veining, Na2 O loss, and ﬂuid inclusions, and by Okamoto et al. (2006) study of pseudotachylyte hosted in a sedimentary rocks within a subduction zone, Southwest Japan. Thus, dry conditions are not a prerequisite for the formation of fault-generated pseudotachylyte. It is also known that the presence of hydrous minerals such as biotite favors pseudotachylyte generation (e.g., Scott and Drever 1953; Allen 1979). It has been postulated that the rock types that most favor pseudotachylyte generation are quartz-rich rocks that contain signiﬁcant amounts of hydrous minerals such as biotite and hornblende; the proportion of melt phase depends largely on the percentage of hydrous minerals within the rock. Quartz-rich rocks are considered to have high shear strength. In addition, the results of high-velocity frictional melting experiments demonstrate that hydrous minerals are conducive to the development of frictional melting during seismic faulting (Spray 1988; Lin and Shimamoto 1998; O’Hara et al. 2006).

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

9.1 Introduction Most of the fault-related pseudotachylyte reported to date is cataclasiterelated (herein termed C-Pt), having formed at shallow depths in brittle-dominated seismogenic fault zones by both frictional melting and crushing mechanisms, as described in previous chapters. Pseudotachylyte has also been described in association with mylonitic rocks (herein termed M-Pt), having formed in deep-level fault shear zones within the semi-brittle to crystalplastic regimes (Sibson 1980a; Passchier 1982, 1984a, b; Hobbs et al. 1986; Koch and Masch 1992; McNulty 1995; Reynolds et al. 1998; Takagi et al. 2000; Lin et al. 2003b, 2005b). In contrast to C-Pt, which is widely reported, M-Pt has only been described from several localities worldwide. In his study of the Outer Hebrides Thrust, Scotland, Sibson (1980a) was the ﬁrst to describe the occurrence of pseudotachylyte within a ductile shear zone. Sibson proposed that the pseudotachylyte formed intermittently in association with seismic slip upon discrete planes within a shear zone that otherwise deformed via crystal-plastic mechanisms. Passchier (1982) described the alternating development of pseudotachylyte and mylonite–ultramylonite within the Saint-Barthkemy Massif, France (Fig. 9.1). Passchier suggested that this alternating deformation pattern reﬂected the passing of the shear zone across the brittle–plastic transition; however, based on observations of plastically deformed pseudotachylyte within the Redbank Shear Zone, Central Australia, Hobbs et al. (1986) proposed that ultramylonite-associated pseudotachylyte can develop entirely within the plastic-dominated regime via plastic instability. They argued that the plastic instability of rocks could lead to seismic slip and the generation of pseudotachylyte within the plastic-dominated regime. Lin et al. (2003b, 2005b) reported large volumes of coexisting C-Pt, M-Pt, and ultramylonite-associated pseudotachylyte (herein termed Um-Pt) in cataclastic and mylonitic rocks within individual shear zones along the Woodroﬀe Thrust, Central Australia, and along the Dahezhen Fault Shear Zone located

178

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.1. Photograph of M-Pt veins and associated mylonite from the SaintBarthkemy Massif, France. Large arrows indicate sense of shear. The coin shown for scale is 2.4 cm across

within an ultrahigh-pressure (UHP) complex in the Qinling-Dabie Shan collisional orogenic belt, Central China. Outcrops along the Woodroﬀe Thrust also reveal granulite-associated pseudotachylytes (herein termed G-Pt) that formed under granulite facies conditions and that are overprinted by C-Pt and M-Pt veins (Lin 2007). Although the formation mechanism of myloniteand granulite-related pseudotachylytes that formed within the crystal-plastic dominated regime remains controversial, there is a general consensus that these rocks formed by frictional fusion during episodes of seismic faulting, as is the case for cataclasite-related pseudotachylyte that forms in the brittledominated regime and that is commonly observed in fault shear zones (e.g., Sibson 1980a; Hobbs et al. 1986; McNulty 1995; Lin et al. 2005b). It is well known that C-Pt veins generally form in the upper crust at shallow depths of <10 km (e.g., Sibson 1975; Toyoshima 1990; Lin 1994a, b; Fabbri et al. 2000). In contrast, M-Pt and G-Pt veins are considered to form within deeper shear zones at greenschist facies conditions (∼300–400◦C), corresponding to depths of 10–15 km assuming an average geothermal gradient of 25◦ C/km (Sibson 1983; Scholz 2002); in addition, other studies (e.g., White 1996; Lin et al. 2003b, 2005b) document the formation of M-Pt vein under amphibolite facies temperatures or higher (>400–600◦C), corresponding to depths of >15–25 km. These two depth boundaries inferred from MPt, 10 km and 15–25 km, are coincident with the upper and lower limits of the brittle–plastic transition zone, lying outside of the ‘normal’ seismogenic zone. It has also been reported that some pseudotachylytes formed within the deep crust at granulite facies conditions of >650–700◦C and 700–800 MPa (Clarke and Norman 1993) and at eclogite facies conditions of 700–800◦C

9.2 Woodroﬀe Pseudotachylytes

179

and 1800–2000 MPa, corresponding to depths of ≥60–70 km (Austrheim and Boundy 1994; Lund and Austrheim 2003). In addition, the results of a seismic inversion study suggest that a melt layer of 31 cm in thickness was generated at a depth of 670 km by frictional melting during the 1994 Bolivian Mw 8.3 earthquake (Kanamori et al. 1998). These studies support the hypothesis that fault-related pseudotachylyte can be generated at various depths ranging from the shallow crust close to the Earth’s surface to deep shear zones in the lower crust and even within the mantle, although the formation mechanisms of deep earthquakes remain poorly understood. In this chapter, to gain an understanding of the deformation mechanisms of pseudotachylyte within the brittle–plastic transition zone, two representative examples are presented of the structural modes and petrologic features of C-Pt, M-Pt, Um-Pt, and G-Pt that coexist within individual fault shear zones that formed in both the brittle-dominated and plastic-dominated regimes. One of the examples is taken from the Woodroﬀe Thrust in the Musgrave Mountains, Central Australia (Camacho et al. 1995; Lin et al. 2005b), while the other is taken from the Dahezhen Shear Zone, located within an ultrahighpressure (UHP) complex in the Qinling-Dabie Shan collisional orogenic belt, Central China (Lin et al. 2002b, 2003b). The structural and textural features of these pseudotachylytes and their host rocks are interpreted in terms of the tectonic history of the fault shear zones, pseudotachylyte formation, earthquake generation, and rupture propagation within fault shear zones that span the brittle–plastic transition in the upper–lower crust.

9.2 Woodroﬀe Pseudotachylytes 9.2.1 Tectonic Setting of the Woodroﬀe Thrust The Woodroﬀe Thrust is an E–W trending structure located within MesoNeoproterozoic metamorphic rocks of the northern part of the Musgrave Block, Central Australia (Fig. 9.2). The Musgrave Block is dominated by the geologically distinct Fregon and Mulga Park subdomains, which are separated by the Woodroﬀe Thrust (Edgoose et al. 1993; Camacho et al. 1995). The Fregon subdomain is composed of granulite facies gneiss derived from interlayered felsic volcanic rocks, while the Mulga Park subdomain consists of amphibolite facies gneiss and porphyritic granite (Gray 1978; Edgoose et al. 1993). The Woodroﬀe Thrust consists of numerous mylonitic shear zones that anastomose around less-deformed or undeformed bodies of gneiss and granite (Bell 1978; Camacho et al. 1995). The main (widest) mylonitic shear zone occurs at the bottom of the sequence of highly deformed rocks (Fig. 9.2) and contains a NE–SW to E–W striking mylonitic foliation and a prominent mineral lineation, plunging 20–30◦ to the south, that is deﬁned by the preferred alignment of clasts. The mylonites generally occur in the amphibolite facies rocks of the Mulga Park subdomain (Camacho et al. 1995) and to a lesser

180

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.2. (a) Geologic map of the central Musgrave Block, Australia. (b) and (c) show detailed geologic maps of the areas indicated in (a). The cross-sections I, II, and III, whose locations are indicated in (b) and (c), are shown in Figs. 9.3, 9.4, and 9.5, respectively. (Modiﬁed from Camacho et al. 2005; Shimamoto and Arai 1997; c Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

9.2 Woodroﬀe Pseudotachylytes

181

degree in the granulite facies rocks of the Fregon subdomain. Ultramylonite is more commonly observed towards the top of the main mylonitic shear zone and is generally interlayered with mylonite and deformed pseudotachylytegeneration zones (Lin et al. 2005b) over a thickness of 300 m through the mylonitic shear zone. The cataclastic rocks occur mainly in the upper part of the shear zone, above the mylonite–ultramylonite zone. Large volumes of pseudotachylyte occur along the Woodroﬀe Thrust (Bell 1978; Camacho et al. 1995), mostly within the upper cataclasite and dispersed throughout the mylonite– ultramylonite zone. 9.2.2 Field Occurrences of the Woodroﬀe Pseudotachylytes The distribution of pseudotachylyte along the Woodroﬀe Thrust and its relationship to fault shear zones are illustrated in three geological sections (Figs. 9.2–9.5; Lin et al. 2005b). Section I is located within the upper part of the mylonite–ultramylonite shear zone, close to the contact with overlying less-deformed granulite rocks of the Fregon subdomain; this section contains C-Pt veins (Fig. 9.3). Section II is located in the eastern Musgrave Ranges near the road between Kulgera and Mulga Park; it contains co-existing C-Pt, M-Pt, Um-Pt, and G-Pt veins (Fig. 9.4). Section III is located on the northwestern margin of Mt. Fraser, close to Mulga Park, and reveals the relationship between the mylonite and M-Pt veins (Fig. 9.5). In all three sections, pseudotachylyte veins are distributed throughout a zone of >300 m thickness, with the base of the shear zone being unexposed (Figs. 9.3–9.5). The thickness of the pseudotachylyte-bearing fault zone observed in these three sections is estimated to be >1.5 km, while the total thickness of the zone is 3–4 km, as shown in the section across the entire thrust zone shown in Fig. 9.2b. The pseudotachylyte-bearing fault zone can be recognized in a Landsat image (Fig. 9.6), in which pseudotachylyte-bearing areas (blue colors) are bounded by granulite (dark-gray). The image reveals a distinct north–south lineation within the granulite, and the width of the pseudotachylyte-bearing zone is 3–4 km, consistent with that inferred from the geological section shown in Fig. 9.2b. The textural and structural relationships between pseudotachylyte veins and wall rocks along the Woodroﬀe Thrust, in addition to vein morphology and host-rock lithology, indicate multiple stages of pseudotachylyte veins, including C-Pt, M-Pt, Um-Pt, and G-Pt veins that formed at diﬀerent times and depths. C-Pt veins are closely associated with cataclasite, which is mainly found in the granulite facies rocks in the upper part of the mylonite– ultramylonite shear zone. C-Pt veins generally cut across zones of mylonite and ultramylonite and overprint M-Pt and Um-Pt veins (Figs. 9.3–9.5), which occur as broad networks of millimeter- to centimeter-scale veins, as with the melt-origin Pt veins described in Chap. 4 (Figs. 4.14 and 4.17).

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9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.3. Geological cross-section I. (a) Sketch showing the relation between the fault plane and the pseudotachylyte-generation zone (b; also see Figs. 4.16 and 4.20) along the Woodroﬀe Thrust, Central Australia. The stereoplot (equal-area lower-hemisphere projection) shows the orientations of fault planes along which pseudotachylyte-generation zones were observed (see Fig. 9.2 for the location of this c site). (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

It is generally diﬃcult to recognize individual pseudotachylyte-generation planes within C-Pt zones, but diﬀuse zones of pseudotachylyte generation are recognized in terms of the sharp boundaries between the veins and host rocks; these zones vary in thickness from ∼2 mm to 10 m (Figs. 9.3–9.5 and 4.17). The generation zones generally contain 10–50 Vol% pseudotachylyte veins, but locally this may reach as high as 70–90 Vol% (Figs. 9.7 and 4.17). Overprinting relationships and the incorporation of older C-Pt fragments within younger overprinting C-Pt veins indicate the occurrence of repeated episodes of C-Pt generation within individual fault zones (Figs. 4.8 and 9.7). The zones that contain M-Pt and Um-Pt are generally located in the middle–upper part of the mylonitic shear zone, bounded above by C-Pt zones and below by C-Pt-bearing mylonite zones (Fig. 9.4). The M-Pt and Um-Pt zones are generally interlayered with mylonite and ultramylonite shear zones and are overprinted by C-Pt injection veins (Fig. 9.8).

9.2 Woodroﬀe Pseudotachylytes

183

Fig. 9.4. Photograph (a) and sketch (b) of cross-section II along the Woodroﬀe Thrust, Central Australia. G-Pt veins co-exist with M-Pt and C-Pt in the shear zone(see Fig. 9.2 for the location of this site) (After Lin et al. 2005b)

It is generally diﬃcult to distinguish mylonite–ultramylonite layers from deformed pseudotachylyte veins (M-Pt and Um-Pt) in the ﬁeld due to strong weathering and the obscuring eﬀects of overprinting deformation; however, a distinction can be made in some fresh outcrops based on the following characteristic features of pseudotachylyte veins: i) injection geometries and cross-cutting relationships; ii) elongation and alignment of fragments within the veins; iii) a sharp boundary between veins and wall/host rocks, and iv) the dark and aphanitic appearance of the veins. The M-Pt veins formed within the mylonite and ultramylonite zones and were subsequently overprinted by ongoing mylonitization (Fig. 9.9). Similar overprinting relationships between Um-Pt veins and mylonite and ultramylonite zones are also observed in the ﬁeld and in hand specimen (Fig. 9.11). Overprinting crystal-plastic deformation probably destroyed most of the evidence of any coseismic brittle deformation and pseudotachylyte formation that occurred in the crystal-plastic dominated region; this may explain the scarcity of mylonite- and ultramylonite-related pseudotachylyte within

184

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.5. Photograph (a) and sketch (b) of cross-section III. (After Lin et al. 2005b). The stereoplot (equal-area lower-hemisphere projection) shows the orientations of the mylonitic foliation (see Fig. 9.2 for the location of this site). (After Lin c et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

fault shear zones worldwide (e.g., Sibson 1980a; Koch and Masch 1992; Lin et al. 2003b, 2005b). The M-Pt and Um-Pt veins contain distinct evidence of ductile deformation, including ﬂattened and aligned fragments of host rocks that were re-oriented parallel to the foliation within the mylonite and ultramylonite, as evidenced from the continuity of the foliation between the host rock and vein fragments (Figs. 9.9 and 9.10). These M-Pt and Um-Pt veins generally cut across the mylonitic foliation, and can locally be traced back to parent veins oriented parallel to the mylonitic foliation. These overprinting structural relationships indicate that repeated pseudotachylyte-generating events occurred within the crystal plastic-dominated shear zone and that the pseudotachylyte veins themselves were mylonitized during ongoing plastic deformation. A similar pattern of repeated pseudotachylyte generation has also been described from the Dahezhen Shear Zone, Central China (Lin et al. 2003b; see below in this chapter). Numerous fragments of host rocks within M-Pt

9.2 Woodroﬀe Pseudotachylytes

185

Fig. 9.6. Landsat image of the Musgrave Mountains region, Central Australia. The white dotted line indicates the boundary between pseudotachylyte-bearing rocks (blue color) and granulite facies gneiss

Fig. 9.7. Photograph of networks of cataclasite-related pseudotachylyte veins (CPt) along the Woodroﬀe Thrust, Central Australia. Note that the old pseudotachylyte veins (Old-Pt) are cut by the younger overprinting veins (Y-Pt). The chisel c shown for scale is 16 cm. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

186

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.8. Photograph of multiple generations of C-Pt (C-Pt1 to C-Pt3) and M-Pt veins along the Woodroﬀe Thrust, Central Australia. The C-Pt veins cut the M-Pt veins. The fragments within the M-Pt veins are crystal-plastically deformed and oriented parallel to the foliation within the surrounding mylonite. The pen shown c for scale is 15 cm long. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

and Um-Pt veins are visible within polished sections cut through hand samples. The fragments generally vary in diameter from 0.1 to 10 mm and show evidence of ﬂow banding and crystal-plastic deformation (Figs. 9.9 and 9.10). Within the rocks observed along the Woodroﬀe Thrust, G-Pt veins occur within granulite facies rocks that are cut by a M-Pt- and Um-Pt-bearing mylonite shear zone and C-Pt veins (Figs. 9.4 and 9.11). It is generally difﬁcult to recognize G-Pt veins in the ﬁeld due to overprinting brittle–plastic deformation and strong weathering of the outcrops. In some fresh outcrops and polished hand samples it is possible to distinguish G-Pt veins from M-Pt and C-Pt veins on the basis of structural and textural features (Fig. 9.11). The boundary between the M-Pt-mylonite shear zone and the host granulite is generally sharp, and is easily recognized by the diﬀerence between the planar foliation within the mylonite and M-Pt and the complex swirling foliation within the granulite facies rocks (Fig. 9.11). Individual G-Pt veins vary in width from several millimeters to several centimeters. The G-Pt veins generally cut the non-planar foliation of the host granulite facies rocks at oblique angles, and most of these veins are not mylonitized and contain similar ﬂexural fabrics to those developed in the host granulite (Fig. 9.11). This indicates that the G-Pt veins formed contemporaneously with metamorphism of the host rocks, under granulite facies conditions, and escaped overprinting mylonitization within the Woodroﬀe Thrust.

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187

Fig. 9.9. Polished hand sample showing foliation (indicated by arrows) developed within M-Pt veins and host mylonite from the Woodroﬀe Thrust, Central Australia. Note that the M-Pt veins cut across the mylonite but contain a foliation that formed during ongoing crystal-plastic deformation. The fragments within the M-Pt veins are plastically deformed and oriented parallel to the foliation within the c surrounding mylonite. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

9.2.3 Microstructures C-Pt Multiple generations of C-Pt veins are commonly observed in thin sections cut from samples from the Woodroﬀe Thrust (Figs. 9.12 and 9.13). The CPt veins are similar to the melt-origin Fuyun pseudotachylytes described in previous chapters: they contain numerous microlites and fragments of quartz, feldspar, and garnet in a ﬁne-grained matrix, with the fragments possessing various shapes ranging from rounded to irregularly embayed (Figs. 9.14 and 9.15). The contacts between C-Pt injection veins and host granitic rocks are generally sharp and locally irregular (Fig. 9.14). The matrix of the veins consists of ﬁne-grained fragments, cryptocrystalline material, and numerous microlites of plagioclase (An30−40 ) and clinopyroxene as well some other minerals (Camacho et al. 1995, Lin et al. 2005b). Microlite morphology varies from simple acicular and lath-like shapes to complex spherulitic and dendritic forms that are mainly composed of plagioclase (Figs. 9.12 and 9.16). These morphologic and compositional features of the spherulites within C-Pt veins are indicative of a formation involving rapid cooling or quenching of a melt, as with the Fuyun pseudotachylyte described

188

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.10. Photographs showing the ﬁeld occurrence (a) of Um-Pt veins from the Woodroﬀe Thrust, Australia, and a polished section cut through a hand sample (b). Note that the Um-Pt veins cut across the foliation (indicated by the long arrows) within the ultramylonite (a). The fragments within the Um-Pt veins are crystal-plastically deformed and oriented parallel to the external foliation. The hand sample shows crystalplastic deformation of the Um-Pt vein. The hammer shown for scale is 35 cm long. c (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

in Chap. 6. The matrix of the veins is typically opaque in thin sections prepared with the conventional thickness of 0.03 mm, but is locally variable in color from colorless to gray, brown, and yellowish-brown. Flow structures within C-Pt veins are deﬁned by thin colored layers and streaks that anastomose around fragments within the matrix (Figs. 9.12 and 9.14). The C-Pt veins are observed to cut M-Pt veins (Fig. 9.13), suggesting that pseudotachylyte-generating events occurred repeatedly over a period of time as the rock mass was exhumed within the fault shear zone from the deeper plastic-dominated regime up through the shallow brittle regime.

9.2 Woodroﬀe Pseudotachylytes

189

Fig. 9.11. Photographs of G-Pt and M-Pt veins within host mylonite from the Woodroﬀe Thrust, Central Australia (a), and a polished section cut through a hand sample (b). Note that the G-Pt veins are cut by the mylonite zone. The coin shown in (a) for scale is 2.4 cm across (After Lin 2007)

M-Pt and Um-Pt Both M-Pt and Um-Pt veins occur parallel to the mylonitic foliation and as cross-cutting injection veins (Fig. 9.17). Where parallel to the foliation, the two types of veins are commonly interlayered (Fig. 9.18a). The injection veins that cut across the mylonitic foliation have undergone crystal-plastic

190

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.12. Photomicrograph showing diﬀerent generations of C-Pt veins from the Woodroﬀe Thrust, Central Australia. The younger vein (Y-Pt) contains numerous spherulitic microlites (Mi) and the overprinting vein contains internal ﬂow strucc tures. Plane polarized light. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

deformation and exhibit a foliation oriented subparallel to mylonite– ultramylonite bands, but they can still be traced back to mylonite-parallel M-Pt veins (Fig. 9.18b). The boundary between the M-Pt veins and the host mylonitic rocks is generally irregular or curved when viewed in the ﬁeld, but sharp when viewed under the microscope (Figs. 9.17–9.19). Fragments of quartz and feldspar within the M-Pt and Um-Pt veins are generally ﬂattened and aligned parallel to the external mylonitic foliation (Figs. 9.17–9.19), indicating crystal-plastic deformation of the fragments subsequent to pseudotachylyte formation. The shapes of the fragments are asymmetrical, indicating a shear sense that is consistent with that of surrounding mylonite and ultramylonite zones (Fig. 9.20). The layered ﬁne-grained matrices of the M-Pt and Um-Pt veins indicate extreme shear ﬂow (Figs. 9.18–9.21). Microlites that range in shape from acicular and lath-like to complex spherulitic shapes are also crystal-plastically deformed, being ﬂattened and re-oriented subparallel to the mylonitic foliation (Figs. 9.20 and 9.21). The M-Pt and Um-Pt veins and mylonitic host rocks are ﬁnely interlayered (Figs. 9.17 and 9.18). The veins cut obliquely across the mylonite foliation and are in turn plastically deformed by ongoing mylonitization, as described above. It is important to emphasize that pseudotachylyte veins that cut across the mylonitic foliation are themselves mylonitized and cut by C-Pt veins. This

9.2 Woodroﬀe Pseudotachylytes

191

Fig. 9.13. Photomicrograph showing the textural relationship between M-Pt and C-Pt veins from the Woodroﬀe Thrust, Central Australia. Note that the M-Pt veins are cut by the C-Pt veins. The fragments within the M-Pt veins are ﬂattened and oriented parallel to the foliation within the mylonitic host rock. Dotted lines indicate the microfaults along which distict oﬀsets are observed. Plane polarized light. (After c Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

Fig. 9.14. Photomicrograph showing ﬂow textures and rounded fragments within the ﬁne-grained matrix of a C-Pt vein from the Woodroﬀe Thrust, Central Australia. Dotted line indicates the boundary between the pseudotachylyte veins and the wall granitic rock. Plane polarized light

192

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.15. Photomicrograph of highly irregular embayed and rounded fragments (white color) within a C-Pt vein from the Woodroﬀe Thrust, Central Australia. Plane polarized light

indicates that the M-Pt veins formed during episodes of seismic faulting and were subsequently subjected to crystal-plastic deformation during aseismic periods. The overprinting deformation occurred within the plastic-dominated region in which the mylonite–ultramylonite formed, whereas the C-Pt veins overprinted the M-Pt veins following uplift and exhumation of the shear zone into the brittle part of the upper crust. G-Pt Textural relationships between C-Pt, M-Pt, and G-Pt veins are shown in Figs. 9.11 and 9.22–9.24. G-Pt veins occur as both simple and complex networks within the granulite host rocks. The veins commonly terminate abruptly against fractures (Figs. 9.22 and 9.23), as observed in the ﬁeld for C-Pt veins. The G-Pt veins generally occur parallel to the foliation within the host granulite and are cut by M-Pt veins (Fig. 9.24). Some ﬁne G-Pt veins are observed in fractures within rock fragments; the veins record minor oﬀset (Fig. 9.23). The foliation within the mylonite and M-Pt veins is oblique (by 10–90◦) to the ﬂexural foliation within the host granulite gneiss and G-Pt veins. This angular relationship and foliation textures indicate a sinistral shear sense (Fig. 9.24b). In contrast to the foliation developed within M-Pt and host mylonite, which shows a uniform shear sense (Figs. 9.17 and 9.18), the foliation

9.2 Woodroﬀe Pseudotachylytes

193

Fig. 9.16. Photomicrographs of spherulitic microlites within a C-Pt vein from the Woodroﬀe Thrust, Central Australia. (b) Close-up photomocrograph of spherulitic textures. Most of the light-colored microlites are made up of plagioclase. Plane polarized light

developed within G-Pt veins and host granulite gneiss is folded into similartype folds (Fig. 9.24a). It is important to note that the G-Pt veins are cut by the mylonite and M-Pt veins but are not crystal-plastically deformed in the way that M-Pt veins are. This suggests that the G-Pt veins formed during episodes of seismic faulting within the plastic ﬂow regime, at deeper levels than that at which the mylonite and M-Pt veins formed. The G-Pt veins were subsequently overprinted by M-Pt veins during alternating aseismic and

194

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.17. Polished section cut through a hand sample collected from the Woodroﬀe Thrust, Central Australia (a) and photomicrographs (b and c) showing the microstructure of the M-Pt veins. (a) The M-Pt veins (dark veinlets) cut obliquely across the mylonitic foliation. (b and c) Fragments within the M-Pt are plastically deformed, ﬂattened, and oriented parallel to the external mylonitic foliation. Note that the mylonitte and M-Pt foliations cut obliquely the boundary between the M-Pt and wall mylonite. Qz: quartz; Pl: plagioclase; B: boundary between the M-Pt vein and host mylonite. (b): Plane polarized light; (c): crossed polarc ized light. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

seismic periods as the rock mass passed across the brittle–plastic transition; this is the period during which the mylonite–ultramylonite and M-Pt formed. The G-Pt veins generally contain fragments of quartz, feldspar, garnet, and pyroxene of various sizes. The fragments exhibit a variety of shapes, including angular, rounded, embayed, and other irregular forms (Fig. 9.25). The fragments range in size from several nanometers to several millimeters when observed under the microscope and via EPMA; in the ﬁeld, they are up to >10 cm across. In contrast to the crystal-plastically deformed fragments found within M-Pt veins and host mylonites in individual outcrops (Fig. 9.19), most of the fragments within G-Pt veins have irregular shapes and show evidence of brittle deformation (Fig. 9.25). The fact that these fragments are

9.2 Woodroﬀe Pseudotachylytes

195

Fig. 9.18. Photomicrographs showing microstructures within Um-Pt (a) and M-Pt (b) veins from the Woodroﬀe Thrust, Central Australia. (a) The branched Um-Pt veins (dark brown areas) developed along the foliation have sharp boundaries with the host ultramylonite. Arrows indicate sense of shear. (b) M-Pt injection veins cut obliquely across the mylonitic foliation. Fragments within the Um-Pt and M-Pt veins are strongly ﬂattened and oriented subparallel to the external mylonitic foliation. c (a): cross polarized light, (b): plane polarized light. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

rounded and irregularly embayed indicates a melting origin for the pseudotachylyte (Lin 1999b). The matrix within G-Pt veins is generally composed of very ﬁne-grained material and microlites, is typically opaque in thin sections of normal thickness (0.03 mm), and is locally variable in color from colorless to gray, brown,

196

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.19. EPMA X-ray maps of Na, Al, and Si (a–c) and compositional map (d) of a M-Pt vein and host mylonite from the Woodroﬀe Thrust, Central Australia. Note that the fragments of K-feldspar (K), quartz (Qz, black), and plagioclase (Pl) within the mylonite and M-Pt veins are plastically-deformed and elongated parallel to the mylonite foliation, indicating a sinistral shear sense (indicated by the black arrows in (c) (After Lin 2007)

and yellow-brown (Figs. 9.22–9.24). Flow structures within the matrix of G-Pt veins are deﬁned by thin colored layers and streaks that anastomose around resistant fragments (Fig. 9.25). Microlite morphology varies from simple acicular and lath-like shapes (Fig. 9.26a) to spherulitic forms (Fig. 9.26b); microlites with complex forms are not found in the granulite host rock. The microlites are composed of quartz, plagioclase (An30−40 ), biotite, ilmenite, and garnet (Table 9.1), and are several nanometers to 5 μm in width and up to 15 μm in length. This size range is generally too small to enable accurate measurements of the chemical composition of the microlites or to identify mineral type. Given the small sizes of the microlites, it is possible that the chemical compositions shown in Table 9.1 may have been contaminated by the adjacent matrix or ﬁne-grained microcrystalline material. A melting origin for the G-Pt veins is indicated by the presence of spherulitic microlites and the microlite mineral assemblage of garnet, ilmenite, plagioclase, and biotite.

9.3 Dahezhen Pseudotachylytes

197

Fig. 9.20. Photomicrograph showing the microstructures of microlites developed around fragments within M-Pt veins from the Woodroﬀe Thrust, Central Australia. Note that the microlites have been reoriented subparallel to the ﬂoliations of pseudotachylyte and the M-Pt vein. Large arrows indicate sense of shear. Plane poc larized light. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

9.3 Dahezhen Pseudotachylytes 9.3.1 Tectonic Setting of the Dahezhen Shear Zone The Dahezhen Shear Zone is located within the Precambrian Tongbei metamorphic terrane (Fig. 9.27), which is one of the major tectonic units within the Qinling-Dabie Shan collisional orogenic belt in Central China (Liu et al. 1993). The Qinling-Dabie Shan metamorphic complex consists of Precambrian gneisses that contain widespread eclogite bands, rare coesite-bearing marble bands, and rare garnet-peridotite lenses. The abundance of eclogite and highpressure (HP) mineral assemblages (Okay et al. 1989; Wang and Liou 1991) indicates that the complex underwent a regional eclogite facies metamorphism of >30 kbar and 600–750◦C (Okay and Seng¨ or 1992). Basement rocks consist mainly of metamorphosed volcanic and sedimentary rocks and gneissic complexes (Figs. 9.27 and 9.28). Gneissic banding dips consistently to the west-northwest and a mineral stretching lineation plunges to the northwest. These tectonolithostratigraphic units are separated by faults and shear zones that strike northwest–southeast, parallel to the strike of the orogenic belt. The Dahezhen Shear Zone is one of the main fault shear zones in this area, marking the boundary between Upper and Lower Proterozoic metamorphic complexes (Figs. 9.27 and 9.28).

198

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.21. Photomicrographs of crystal-plastically deformed spherulitic microlites (a) and fragments (b) within M-Pt veins from the Woodroﬀe Thrust, Central Australia. Note that the fragments are elongate and oriented subparallel to the foliation within the M-Pt veins. Large arrows indicate sense of shear. Plane polarized light. c (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

9.3.2 Field Occurrence of the Dahezhen Pseudotachylytes Pseudotachylytes within the Dahezhen Shear Zone co-exist with mylonitic and cataclasitic rocks. The pseudotachylyte veins occur at the margin of the shear zone, where the mylonite is sharply bounded by orthogneiss and quartz-monzonitic gneisses (Figs. 9.28–9.30). Stretching lineations within the mylonite plunge to the northwest; this contrasts with the plunge directions of other gneisses in the area (Fig. 9.28). The boundary between the mylonite

9.3 Dahezhen Pseudotachylytes

199

Fig. 9.22. Photomicrograph of G-Pt veins (dark areas) from the Woodroﬀe Thrust, Central Australia. The G-Pt veins were injected into the granulite rock. Qz: quartz. Plane polarized light (After Lin 2007)

zone and the surrounding gneisses dips moderately to the southwest, while the gneissic foliation dips moderately to the northwest. The mylonite zone is overprinted by late-stage cataclastic deformation. The stretching lineation and foliation within the mylonite are deﬁned by the alignment of ﬂattened fragments, the preferred orientation of biotite and feldspar, and the alignment of quartz- and feldspar-rich domains within the rock mass (Lin et al. 2003b). The asymmetric shape of feldspar and hornblende fragments within the mylonite indicates a sinistral sense of shear, meaning that the fault (or shear zone) is a thrust with a component of left-lateral displacement (Fig. 9.30). As with the Woodroﬀe Thrust, the Dahezhen Shear Zone contains two types of pseudotachylyte veins: C-Pt and M-Pt. The C-Pt veins cross-cut the M-Pt veins and the mylonitic foliation, whereas the M-Pt veins are overprinted by the mylonite (Fig. 9.29). M-Pt veins occur as individual veins along distinct fault planes (fault veins) or as injected vein networks (injection veins). Fault veins that occur along distinct shear planes have a high length-to-width ratio and can be traced along strike for several tens of meters; they range in thickness from several mm to 20 cm. In contrast, the injection veins are generally discontinuous and can only be traced over short distances (Fig. 9.29); they branch obliquely from fault veins, commonly intrude dilatant fractures, and terminate abruptly. For some thick veins, the color of the vein varies from dull colors at the vein margin to brighter colors in the center. In contrast to the C-Pt veins, M-Pt veins are observed at the relatively few locations, which are interlayered with mylonite and ultramylonite (Fig. 9.30). Figure 9.31 shows the typical texture of M-Pt veins, as observed on the

200

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.23. Photomicrographs of G-Pt veins from the Woodroﬀe Thrust, Central Australia. The G-Pt veins were injected into and around the plagioclase fragments. Pl: plagioclase, Qz: quartz. (a): plane polarized light, (b): cross polarized light (After Lin 2007)

polished X–Z surface of a hand sample. The mosaic structures evident within M-Pt veins indicate that the veins were injected into the mylonite and subsequently overprinted by ongoing mylonitization. A number of M-Pt veins cut the mylonitic foliation (Figs. 9.30 and 9.31) and can be traced back to their parent foliation-parallel M-Pt veins. An M-Pt fault vein that is 3–4 cm

9.3 Dahezhen Pseudotachylytes

201

Fig. 9.24. Photomicrographs showing the textural relationship between M-Pt and G-Pt veins from the Woodroﬀe Thrust, Central Australia. The M-Pt vein cuts the G-Pt vein and the foliation with the host granulite at angles of 40–90◦ . (a) Both the G-Pt veins and host granulite are aﬀected by crenulations. (b) Interlayered GPt veins and foliated granulite are cut by the M-Pt vein in the upper part of the image. Fragments within the M-Pt vein are strongly elongated and oriented oblique to the contact between the G-Pt and M-Pt veins, indicating sinistral displacement (indicated by large arrows). Plane polarized light (After Lin 2007)

202

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.25. EPMA-derived compositional images of G-Pt veins from the Woodroﬀe Thrust, Central Australia. Small white areas are maﬁc microlites. Plagioclase (Pl), quartz (Qz), and pyroxene fragments show embayed, rounded (a), and elongate (b) shapes set in the ﬁne-grained matrix of the vein. Mqz: silica (quartz) composition of matrix, Py: pyroxene fragment (After Lin 2007)

9.3 Dahezhen Pseudotachylytes

203

Fig. 9.26. EPMA-derived compositional images showing the microstructures of microlites developed within G-Pt veins from the Woodroﬀe Thrust, Central Australia. (a) Maﬁc minerals formed as microlites (white colors) around a plagioclase fragment (Pl). (b) Spherulitic microlites (Sp) and microlites of maﬁc minerals (Maf ) with the chemical compositions of biotite, ilmenite, and garnet (see Table 9.1 for details). Pm: plagioclase microlite (After Lin 2007)

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9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

thick occurs adjacent to the mylonite and is in sharp contact with M-Pt injection veins (Fig. 9.31). All of these M-Pt veins are crystal-plastically deformed and contain foliations deﬁned by ﬂattened and plastically deformed fragments (Fig. 9.30). This suggests that repeated pseudotachylyte-generating events (seismic faulting) occurred in the plastic-dominated region and were overprinted by ongoing mylonitization during periods of aseismic deformation. The fragments within M-Pt veins vary in diameter from several tens of microns to several centimeters and deﬁne meso- and micro-scale ﬂow banding within the ﬁne-grained matrix. The asymmetric shape of the deformed fragments indicates a sinistral shear sense, which is consistent with the shear sense recorded in the host mylonite (Figs. 9.30 and 9.31). The foliation developed within M-Pt veins that are overprinted by mylonitization are generally oriented parallel to the foliation within the bounding mylonite zone (Figs. 9.30 and 9.31). Lineations within M-Pt veins generally plunge moderately to the northwest, as with those within the host mylonite. 9.3.3 Microscopy and Chemical Composition C-Pt Microstructurally, C-Pt veins within the Dahezhen Shear Zone are characterized by textures that are typical of melt-origin pseudotachylytes (see previous Chapters), including microlites with varied morphologies, rounded and irregularly embayed fragments, and the evidence of ﬂow structures within the ﬁne-grained matrix (Fig. 9.32). The contacts between the C-Pt veins and host rocks are generally sharp, but locally are highly irregular. The matrix is generally composed of very ﬁne-grained material and microlites of An-rich plagioclase (Table 9.2). The shapes of microlites vary from acicular and lath-like to dendritic (Fig. 9.32), as with those described in Chap. 7. The geometry and chemical compositions show that these microlites which are similar to those of plagioclase fragments derived from host rock in chemical composition (Table 9.2, columns Fra1-Fra2) formed from rapid cooling or quenching of a melt, as reported by Lin (1994b). The matrix is typically opaque and locally variable in color from colorless to gray, brown, and yellowish-brown when viewed under the microscope. Fragments within the matrix consist of quartz, feldspar, garnet, and hornblende, displaying a variety of shapes such as rounded, embayed, and other irregular forms (Lin et al. 2003b). M-Pt M-Pt veins occur as simple veins oriented parallel to the mylonitic foliation and as injection veins oriented perpendicular or oblique to the foliation (Figs. 9.33–9.36). The mylonite is mainly composed of quartz, hornblende, plagioclase, K-feldspar, muscovite, biotite, and garnet. Recrystallized grains

67.19 0.25 17.58 1.05 0.05 0.0 2.10 8.77 2.26 0.10

99.26

SiO2 TiO Al2 O3 FeO* MnO MgO CaO K2 O Na2 O P2 O5

Total

55.27 0.0 28.34 0.05 0.01 0.0 11.14 0.03 4.18 0.0

Pl1

99.02 100.52

65.82 0.01 17.81 0.04 0.01 0.0 0.54 13.94 1.11 0.02

Kf2

99.30

63.66 0.0 24.22 0.07 0.01 0.0 6.00 0.04 6.47 0.05

Pl2

98.38

52.21 0.41 17.40 14.99 0.07 0.0 9.10 0.37 3.82 0.01

Gar1

96.28

38.17 0.04 23.10 12.06 0.10 0.0 22.83 0.02 0.02 0.03

Gar2

98.59

48.40 0.77 24.43 10.67 0.07 0.0 10.08 1.37 2.77 0.03

Gar3

99.45

0.01 50.88 0.0 38.42 10.14 0.0 0.0 0.0 0.0 0.0

Ilm1

93.53

2.14 49.25 1.29 32.83 7.60 0.0 0.13 0.27 0.06 0.0

Ilm2

(Host * and Max* data are from XRF analyses undertaken by Camacho et al. 1995)

Kf1

Wt%

91.29

45.68 2.32 12.34 23.22 0.26 0.0 0.35 6.61 0.45 0.06

Bio1

93.24

37.36 0.43 25.68 22.34 0.03 0.0 0.05 7.08 0.18 0.09

Bio2

95.01

0.0 0.42 0.01 94.50 0.03 0.01 0.0 0.0 0.03 0.01

Mag1

92.99

2.15 0.48 0.80 89.15 0.01 0.0 0.19 0.12 0.05 0.04

Mag2

99.47

73.91 0.31 13.48 1.44 0.06 0.65 1.54 4.87 3.21 —

Host*

99.89

63.03 0.77 18.05 5.11 0.24 2.05 2.95 2.92 4.77 —

Max*

Table 9.1. Representative chemical compositions of microlites within G-Pt veins, host felsic granulite, and the matrix of C-Pt veins from the Woodroﬀe Thrust, Australia. Kf1-2: K-feldspar; Pl1-2: plagioclase; Gar1-3: garnet; Ilm1-2: ilmenite; Bio1-2: biotite; Mag1-2: magnetite (analyzed using EPMA). Host*: average chemical composition of the host felsic granulite; Max*: average composition of the matrix within C-Pt veins. FeO* = Total Fe as FeO

9.3 Dahezhen Pseudotachylytes 205

206

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.27. Tectonic map of the Qinling-Dabie Shan collisional orogenic belt, Central China. Q-D belt: Qinling-Dabie Shan collisional orogenic belt (After Lin et al. 2003b)

of quartz and feldspar within the mylonite matrix are typically 2–20 μm in diameter. The mineral assemblage suggests that the host mylonite was deformed at amphibolite facies conditions (Lin et al. 2003b). Most of the feldspar and hornblende porphyroclasts have ‘ﬁsh’ geometries (Fig. 9.33). S-C (and C’) fabrics are well developed in the mylonite, both at the meso- and micro-scale (Figs. 9.30 and 9.33). S-surfaces within quartz-rich domains are deﬁned by the preferred orientation of quartz grains; in other domains, they are deﬁned by aligned biotite, amphibole, and feldspar porphyroclasts. C-surfaces are oriented parallel to the boundaries of mineralogical domains and are deﬁned by the ‘tails’ of asymmetric biotite, amphibole, and feldspar ‘ﬁsh’ as well as ﬁnegrained mineral aggregates. C’-surfaces are inclined in the opposite direction to the S-surfaces, which are deﬁned by micro-shear bands or microcracks. The

9.3 Dahezhen Pseudotachylytes

207

Fig. 9.28. (a) Geological map of the Tongbei region, Qinling-Dabie Shan collisional orogenic belt Central China, and (b) sketch of the boundary zone between mylonite– pseudotachylyte and gneiss (Modiﬁed from Lin et al. 2003b)

S-C fabrics indicate a sinistral shear sense, in agreement with ﬁeld observations (Fig. 9.33). A representative sample shown in Fig. 9.34 shows a M-Pt injection vein that cuts the mylonitic foliation and is itself deformed by a mylonite– ultramylonite band. The vein can be traced back to a foliation-parallel M-Pt vein. M-Pt and mylonite are interlayered at the meso- and micro-scales (Figs. 9.31 and 9.33). The boundary between the veins and host mylonite is generally irregular and curved, but locally sharp. Extreme shear ﬂow of

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9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.29. Photograph of C-Pt veins within the Dahezhen Shear Zone, Central China. The coin shown for scale is 2.4 across (After Lin et al. 2003b)

the M-Pt is indicated by elongated porphyroclasts and layering within the ﬁne-grained matrix (Figs. 9.33–9.36). In contrast to the C-Pt veins, most of the fragments within the M-Pt are elongate and ﬂattened along the foliation plane. This indicates that the M-Pt veins were overprinted by ongoing mylonitization. Some of the M-Pt veins are cut by younger C-Pt veins (Fig. 9.35), suggesting that pseudotachylytegenerating events occurred repeatedly as the rock mass passed from the plastic-dominated to brittle regimes, as with the example described above from the Woodroﬀe Thrust. Some of the fragments within the M-Pt veins are

9.3 Dahezhen Pseudotachylytes

209

Fig. 9.30. Photographs of M-Pt veins and host mylonite within the Dahezhen Shear Zone, Central China. (a) M-Pt veins (arrows) injected into host mylonitic rocks and subsequently overprinted by cataclasite. (b) Interlayered M-Pt and mylonite. The asymmetric porphyroclast fabrics indicate a sinistral shear sense (indicated by large arrows). The marker pen shown in (a) and the coin shown in (b) for scale are 16 cm long and 2.4 across, respectively(After Lin et al. 2003b)

sub-rounded and elliptical, contain irregular embayments, and have a roundness of >0.4, indicating a melting origin. The matrix within M-Pt veins is composed of ﬁne-grained material, including microcrystalline material that is too small to identify by optical microscope; however, EPMA X-ray element mapping clearly shows the different mineral types and their associated microstructures (Fig. 9.36). The foliation developed within M-Pt veins and host mylonite is characterized by crystal-plastically deformed K-feldspar, plagioclase, and quartz crystals, as well as aligned bands of microcrystalline quartz and plagioclase (Fig. 9.36). These textural modes indicates that both quartz and feldspar are dynamically recrystallized and aligned subparallel to the shear bands. All of the quartz and K-feldspar fragments and most of the plagioclase fragments within the M-Pt veins are strongly ﬂattened along the foliation plane, generally with aspect ratios of 10–20:1 or higher (Figs. 9.33–9.36). The fact that M-Pt veins are overprinted by mylonitization and C-Pt veins, in combination with the occurrence of crystal-plastically deformed fragments within the veins, indicates that they formed during rapid seismic slip and were subsequently crystal-plastically

210

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.31. Sketch of an X–Z section through a hand sample showing M-Pt and mylonite from the Dahezhen Shear Zone, Central China. A network of M-Pt veins is overprinted by ongoing mylonitization. Porphyroclasts within the M-Pt are plastically deformed and elongate subparallel to the foliation within the mylonite. Large arrows indicate sense of shear (After Lin et al. 2003b)

deformed during aseismic deformation within the plastic-dominated region. Mylonitization also occurred within the plastic-dominated region, while the C-Pt veins formed later within the brittle-dominated region during uplift of the rock mass. The M-Pt veins are greenish-yellow in the Al-element map shown in Fig. 9.36 and are thereby easily distinguished from the mylonite. Table 9.3 lists the representative chemical composition of the matrix within M-Pt veins, including that of microcrystals. M-Pt veins are chemically heterogeneous;

9.3 Dahezhen Pseudotachylytes

211

Fig. 9.32. Photomicrographs of C-Pt veins from the Dahezhen Shear Zone, Central China. (a) Network of C-Pt veins injected into cataclasite. Cross polarized light. (b) C-Pt injection veins. Plane polarized light (After Lin et al. 2003b)

however, the mylonitic layers of plagioclase, K-feldspar, and quartz are generally homogeneous, as shown in Fig. 9.36. Within individual M-Pt veins, SiO2 concentrations vary from 36 to 80 wt% and FeO + Mg + Mn varies from 1 to 30 wt% (Table 9.3). There is also a distinct contrast in chemical composition between the ﬁne-grained matrix and the host mylonite (Fig. 9.36). The compositions of the microlites include high-Ti biotite and high-An plagioclase (see M4 and M5 in Table 9.3). These are probably primary microlites that formed from the melt, as reported by Lin (1994b).

212

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Table 9.2. Representative chemical compositions of plagioclase microlites (Pl1–Pl4) and plagioclase fragments (Fra1 and Fra2) within C-Pt veins from the Dahezhen Shear Zone, Central China (analyzed using EPMA) Wt%

Pl1

Pl2

Pl3

Pl4

Fra1

Fra2

SiO2 TiO Al2 O3 FeO* MnO MgO CaO K2 O Na2 O P2 O5

61.57 0.56 20.70 3.91 0.13 1.31 3.19 2.72 5.83 0.02

63.39 0.44 20.62 2.63 0.39 0.52 2.88 2.59 6.58 0.00

52.99 0.39 28.32 0.96 0.03 0.17 12.01 0.48 3.57 0.01

52.59 0.29 29.00 1.93 0.09 0.06 12.25 0.44 3.22 0.01

63.01 0.03 22.51 0.12 0.03 0.03 6.03 0.10 8.01 0.02

66.33 0.00 18.14 0.01 0.00 0.00 0.10 13.85 1.50 0.01

Total

99.94

100.04

98.93

99.88

99.89

99.94

FeO* = Total Fe as FeO (Data from Lin et al. 2003b)

9.4 Discussion 9.4.1 Formation Mechanisms of Large Volumes of Pseudotachylytes Seismic and geological data reveal that most large-magnitude intracontinental earthquakes appear to occur on pre-existing mature active faults and that these faults control the temporal and spatial distribution of displacement and rupture processes (e.g., Yeats et al. 1997; Lin et al. 2003a). Long-lived active faults that record hundreds of kilometers of accumulated displacement are considered to represent repeated episodes of seismic slip associated with thousands of large-magnitude earthquakes along the main fault plane. Previous studies reveal that C-Pt forms by seismic slip upon fault planes in the upper crust, generally at depths of <10 km (e.g., Sibson 1975; Toyoshima 1990; Lin 1994a, b; Fabbri et al. 2000), and that M-Pt forms in the greenschist facies or even lower grades, corresponding to temperatures of <300–400◦C and depths of <10–15 km assuming an average geothermal gradient of 25◦ C/km (e.g., Sibson 1983; Hobbs et al. 1986). The depth of formation of M-Pt is generally coincident with the lower depth limit of the seismogenic zone, as shown in Fig. 9.37. The formation of pseudotachylyte has also been documented during coseismic landsliding at depths of <40 m during the 1999 Mw 7.6 Chi-Chi (Taiwan) earthquake (Lin et al. 2001a; see Chap. 11 for details). High-velocity frictional melting experiments also demonstrate that frictional melting can occur at conditions of low normal stress, even less than 1 MPa, corresponding to

9.4 Discussion

213

Fig. 9.33. Photomicrographs of deformed M-Pt veins from the Dahezhen Shear Zone, Central China. (a) M-Pt veins interlayered with mylonite. Fragments within the veins are crystal-plastically deformed. (b) S-C fabrics developed within the mylonitized pseudotachylyte vein (M-Pt). The geometry of the S-C fabrics indicates a sinistral sense of shear (indicated by large arrows). S: S-surface; C: C-surface. (a) Cross polarized light, (b) plane polarized light. (After Lin et al. 2003b)

depths of <30 m (Lin 1991; Lin 1994, 1998). The above geologic, seismic, and experimental data indicate that melt-origin pseudotachylyte can form at various depths in the crust, from the near-surface to within deep-level seismogenic fault zones.

214

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.34. Photomicrograph (a) and accompanying sketch (b) of an M-Pt vein within mylonite from the Dahezhen Shear Zone, Central China. Note that the M-Pt vein cuts the mylonitic foliation and is itself sheared and has been transformed to mylonite. (a): Plane polarized light (After Lin et al. 2003b)

As discussed in Chap. 3, quartz, which is one of the major rock-forming minerals in granitic rocks, is a sensitive indicator of the transition between brittle and plastic deformation within a fault shear zone. Quartz fragments within C-Pt veins from both the Woodroﬀe Thrust and Dahezhen Shear Zone show no explicit evidence of dynamic recrystallization or crystal-plastic deformation. This suggests that the C-Pt veins formed within the brittle-dominated regime at temperatures of <300◦C. A melt origin for these pseudotachylytes is indicated by ﬂow structures within C-Pt veins and the presence of microlites that are similar to those found in igneous rocks and melt-origin pseudotachylytes as described in Chap. 6. For both the Woodroﬀe (Camacho et al. 1995; Lin et al. 2005b) and the Dahezhen pseudotachylytes (Lin et al. 2003b), the matrix within C-Pt veins is generally deﬁcient in silica relative to the host rocks. This is indicative of a melt origin and is considered to reﬂect the preferential melting of minerals with low melting temperatures during pseudotachylyte formation (e.g., Lin 1994a; Camacho et al. 1995; Lin and Shimamoto 1998). It is diﬃcult to generate large volumes of pseudotachylyte, up to 70–90 Vol% of the generation zone across a wide zone of >3.0 km in the case of the Woodroﬀe Thrust, during a single seismic faulting event because of the limited

9.4 Discussion

215

Fig. 9.35. Photomicrograph showing the microstructural relationship between MPt, mylonite, and C-Pt veins from the Dahezhen Shear Zone, Central China. M-Pt veins interlayered with mylonite are cut by C-Pt veins. Note that the M-Pt vein is plastically deformed and that fragments within the vein are elongate subparallel to the foliation. Plane polarized light (After Lin et al. 2003b)

slip that takes place during a large-magnitude earthquake. Seismic inversion studies suggest that even the large-magnitude (Mw 8.3) Bolivian (1994) earthquake only produced a melt layer of 31 cm thickness at a depth of 670 km (Kanamori et al. 1998). The empirical relation between the thickness of a pseudotachylyte vein and the amount of fault slip dictates that a slip increment of 5 m would give rise to a uniform layer of melt of just 1 cm in thickness (Sibson 1975). The amount of coseismic displacement arising from a large (M >7) earthquake is generally <10 m (Yeats et al. 1997; Lin et al. 2001b) which produces a uniform melt layer of <2 cm in thickness, although more than 10 m of displacement has been recorded in the past (Lin and Lin 1998; Lin et al. 2002a, 2003a). The overprinting of pseudotachylyte veins, in which fragments of the older vein are included in the younger vein (as observed in the Woodroﬀe C-Pt pseudotachylytes), indicates the occurrence of multiple pseudotachylyte-generating events within the same part of the fault zone. The overprinting of successive generations of C-Pt and M-Pt veins also indicates that pseudotachylyte formed during repeated seismic faulting events in both the brittle- and crystal plastic-dominated regimes within the same fault shear zone. On this basis, it is concluded that the large volumes of pseudotachylyte developed along the Woodroﬀe Thrust reﬂect multiple large-magnitude seismic

216

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.36. EPMA X-ray Al map of M-Pt veins and host mylonite from the Dahezhen Shear Zone, Central China. Note that the porphyroclasts of K-feldspar (red colors), quartz (black), and plagioclase (pink) within the mylonite and M-Pt veins are plastically deformed and elongate subparallel to the foliation within the mylonite (After Lin et al. 2003b)

faulting events that occurred at various depths from the near-surface to deeplevel fault zones within the lower crust (Lin et al. 2005b). 9.4.2 Conditions of Formation of the Dahezhen and Woodroﬀe M-Pt Veins Dahezhen M-Pt The Qinling-Dabie Shan collisional orogenic belt formed during the Triassic collision of the North China and Yangtze continental blocks (Seng¨or 1985). The exposure of UHP rocks within the orogenic belt implies that the collision event led to their exhumation. Several tectonic models have been proposed to explain the exhumation of these coherent continental high-pressure (HP) metamorphic terrains within the collisional orogen: i) thrusting and concomitant erosion (e.g., Rubie 1984; Hs¨ u 1991), ii) underplating and extension (Platt 1986), iii) Triassic subhorizontal extrusion and Cretaceous magmatic doming (Maruyama et al. 1994), and iv) Triassic indentation, subvertical extrusion, and Cretaceous plate-margin transtension (Hacker et al. 1995).

9.4 Discussion

217

Table 9.3. Representative chemical compositions of the matrices (M1–M3) and microlites (M4 and M5) within M-Pt veins and a hornblende fragment (Hor1), plagioclase (Pl1 and Pl2), K-feldspar (K1), and biotite (Bi1) layers within mylonite from the Dahezhen Shear Zone, Central China Wt%

M1

SiO2 78.71 TiO 0.40 Al2 O3 10.95 FeO* 3.06 MnO 0.00 MgO 1.24 CaO 1.39 K2 O 1.80 Na2 O 1.98 P2 O5 0.02 Total 99.57

M2

M3

M4

M5

Hor1

Pl1

Pl2

80.30 0.30 11.09 1.45 0.02 0.70 1.99 1.96 2.15 0.01 99.96

67.13 0.66 15.85 6.36 0.13 2.07 0.36 3.23 3.18 0.00 98.97

36.22 6.54 18.45 16.63 0.08 13.61 0.30 5.16 0.30 0.11 97.39

51.71 0.03 25.95 0.06 0.01 0.03 10.99 0.50 10.68 0.02 99.98

39.16 3.74 13.62 13.37 0.27 10.41 10.39 1.39 2.38 0.15 94.88

52.84 57.89 0.23 0.05 26.13 25.79 1.00 0.04 0.23 0.12 0.11 0.06 6.23 7.25 1.44 0.78 10.85 7.45 0.00 0.01 99.06 100.24

K1

Bi1

63.95 0.00 18.25 0.00 0.07 0.00 0.02 17.16 0.78 0.01 99.44

38.76 9.34 16.45 11.64 0.05 14.35 0.75 5.16 1.30 0.12 97.92

FeO* = Total Fe as FeO (After Lin et al. 2003b)

Approximately 265 km of post-collisional thrusting beneath the UHP metamorphic rocks has been documented within the belt (Okay and Seng¨ or 1992). The underthrust lower plate was exhumed by the southward propagation of thrust planes and consequent incorporation of the subducted plate into the hanging wall (Okay and Seng¨ or 1992; Nie et al. 1994). This shows that major intracontinental thrusting concomitant with erosion and back-thrusting at the rear of the orogen were the main factors in the exhumation of the UHP metamorphic rocks (Okay and Seng¨or 1992). As stated above, the Dahezhen Shear Zone represents an example of these major intracontinental thrusts. Within the Dahezhen Shear Zone, the transition from crystal-plastic to brittle deformation is marked by a succession of deformation styles: from HPUHP metamorphic rocks to mylonite, M-Pt within a ductile shear zone, and ﬁnally cataclasite and C-Pt within a brittle fault zone. This sequence reveals that seismic faulting occurred repeatedly in the shear zone within both the brittle-dominated and crystal plastic-dominated regimes. Field relations and meso- and micro-structures suggest that the mylonite and multi-stage MPt veins formed within the crystal plastic-dominated regime and were later overprinted by cataclasis and C-Pt within the brittle-dominated regime during uplift and exhumation of the shear zone. There is no evidence of dynamic recrystallization or crystal-plastic deformation within fragments in the C-Pt and related cataclastic rocks. This indicates that the C-Pt formed at shallow depths within the brittle domain at <300◦C. In contrast, the mylonite and M-Pt are characterized by layers of ﬁne-grained dynamically recrystallized quartz, feldspar, and amphibole that deﬁne the foliation within mylonite along with plastically deformed biotite, feldspar, and amphibole porphyroclasts.

218

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

Fig. 9.37. Schematic model of the propagation of seismic slip within a large-scale fault zone developed in continental crust. The proﬁle of shear resistance (τ) with depth is arbitrary, given a geothermal gradient of 25◦ C/km (after Sibson 1983 and Scholz 2002). Repeated cycles of earthquake nucleation occur at the base of the seismogenic zone (indicated by the catﬁsh) within the brittle-dominated regime. During each rupture event, seismic slip propagates down into the plastic-ﬂow regime, even into the lower crust. Over time, large volumes of M-Pt and Um-Pt form as veins within the plastic-dominated regime; these veins are ﬁrst overprinted by subsequent mylonitization and then by cataclasis and the formation of C-Pt in the brittledominated seismogenic zone following uplift and exhumation. The long-lived and repetitive nature of seismic events along the fault zone explains the coexistence of both C-Pt and Um-Pt veins and results in large accumulated displacements and large volumes of pseudotachylyte. The catﬁsh represents the focus area of large earthquakes. According to Japanese and China folklore, earthquakes occur when a c giant underground-dwelling catﬁsh thrashes about. (After Lin et al. 2005b). 2007, with kind permission from Elsevier Science Ltd

The crystal-plastic deformation of feldspar fragments within M-Pt veins indicates a minimum deformation temperature of 450◦ C (Simpson 1985; Carter and Tsenn 1987; Scholz 1988). The observed deformation textures, in combination with the mineral assemblage of hornblende, plagioclase,

9.4 Discussion

219

K-feldspar, quartz, muscovite, biotite, and garnet, suggests that the mylonite and M-Pt veins formed under amphibolite facies conditions corresponding to temperatures of 400–650◦C. The anorthite component of recrystallized plagioclase layers within the mylonite and M-Pt veins varies from An37 to An51 (see column Pl1 in Table 9.3). On the basis of the Al2 O3 content of hornblende (see Plyusinina, 1982), the temperature and pressure conditions of the mylonite and related M-Pt veins, as determined from plagioclase- and hornblende-bearing assemblages under conditions of chemical equilibrium, are estimated to be 400–800 MPa and 490–570◦C (see column Hor1 in Table 9.3). Taking account of the 50◦ C uncertainty involved in the models of Spear (1980) and Plyusinina (1982), we arrive at a minimum estimated temperature of 400–420◦C and maximum of 620– 650◦C (Lin et al. 2003b). This corresponds to depths of 20–33 km assuming an average geothermal gradient of 20◦ C /km for the intracontinental crust in the study area (Cong and Wang, 1995). Accordingly, it is inferred that the M-Pt veins within the Dahezhen Shear Zone formed at a depth of 20–33 km. The coexistence of the two types of pseudotachylyte veins reveals that pseudotachylyte-producing seismic faulting events occurred repeatedly at different crustal levels within the shear zone. These seismic events occurred during rapid thrust-related exhumation of UHP metamorphic rocks within the Qinling-Dabie Shan collisional orogenic belt. As documented above, the Dahezhen M-Pt formed at estimated conditions of 400–650◦C and 400–800 MPa, corresponding to the crystal plastic-dominated regime and deeper than the base of the ‘normal’ seismogenic zone (Fig. 9.37). Woodroﬀe M-Pt and G-Pt In the case of the Woodroﬀe Thrust, the co-existence of large volumes of CPt and interlayered M-Pt and related mylonite–ultramylonite indicates cyclic mylonitization and pseudotachylyte formation within the crystal plasticdominated regime, deep within the fault shear zone. As with the Dahezhen M-Pt, most of the quartz and feldspar fragments within M-Pt and U-Pt veins along the Woodroﬀe Thrust are crystal-plastically deformed, although fragments of garnet within M-Pt veins remain undeformed. In contrast to the C-Pt, the mylonite–ultramylonite, M-Pt, and Um-Pt are characterized by dynamically-recrystallized ﬁne-grained quartz and feldspar. These textural features, in combination with the observed mineral assemblages, suggest that the mylonitic rocks, M-Pt, and Um-Pt formed under amphibolite facies conditions (450–600◦C). This corresponds to depths of 25– 30 km assuming an average geothermal gradient of 20◦ C /km. The fact that the foliation within M-Pt and Um-Pt veins is generally oriented parallel to that within the mylonite shear zone indicates that the mylonitic rocks and M-Pt and U-Pt veins formed contemporaneously under the same shear-stress and P-T conditions.

220

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

The structural and textural features of the fault rocks and various types of pseudotachylyte developed along the Woodroﬀe Thrust (as described above) reveal the following: i) that G-Pt veins co-exist with M-Pt, Um-Pt, and C-Pt veins, along with cataclastic and mylonitic rocks, within individual fault outcrops, ii) that the ﬂexural fabrics developed within G-Pt veins are similar to those observed in the host granulite gneiss, and iii) that the G-Pt veins are cut by younger M-Pt veins and mylonite shear zones. In turn, these observations suggest that: i) the G-Pt veins formed contemporaneously with metamorphism of the host granulite facies rocks, prior to the formation of M-Pt and associated mylonitic rocks; and ii) the G-Pt veins were dynamically overprinted by younger M-Pt veins during coseismic faulting and aseismic mylonitization; this occurred within the brittle–plastic transition zone, followed by overprinting by C-Pt and associated cataclastic rocks within the brittle-dominated regime in the upper crust. The mylonite–ultramylonite, M-Pt, and Um-Pt are characterized by dynamically recrystallized ﬁne-grained quartz and feldspar. This suggests that they formed under amphibolite facies conditions (Lin et al. 2005b). That fact that K-feldspar fragments within those mylonitic rocks associated with MPt are crystal-plastically deformed and dynamically recrystallized indicates temperatures of up to ∼650◦ C (Tullis and Yund, 1991). This corresponds to depths of ∼30–35 km assuming an average geothermal gradient of 20◦ C /km for continental crust. The petrologic and structural features described above indicate that the G-Pt veins formed in the deeper plastic-dominated regime under granulite facies conditions of ∼8 kbar and 650–700◦C (Camacho et al. 1997, 2001), corresponding to depths of >∼35 km, deeper than those at which the M-Pt and mylonitic rocks formed. In summary, the G-Pt veins formed within the crystal plastic-dominated regime of the lower crust under granulite facies conditions; these were subsequently overprinted by M-Pt and C-Pt veins during exhumation and uplift of the granulite facies rocks into the shallow brittle-dominated regime of the upper crust. Propagation of Seismic Slip from the Brittle to Plastic Deformation Regimes M-Pt veins reported in previous studies generally co-exist with C-Pt veins and cataclastic and mylonitic rocks within individual outcrops along a fault shear zone, as documented along the Woodroﬀe Thrust and the Dahezhen Shear Zone. The joint occurrence of these rocks indicates that M-Pt veins and associated mylonitic rocks form in the ductile regime and are dynamically overprinted by C-Pt veins and cataclastic deformation within the brittle regime. These features record either a continuous and long-lived history of faulting or the discontinuous reactivation of pre-existing faults alternating with long periods of seismic quiescence. These repeated pseudotachylytegenerating events occurred during the exhumation and uplift of the rock mass

9.4 Discussion

221

from the deep-level crystal-plastic-dominated regime to shallow depths within the brittle-dominated regime. Previous studies of mylonite and associated M-Pt veins within fault shear zones have led to the proposal of three primary deformation paths that explain the synchronous development of M-Pt and mylonite: i) cyclic generation of pseudotachylyte within deforming mylonite as a thrust block moves through the base of the seismogenic zone (e.g., Passchier 1982; Koch and Masch 1992); ii) downward propagation of dynamic fractures, which initiated within the seismogenic zone, to the plastic-dominated regime (Sibson 1980a; Scholz 2002); and iii) development of plastic instability in the plasticdominated regime, leading to the development of temperature transients and associated melting (Hobbs et al. 1986; White 1996). In models i) and ii) above, the pseudotachylyte-generating events occur within the seismogenic zone, whereas in model iii) the events occur at depths greater than those of the ‘normal’ seismogenic zone. Both models i) and iii) dictate that the ambient temperature of the shear zone within which M-Pt is generated is less than 300–400◦C (Sibson 1983; Hobbs et al. 1986), corresponding to greenschist facies conditions. In terms of model iii), a necessary condition for plastic instability in the plastic regime at a given strain rate is that the temperature is below the critical values of <400◦C at which transient work hardening is equal to the product of stress relaxation associated with thermal ﬂuctuation and heat generated by shearing (Hobbs et al. 1986). The cyclic brittle–plastic formation of mylonite-hosted pseudotachylyte is considered to occur at depths greater than that of the “normal” seismogenic zone of <10–15 km (Hobbs et al. 1986; McNulty 1995; White 1996; Lin et al. 2003b, 2005b). Hobbs et al. (1986) investigated M-Pt within the Red Bank Shear Zone, Central Australia, and suggested that plastic instability in the plastic-dominated regime led to the development of thermal transients and associated melting and generation of M-Pt. White (1996) inferred that M-Pt along the Outer Hebrides Thrust, Scotland, formed at amphibolite facies conditions of ∼500◦C and 140–210 MPa, corresponding to depths of ∼25 km assuming a geothermal gradient of 20◦ C/km; this was based upon the interpretation that the M-Pt formed by plastic instability (Hobbs et al. 1986). In contrast to the plastic-instability model, the above propagation model (model ii) is more appropriate in explaining the coexistence of M-Pt and C-Pt (which developed in both the brittle- and plastic-dominated regimes) within individual fault shear zones, as well as being more appropriate in explaining their formation mechanisms (Lin et al. 2005b). Plastic instability is unlikely to have been an important process in the formation of pseudotachylyte along the Woodroﬀe Thrust and the Dahezhen Shear Zone because of the high temperatures (>400–650◦C) of pseudotachylyte formation in these areas and the relatively low resistance to large-magnitude earthquakes within the regime dominated by crystal-plastic deformation (Lin et al. 2003b, 2005b). In this deep regime, the shear strain energy would be consumed by the continuous crystal-plastic deformation and dynamical recrystallization of quartz and

222

9 Formation of Pseudotachylyte in the Brittle and Plastic Regimes

feldspar within the fault shear zone. The greatest resistance and distortional strain energy within the fault shear zone are inferred to be concentrated at or near the base of the seismogenic zone as shown in Fig. 9.37. Accordingly, it is diﬃcult to accumulate the shear strain required to trigger large-magnitude earthquakes solely by plastic instability in the regime dominated by crystalplastic deformation. In this part of the crust, the shear resistance is much less than that within a continental fault zone in the brittle-dominated regime (Fig. 3.37). It is well known that large-magnitude earthquakes occur repeatedly along long-lived mature active faults and that they nucleate near the base of the seismogenic zone at <10–15 km depth within continental crust. Although it is generally assumed that aftershocks delineate the rupture zone and depth extent of the main shock, they do not provide a reliable indication of the regime within which seismic rupture associated with large-magnitude earthquakes is able propagate downward and upward within the crust (Scholz 2002). In fact, a previous seismic inversion study reveals that seismic slip associated with large-magnitude earthquakes is able to propagate downward to greater depths in the crust than that aﬀected by related aftershocks (Kikuchi 2003). Calculations based on laboratory-derived experimental data indicate that large strike-slip earthquakes occur around the nucleation depth and above, but seismic rupturing can propagate downward to greater depths (Tse and Rice 1986). Modern seismologic techniques make it possible to accurately deﬁne the focus location and maximum depth of seismic rupturing (Kikuchi and Kanamori, 1996; Kikuchi, 2003). Seismic inversion results reveal that most of the recent Mw >7.5 strike-slip earthquakes occurred at a focus depth of <10–20 km but that associated seismic slip propagated to depths of up to 20–40 km (Kikuchi 2003). An example of such an earthquake is the 2001 Kunlun Mw 7.8 earthquake (Tibet), which generated a maximum sinistral displacement of 16.3 m along a surface rupture zone of 400–450 km in length (Lin et al. 2002a, 2003a). Rupture initiated at 17 km depth and propagated downward to around 40 km (Lin et al. 2003a). Similar examples of downward rupture propagation have also been documented from large thrust-related earthquakes such as the 1999 Taiwan Mw 7.6 earthquake (focus depth of 7 km, propagation to 20 km), the 2001 India Mw 7.6 earthquake (focus depth of 10 km, propagation to 40 km), and the 2001 Peru Mw 8.2 earthquake (focus depth of 30 km, propagation to 50 km) (Kikuchi 2003). These seismic inversion results demonstrate that both large strike-slip and thrust-related earthquakes are able to rupture the entire seismogenic zone and that coseismic slip not only propagates upward to the Earth’s surface where displacement can be directly observed, but also downward deep into the lower crust where deformation occurs dominantly within the crystal-plastic regime (Fig. 9.37). On this basis, it is proposed that M-Pt veins form within the plasticdominated regime by frictional melting resulting from the downward propagation of seismic slip associated with large-magnitude earthquakes that

9.4 Discussion

223

nucleated at the base of the brittle-dominated seismogenic fault zone (Lin et al. 2003b, 2005b). Earthquake data reveal that large-magnitude intracontinental earthquakes (M >5) generally occur at depths of <15 km (Sibson 1983; Hill et al. 1990; Kikuchi 2003); this supports the idea that seismic slip propagates downward to deep levels within the fault zone. During aseismic periods, pseudotachylyte veins that formed during periods of seismic faulting within the plastic-dominated regime are overprinted by ongoing mylonitization at deep levels within the fault shear zone. Although previous studies have reported the formation of pseudotachylyte in the lower crust at granulite–eclogite facies conditions of >700◦ C, 700– 1900 MPa, and >60 km depth (Clarke and Norman 1993; Austrheim and Boundy 1994; Lund and Austrheim 2003), this can be explained by pseudotachylyte formation associated with deep earthquakes that occurred at depths greater than the lower limit of continental crust (Lund and Austrheim 2003). This suggests that such pseudotachylytes are the products of deep-level seismicity but are not related to faulting in the plastic-dominated part of the shear zone that occurs within the normal seismogenic zone within continental crust. Such pseudotachylyte veins are not overprinted by brittle–plastic deformation. This is in marked contrast to pseudotachylyte observed along the Woodroﬀe Thrust, where G-Pt veins are overprinted by M-Pt, Um-Pt, and C-Pt veins and associated cataclastic and mylonitic rocks within individual fault zones. The important question here is whether the G-Pt veins formed in association with earthquakes that nucleated within the plastic-dominated regime under granulite facies conditions. The strength proﬁle of continental crust reveals that the highest resistance and distortional strain energy within a fault shear zone are concentrated at or near the base of the seismogenic zone (Sibson 1983), as shown in Fig. 9.37. The shear strain energy within fault shear zones at this depth is consumed by the continuous crystal-plastic deformation and dynamical recrystallization of quartz and feldspar under granulite facies conditions. Therefore, within continental fault zone it is diﬃcult to accumulate suﬃcient shear strain solely on the basis of plastic instability to trigger large-magnitude earthquakes in the plastic ﬂow regime where the fault shear resistance is much less than that at the base of brittle-dominated regime. The Woodroﬀe Thrust zone, within which G-Pt veins are developed, is 3–4 km thick (Fig. 9.2). This strongly suggests that the G-Pt veins formed at deep levels within the fault shear zone under granulite facies conditions due to the downward propagation of seismic slip from the hypocenter located in the brittle-dominated seismogenic zone (Fig. 9.37). The formation mechanism of the G-Pt is comparable to that of M-Pt, as proposed by Lin et al. (2005b). The modiﬁed fault model shown in Fig. 9.37 states that the repeated seismic rupturing and propagation of slip that occurs during large-magnitude earthquakes along mature faults is able to explain the large accumulated displacements observed along the faults and the coexistence of large volumes of C-Pt, M-Pt, Um-Pt, and G-Pt veins and associated fault rocks that formed within

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individual fault zones at diﬀerent depths from the brittle-dominated regime in the upper crust down to the brittle–plastic transition zone and ﬁnally into the plastic ﬂow regime of the lower crust. The downward propagation of seismogenic rupture into the predominantly aseismic regime may account for the generation of deep-level pseudotachylyte within deep fault shear zones.

10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

10.1 Introduction One of the major goals of observing fault rocks as part of paleoseismological studies is to establish criteria that can be used to distinguish seismic slip from aseismic creep in fault zones. In addition to the melt-origin pseudotachylyte described in previous chapters, the following rock types are also considered to record fossil earthquakes, i.e. seismic slip at shallow depths within fault zones: veinlet cataclastic rocks such as fault gouge (Lin et al. 1994, 1998b; Lin 1996), clastic-type pseudotachylyte (Philpotts 1964; Lin 1996, 1997a; Shigetomi and Lin 1999; Kano et al. 2004), fault breccia (Sibson 1986), and veins of cataclasite (Chester and Chester 1998) that occur as both simple veins and complicated networks within fault zones. Veinlet cataclastic rocks share some of the characteristics of melt-origin pseudotachylyte veins that are injected into fractures, although some resemble sedimentary breccias and are cemented by a ﬁne-grained matrix. The matrices of clastic-type pseudotachylyte veins comprise ﬁne-grained fragments of the host rock that are nanometer to sub-millimeter in scale, with only a small amount of amorphous material such as glass and without typical melting textures such as microlites, vesicles, amygdules, or ﬂow structures. As with meltorigin pseudotachylyte veins, the clastic-type pseudotachylyte veins also occur as simple veins and vein networks injected into the country rocks and are dark and aphanitic in appearance (e.g., Lin 1996, 1997a; Kano et al. 2004). To distinguish them from melt-origin pseudotachylyte, clastic-type pseudotachylytes are termed crushing-origin pseudotachylyte (Lin 1996, 1997a). These rocks are generated by rapid comminution and injection during seismic faulting and are therefore also considered to represent fossil earthquakes, as with melt-origin pseudotachylyte (e.g., Lin et al. 1994; Lin 1996, 1997a; Kano et al. 2004). The structural modes of non-primary cohesive or incohesive cataclastic injection veins, including crushing-origin pseudotachylyte, fault gouge, and some calcite veins that occur along fault-fracture networks, suggest that the associated fault-fracture networks are related to the propagation of dynamic

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rupture and fault movement during incremental coseismic displacement or rapid slip (Lin 1996, 1997a). Although many studies have demonstrated that veinlet cataclastic rocks are related to seismic faulting, a lack of ﬁeld and petrologic evidence means that it remains disputed as to whether they actually preserve a record of seismic slip (Cowan 1999). This chapter presents a number of representative examples of veinlet cataclastic rocks, including crushing-origin pseudotachylyte, fault gouge veins, and crack-ﬁll veins, that provide evidence of seismic faulting within fault zones. Their formation mechanisms are discussed along with their signiﬁcance in terms of seismic faulting events. This study of the formation mechanism of deformation-related cataclastic veins attempts to bridge the gap between theoretical simulations, frictional experiments, and ﬁeld observations in terms of investigating the seismic records of seismogenic fault zones.

10.2 Occurrence of Crushing-Origin Pseudotachylyte and Cataclastic Veins 10.2.1 Crushing-Origin Pseudotachylyte Typical crushing-origin pseudotachylyte veins have been reported from active fault zones in Japan such as the Iida–Matsukawa Fault (Lin 1989, 1996; Lin et al. 1994a, b), the Nojima Fault, which triggered the 1995 Kobe Mw 7.2 earthquake (Shigetomi and Lin 1999; Lin 2001), and the Itoigawa–Shizuoka Tectonic Line Active Fault Zone (Kano et al. 2004). As with melt-origin pseudotachylyte veins, crushing-origin veins occur both as simple veins (fault veins) along primary fault planes (Fig. 10.1) and complex vein networks (injection veins) within country rocks (Fig. 10.2). Injection veins are generally developed within fault-fracture zones composed of cataclasite and fault breccia; these are generally located between several centimeters and several meters from the source fault plane and can locally be traced back to the parent generation plane upon which continuous fault veins are observed. Injection veins are locally preserved within fault zones as pocked pools of >10 cm in diameter and thin isolated lenses of 1–20 mm in width, accompanied by cataclastic rocks (Fig. 10.2). Pseudotachylyte veins may be oﬀset along fractures due to faulting (Fig. 10.3), and dark and aphanitic pseudotachylyte fragments are commonly found within the fault breccia–gouge zones that occur along primary fault planes (Fig. 10.4). The structural features of these oﬀset and fragmented pseudotachylyte veins indicate that seismic rupturing events occur repeatedly upon individual fault planes, as is also indicated by overprinting melt-origin pseudotachylyte veins (see Chaps. 4 and 9). The contacts between injection veins and host cataclastic rocks are generally sharp, although locally they may be gradational. Individual veins vary in width from several millimeters to several

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Fig. 10.1. Fault vein of pseudotachylyte (Pt) along the main fault plane of the c Iida–Matsukawa Fault, Japan. (After Lin 1996). 2007, with kind permission from Elsevier Science Ltd

centimeters, with most being 3–10 mm; fault veins can generally be traced for up to several meters along the fault plane. The granite-hosted pseudotachylyte veins described in this chapter are generally characterized by a compact and aphanitic appearance, are dark-brown to black in color, and locally show a vitreous luster that is similar to that of melt-origin pseudotachylytes. On fresh outcrop cuttings, pseudotachylyte veins may have a protuberant appearance, occurring as thin dikes that stand out from the host cataclasite (Fig. 10.3). This mode of occurrence indicates that the veins are denser and harder than the host cataclasite. Based on the structural relationship between source faults and pseudotachylyte veins, it is sometimes possible to infer the timing of the formation of crushing-origin pseudotachylyte veins. Such an example has been reported from the Itoigawa–Shizuoka Tectonic Line Active Fault Zone

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Fig. 10.2. Network of pseudotachylyte (Pt) veins (indicated by arrows) within granitic cataclasite in the hanging wall of the Itoigawa–Shizuoka Tectonic Line Active Fault Zone at Tozawa site, Japan. (a) Complex network veins. Large arrows indicate the fault plane (F), along which the granitic rock is thrust over alluvial deposits. (b) Close-up view of injection veins. The interval of the white grid lines shown in (a) which are used for sketching the section is 50 cm

(Kano et al. 2004), where granitic rocks are thrust over alluvial deposits (Fig. 10.2). At this site, pseudotachylyte veins occur in the hanging wall of the fault as networks of injection veins within granite-hosted cataclasite, as well as fault veins along the main fault (Fig. 10.2). The fault veins are bounded

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Fig. 10.3. Photograph of a pseudotachylyte injection vein (Pt) within granitic cataclasite within the Iida–Matsukawa Fault zone, Japan. The injection vein (indicated c by the arrows) is oﬀset along a fault. (After Lin 1996). 2007, with kind permission from Elsevier Science Ltd

below by the alluvial deposits of the footwall and are locally injected into these deposits. The intrusive mode of the injection veins within the alluvial deposits suggests that the veins were injected into coseismic ruptures that formed during recent seismic faulting events. Trenches excavated across the Itoigawa–Shizuoka Tectonic Line Active Fault Zone at a site located close to this pseudotachylyte outcrop reveal that the most recent seismic faulting event occurred between 1550 ± 70 and 2350 ± 60 years B.P. with a displacement of 1–1.2 m along the main fault plane (Touda et al. 2000; Tachikawa 2002). Based on the structural mode of the injection veins and 14 C dating of the alluvial deposits, it is inferred that at least some of the pseudotachylyte veins formed during the most recent seismic faulting events, during the past 1500–2500 years, at depths of <10 m (Kano et al. 2004).

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Fig. 10.4. Photograph of a pseudotachylyte (Pt) fault vein bounded by cataclasite (to the left) and a fault breccia zone (to the right) that contains fragments of pseudotachylyte within the Iida–Matsukawa Fault, Japan

10.2.2 Fault-Gouge Injection Veins Fault gouges are generally produced along primary fault planes as the result of incremental fault slip during recent faulting events along active fault zones. Over the past two decades, fault gouge veins and vein networks have also been reported from active fault zones located far from the main fault plane (e.g., Lin 1989, 1996; Maruyama and Lin 2004). Typical fault-gouge veins and vein networks have been reported from active fault zones in Japan such as the Iida– Matsukawa Fault (Figs. 10.5 and 10.6), which contains crushing-origin pseudotachylyte veins (Lin 1996), the Rokko Fault (Lin et al. 1998b, 2001d), and the Arima-Takatsuki Tectonic Line (Maruyama and Lin 2004). Fault gouge veins hosted within granitic rocks generally show varied colors ranging from gray to brown-gray, dark-gray, yellow-reddish brown, and dark brown. The contacts between gouge injection veins and host cataclastic rocks are generally sharp, with no distinct oﬀset across the contacts and no evidence of continuous shear fabrics across the contacts. Where observed along the Iida–Matsukawa

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Fig. 10.5. Photograph of a fault vein along the main fault plane (F) and injection vein of fault gouge within foliated cataclasite in the Iida–Matsukawa Fault zone, c Japan. The pen shown for scale is 15 cm long. (After Lin 1996). 2007, with kind permission from Elsevier Science Ltd

Fault, fault-gouge injection veins are weakly consolidated but are generally more strongly indurated than the host cataclastic rocks in which they occur (Lin 1996, 1997a). The above observations indicate that fault gouge veins are intrusive in origin and are generated upon the main fault plane, as with pseudotachylyte injection veins. The following structural features are common to fault gouge veins and crushing-origin pseudotachylyte injection veins: i) the veins generally occur as simple veins and complex networks within fault-fracture zones in association with cataclastic rocks (Figs. 10.5 and 10.6); ii) the injection veins can generally be traced back to the parent fault vein along the main fault plane (Fig. 10.6a); iii) the veins are locally preserved as isolated lenses within faultdamage zones; iv) the veins terminate sharply against fractures at distances of several centimeters to several meters from the main fault plane (Figs. 10.5 and 10.6); and v) the veins vary in width from several millimeters to >10 cm, with most being 3–10 mm.

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Fig. 10.6. Injection veins of fault gouge within the Iida–Matsukawa Fault zone, Japan. (a) Gouge injection veins are connected to the fault vein that occurs along the fault plane (F, indicated by large arrow). (b) The gouge injection vein terminates sharply within cataclasite. The pencil shown for scale is 15 cm long. (After Lin 1996). c 2007, with kind permission from Elsevier Science Ltd

10.2.3 Layered Fault Gouge and Pseudotachylyte Veins Fault gouges generally occur as a thin layer or ﬁlm along the main fault plane. Interlayered fault gouge, crushing-origin pseudotachylyte, and fault breccia have been reported from the Nojima Fault, Japan (Fig. 10.7; Shigetomi and Lin 1999; Lin 2001). In this case, thin and hard pseudotachylyte veins are characterized by sandwich structure, which consists of interlayered weak to

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Fig. 10.7. Polished hand sample showing interlayered pseudotachylyte (Pt) and fault gouge veins from the Nojima Fault, Japan. F: Fault plane. (After Lin 2001). c 2007, with kind permission from Elsevier Science Ltd

unconsolidated fault gouge, breccia layers varying in color from yellowishand pinkish-gray to dark gray, and dark consolidated pseudotachylyte veins. Figures 10.7 and 10.8 and show the typical occurrence of sandwich structure that consists of pseudotachylyte fragments, fault gouge, and pseudotachylyte veins. The ﬁgure shows at least 14 thin gouge layers interlayered with 4 pseudotachylyte veins within a 13 cm wide hand specimen (Fig. 10.8). Sharp boundaries between the fault gouge and pseudotachylyte veins are one of the characteristic features of these kinds of interlayered veins (Fig. 10.7). The pseudotachylyte veins are generally dark, compact, and aphanitic; they are harder than the related fault gouge layers and contain small numbers of ﬁnegrained host rock fragments that can be recognized with the naked eye. The gouge layers are easily distinguished from the dark and aphanitic pseudotachylyte veins on the basis of color and hardness. Dark fragments of pseudotachylyte are involved within the fault gouge layers (Figs. 10.7 and 10.8). Individual pseudotachylyte veins are generally less than 1 cm in width, and little or no mixing is observed between diﬀerent layers. A foliation is well developed within both the pseudotachylyte and fault gouge layers; the foliation

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Fig. 10.8. Sketch of a polished section through a hand sample collected from the Nojima Fault, Japan. The pseudotachylyte veins are interlayered with fault gouge and fault breccia zones. Numerous pseudotachylyte fragments are involved in the interlayered fault gouge layers (After Lin et al. 2001d)

is deﬁned by layers of contrasting color, the preferred orientation of visible fragments, and microscale shear bands (Figs. 10.7 and 10.8). 10.2.4 Crack-Fill Veins Fault-fracture (crack) networks are important structural features within hydrocarbon systems, as they act as ﬂuid-ﬂow conduits during hydrocarbon migration and/or reservoir production. Such networks can change in character from zones of high to low permeability due to cementation and/or pore collapse (e.g., Antonellini and Aydin 1995; Caine et al. 1996). The permeability of fault-fractures also varies as a consequence of seismic deformation (Sibson 1990). For this reason, the veinlet material that ﬁlls fault-fracture networks is also considered to be a kind of fossil earthquake, representing repeated seismic faulting events within seismogenic fault zones (Lin et al. 2001d, 2003c; Uda et al. 2001). Typical fracture (crack)-ﬁll network veins, composed mainly of ﬁne-grained material and calcite crystals, have been reported from both surface outcrops

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of the Nojima Fault and drill cores through the fault (Lin et al. 2001d, 2003c, 2007; Uda et al. 2001). A program of drilling through the Nojima Fault was conducted soon after the 1995 Mw 7.2 Kobe earthquake. The drilling was jointly undertaken by the Geological Survey of Japan, the Nation Institute for Earth Science and Disaster Prevention of Japan, and the Disaster Prevention Research Institute, Kyoto University. As deep drill cores are free of the physical and chemical weathering that aﬀects rocks close to the surface, the project provided the rare opportunity to study a fault zone at depth, soon after seismic movement. The structural features of the crack-ﬁll veins described in this section were mainly observed from two drill cores of 500 m and 1800 m in length (depth) that penetrated the Nojima Fault zone in the northern part of Awaji Island, Japan (Fig. 10.9). Most of the observed fractures are ﬁlled by ﬁne-grained material, clasts of the host rock, and calcite crystals. The fractures are developed within both fault-fracture networks and undeformed granitic country rocks at depths of up to >1800 m (Lin et al. 2003c). The crack-ﬁll material is generally unconsolidated to weakly consolidated and is grayish-white to gray-brown and dark brown in color; it has the appearance of clay–mud or wet soil (Fig. 10.10). These veins vary in width from several microns to several centimeters, with most being 1–5 mm. Numerous yellowish-brown hairline cracks are observed on polished sections of core samples (Fig. 10.11). The boundaries between the crack-ﬁll veins and the granitic host rocks are generally sharp, lacking any signs of a shearingrelated foliation. The above structural features suggest that the cracks are

Fig. 10.9. Simpliﬁed geological map of northern Awaji Island, Japan, showing the locations of the 500 m, 800 m, and 1800 m drill holes (black squares). The crosssection shows the 500 m and 1800 m drill holes through the Nojima Fault (Modiﬁed from Mizuno et al. 1990; Rgafj 1991; Lin and Uda 1995; Lin et al. 2001d, 2007)

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Fig. 10.10. Crack-ﬁll material within drill core at depths of 1344.5 m (a) and 1522.0 m (b) from the Nojima Fault, Japan. The crack-ﬁll material is unconsolidated. c (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

typical extensional fractures that act as ﬂuid conduits and that the crack-ﬁll veins did not form by in situ shearing along the cracks; rather, they formed via the deposition of ﬁne-grained material that was previously dissolved within or ﬂoating within an externally derived ﬂowing ﬂuid. The fault-fracture networks also contain calcite veins that are whitegray to brownish-gray in color; the veins are several millimeters to 2 cm in width (Fig. 10.12). Numerous yellow-brown to red-brown open cracks are also observed, being similar in appearance to the crack-ﬁll veins. There are no recognizable fault-shearing-related textures within the granitic rocks

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Fig. 10.11. Photographs of typical crack networks observed in polished sections of drill core recovered from depths of 1695.10 m (a), 1692.90 m (b), 1694.60 m (c), and 1605.70 m (d) from the Nojima Fault, Japan. (a–c): The fractures are generally ﬁlled by brown ﬁne-grained material. (d): The host granite is partially iron-oxidized c around the fractures. (After Lin et al. 2007). 2007, with kind permission from Elsevier Science Ltd

adjacent to the cracks, although hairline cracks are observed. This suggests that the extensional cracks developed by dilatation under tensional stress and were subsequently oxidized by circulating oxygen-rich ﬂuids (Lin et al. 2003c).

10.3 Petrologic Characteristics of Veinlet Cataclastic Rocks 10.3.1 Microstructures of Veinlet Cataclastic Rocks Crushing-origin Pseudotachylyte Crushing-origin pseudotachylyte veins consist mainly of ﬁne-grained material that has the optical characteristics of glass, glassy material that is generally too ﬁne-grained to observe under the microscope, and randomly oriented fragments of quartz and feldspar (Fig. 10.13). SEM images reveal that the

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Fig. 10.12. Calcite and crack-ﬁll veins at depths of 535.1 m (a) and 1334.8 m (b) within the 1800 m drill core, Nojima Fault, Japan. Note the coexistence of ﬁnegrained brown crack-ﬁll material and white calcite veins. (After Lin et al. 2003c). c 2007, with kind permission from Elsevier Science Ltd

pseudotachylyte matrix is mostly composed of ﬁne-grained angular fragments and minor sub-angular to sub-rounded fragments similar to those found in interlayered fault-gouge veins and host cataclastic rocks (Fig. 10.14). These fragments vary in size from several nanometers to several millimeters, with most being 1–100 μm. Distinct foliations and layering are observed in the interlayered pseudotachylyte and fault gouge veins; these structures are deﬁned by gray, brown, dark-brown and dark layers and fragments oriented subparallel to the boundaries between diﬀerent layers (Figs. 10.14 and 10.15). Overprinting structures can also be observed under the microscope. These include younger veins cutting older veins and the inclusion of numerous fragments of preexisting pseudotachylyte within younger veins (Figs. 10.15 and 10.16). These overprinting relationships demonstrate that multiple pseudotachylyte-forming

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Fig. 10.13. Photomicrograph of crushing-origin pseudotachylyte vein (Pt) bounded by granitic cataclasite within the Iida–Matsukawa Fault zone, Japan. Planec polarized light (After Lin 1996). 2007, with kind permission from Elsevier Science Ltd

Fig. 10.14. SEM-EBS (electron back-scatter ) image showing interlayered pseudotachylyte (Pt) and fault gouge veins from the Nojima Fault, Japan

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Fig. 10.15. Photomicrograph showing microstructures of the interlayered pseudotachylyte (Pt) and fault gouge veins in the Nojima Fault, Japan. Note that numerous pseudotachylyte fragments are involved in the fault gouge veins. Plane-polarized light

events occurred within individual fault zones, as is also indicated by overprinting relationships between melt-origin pseudotachylyte veins. The interfaces between the pseudotachylyte veins and host cataclasite or fault gouge layers are generally sharp, but locally they are obscure, even when viewed under the microscope. Biotite fragments present within the host granitic cataclastic rocks are not observed in the pseudotachylyte and fault gouge veins when viewed under the microscope. TEM analyses of pseudotachylyte veins from the Iida–Matsukawa Fault, Japan, reveal that the ﬁne-grained matrix with the veins contains little or no amorphous material and a lack of distinct melting textures, although it contains numerous ultra-ﬁne crystalline fragments (Ozawa and Takizawa, 2007). TEM analysis also reveals that the nanometer-sized fragments within the

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Fig. 10.16. Photomicrographs of overprinting crushing-origin pseudotachylyte veins from the Iida–Matsukawa Fault, Japan. (a) A younger pseudotachylyte vein (Y-Pt) cuts across an older vein (Old-Pt) and cataclasite. (b) Fragments of pre-existing pseudotachylyte within an overprinting vein (Pt) and fault microbreccia zone. Planec polarized light. (After Lin 1996). 2007, with kind permission from Elsevier Science Ltd

ﬁne-grained matrix are generally angular to subangular, although the superimposition of diﬀuse ring patterns associated with biotite spacings and diﬀraction spots limit the analysis of the ﬁne-grained matrix (Ozawa and Takizawa, 2007). These micron- to nanometer-scale structures indicate the existence of a small amount of non-crystalline (amorphous) material within

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crushing-origin pseudotachylyte veins; this material probably formed by strong comminution, without melting, during seismic faulting (see Sect. 10.4 below for further discussion on this topic). Fault Gouge Microstructurally, fault-gouge injection veins observed along the Iida– Matsukawa Fault, Japan, have features in common with the pseudotachylyte injection veins that occur within the same outcrops (Figs. 10.14 and 10.15). A foliation and layering deﬁned by gray, brown, dark-brown, and black layers are also observed in the fault gouge layers interlayered with fault veins and pseudotachylyte veins; these structures are not observed within injection fault-gouge veins (Figs. 10.14 and 10.15). These fault veins are yellowishbrown to brown in color when viewed under plane polarized light and dark to dark brown when viewed under crossed polarized light. They are mainly composed of ﬁne-grained angular fragments that are generally larger than 2– 3 μm in size (Fig. 10.17). Grain-size analyses of the fragments reveal similar size-distribution patterns for both injection and fault veins of pseudotachylyte and fault gouge, but there is a distinctly diﬀerent distribution pattern between the host cataclasitic rock and those of the pseudotachylyte and gouge veins (Fig. 7.8). These similarities in the size-distribution patters of the injection veins and fault veins of pseudotachylyte and fault gouge may indicate that they be formed coevally during seismic faulting (see Sect. 10.4).

Fig. 10.17. SEM-EBS image showing the microstructurs of fault gouge from the Iida–Matsukawa Fault, Central Japan

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Crack-ﬁll Veins Most of the minerals and clasts that inﬁll the intensively fractured area are altered and iron-oxidized, and most are cemented by calcite and ﬁne-grained material that contains numerous fragments of the wall rock (Fig. 10.18). Zoned structures are locally developed within the fracture-ﬁll veins, which are characterized by ﬁne-grained clasts of the granitic host rock and inter-zoned veins of calcite (Fig. 10.19). Fragments of calcite and ﬁne-grained clasts occur within these zoned veins. The veins are in part cut and deformed by faults and cemented by younger calcite veins and other ﬁne-grained material. Zoning structures within the crack-ﬁll calcite veins are characterized by inter-zoning layers of calcite and ﬁne-grained clasts of the host granite; the diﬀerent layers

Fig. 10.18. Photomicrographs of inﬁlled fault-fracture veins from the Nojima Fault, Japan. (a and b): Cracks ﬁlled by calcite veins (Ca) within drill core from a depth of 1695.10 m. (c): Zoned fracture-ﬁll veins composed of ﬁne-grained carbonate and clasts of host granite; depth, 1692.90 m. Ca1–Ca3: crack-ﬁll calcite veins. F-V: crackﬁll vein of ﬁne-grained material. Ca-clast: calcite clasts that pre-dated the calcite veins and were fracture-brecciated during later brittle deformation. (d): Crack-ﬁll calcite veins (Ca) observed at the same depth as those shown in (c) are cut by late-stage fracture-ﬁll veins (F–V) that consist of ﬁne-grained material including calcite and clasts of the host granite. Crossed polarized light. (Modiﬁed from Lin et al. 2007)

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Fig. 10.19. Photograph (a) and photomicrographs (b–d) of void cavities within a calcite vein and zoning within calcite veins at depths of 384.10 m (a–b) and 463.0 m (c–d) within the 1800 drill core, Nojima Fault, Japan. Note that the open crack is not completely ﬁlled by ﬁne-grained and carbonate material; rather, it occurs as an elongate lens-shaped cave structure within the calcite vein (c–d). Zoning within the calcite veins is characterized by colored laminations oriented parallel to the cavity wall (b–d). (b, d): Crossed polarized light, (c): plane polarized light. (Modiﬁed from c Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

are distinguished by their contrasting color (Fig. 10.19). These zoned layers are oriented subparallel to the walls of calcite veins, which are considered to have formed during multiple stages of ﬂuid ﬂow (Lin et al. 2003c). Locally, unﬁlled cavities act as active conduits of ﬂuid ﬂow (Fig. 10.19a, b). When observed in hand samples, the boundaries between crack-ﬁll veins and host granite are generally sharp. The above textural features indicate that repeated episodes of vein formation and deformation occurred within the Nojima Fault zone. 10.3.2 Powder X-ray Diﬀraction Analysis of Veinlet Material Veinlet cataclastic rocks such as fault gouge are generally composed of very ﬁne-grained material that is diﬃcult to identify under the optical microscope. Accordingly, powder X-ray diﬀraction is commonly used to identify the mineral types within veinlet material; these data are useful in comparing the

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petrologic properties of the veins and their host rocks and in assessing their origin and deformation mechanisms. Crushing-origin Pseudotachylyte and Fault Gouge Veins Powder X-ray diﬀraction data reveal that the crushing-origin pseudotachylytes described above are mainly composed of crystalline material with little or no non-crystalline material (Lin et al. 1994; Lin 1996, 1997a; Kano et al. 2004). X-ray diﬀraction spectra for samples from the Iida–Matsukawa Fault (Fig. 10.20) and the Nojima Fault, Japan (Fig. 10.21), show similar diﬀraction patterns among crushing-origin fault veins, injection veins of pseudotachylyte, fault gouge and injection gouge veins, and the host granitic cataclasites. For granite-hosted pseudotachylyte veins, it is commonly found that mica peaks such as biotite can only be recognized within the host granitic rocks and that the peaks of minor clay mineral such as chlorite and montmorillonite, which may have formed from the alteration of mica, are only recognized in the pseudotachylyte and gouge spectra (Figs. 10.20 and 10.21). TEM-derived lattice fringe images reveal that structural changes to amorphous-like material occur within deformed mica fragments of several tens of nanometers in size; these changes are considered to reﬂect stacking rather than melting (Ozawa and Takizawa 2007). The powder X-ray diﬀraction spectra of such samples (Figs. 10.20 and 10.21) do not possess the wide diﬀraction bands of non-crystalline material at low diﬀraction angles (2θ) of 12–40˚ such as those found in analyses of glassy pseudotachylyte (Fig. 5.16). Instead, the diﬀraction patterns reveal that the rock-forming minerals within crushing-origin pseudotachylyte veins are the same as those of the fault gouge veins and related wall rocks (Figs. 10.20 and 10.21). This indicates that the material in the cataclastic veins within the Nojima Fault zone originated from the granitic host rocks. Crack-ﬁll Veins Networks of crack-ﬁll veins are generally heterogeneous, even within individual veins; therefore, to compare the vein-forming minerals with those of the host rocks, it is necessary to analyze a large number of vein samples. In previous studies, more than 60 samples of crack-ﬁll veins (Table 10.1) collected from both outcrops and drill core from varying depths along the Nojima Fault were analyzed via powder X-ray diﬀraction (Lin et al. 2001d, 2003c, 2007; Uda et al. 2001). The obtained diﬀraction spectra reveal that the ﬁne-grained crack-ﬁll material is mainly composed of laumontite, siderite, and calcite, while typical granite-forming minerals such as quartz, feldspar, and biotite are also included in the veins as ﬁne-grained clasts (Fig. 10.22; Table 10.1). The mineral assemblages identiﬁed from the diﬀraction spectra indicate an external source for the ﬁne-grained crack-ﬁll material and calcite veins, probably related to seismic faulting (see Sect. 10.4 below).

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10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

Fig. 10.20. Powder X-ray diﬀraction spectra for pseudotachylyte veins and granitic host rock from the Iida–Matsukawa Fault, Japan. (a): standard sample of quartz, (b): granitic host rock, (c): fault vein of pseudotachylyte, (d): injection vein of pseudotachylyte. Qz: quartz; Pl: plagioclase; Bi: biotite; Mo: montmorillonite; ch: chloc rite. (Modiﬁed from Lin 1996). 2007, with kind permission from Elsevier Science Ltd

Fig. 10.21. Powder X-ray diﬀraction spectra for pseudotachylyte veins (a and b), fault gouge (c–e), and granitic host rock (f) from the Nojima Fault, Japan. (a and b): interlayered pseudotachylyte veins, (c and d): interlayered fault gouge zones and pseudotachylyte veins, (e): gouge within drill core at a depth of 388.42 m, (f): granitic host rock. Qz: quartz; Pl: plagioclase; Bi: biotite; Sm: smectite; Sd: siderite: (Modiﬁed from Lin et al. 2001d)

Depth (m)

103.50 113.15 215.10 316.00 422.90 533.20 786.85 787.60 788.75 790.20 790.20 901.40 1276.80 1279.20 1284.20 1288.70 1294.80 1299.40 1303.55 1308.70 1314.40 1334.80

Sample no.

1 3 4 5 7 8 11 12 13 14 15 17 20 21 22 23 24 25 26 27 28 29

Pla. ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ –

Qz

◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ – ◦ ◦ ◦ – – – – – – – – – – – ◦ – – – – – – – – –

Ort. – ◦ – – ◦ – – – – – – – – – – – – – –

Bio. ◦ – ◦ – ◦ – – – – – – – – – – – – – –

Horn.

– – – – – – – – – – – – – – –

– – – ◦

Chl. ◦ – – – – – – – – – – – – – – – – – – – – –

Smec. – – – – – – – – – – – – ◦ – ◦ – – – – – ◦

Lau. – – ◦ – – – – – – – – – – – – –

Sti.

– – – – – – – – – – ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦

Cal.

– – ◦ – – ◦ – – ◦ ◦ – ◦ ◦ ◦ ◦

Sid.

Table 10.1. Mineral assemblages of clay-like crack-ﬁll material identiﬁed by power X-ray diﬀraction analysis. (Modiﬁed from Lin c et al. 2003c). 2007, with kind permission from Elsevier Science Ltd 10.3 Petrologic Characteristics of Veinlet Cataclastic Rocks 247

Depth (m)

1390.10 1403.35 1416.30 1428.80 1437.00 1454.80 1469.25 1477.85 1495.95 1505.85 1507.20 1512.85 1518.30 1522.00 1528.30 1550.15 1555.30 1555.90 1567.75 1567.75 1574.15 1578.25 1597.10

Sample no.

30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52

Pla. ◦ ◦ – ◦ ◦ – ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦

Qz

◦ ◦ ◦ ◦ ◦ ◦ – ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦ ◦

– – – – – – – – – – – – – – – – – –

Ort. – – – – – – – – – – – – – – – – – – – – – –

Bio. – – – – – – – – – – – – – – – – – – – – – – –

Horn.

Table 10.1. Continued

– – – – – – – – – – – – – – – – – – – – – –

Chl. – – – – – – – – – – – – – – – – – – – – – – –

Smec. – – – – – – – ◦ – ◦ – – – – ◦ – – – – – – – –

Lau. – – – – – – – – – – – – – – – – – – – – – – –

Sti. – ◦ – ◦ ◦ – ◦ ◦ – – – ◦ – ◦ – ◦ – – – ◦ ◦ –

Cal.

– ◦ ◦ ◦ ◦ ◦ – – ◦ ◦ ◦ – – ◦ – – ◦

Sid.

248 10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

53 1605.10 ◦ – – – – – – ◦ ◦ 54 1624.15 ◦ ◦ – – – – – – – ◦ 55 1627.25 ◦ ◦ – – – – – – – ◦ 56 1635.45 ◦ – – – – – – – – 57 1650.55 ◦ ◦ – – – – – – – – 58 1658.35 ◦ – – – – – – – – ◦ 59 1673.15 ◦ ◦ – – – – – – – – – 60 1700.15 ◦ ◦ – – – – – – ◦ 61 1708.00 ◦ – – – – – – – – – Qz: quartz, Pla: plagioclace, Ort: orthoclase, Bio: biotite, Horn: hornblende, Chl: chrolite, Smec: smectite, Lau: laumontite, Sti: Stilbite, Cal: calcite, and Sid: siderite. : strong crystal peak, ◦ : intermediate crystal peak, : weak crystal peak, –: no crystal peak or the crystal peak is too weak to be recognized.

10.3 Petrologic Characteristics of Veinlet Cataclastic Rocks 249

250

10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

Fig. 10.22. Powder X-ray diﬀraction spectra for ﬁne-grained crack-ﬁll material taken from the granitic host rock (a), bentonite (b), fault gouge collected from a fault outcrop (c), 500 m drill core (d), and 1800 m drill core (e–g), Nojima Fault, Japan. Qz: quartz; Pl: plagioclase; Bi: biotite; Ca: calcite; Sm: smectite; Sd: siderite; Lmt: laumontite. The bentonite sample was used to prevent collapse of the drill hole during drilling; its diﬀraction peaks diﬀer to those of the pseudotachylyte veins, fault gouge zones, crack-ﬁll veins, and granitic host rock. (Modiﬁed from Lin et al. 2003c). c 2007, with kind permission from Elsevier Science Ltd

10.3.3 Chemical Composition Data and Isotope Analyses Chemical and isotopic analyses are generally used to infer the origin of ﬂuid within fault-related fracture-ﬁll veins (Kerrich et al. 1984; Janssen et al. 1997; Kharaka et al. 1999; Lin et al. 2003c). The results of XRF analyses of samples collected from the Nojima Fault, Japan, reveal that the rock-forming mineral assemblages of crushing-origin pseudotachylyte injection veins are similar to those of fault gouge veins and the granitic host rocks (Table 10.1). This similarity indicates that both the crushing-origin pseudotachylyte and the fault gouge veins originated from the same granitic host rocks.

10.3 Petrologic Characteristics of Veinlet Cataclastic Rocks

251

Fig. 10.23. δ 18 O values versus depth for the 1800 m drill core, Nojima Fault, Japan. The data of Ueda et al. (1999) were obtained from the Hirabayashi core (see c Fig. 10.9 for locality). (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

Fig. 10.24. δ 13 C values versus depth for the 1800 m drill core, Nojima Fault, Japan. c (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

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10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

Isotope analyses of calcite-rich crack-ﬁll and calcite veins extracted from various depths within drill cores from the Nojima Fault reveal that the δ 18 O and δ 13 C values of carbonate samples vary from 21.0%◦ to 27.7%◦ and from –4.6%◦ to –18.5%◦ , respectively, and that zoned calcite samples have diﬀerent values that range from 19.3%◦ to 24.8%◦ (δ 18 O) and –5.9%◦ to –11.3%◦ (δ 13 C), respectively (Figs. 10.23 and 10.24; Table 10.2). Based on these results, it is inferred that the calcite zones with contrasting isotopic compositions formed from diﬀerent source ﬂuids derived from surface waters (Lin et al. 2003c; see Sect. 10.4 below).

10.3.4 Age Data for Crack-ﬁll Veins The presence of zoned crack-ﬁll veins that contain carbonate and unconsolidated ﬁne-grained material indicates that they formed over the course of multiple seismic-related events in recent geological time, as discussed below. It is therefore possible to use the 14 C ages of samples of carbonate material and calcite veins to determine the timing of vein-forming events. In such cases, samples of pure calcite veins are generally analyzed. Lin et al. (2003c) reported that the 14 C ages of fracture-ﬁll calcite veins along the Nojima Fault range from 35,000 to 58,430 years B.P. (Fig. 10.25; Table 10.3). These results demonstrate that at least some of the crack-ﬁll veins formed during the Late Pleistocene, possibly related to recent seismic events along active faults (see Sect. 10.4 for more details).

Table 10.2. Isotope compositions of carbonate samples from the 1800 m core, c Nojima Fault, Japan. (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd 13

Sample No.

Sample depth (m)

δ

1 2 3 4 5 6 7 8 9 10 11 12 13 14

348.10 786.85 1267.70 1294.80 1299.40 1323.60 1340.05 1458.50 1477.85 1484.40 1522.00 1575.85 1624.15 1700.15

–8.372±0.049 –4.575±0.021 –14.458±0.014 –12.479±0.017 –10.890±0.018 –11.346±0.036 –14.925±0.045 –8.456±0.020 –9.538±0.042 –8.539±0.030 –9.158±0.027 –12.831±0.024 –14.637±0.040 –18.487±0.015

C (%◦ PDB)

δ

18

O (%◦ SMOW)

22.230±0.050 20.855±0.034 27.668±0.051 20.321±0.057 20.543±0.039 20.990±0.057 25.659±0.038 21.896±0.051 22.297±0.067 21.839±0.050 20.899±0.063 22.659±0.032 22.832±0.049 24.128±0.022

10.4 Discussion on the Formation Mechanisms of Veinlet Cataclastic Rocks

253

Fig. 10.25. 14 C ages of carbonate veins plotted versus depth for the 1800 m core, Nojima Fault, Japan (see Table 10.3 for detail depths of samples). (After Lin c et al. 2003c). 2007, with kind permission from Elsevier Science Ltd Table 10.3. 14 C dates for carbonate samples from the 1800 m core, Nojima Fault, c Japan. (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd Sample no.

Depth (m)

1 2 3 4 5 6 7 8 9 10

258.03 389.40 1267.75 1318.40 1323.60 1340.10 1484.40 1575.90 1624.15 1665.20

14

C age (years B.P.)

38,132±318 52,94±1,279 58,431±1,535 36,510±264 34,955±232 45,660±573 54,949±1,491 49,787±625 50,872±991 39,241±367

10.4 Discussion on the Formation Mechanisms of Veinlet Cataclastic Rocks 10.4.1 Formation Mechanism of Amorphous Material Within Veinlet Cataclastic Rocks Crushing-origin pseudotachylyte veins are similar to melt-origin pseudotachylyte veins in terms of their dark color, dense and aphanitic appearance, and occurrence as both simple veins and irregular networks intruded into country

254

10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

rocks. Despite these similarities, the crushing-origin pseudotachylyte veins described in this chapter are devoid of any evidence of frictional melting such as variously shaped microlites, vesicles, amygdules, and glass material within the matrix. These petrological and textural characteristics indicate that both fault gouge and crushing-origin pseudotachylyte veins formed mainly by crushing, without frictional melting. Pocked glassy or amorphous material can sometimes be observed in fault gouge and crushing-origin pseudotachylyte veins; this is considered to have formed by comminution rather than frictional melting. This interpretation is made because glassy or amorphous material that forms via the cooling of a melt generally has sharp boundaries and uniform contrast in bright images; in contrast, these features are not observed in the amorphous material found within pseudotachylyte veins from along the Iida–Matsukawa Fault (Ozawa and Takizawa 2007). TEM images of Iida–Matsukawa pseudotachylyte reveal that biotite fragments of just several tens of nanometers in size are fragmented and deformed, with lattice spacings of 1 nm (Ozawa and Takizawa 2007). These biotite fragments show amorphous TEM diﬀraction features that are explained by mechano-chemical eﬀects associated with the intense fragmentation and deformation that accompanied comminution (Ozawa and Takizawa 2007). Natural non-crystalline materials such as glass are known to form not only by the rapid cooling of melts but also by alteration (e.g., Henley and Ellis 1983), rapid comminution (Kieﬀer et al. 1976), and solid-state shocking associated with impact events (e.g., Chao 1968; Christie et al. 1973). Experimental results show that amorphous materials can also be produced within fault gouge, even during events with small displacements (e.g., Friedman et al. 1974; Logan et al. 1981; Yund et al. 1990). These studies of natural and experimental fault-related materials reveal that comminution associated with rapid faulting is probably a major formation mechanism of the small, pocked, amorphous material that is found in fault gouge and crushing-origin pseudotachylyte veins. 10.4.2 Coseismic Fluidization of Fine-grained Material Within Fault Zones The intrusive modes and petrologic features of crushing-origin pseudotachylyte and vein networks of fault gouge reveal they formed via the injection of ﬁne-grained material derived from the source generation (fault) plane; this material contains little or no amorphous or glassy material. The question is whether these injection veins formed by i) the rapid spray-like intrusion of ﬁne-grained fragments during seismic slip, ii) the slow deposition of clay minerals or ﬁne-grained fragments transported by hydrothermal ﬂuids, or iii) slow intrusion along fractures during aseismic periods. If the crushing-origin pseudotachylyte and fault-gouge injection veins described above formed by slow ﬂow during aseismic periods, it would be diﬃcult

10.4 Discussion on the Formation Mechanisms of Veinlet Cataclastic Rocks

255

to intrude ﬁne-grained solid fragments into sub-millimeter to centimeter-scale cracks over a distance of several meters or more from the source fault plane without the presence of water. If the injection veins formed via the slow ﬂow of groundwater, we would expect to ﬁnd: i) contrasting chemical compositions among the injection veins, fault veins, and granitic host rock due to dissolution-related sorting of minerals within the ﬂuid, and ii) greater amounts of ﬁne-grained fragments in the injection veins than in the fault veins because clay materials and ﬁne-grained materials would have been sorted and preferentially transported into narrow cracks (<1 cm in width) located 1–2 m from the source fault plane. However, the chemical compositions, grain-size data (see Fig. 7.8), and X-ray diﬀraction data described above show no indication of this type of sorting. These data do, however, indicate a common source material for the fault veins and injection veins of pseudotachylyte and fault gouge (i.e. the host rock); this ﬁnding supports the proposal that the injection veins of both crushing-origin pseudotachylyte and fault gouge formed by the rapid injection of ﬁne-grained fragments during episodes of seismic slip. The mechanism of ﬂuidization has been proposed to explain the rapid injection of ﬁne-grained material within fault zones during seismic faulting (Lin 1996, 1997a, 2006b). The Glossary of Geology (Bates and Jackson 1980) deﬁnes ﬂuidization as “the mixing process of the whole ﬂows like a liquid, e.g. the formation of an ash ﬂow or nuee ardente during a volcanic eruption.” The term was ﬁrst used in explaining geological phenomena by Reynolds (1954). Although the term ﬂuidization is applied speciﬁcally to a gas–solid system, it is also applicable to a suspension of solid particles within an upwardﬂowing stream of liquid that has a lower density than that of the particles. Geological examples of intrusive ﬂuidized systems are characterized by network veins within dikes, breccias in which many of the fragments derived from the adjacent wall rocks are rounded as if sandblasted, and mechanical hybrids (Reynolds 1954). Other geological examples of intrusive veins are associated with seismic faulting, such as melt-origin pseudotachylyte veins (Lin 1991, 1994a), and liquefaction veins that form during large earthquakes (e.g., Lin 1997c, 2006a); these also occur as simple veins and vein networks injected into the host rocks. It has also been observed that typical melt-origin pseudotachylyte veins such as the Fuyun and Outer Hebrides pseudotachylytes contain 60–70 Vol% fragments of various sizes (Figs. 5.3, 7.3 and 7.4). This ﬁnding suggests that the ﬁne-grained fragments mixed with the melt were injected into the void spaces generated during seismic faulting by the rapid intrusion-like spraying of a gas–solid–liquid system, much like a pyroclastic ﬂow rather than the slow ﬂow of a melt liquid. The substantial cavities into which this material is injected may form in association with seismic slip in strong rocks at depths greater than several kilometers (Sibson,1986). The cavities, which act as transitory low-pressure

256

10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

channels, are suitable sites for the rapid passage of ﬂuidized particles associated with rapidly developed ﬂuid-pressure gradients generated in dilational jogs during rupture arrest (Sibson 1986). It is important to realize that the bubble phase of gas–solid systems has no counterpart in liquid–solid systems; consequently, turbulently expanded beds are speciﬁc to gas–solid systems. This fact is of importance to geologists because the recognition of turbulent expansion and generation of rock fragments that remain close to the source rock, together with a lack of sorting of the fragments and the possible presence of druses indicates that the active medium was a gas rather than a liquid. A similar process to that of the ﬂuidization of solid particles by gas is the method of painting by spraying (Reynolds 1954). In summary, it is possible that injection veins of both melting- and crushingorigin pseudotachylyte and fault gouge form via the seismic ﬂuidization of ﬁne-grained materials in a gas–solid–liquid system within seismogenic fault zones during large earthquakes (Lin 1996, 1997a, 2006b). 10.4.3 Repeated Events of Seismic Slip The contacts between melt-origin pseudotachylyte and cataclastic veinlet material are sometimes gradational; this demonstrates that some of the veins composed of cataclastic material also formed in association with seismic faulting and might therefore represent paleoseismic rupture zones, possibly those of very large earthquakes, with or without accompanying molten material (Rowe et al. 2005). The layered structures with sharp boundaries that occur between fault gouge and pseudotachylyte veins along the Nojima Fault, Japan, are considered to potentially record the mode of seismic faulting during the formation of the gouge and veins (Shigetomi and Lin 1999). As described above, a 13cm-wide fault gouge–pseudotachylyte zone from the Nojima Fault contains at least 18 distinct thin layers with sharp boundaries between them. Geological evidence indicates that the Nojima Fault is currently active as a seismic source fault and that seismic faulting events occurred repeatedly on this fault during the Holocene (Lin and Uda 1996); despite this, creep movement is not observed at the Earth surface above the Nojima Fault, not even following the 1995 Mw 7.2 Kobe earthquake (Lin 1996). It is suggested that each layer within the layered structure records at least one seismic faulting event, as it is diﬃcult to generate distinct multi-layers with sharp boundaries during a single seismic faulting event (Shigetomi and Lin 1999). The layering structures within fault gouge veins also indicate that seismic slip occurred within a narrow zone of less than several millimeters in width. Lin et al. (1998b, 2003c) reported that numerous small calcite veins that ﬁll microcracks as networks within the fault gouge zone of the Nojima Fault were disrupted within a narrow zone (several millimeters in width) along the master fault plane at the time of the distinct slip that accompanied the 1995 Mw 7.2 Kobe earthquake. Along this disrupted zone, the electron spin

10.4 Discussion on the Formation Mechanisms of Veinlet Cataclastic Rocks

257

resonance (ESR) signals decrease sharply across a zone of less than 3 mm in width (Fukuchi and Imai, 1998; Fukui, 2003); this is explained by annealing associated with frictional heating generated during seismic slip (Fukuchi and Imai, 1998). Likewise, analyses of the meso- and micro-scale structures of fault gouge layers developed along the Chelungpu Fault, Taiwan, which triggered the 1999 Mw 7.6 Chi-Chi earthquake, reveal that coseismic slip during the Chi-Chi earthquake was restricted to a narrow zone along the main fault plane of <2–3 mm in width (Lin et al. 2005a). The results of theoretical simulations (that assume a frictional fault plane with zero thickness) and geophysical observations also support the proposal that zones of coseismic frictional melting generated upon fault planes are generally no thicker than 1 cm (McKenzie and Brune 1972; Sibson 1973; Richards 1976; Cardwell et al. 1978). Laboratory experiments of unstable slip demonstrate that slip occurs in a narrow fault zone at upper crustal levels, even when in the presence of fault gouge (e.g., Engelder et al. 1975; Marone and Kilgore 1993; Beeler et al. 1996). High-velocity frictional experiments (e.g., Spray 1987, 1995; Lin and Shimamoto 1998) conducted under conditions that replicate seismic faulting at shallow depths reveal that frictional melting occurs in narrow zones of less than 2–3 mm in width. These frictional experiments were designed to generate frictional melting and sliding along a contact surface or a thin layer of simulated gouge. Based on the results of the above studies and the structural modes of veinlet cataclastic rocks, Lin et al. (2005a) concluded that seismic slip during moderate- to large-magnitude earthquakes occurs along fault zones that are narrower than 2–3 mm. It is therefore interpreted that the 13-cm-wide interlayered zone of gouge and crushing-origin pseudotachylyte observed within the Nojima Fault zone records at least 18 seismic events (Shigetomi and Lin 1999). 10.4.4 Repeated Coseismic Inﬁltration of Surface Water into Deep Fault Zones Source of Vein-forming Fluids Within Fault Zones A large proportion of the brittle deformation induced by an increment of slip on a master fault may occur well away from the principle slip surfaces (Sibson 1986). Indeed, coseismic surface ruptures, including numerous extensional cracks, were generated across a wide zone of up to 10 km during the 2001 Mw 7.8 Kunlun (Tibet, China) earthquake, even though the primary deformation was restricted to a narrow zone of <10 m along the pre-existing active fault zone (Lin et al. 2002a, 2003a; Lin and Nishikawa 2007). This indicates that during large seismic faulting events, networks of fractures and extensional cracks can be generated across a wide zone of 1–10 km about the main fault plane. Seismic faulting not only aﬀects the fault-fracture architecture, it also leads to major changes in permeability that in turn signiﬁcantly impact upon the magnitude and pattern of ﬂuid ﬂux and solute transport both within and

258

10 Crushing-Origin Pseudotachylyte and Veinlet Cataclastic Rocks

about fault zones. It is therefore proposed that the fault-fracture networks related to veinlet material along the Nojima Fault, Japan as described above, also form from repeated coseismic faulting (Lin et al. 2003c). Evidence for ﬂuid ﬂow and sealing along fault and fracture networks within the Nojima Fault is derived from three sources (Lin et al. 2003c). First, fractures are ﬁlled with carbonate cement, commonly with a banded pattern, and carbonate-ﬁlled veins are concentrated in areas close to faults. Second, iron-oxide alternation is present along most of the faults and veins. Thirdly, calcite-cemented ﬁne-grained clasts, which originated from the host rocks and were crushed by faulting in the core and damage zones, are deposited along the fractures. The lack of distinct shear fabrics adjacent to those fractures ﬁlled by ﬁne-grained clasts indicates that the ﬁne-grained material formed within fault zones and was subsequently transported by ﬂuid ﬂow. The Nojima Fault is characterized by networks of sub-faults, extensional fractures, and microcracks that are mostly ﬁlled by multi-stage calcite veins, calcitecemented breccia, ﬁne-grained clasts of the host rock, and, locally, iron-oxide alternation of the host rock (Lin et al. 2001d, 2003c; Uda et al. 2001). The presence of crack-ﬁll vein networks of calcite and ﬁne-grained carbonate material and their very nature indicates that they formed from multiple ﬂuid-ﬂow events that permeated the fault-fracture networks. Isotopic data and structural features indicate that the crack-ﬁll calcite veins and ﬁne-grained calcite-cemented material formed via the repeated inﬁltration of O2 - and CO2 -bearing surface waters, including meteoric water and seawater, during seismic faulting events of the Late Pleistocene and Holocene (Lin et al. 2003c). The presence of ﬁne-grained crack-ﬁll material, calcite veins, and yellow-brown cracks within the Nojima Fault indicates the circulation of ﬂuids rich in Ca, CO2 , and O2 . Chemical and isotopic analyses of carbonate veinlets within fault zones are commonly used to infer the origin of ﬂuids within fault zones (Kerrich et al. 1984; Janssen et al. 1997; Ueda et al. 1999; Kharaka et al. 1999). Figure 10.26 shows δ 18 O isotopic data for ﬂuids in chemical equilibrium with calcite at various depths within the 1800 m core drilled through the Nojima Fault; these values were calculated using the equilibrium equations proposed by O’neil et al. (l969) and for a geothermal gradient of 30◦ C/km (Yamano and Goto 1998). The δ 18 O values of the carbonate veins are consistent with those of seawater and typical Japanese meteoric water, but diﬀerent to that of mantle-derived water (Fig. 10.27). Similar δ 18 O values (SMOW, Standard Mean Ocean Water), ranging from 19 to 26, have also been reported for the North Anatolian Fault Zone, Turkey; these are interpreted to have originated from ﬂuids within adjacent limestone (Janssen et al. 1997). Negative δ 18 O values of –3 to –4 %0 recorded at 500–540◦C have been reported for ﬂuids within a fault shear zone at Lagoa Real, Brazil; these values are interpreted to reﬂect the dewatering of sedimentary-related brines that originated from an earlier recharge of Proterozoic sedimentary-basin aquifers by meteoric water (Kerrich et al. 1984).

10.4 Discussion on the Formation Mechanisms of Veinlet Cataclastic Rocks

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Fig. 10.26. Depth proﬁles of δ 18 O values for seawater (SMOW), meteoric water (data from Matsuo 1997) and δ 18 O values for carbonate veins analyzed in the present study for ﬂuids in chemical equilibrium with the carbonate material, using c temperatures measured within the drill hole. (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

Fig. 10.27. δ 13 C values (PDB) of typical mantle-derived carbon, marine plants, sedimentary rocks, atmospheric CO2 (data from Matsuo 1997), and crack-ﬁll carbonate veins measured by Lin et al. (2003c). The δ 13 C values of carbonate material are similar to those of atmospheric CO2 , marine plants, and sedimentary rocks, but diﬀer from those for mantle-derived carbon and deep-ocean carbonate. (After Lin c et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

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The δ 18 O-rich ﬂuid within the Nojima Fault Zone probably reﬂects the inﬁltration of surface water and progressive interaction between the inﬁltrating ﬂuid and partially altered granitic host rock. The δ 13 C values of carbonate material within the Nojima Fault Zone are similar to those of atmospheric CO2 , marine plants, and sedimentary rocks, but diﬀerent from those of mantlederived carbon and deep-ocean carbonate (Fig. 10.27). Marine-origin ﬂuids with δ 13 C values of –2 to 4.5 (PDB, a standard of comparison in determining the isotopic composition of carbon and oxygen) have been reported from carbonate veins developed within the North Anatolian Fault Zone (Janssen et al. 1997), indicating that the veins formed from ﬂuids that originated from seawater, possibly mixed with meteoric water. The diﬀerences in the mineral assemblages of the ﬁne-grained crack-ﬁll material, which includes calcite, and the fault gouge indicates that the crack-ﬁll material was not generated within the cracks during shearing of the granitic host rock. Siderite generally forms under conditions of low Eh (oxidation potential) (Garrels and Christ 1965). The groundwater in the area of the Nojima Fault is rich in HCO3 (Sato and Takahashi, 1997), and high concentrations of CO2 and H2 were detected in the 1800 m drill hole (Arai et al. 1998). This CO2 -rich ﬂuid is not mantle- or magma-derived, but probably formed via the chemical reaction of meteoric O2 and CO2 within the groundwater in the upper zone, above the stable groundwater level (Arai et al. 1998). It is therefore proposed that the ﬁne-grained crack-ﬁll material and brown-colored open cracks that contain siderite (FeCO3 ) and calcite (CaCO3 ) formed via the reaction of Ca and HCO3 or CO2 , which originated from surface waters (meteoric and seawater) that circulated within the fault zone. Cracks that form at shallow depths are generally weathered, eroded, and oxidized by oxygen-rich groundwater, including meteoric and seawater. The brown-colored open cracks observed at shallow levels around the stable groundwater level, which generally occurs at depths of <100 m in Japan, and which are typically ﬁlled by ﬁne-grained sediment and calcite material, generally formed by the oxidization of Fe because of the oxygen-rich nature of the groundwater that circulates within the fault zone (Uda et al. 1999). In general, cracks within deep fault zones remain unoxidized due to anoxic conditions at depth. The distribution depth of brown cracks is commonly used in hydrogeologic investigations as an indicator of the stable groundwater level; however, open brown-colored cracks are observed at all depths within the 1800 m core extracted from the Nojima Fault, even though the stable groundwater level at the drill site occurs at less than 6 m depth (Uda et al., 1999). This ﬁnding indicates that considerable volumes of oxygen-rich water were drawn into the fault zone to depths of >1800 m, thereby oxidizing the open cracks. Calcite material usually occurs in reactivated sections of cracks as a coating on earlier crack-ﬁll carbonate material. The δ 18 O and δ 13 C values of zoned carbonate veins indicate that the colored laminae formed from ﬂuids derived from varied sources, mainly from surface water (seawater and meteoric water)

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containing carbon-rich material. The presence of unconsolidated ﬁne-grained crack-ﬁll material suggests that at least some of the veins formed during repeated events that occurred during recent geologic time. The co-existence of unconsolidated clay-like crack-ﬁll material and consolidated calcite veins reveals that they formed from a multi-stage process of repeated inﬁltration events. The newly formed carbon material within ﬂowing water was probably contaminated by or mixed with pre-existing ﬂuids containing dead carbon before being deposited on the pre-existing carbonate material. It is impractical to separate the diﬀerent generations of carbonate material for 14 C dating; therefore, the 14 C age determined for the carbonate material is probably a mixed age that reﬂects multiple stages of vein formation; such ages are generally older than the age of the most recently deposited carbonate material (Table 10.3). It is proposed that the deposition of carbonate material in at least some of the open cracks took place repeatedly during the past 30–60 ka, primarily via the inﬁltration of meteoric water and seawater into the Nojima Fault zone (see below for details). The colored lamellae (zoning textures) within calcite veins also indicate an origin associated with multiple episodes of ﬂuid ﬂow, thereby supporting the proposal that the variable 14 C ages determined for the calcite veins represent repeated events of carbonate formation. Coseismic Inﬁltration of Surface Water Deep into the Active Fault Zone The accumulated evidence suggests that CO2 - and O2 -rich surface waters, including meteoric water and seawater, repeatedly inﬁltrated the Nojima Fault zone, Japan. The groundwater level measured within drill holes that penetrate the Nojima Fault zone is <6 m (Uda et al. 1999), while the depth of the sea ﬂoor around Awaji Island is <150 m (Lin and Uda 1996b). Thus, it is unlikely that groundwater containing meteoric water and seawater inﬁltrated to depths of >1800 m along the Nojima Fault solely due to seasonal ﬂuctuations in groundwater levels. The question remains as to how surface waters inﬁltrated deep into the fault zone. One consequence of the dilatancy/ﬂuid-diﬀusion mechanism that operates during shallow earthquakes is that considerable volumes of ﬂuid are rapidly redistributed in the crust following seismic faulting, as illustrated by outpourings (Sibson 1975) and lowered groundwater levels along fault traces (Osada et al. 1997; Touda et al. 1995; Huang et al. 1999). Fluctuations in groundwater levels during earthquake events have been reported previously (Touda et al. 1995; Huang et al. 1999). Outpourings and falling groundwater levels were described from many localities along the Nojima Fault zone during the 1995 Kobe earthquake (Touda et al. 1995); the groundwater level measured within boreholes 4 months after the earthquake was about 2 m lower than that measured before the earthquake (JSEG 1999). On the day before the 1999 Chi-Chi Mw 7.6 (Taiwan) earthquake, groundwater levels along the

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Fig. 10.28. Suction pumping model for a seismic fault. The mechanics of seismic faulting leads to vacuum suction pumping and a high groundwater potential that draws meteoric water and seawater deep into the fault zone. The groundwater spouts from the surface trace or inﬁltrates the fault zone due to changes in potential during large earthquakes (b). It may take a long time to restore the potential of the groundwater (c) to the state prior to the earthquake (a). Recurrent large-scale seismic faulting events lead to the repeated inﬁltration of meteoric water and seawater c deep into fault zones. (After Lin et al. 2003c). 2007, with kind permission from Elsevier Science Ltd

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footwall of the fault fell by up to 11 m along the 100 km Chelungpu surface rupture zone along the pre-existing Chelungpu Fault (Lin et al. 2001b); many outpourings were also reported (Huang et al. 1999). These observations indicate that ﬂuctuations in groundwater potential in and around fault zones result from crustal deformation during large-magnitude earthquakes. Given the above, a fault suction pumping model is proposed to explain the downward inﬁltration of surface waters deep into the fault zone to generate ﬁne-grained crack-ﬁll material and calcite veins and to oxidize and weather open cracks (Fig. 10.28; Lin et al. 2003c). In contrast to workings of the seismic pumping model of Sibson (1975), observations along the Nojima Fault zone reveal that the carbonate materials deposited deep within the fault zone originated from meteoric water and seawater. This suggests that a suction mechanism led to rapid changes in groundwater levels that in turn promoted the inﬁltration and circulation of meteoric water and seawater deep into the fault zone. The dilatancy/ﬂuid-diﬀusion model of Scholz et al. 1973) provides an explanation of the intermittent ﬂow of hydrothermal ﬂuids in and around fault zones, and suggests that seismic faulting acts as a pumping mechanism whereby individual earthquakes are capable of moving signiﬁcant quantities of ﬂuids from one crustal environment to another (Sibson 1975). It is also proposed that the mechanics of seismic faulting perform the role of a vacuum-like suction pump, leading to a high groundwater potential and drawing meteoric water and seawater deep into the fault zone (Fig. 10.28). Topographically, vertical coseismic displacements are a further important factor in terms of rapidly generating a high groundwater potential across a fault plane. A number of laboratory and ﬁeld studies demonstrate that seismogenic faulting is strongly aﬀected by the mechanical and chemical interaction of ﬂuids and rocks (Evans and Chester 1995; Tsunogai and Wakita 1995). These interactions are probably responsible for the discharge phenomena commonly observed within epicentral areas. Following a strong earthquake, fractured zones are restored to their prior strength and groundwater levels are restored to their usual stable level from the funnel-shaped pattern that developed during the earthquake (Fig. 10.28). Accordingly, during the long restoring process following a seismic event, O2 - and CO2 -bearing groundwater is drawn deep into the fault zone by vacuum pumping or a high groundwater potential, as shown in Fig. 10.28. In this way, recurring large-magnitude seismic faulting events potentially lead to the repeated inﬁltration of meteoric water and seawater deep into fault zones.

11 Landslide-related Pseudotachylyte

11.1 Introduction Natural glass is well known to be produced by a number of diﬀerent geological processes, including the rapid cooling of magma, seismic-related frictional melting upon a fault plane (as described in previous chapters), and impact-related shocking (e.g., Chao 1968; Christie et al. 1973). Molten material has also been reported from the glide planes beneath landslides (e.g., Marsh et al. 1985; Lin et al. 2001a). A well-known landslide-generated pseudotachylyte from the Langtang region of the Nepalese Himalayas (Marsh et al. 1985) was initially interpreted to represent frictional fusion along a tectonic thrust (Scott and Drever 1953). Scott and Drever (1953) termed this vesicular glass layer hyalomylonite because they considered that the formation of glass in this case was the end-product of typical mylonitization processes (or “ultra-mylonisation”). The authors deﬁned mylonitization as “Any rock, layer or matrix in which there has been a considerable amount of ﬁne microbrecciation under tectonically determined cataclastic condition”. Based on our current knowledge of fault rocks, such melt-origin glass is considered to form by frictional melting during rapid seismic faulting or gliding upon a slip plane over a short period of time ranging from several seconds to a minute; the origin of such rocks is not considered to be directly related to mylonitization. To avoid any potential confusion in nomenclature, the term pseudotachylyte, as deﬁned in Chap. 2, is herein used in place of hyalomylonite. Scott and Drever (1953) described the ﬁeld occurrence and petrologic characteristics of molten material within the Langtang Himalaya area and concluded that the glass formed via frictional heating upon a slip (glide) plane within crystalline rocks. A second example of landslide-origin molten material has been reported from the K¨ ofels landslide, Austria (Marsh et al. 1985); the volume and area of the landslide in this case are similar to those of the

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Langtang landslide. Both the Langtang and K¨ ofels landslide-generated pseudotachylytes are mainly composed of molten material that is characterized by typical melting textures such as vesicular glass, microlites, bubbles, ﬂow structures, and rounded and embayed fragments (Scott and Drever 1953; Masch et al. 1985). The most recent example of landslide-generated pseudotachylyte is that reported from the Chiufener-Shan landslide, which occurred during the 1999 Chi-Chi (Taiwan) Mw 7.6 earthquake, producing a thin millimeter-scale layer of molten material (pseudotachylyte) upon the glide plane of the landslide (Lin et al. 2001a). This is the ﬁrst discovery of molten material that formed in association with an earthquake-related landslide. These types of pseudotachylyte are considered to represent a type of “fossil earthquake” because large-scale landslides commonly occur in association with the strong groundmotions that accompany large earthquakes (Lin et al. 2001a). This chapter focuses on two representative examples of landslide-generated pseudotachylytes (from the Langtang Himalaya area, Nepal, and ChiufenerShan, Taiwan) and their conditions of formation.

11.2 Occurrences of Landslides and Related Pseudotachylytes 11.2.1 Langtang Himalaya Landslide and Related Pseudotachylyte The Langtang Himalaya landslide is located north of Kathmandu, Nepal, close to the border between China and Nepal (Fig. 11.1). The area in which the landslide occurred is bordered by the Langtang Glacier to the west, the Phrul Rangjen Glacier to the east, and the Langtang River to the south (Fig. 11.1). At this site, a pseudotachylyte vein can be observed within the boundary zone between the overlying basement rocks (northern side) and underlying alluvial deposits (southern side) in the frontal part of the landslide, close to the Langtang River (Figs. 11.2 and 11.3). Scott (1953) ﬁrst discovered this dark, vesicular, glass vein, which is accompanied by brecciated host rock, and suggested that it formed by frictional fusion during recent movement along the boundary zone, which was interpreted as a tectonic thrust. Subsequent ﬁeld investigations, however, reveal that the glass vein is more likely to have formed in association with the landslide itself (Masch and Preuss 1977; Masch 1979; Masch et al. 1985). A Landsat image and topographic map of the area reveal topographic features that are indicative of landslide-related landforms (Fig. 11.1). It is estimated that the landslide was displaced horizontally by about 2 km to the southwest, with the present volume of the landslide being ∼3 km3 , covering an area of 14 km2 (Fig. 11.1). Masch et al. (1985) documented a distinct diﬀerence in the deformation modes of the transported mass and autochthonous rocks. The transported rocks are mainly fault gouge, breccia, and fractured rocks, all derived

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Fig. 11.1. Landsat image showing the topographic features of the Langtang Himalaya landslide, Nepal (image sourced from Google). The dashed line indicates the extent of the landslide area, while the arrow shows the movement direction of the landslide

Fig. 11.2. Photograph of the glide plane (F, indicated by arrows) of the Langtang Himalaya landslide, exposed in the Langtang Valley, Nepal. Photograph courtesy K. Arita

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Fig. 11.3. Photographs of the Langtang pseudotachylyte vein (Pt). The vein varies in thickness from ∼2 to 20 cm along the glide surface. (b) Enlargement of the central area within (a). The hammer shown in (a) for scale is 35 cm long. The boundaries between the vein and the wall rock are generally sharp. Photograph courtesy K. Arita

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from migmatized biotite gneiss and schist (Scott and Drever 1953; Masch et al. 1985); in contrast, the autochthonous rocks are massive and relatively uniform layered biotite gneiss (Scott and Drever 1953) that grades upward into a sequence of interlayered biotite gneiss and granite of the Langtang migmatite zone (Arita et al. 1973). The glide surface of the landslide dips to the southwest at 15–30◦ (Fig. 11.2; Masch et al., 1985). The pseudotachylyte vein at the base of the landslide is developed upon the glide surface; it varies in thickness from several millimeters to ∼20 cm, with local pockets reaching 60 cm in thickness (Fig. 11.3; Masch et al. 1985). The vein is mainly composed of glassy material and numerous ﬁne-grained fragments of the host rock, giving the vein appearance of a fresh vesicular basalt dike (Fig. 11.3b). A polished section cut through a hand sample of the vein reveals vesicular textures (Fig. 11.4). Most of the vesicles are elongate (aspect ratios of 10:1) and aligned subparallel to ﬂow lamellae (Fig. 11.4b). The widths of the elongate vesicles are generally less than 2–3 mm, with lengths up to >2 cm. The volume percent of vesicles within the vein increases from ∼2 to 30 Vol% from the margins to the center of the vein, as observed in glass pseudotachylyte veins from the Fuyun Fault, China (see Chap. 4 for details). In contrast, the volume percentage of fragments increases from ∼2 Vol% in the center of the vein to >50 Vol% at the margins (Fig. 11.4c). Flow of the melt is indicated by ﬂow structures deﬁned by thin colored layers, trails of oriented fragments, and elongate vesicles. The visible fragments of the host rock are mostly angular to sub-angular, but rounded and irregularly embayed fragments are also observed. The ﬁeld occurrence of the vein and the nature of meso-scale textures indicate that the vein originated from a primary melt formed by frictional fusion within the narrow glide zone of the landslide. 11.2.2 Chiufener-Shan Landslide and Related Pseudotachylyte The Chi-Chi Mw 7.6 earthquake struck west-central Taiwan on 21 September 1999, resulting in 2300 deaths and widespread damage. The quake produced a 100 km-long coseismic surface rupture zone along the boundary between a mountain range and adjacent basins (Lin 2001b, c). Numerous landslides were generated by the strong ground motions associated with the earthquake, with as many as 2365 landslides occurring over an area of 14,300 ha (Wang et al. 2000). Large-scale landslides of up to 0.8–1.0 × 108 m3 in volume occurred at Chiufener-Shan, near the epicenter of the quake, and at Tsaoling, along the southern part of the coseismic surface rupture zone (Fig. 11.5). The area in which the Chiufener-Shan landslide occurred is cut by ﬁve major active faults: the Shuiliken, Tamoupu-Hsuangtung, Chelungpu, and Changhua Faults, which dips moderately to the east, and the Tuntzechiao Fault, which dips moderately to the northwest (Fig. 11.5). The Chelungpu

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Fig. 11.4. Photographs of a polished section cut through a hand sample of the Langtang pseudotachylyte. Vesicles in the central part of the vein are elongated subparallel to the vein margins. (b): Enlarged view of the central part of the vein. (c): Enlarged view of the vein margin, showing numerous light-gray fragments. Sample courtesy K. Arita

Fault marks the geological and topographical boundary between a mountain range of Neogene sediments and adjacent Quaternary basin-ﬁll deposits (Fig. 11.5). The basement in this area consists mainly of weakly consolidated interbedded Pleistocene–Pliocene shale, mudstone, siltstone, and sandstone. The Chi-Chi earthquake produced a 100 km-long surface rupture zone along the pre-existing Chelungpu Fault; with maximum displacements of 11.1 m in the horizontal and 7.5 m in the vertical were recorded along the northern section of the fault (Lin et al. 2001c).

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Fig. 11.5. Geological map of west-central Taiwan, showing the locations of the Chiufener-Shan and Tsaoling landslides. F: Fault. (Modiﬁed from Central Geological Survey of Taiwan 1985; Hu et al. 1997; Lin et al. 2001a, b, c)

A large-scale landslide occurred close to the earthquake epicenter at Chiufener-Shan, between the Shuiliken and Tamoupu-Hsuangtung Faults (Fig. 11.5). Ground ﬁssures formed along these faults in the area close to the landslide, but no evidence of fault displacement was recorded (Lin et al. 2001a). The failure surface of the landslide developed within Miocene– Pliocene shale and siltstone along a bedding surface that dips to the southeast at 25–28◦ (Fig. 11.6). The landslide displaced an estimated 0.8 × 108 m3 of material by up to 2 km (Yu et al. 1999), making it comparable to the Langtang Himalaya landslide in terms of movement distance. The ﬁeld occurrence of the Chiufener-Shan landslide is reported in detail by Lin et al. (2001a).

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Fig. 11.6. Geologic cross-section through the Chiufener-Shan landslide, Taiwan. The topography is shown prior to (dotted line) and following (solid line) the earthquake. The landslide failure surface developed along a bedding surface within interbedded shale and siltstone (Modiﬁed from Lin et al. 2001a)

The transported material was broken into numerous blocks and fragments of varying sizes that were rotated as they slid downslope and came into contact with the ground surface. The landslide debris blocked rivers to form temporary lakes. Several houses and trees located upon blocks within the landslide were transported in excess of 2 km without collapse, demonstrating that parts of the landslide moved relatively smoothly, without causing significant damage to surface features. Thirty-ﬁve people died when a small village was buried under the landslide. Survivors of the landslide reported hearing a bomb-like sound during the earthquake, accompanied by a strong jolt (Lin et al. 2001a). The second large coseismic landslide occurred at Tsaoling, close to the southern end of the Chelungpu surface rupture zone (Fig. 11.5), displacing an estimated 1.0 × 108 m3 of material (Hung et al. 2000). Thirty-six people who lived upon the hillside were transported a distance of 3.5 km upon the landslide, with just seven surviving. Survivors reported that the landslide occurred immediately following strong ground motions (associated with the Chi-Chi earthquake). The results of ﬁeld investigations carried out immediately after the earthquake, in combination with eyewitness accounts, conﬁrm that both of the large-scale Chiufener-Shan and Tsaoling landslides were generated coseismically and resulted from strong ground motions associated with the 1999 Chi-Chi earthquake (Lin et al. 2001a). Molten material (pseudotachylyte) with the appearance of melt-bearing pseudotachylyte was discovered within interbedded shale and siltstone on the glide plane of the Chiufener-Shan landslide (Fig. 11.7). The pseudotachylyte occurs as a thin layer upon the landslide failure plane and as veins injected into cracks running obliquely from the plane (Fig. 11.8). The pseudotachylyte varies in thickness from 1 to 10 mm and is locally preserved in curved sections of the glide plane, accompanied by prominent slickenside striations. In freshly exposed sections, the pseudotachylyte is dark-brown to black in color, with a vitreous or pearly luster. The vein contains fragments of the sedimentary host rock cemented within a dark ﬁne-grained matrix.

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Fig. 11.7. Glide plane of the Chiufener-Shan (Taiwan) landslide. The glide plane is a bedding plane within interbedded shale and sandstone. Striations upon the glide plane dip to the southeast at 25–28◦ . View to the southeast. Note that the landslide debris blocked rivers to form the temporary lake (After Lin et al. 2001a)

Fig. 11.8. Thin layers of brown-black pseudotachylyte (Pt) that formed upon the glide plane of the Chiufener-Shan landslide, Taiwan (After Lin et al. 2001a)

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11.3 Petrographic Characteristics of Landslide-related Pseudotachylytes 11.3.1 Petrography of the Langtang Himalaya Pseudotachylyte The petrologic characteristics of the Langtang pseudotachylyte have been reported in detail by Scott and Drever (1953) and Masch and Preuss (1977). Masch et al. (1985) conﬁrmed the presence of glass in the pseudotachylyte on the basis of TEM analysis. When viewed under crossed polarized light on an optical microscope, the pseudotachylyte matrix shows the extinction characteristics of volcanic glass, being pale-gray, gray-brown, yellowish-brown, and dark brown in color (Figs. 11.9 and 11.10). The percentage of glass matrix within the vein varies from 20 to 80%, with typical values of 50–60%. The refractive index of the glass ranges from 1.492 to 1.508, indicating a silica-rich composition (Scott and Drever 1953). To check the refractive index of the glass, Scott and Drever (1953) melted a sample to produce an artiﬁcial glass; the authors then conﬁrmed that the refractive index of the glass was 1.495 (± 0.003). Colorless glass locally envelops fragments and penetrates to the interior of fragments (Fig. 11.9), as observed in the Fuyun pseudotachylyte (Fig. 5.5). Translucent and transparent patches of glass are observed within the generally yellowish-brown to dark-brown glass matrix (Fig. 11.9). TEM analyses

Fig. 11.9. Photomicrograph of the glass matrix and fragments within pseudotachylyte from Langtang Himalaya, Nepal. Transparent and pale to dark glass surrounds fragments and locally penetrates into the fragments (Fra). Flow: Flow structure within the glass matrix. TG: transparent glass. Plane polarized light

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Fig. 11.10. Photomicrograph of ﬂow structures within vesicular matrix from the Langtang pseudotachylyte. The elongate vesicle (Ve) has an aspect ratio of 1:10. ﬂow: Flow structures which are deﬁned by colored bands and fragments (frag) reoriented subparallel to the ﬂow streaks. Plane polarized light

of these patches yield diﬀraction patterns that are typical of non-crystalline material: the patterns lack Bragg contours and crystal defects such as dislocations, and crystalline diﬀraction patterns are absent (Masch et al. 1985). This provides indisputable proof of the occurrence of a large amount of glass material within the Langtang pseudotachylyte. Flow structures are also apparent under the microscope. These are generally deﬁned by translucent and transparent ﬂow streaks of varying color, the glass-rich matrix, and fragment-rich bands (Figs. 11.9 and 11.10). The glass matrix contains abundant quartz and feldspar fragments. The fragments are locally concentrated around pale to dark patches of glass matrix (Fig. 11.11). No biotite fragment can be recognized within the pseudotachylyte, even though biotite is one of the main rock-forming minerals within the host biotite gneiss and granite. The lack of biotite in the pseudotachylyte is considered to reﬂect selective melting (Scott and Drever 1953), as occurs within melt-origin pseudotachylyte via the preferential melting of maﬁc minerals that have relatively low melting points (Lin 1994a). Most of the fragments are angular to sub-angular, but smaller fragments are more rounded and embayed than larger ones (Figs. 11.9–11.11). Vesicles are generally elliptical and spherical in shape (Fig. 11.10), and some amygdaloidal structures are ﬁlled with ﬁne-grained carbonate material (Fig. 11.12). Microlites are also observed in the glass-rich matrix; they are generally too small to identify in terms of mineral type and most cannot be identiﬁed

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Fig. 11.11. Photomicrograph of a dark patch of glass matrix within the Langtang Himalaya pseudotachylyte. Numerous fragments (Fra) are concentrated around the margins of the patch. Plane polarized light

Fig. 11.12. Photomicrograph of microstructures within microlitic amygdaloidal matrix within the Langtang Himalaya pseudotachylyte. The microlites (Mi) are mainly lath-like and acicular feldspar. An amygdaloidal structure (Amy) is ﬁlled with ﬁne-grained carbonate material. Crossed polarized light

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from powder X-ray diﬀraction spectra (Scott and Drever 1953). The microlites have relative simple morphologies, including granular, acicular, lath-like, and skeletal (Fig. 11.12).

11.3.2 Petrography of the Chiufener-Shan Pseudotachylyte The Chiufener-Shan pseudotachylytes contain distinctive textures that are indicative of a melt origin; these include a glassy matrix, vesicles, rounded and embayed fragments, and ﬂow structures (Fig. 11.13), as also observed in the Langtang Himalaya pseudotachylyte and the melt-origin Fuyun pseudotachylyte described in previous chapters. Injection veins are also observed under the microscope; these branch from parent veins that formed on the glide plane of the landslide (Fig. 11.14). The matrix is mainly composed of glassy material and varies in color from pale yellow-brown to black when viewed under the microscope; it has the optical characteristics of volcanic glass. Zones of transparent glass surround quartz and feldspar fragments. Vesicles within the pseudotachylyte are circular to elliptical in shape, and range in diameter from ∼2 to 200 μm (Figs. 11.13). Locally, the proportion of vesicles within the glassy matrix reaches 20%. Fragments of the host sedimentary rock are varied in shape, being angular, rounded, and embayed, as well as possessing other irregular outlines. Flow structures are deﬁned by ﬂow streaks, oriented clasts, and layers of contrasting color (Fig. 11.13).

Fig. 11.13. Photomicrograph of microstructures within the Chiufener-Shan pseudotachylyte, Taiwan. The glass matrix contains prominent ﬂow streaks and irregularly shaped fragments. Numerous elongate vesicles (Ve) are aligned subparallel to ﬂow streaks. Plane polarized light (After Lin et al. 2001a)

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Fig. 11.14. Photomicrograph of microstructures within the Chiufener-Shan pseudotachylyte. A dark-brown to black pseudotachylyte vein is injected obliquely into the wall rock oblique from the glide plane. Plane polarized light. (After Lin et al. 2001a)

Fig. 11.15. Powder X-ray spectra of Langtang pseudotachylytes (a and b) and obsidian (c). (a): Vesicular glass from the center of a pseudotachylyte vein (shown in Fig. 11.4b). (b): Marginal zone of a pseudotachylyte vein composed almost entirely of fragments of the host rock (shown in Fig. 11.4c). (c) Obsidian sample from Wadatoge, Central Japan

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11.3.3 Glass Contents of the Observed Pseudotachylytes Powder X-ray diﬀraction patterns for the Langtang Himalaya and ChiufenerShan pseudotachylytes show a broad band of 2θ values, ranging from 12 to 42◦ (Figs. 11.15 and 11.16); these patterns are similar to those for obsidian (Fig. 11.15) and the Fuyun glass pseudotachylyte (Fig. 5.16b; see Chap. 5). The intensities of the quartz and plagioclase peaks for the pseudotachylyte are weaker than those for the host sedimentary rock (Figs. 11.15 and 11.16). The obtained spectra reveal that glass is present in both the Langtang and Chiufener-Shan pseudotachylytes and that some of the quartz and plagioclase crystals were melted within the pseudotachylyte. The diﬀraction patterns for the Chiufener-Shan samples contain peaks that are coincident with those for the host sedimentary rock, suggesting that the material within the pseudotachylyte veins was derived from the host rock. The diﬀraction patterns indicate glass contents of approximately 40–50 wt% and 50–60 wt% in the Langtang and Chiufener-Shan pseudotachylyte veins, respectively (Figs. 11.15 and 11.16). These values are comparable with those for glassy-type pseudotachylytes shown in Fig. 5.16. The glass

Fig. 11.16. Powder X-ray spectra of calibration materials (a–c), samples of Chiufener-Shan pseudotachylyte, Taiwan (d and e), and the host sedimentary rock (f). The calibration materials have glass contents of 100 (a), 50 (b), and 25 wt% (c), with the remainder of the material being powdered host rock. The pseudotachylyte veins (d and e) yield diﬀraction patterns similar to those of the glassy calibration material (a–c). On the basis of comparison with the calibration substances, the glass content of the pseudotachylyte veins (d and e) is estimated to be 40–50 wt%. Injection Pt: Injection pseudotachylyte vein; Fault-Pt: fault pseudotachylyte vein; Qz: quartz; Pl: plagioclase (After Lin et al. 2001a)

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11 Landslide-related Pseudotachylyte

content of the Langtang pseudotachylyte vein estimated from the diﬀraction data is consistent with that estimated by Scott and Drever (1953).

11.4 Discussion of the P-T Conditions during the Formation of Landslide-related Pseudotachylyte Both the Langtang and Chiufener-Shan pseudotachylytes contain typical melting textures that are similar to those found in fault-related pseudotachylytes that form by frictional melting. Field investigations carried out immediately after the 1999 Chi-Chi earthquake, in combination with petrological evidence, reveal that the Chiufener-Shan pseudotachylyte formed by frictional melting associated with a coseismic landslide initiated by the strong ground motions that accompanied the 1999 Mw 7.6 Chi-Chi (Taiwan) earthquake (Lin et al. 2001a). Although there are no historical records of large earthquakes in the area of the Langtang Himalaya landslide, it is inferred that the Langtang pseudotachylyte also resulted from frictional heating on the failure plane of the landslide and that the slide occurred during a large earthquake in the late Holocene. This interpretation is based on the pristine geomorphology of the landslide and the nature of alluvial deposits that form the lowest terraces along the Langtang River. The average current-day thickness of the Langtang Himalaya landslide deposit is calculated to be about 200 m, although topographical features indicate that the original thickness was in the order of 600–800 m (Masch et al. 1985). In contrast, the average thickness of transported blocks within the Chiufener-Shan landslide is estimated to be approximately 40 m (Fig. 11.6; Lin et al. 2001a). Considering the self-lubrication eﬀects of the fused material (Erismann 1986), the lithostatic pressures upon the glide plane of the Langtang Himalaya and Chiufener-Shan landslides are estimated to have been <30 MPa (<800 m in thickness) and <1.5 MPa (<40 m in thickness), respectively (Lin et al. 2001a). This result demonstrates that frictional melting upon slip planes can form at depths shallower than those estimated from well-known fault-related pseudotachylytes (e.g., Sibson 1975; Lin 1994a). The results of high-velocity frictional experiments (e.g., Lin and Shimamoto 1998) and observations made during drilling (e.g., Killick 1990) indicate that the experimental conditions were very similar to those of the landslide-generated pseudotachylytes, suggesting that frictional melting occurs under conditions similar to those of brittle faulting at shallow depths of ∼30–1000 m. These previous studies of natural and artiﬁcially generated pseudotachylytes reveal that frictional melting is readily generated at depths of <1 km, even as shallow as 30 m. It appears that the slip rate plays an important role in the generation of frictional melting upon the friction plane. The melting temperature during pseudotachylyte formation is commonly estimated from geothermometry analyses of microlites (e.g., Toyoshima 1990; Lin 1994a, b; see Chap. 8) and the chemical composition of the glass matrix

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281

(e.g., Lin 1994a; Lin and Shimamoto 1998), assuming melting at chemical equilibrium (see Chap. 8). Petrologic features within both the Langtang Himalaya and Chiufener-Shan pseudotachylytes indicate that feldspar and quartz crystals from the host rock were melted during formation of the pseudotachylyte. This means that the friction-related melting temperature within the landslide exceeded the melting temperatures of quartz and feldspars under a low normal stress of <30 MPa. The melting point of quartz under a normal stress of 30 MPa is ∼1600◦C (see Fig. 8.6; Kennedy et al. 1962), while that of feldspar varies from 1100 to 1550◦C; thus, the minimum estimated temperatures for the Langtang and Chiufener-Shan pseudotachylytes are 1100–1600◦C, which overlap with the melting points of quartz and feldspar. In these cases, frictional melt formed from the preferential melting of low-melting-point minerals under conditions of chemical non-equilibrium and without chemical reaction (e.g., Lin 1991, 1994a; Spray 1993; Lin and Shimamoto 1998); accordingly, the estimated temperatures are likely to be much lower than the actual temperatures.

12 Experimentally Generated Pseudotachylyte

12.1 Introduction The preceding chapters considered the properties of melt-origin pseudotachylyte, whose origin is demonstrated unequivocally by the presence of as much as 90% glass matrix within the Fuyun pseudotachylyte veins, vesicles and amygdules, high-temperature microlite minerals of varied morphologies, and rounded and irregularly embayed fragments within pseudotachylytes hosted within a variety of host rocks. The results of high-velocity frictional melting experiments (e.g., Spray 1987, 1988, 1993; Lin 1991; Lin et al. 1992; Lin 1994, 1998) and drilling projects (e.g., Bowen and Aurousseau 1923; Killick 1990; Kennedy and Spray 1992) also demonstrate that frictional melt can be generated upon a slip plane under conditions similar to those of seismic faulting at shallow depths, even at depths of less than 30 m (Lin 1991; Lin and Shimamoto 1998). It is important to determine the physical conditions that accompany the formation of frictional melt during seismic slip and the nature of related chemical processes; i.e., whether frictional melting occurs under conditions of chemical equilibrium. This is important in terms of the following factors: 1) understanding the mechanical mechanisms and melting processes involved in the generation of fault-related pseudotachylyte, 2) determining the temperature of the molten zone during frictional melting, and 3) assessing the thermal evolution and eﬀect of melting on the mechanics of seismic slip within a seismogenic zone during fault movement. An understanding of the chemical and physical processes of frictional melting is essential for the correct application and explanation of laboratory data in terms of ﬁeld examples and mesostructures and microstructures, as well as in explaining compositional variations within natural fault rocks such as fault-generated pseudotachylyte. Accurate estimates of frictional temperature are critical in estimating shear stress and assessing the process of frictional heating on a fault plane in terms of the vein thickness–displacement relation obtained for fault-generated pseudotachylyte (Sibson 1975). Such data provide important constraints on

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laboratory experiments that investigate dynamic fault strength during seismic fault movement. The frictional temperature attained during the frictional formation of melt-origin pseudotachylyte is commonly estimated from the chemical stability ﬁelds of minerals that grew within the pseudotachylyte veins and that possess textures indicative of rapid growth, such as dendritic microlites, assuming that melting occurred under conditions of chemical equilibrium (e.g., Toyoshima 1990; Lin 1991, 1994a, b; Magloughlin and Spray 1992); however, the assumption of chemical equilibrium may be invalid. In such a case, the temperature distribution within a molten zone is likely to be relatively heterogeneous. The dynamic fault strength during frictional melting is primarily controlled by the growth of a molten layer and its viscosity, with the latter being strongly dependent on melt temperature (Spray 1993). The pattern of thermal evolution within a seismic fault zone can diﬀer markedly between equilibrium melting, in which the temperature remains at the melting temperature for a suitably long period, and non-equilibrium melting, which involves overstepping of the melting temperature and a heterogeneous temperature distribution during an instantaneous period of generally less than 1 minute during an individual large-magnitude earthquake. Frictional melting during seismic fault movements generally takes place over a period of ∼2–20 seconds, and it is uncertain as to whether chemical equilibrium is attained during such rapid frictional melting. In an attempt to solve the problems documented above, a number of studies have conducted a series of high-velocity frictional melting experiments over the past two decades (e.g., Spray 1987, 1988, 1993; Lin 1991; Lin 1994, 1998; Shimamoto and Lin 1994; Tsutsumi and Shimamoto 1994, 1996, 1997a, b; Hirose and Shimamoto 2003, 2005; O’Hara et al. 2006; Di Toro et al. 2006). These experiments have successfully simulated the geometric modes and properties of molten material, as well as the formation processes of fault-generated pseudotachylyte veins. The experiment results demonstrate that frictional melting is a non-equilibrium chemical process (e.g., Lin and Shimamoto 1998) and suggest that frictional melting acts to reduce fault strength during seismic faulting (Hirose and Shimamoto 2005; Di Toro et al. 2006). This chapter focuses on uniaxial high-velocity frictional melting experiments and the chemical and physical processes of frictional melting during seismic faulting within seismogenic zones.

12.2 High-Velocity Frictional Experiments 12.2.1 Test Equipment and Experimental Conditions With the aim of understanding the high-velocity frictional properties of rocks within seismic fault zones, Professor T. Shimamoto of The University of Tokyo, Japan, designed a rotary-shear high-velocity frictional test machine that was installed at the Earthquake Research Institute, University of Tokyo

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Fig. 12.1. Photograph (a) and sketch (b) of the high-velocity frictional testing machine 1: a servo-motor, 2: torque, 3: torque gage, 4: electromagnetic clutch, 5: cylindrical specimen of 25 cm in diameter, 6: actuator for applying axial force, 7: water reservoir, 8: motor controller

(Fig. 12.1). This test machine was ﬁrst used for frictional melting experiments in 1991, although at this time it was impossible to make accurate measurements of torque and other mechanical data (Lin 1991). After testing and improving the system over several years, it became possible to obtain mechanical data such as torque, frictional coeﬃcient, axial-shortening amount, normal

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stress and shear stress, frictional melt temperature, and displacement; these values could be accurately measured during experiments to quantitatively assess the dynamic behavior of a simulated fault at high velocities and under large displacements (Shimamoto et al. 2006). The test machine is simple in its construction and is easy to operate, making it one of the most successful experimental rock-deformation machines in the world in terms of studying high-velocity deformation mechanisms and the rheology of fault rocks. The experimental results obtained using this machine have been published in a number of studies (e.g., Lin 1991 1992; Lin 1994, 1998; Ohtomo and Shimamoto 1994; Shimamoto and Lin 1994; Shimamoto and Tsutsumi 1994, Tsutsumi 1994, 1999; Tsutsumi and Shimamoto 1994, 1996, 1997a, b; Nakamura et al. 2002; Hirose 2002; Hirose and Shimamoto 2003, 2005; Di Toro et al. 2006; O’Hara et al. 2006; Han et al. 2007). All experiments to date on frictional melting have been performed under dry conditions because the test machine lacks a hydraulic pressure system; however, a new and improved rotary-shear high-velocity testing machine has been produced equipped with a hydraulic pressure apparatus. This new machine was installed in Shizuoka University, Japan, early in 2007 (Fig. 12.2). The new test machine contains two rotary-shear high-velocity testing devices: one with the same capacity as the original testing machine (see below for details), and another that contains a hydraulic friction apparatus with a controlling device for high pore-water pressure attached to the uniaxial

Fig. 12.2. Photograph of the new rotary-shear low-high velocity testing machine equipped with a hydraulic pressure apparatus 1: a servo-motor and gear system, 2:one-axis rotary-shear specimen box, 3: torque and axial force gauges, 4: hydraulic friction apparatus with a controlling device for high pore-water pressure, 5: axial force and torque gauges, 6: an oil pressure pump, 7: pressure generator system, 8: data recording system. This machine has been installed in Shizuoka University, Japan in the early of 2007

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rotary-shear high-velocity friction device (Fig. 12.2); this device has a wide range of slip rates, from <10 cm/yr to 10 m/s. The new machine enables high-velocity friction experiments on faults with supercritical pore-water pressure and low rates of shear deformation; these experiments can be used to simulate both seismic and aseismic crystal plastic deformation processes of mynolite-related pseudotachylyte veins within fault zones where high pore-ﬂuid pressure is present. The two devices are set together with a single rotary axis that is powered by a single servo-motor (Fig. 12.2). Preliminary experiments under high hydraulic pressure conditions are currently being carried out within the Laboratory of Earthquake Geology at Shizuoka University, Japan. The results of earlier experiments carried out using the uniaxial testing machine under dry conditions are described below. A pair of cylindrical specimens of 25 mm in diameter and 50 mm in length is used for high-velocity rotary-shear frictional melting experiments (Fig. 12.3), with the circular interface of the two specimens being set as the slip plane (Fig. 12.4b). Hollow cylindrical specimens are also used in some experiments, as the reduced contact area acts to increase the normal stress (Fig. 12.3). High-velocity slip is attained by rotating one of the cylindrical specimens via a 75 kW AC servo-motor (left-hand side of Fig. 12.4b) while the other specimen is kept stationary (right-hand side of Fig. 12.4b). The machine is capable of producing slip rates of up to 2 m/s, corresponding to 1,500 rpm under an axial load of up to 1000 kg. A disadvantage of the rotary-shear testing machine is

Fig. 12.3. Photograph of hollow-cylindrical (a) gabbro and cylindrical (b) coursegrained and (c) ﬁne-grained granite specimens used for high-velocity rotary-shear frictional melting experiments. The samples are 25 mm in diameter and 50 mm in length

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Fig. 12.4. Photographs of the specimen assembly in the high-speed frictional testing machine (a and b) and sparks with cherry red melt ejected from the interface (white arrow) of two cylinder specimens during a high-velocity frictional experiment involving gabbro (c). The circular interfaces of two specimens are in contact as the slip plane (b). Progress at the sliding interface during the frictional experiments is observed through a transparent window cover

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that the slip rate varies with the radial position, i.e., for solid, cylindrical specimens the rate is zero in the center and reaches a maximum at the outer edge. To assess the total frictional work, we generally use an equivalent slip rate (Veq), whereby Veq multiplied by the area of sliding provides the rate of frictional work (Shimamoto and Tsutsumi 1994; Lin and Shimamoto 1998; Shimamoto et al. 2006). On the sliding area of a cylindrical specimen with a diameter r, the total area of the circle is ds = 2π rdr and the slip velocity (v) at a point with the distance (radius) r is 2π rR, where R is the revolution rate of the motor; therefore, the frictional work is given by dw = vτds = (2π r)2 τRd r = 4π2 r2 τRd r

(12.1)

where τ is shear stress. Assuming the same frictional coeﬃcient over the entire frictional surface, the frictional work W is given by W = dw = 4π 2 τR r2 dr = 4/3π 2 τR(r2 2 − r1 2 )

(12.2)

Here, the frictional area S is deﬁned as S = π(r2 2 − r1 2 )

(12.3)

for a cylindrical specimen of radius (r = r2 , r1 = 0). Thus, Veq is given by Veq = 4π R(r2 2 + r2 r1 + r1 2 )/3(r1 + r2 ) = 4/3π r R

(12.4)

At a rotation rate of 1500 rpm for a cylindrical specimen of 25 mm in diameter, the maximum Veq on the circular frictional area is up to 2.6 m/s. The test equipment can be used to apply normal stresses up to approximately 100 MPa, close to the typical uniaxial strength of crystalline rocks, provided that anvil-shaped specimens are used. As it turns out, however, it is only possible to apply normal stresses of <10 MPa because of the severe thermal fracturing of unconﬁned specimens upon frictional heating. Temperature measurements and calculations reveal that the temperature of frictional melt can exceed 1100◦ C when using gabbroic specimens (Tsutsumi and Shimamoto 1997b). The uniaxial strength of granite is reduced by almost two orders of magnitude at such high temperatures (Bauer and Johnson 1979). The reduction in the uniaxial strength of rocks due to thermal fracturing is even more severe during high-velocity frictional melting experiments because of the high rate of frictional heating (Ohtomo and Shimamoto 1994). Thermal fracturing is particularly pronounced in quartz-bearing rocks because quartz undergoes α–β transformation, which is accompanied by substantial volume change; accordingly, frictional melting experiments commonly make use of ﬁne-grained gabbroic samples that contain little quartz, with any quartz being largely free of visible cracks. High-velocity frictional melting experiments are performed at room temperature, under dry conditions, and with a revolution rate of up to 1500 rpm. Experiments are generally conducted for 20–40 seconds for gabbroic

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rocks and 10–20 seconds for granite at equivalent slip rates of 0.75–1.90 m/s and normal stress of ≤1.5 MPa. As high values of normal stress cannot be applied in these tests, we generally increase the total displacement to several tens of meters such that the total frictional work corresponds to that of much smaller displacements at higher normal stress. Because of the low normal stresses applied to the simulated faults, in terms of the frictional work the high-velocity frictional melting experiments with displacements as large as 50–100 m correspond approximately to those with displacements of the order of 1 m under normal stresses of 100 MPa. These displacements are much larger than those recorded during individual large-magnitude earthquakes (maximum values of approximately 10 m upon a single slip plane). The frictional energy expended per unit area in the high-velocity frictional melting experiments under a slip rate of ∼1 m/s and a large displacement of up to 50 m is very similar to that of a mid-crustal earthquake (Table 12.1; O’Hara et al. 2006). The experimental conditions are therefore comparable to those during a large earthquake at shallow depths, at least in terms of velocity and stress drops (e.g., Kanamori and Stewart 1976; Kikuchi 1995). 12.2.2 Experiment Samples and Procedures The high-velocity frictional melting experiments are conducted using samples of gabbro; granite; pairs of albitite–anorthosite and albitite–quartzite, which have eutectic relations (Bailey 1976); and anorthosite–anorthosite pairs, which consist mostly of plagioclase with an anorthite composition of 70–75%, which is a well-known solid solution of albite and anorthite (Deer et al. 1992). The gabbro samples consist of two diﬀerent grain sizes that are 1 and 2 mm in diameter; for convenience, these are herein termed ﬁne-grained and medium-grained, respectively. Both types of gabbro samples consist mainly of plagioclase, pyroxene, hornblende, and biotite. The gabbro samples are used to duplicate the melting process in quartz-poor rocks that are rich in hydrous minerals; these are suitable for frictional melting experiments and demonstrations for students because the quartz-poor nature of the specimens means that they are not destroyed during the experiment. Ultramaﬁc rock Table 12.1. Frictional energy expended per unit area from high-velocity frictional melting experiments on coal gouges and from a mid-crustal earthquake. The coeﬃcient of friction μ is set to 0.6 Event

Shear stress

Displacement

Frictional work

Earthquake Experiment

τ= 66 MPa τ= 0.6 MPa

D = 1m D = 50 m

66 × 106 J/m2 30 × 106 J/m2

(Data from O’Hara et al. 2006)

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hosted pseudotachylytes are also found in natural fault zones (Brandl and Reimond 1990, Obata and Karato 1995). Most fault-related pseudotachylyte occurs within granitic rocks; therefore, granite samples are used to simulate the frictional melting process and the formation mechanisms of quartz-rich quartzo-feldspathic rocks within continental fault zones. The granite samples used in the experiments can be divided into three types on the basis of their average grain size: 1, 2, and 4 mm, which are herein termed ﬁne-grained, medium-grained, and coarse-grained samples, respectively. The granite samples consist mainly of K-feldspar, plagioclase, quartz, pyroxene, biotite, and magnetite, and contain visible microcracks, particularly in course-grained quartz crystals. Frictional melting experiments using granite specimens are generally diﬃcult to perform because of the rapid destruction of specimens caused by the formation of heating cracks within quartz crystals. Based on the eutectic and solid solution properties of minerals, albitite, quartzite, and anorthosite samples are used in experiments that are intended to shed light on the chemical processes of frictional melting (Lin 1991; Lin and Shimamoto 1998). Albitite samples are obtained from Itoigawa, Niigata Prefecture, Japan, with quartzite sourced from Brazil (exact locality unknown), and anorthosite from the Outer Hebrides, Scotland, where well known meltorigin pseudotachylyte is found. The quartzite and anorthosite samples are generally stiﬀ and largely free of visible cracks, but albitite blocks contain numerous visible cracks that must be avoided in extracting cores from the specimens. The choice of samples for quartz is diﬃcult because of thermal fracturing; accordingly, large crystals of synthetic quartz, chert blocks, and natural quartz veins are commonly used. The albitite samples are not pure albite: they also contain considerable amounts of quartz; however, they are largely free of other minerals, meaning that the experimental conditions are not signiﬁcantly complicated by contamination. As the albitite contains some quartz crystals but the quartzite is largely free of albite, the sliding that occurs during the experiments conducted using albitite–quartzite pairs is primarily albite-on-quartz and quartz-on-quartz (not albite-on-albite). This precludes the melting of albite due to albite-on-albite sliding. The anorthosite from Northwest Scotland (Macaudi`ere and Brown 1982) consists mainly of plagioclase (more than 93%); consequently, the sliding that occurs during anorthosite–anorthosite experiments is mainly plagioclase-on-plagioclase sliding. Cylindrical specimens are obtained by coring; the end faces of the cores are polished using a silicone-carbide grinding wheel to ensure that the specimen ends are normal to the cylinder axis (Fig. 12.3). The specimens are washed for approximately 20 minutes with an ultrasonic cleaner and dried at room temperature. Progress at the sliding interface during frictional experiments can be observed through a transparent window cover, as shown in Fig. 12.4.

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12.2.3 High-Velocity Frictional Properties Experiments involving the high-velocity frictional melting of rocks demonstrate that frictional melting is readily initiated during seismic faulting, even under a low normal stress of 1 MPa (Lin 1991; Lin and Shimamoto 1998). High-velocity frictional properties and the mechanical behaviors of rocks have also been observed during frictional melting experiments carried out on the rotary-shear high-velocity frictional testing machine (Tsutsumi and Shimamoto 1997a; Hirose and Shimamoto 2005; O’Hara et al. 2006). Frictional melting phenomena are clearly observed during high-velocity frictional melting experiments of gabbro samples. The simulated fault (sliding interface) between gabbro specimens generally begins to spark and produce dust within several seconds of the initiation of fault motion; the interface changes to a cherry-red color after 5–7 seconds and produces frictional melt.

Fig. 12.5. Photograph of the frictional melting zone where the melt was squeezed out of the shear zone

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293

Cherry-red run products are then extruded from the sliding surface (Figs. 12.4 and 12.5). Simulated faults in granite, quartzite–albitite, and anorthosite generate sparks and dust in the ﬁrst 5–6 seconds of the experiment; however, the specimens generally fracture and disintegrate after no more than 10 seconds. High-velocity frictional melting experiments are much more diﬃcult to perform for granite, quartzite, and albitite pairs than for gabbro, as granite and quartzite contain numerous quartz crystals that undergo α–β transition during frictional heating, resulting in thermal fracturing (Ohtomo and Shimamoto 1994). This problem commonly results in the granite experiments being terminated within <10 seconds. The host granite commonly fractures into pieces, meaning that it is very diﬃcult to generate injection veins of pseudotachylyte similar to those generated in gabbro experiments. Simulated faults that consist of an albitite–quartzite pair began to emit sparks after just several seconds from the onset of the experiment; they turn bright reddish-orange after approximately 10 seconds. Frictional melt then begins to form and is extruded for approximately 4–5 seconds before the specimen breaks into fragments and the run is terminated. Simulated faults that consist of anorthosite–anorthosite pairs emit bright yellow sparks and turn bright reddish-orange after 7–8 seconds; the experiments continue with frictional melting for approximately 30 seconds, without breakage of the specimen. The total circumferential displacement is somewhat less than 30 m for albitite–quartzite pair and approximately 80 m for anorthosite pairs. These observations reveal that frictional melt can be produced over short periods of less than several seconds.

12.3 Microstructures of Experimentally Generated Pseudotachylyte 12.3.1 Textures of the Fault Shear Plane To enable the observation of microstructures upon the experimental shear plane, the cylindrical specimens are separated as soon as the experiment ﬁnishes, before the frictional surface is bonded by rapid cooling of the frictional melt. Figure 12.6 shows a representative example of the rotary-shear frictional surface of a gabbro sample. The fault surface is generally uneven and contains ﬂow or drag structures with numerous striations or grooves oriented parallel to the slip direction (Figs. 12.6 and 12.7a) although the boundaries are gradual. The distribution of melt on the frictional surface is generally heterogeneous. The circular frictional surface can be divided into three ring-shaped zones on the basis of texture: the central (A), middle (B), and outer zones (C) (Fig. 12.6). The production rate of frictional heat is small in the central zone because of the low slip rate close to the centre of the circular frictional surface. In the outer zone, frictional heat is rapidly lost to the air, although the slip rate

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Fig. 12.6. Photograph of the frictional melting surface of a gabbro specimen. The melt distribution on the frictional surface can be divided into three ring-shaped zones on the basis of textural features: central zone (A), middle zone (B), and outer zone (C). The central and outer zones (A, C) are areas in which fragments accumulated and were cemented by small amounts of melt that locally show a glassy luster. The black-colored middle zone (B) occupies approximately half the radius of the circular surface and typically possesses a morphology that is similar to that of fresh lava, including vesicles, cavities, and the surface roughness of a ploughed ﬁeld. (After c Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

is suﬃciently high to generate frictional heat; consequently, the temperature is highest in the middle zone (B), which is insulated by the adjacent zones (Lin 1991; Tsutsumi and Shimamoto 1994; Lin and Shimamoto 1998). The central and outer zones (A, C) are the areas in which ﬁne fragments accumulate, cemented by small amounts of frictional melt that locally shows a glassy luster. The black-colored middle zone (B) occupies approximately half of the radius of the circular surface and is typically similar in morphology to fresh lava, exhibiting vesicles, cavities, and a surﬁcial roughness that resembles a ploughed ﬁeld (Fig. 12.7). These textures upon the frictional surface indicate that frictional milling is the main mechanism that operates at the shear surface prior to melting. The vesicles vary in diameter from several microns to 1 mm and are generally circular in shape, although in places they are elliptical in shape, with the long axis oriented parallel to the slip direction (Fig. 12.7). 12.3.2 Vein Geometry and Texture of Molten Material The microstructures of experimental fault shear zones are observed in thin sections cut perpendicular to cylindrical specimens that are bonded by the

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Fig. 12.7. SEM images of the frictional melting textures generated on the surface of a gabbro specimen during a high-velocity frictional experiment. (a): Flow structures generated upon the frictional surface. (b): Cavernous textures developed upon the c frictional surface. (After Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

frictional melt. Figure 12.8 shows the typical textural features of a frictional shear zone. The generation zone of frictional melt is approximately 50 μm in width, and injection veins vary in width from ∼2 to 50 μm. The boundaries between the generation zone of frictional melt and the walls of the host gabbro are generally sharp and straight, but locally irregular where injection veins branch from the main generation surface (Fig. 12.8).

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Fig. 12.8. SEM images of experimentally generated pseudotachylyte veins. (a and b): Fault vein and network veins injected into the host rock. The thin sections were cut normal to the shear zone and parallel to the sliding direction. (After Lin and c Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

The injection veins are highly irregular in geometry and narrow in width, even less than one micron, and are generally much thinner than the generation zone; most can be traced back to the generation zone (Fig. 12.8). The injection veins are generally thicker at the branching point and gradually decrease in width away from the shear plane. The geometric features of the generation zone and injection veins are similar to those observed in

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natural pseudotachylyte veins, as described in Chap. 4 (Fig. 4.10a). Both the injection veins and generation zone contain numerous rock fragments (Fig. 12.8). The above textural features of experimentally generated pseudotachylyte veins demonstrate that the geometry and structural modes of

Fig. 12.9. SEM images of the melting textures within experimentally generated pseudotachylyte. (a): Vesicular texture. Note that most of the fragments are rounded. (b): Enlargement of image (a). Pl: plagioclase fragment; Py: pyroxene fragc ment. Ve: Vesicle. (After Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

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12 Experimentally Generated Pseudotachylyte

natural pseudotachylyte veins can be successfully simulated in high-velocity frictional melting experiments. During frictional experiments, run products including molten material and fragments are commonly squeezed out and splashed from the fault shear zone (Fig. 12.4). The irregular shapes of these products resemble hairs and small pillows, and they contain numerous ﬁne fragments of the host rock. Pseudotachylyte generated experimentally from gabbro comprises mainly molten material and sub-angular to rounded fragments of pyroxene and feldspar derived from the host gabbro (Figs. 12.8 and 12.9); no biotite or hornblende fragments are observed, even though these are the main rock-forming minerals of the host gabbro. Although these experiments are conducted under dry conditions in the absence of external water, numerous vesicles are observed in the molten layer (Fig. 12.9). These vesicles are generally circular in shape, but some are elliptical, and most are less than 20 μm in diameter. Such vesicles represent the traces of bubbles produced from hydrous minerals such as biotite and hornblende that were preferentially melted during frictional melting (Lin 1991; Lin and Shimamoto 1998). Figure 12.10 shows a circular run product of a ﬁne-grained granite sample that is connected by molten material onto fragments of the host granitic rock. The circular run product was thrown out from the holder when the cylindrical specimens broke during the experiment. Biotite crystals were fused and ﬂowed parallel to the shear direction on the slip surface, forming black bands

Fig. 12.10. Photograph of a broken ring-shaped melt product that was ejected from the simulated fault during an experiment involving a ﬁne-grained granite specc imen. (After Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

12.3 Microstructures of Experimentally Generated Pseudotachylyte

299

Fig. 12.11. Photograph of run products generated from a granite specimen. The maﬁc mineral (biotite, black area indicated by arrows) was fused onto the simulated c fault surface. (After Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

(Figs. 12.11 and 12.12). The transparent and translucent hair-like products are elongate and aligned parallel to the shear direction. The molten material derived from the granite sample mainly comprises angular and subangular to sub-rounded quartz and feldspar fragments within a heterogeneous glassy

Fig. 12.12. SEM image of fused material (F-M, white area in the center of the image) from biotite crystal upon the frictional surface of a granite specimen. (After c Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

300

12 Experimentally Generated Pseudotachylyte

matrix (Fig. 12.12). The cemented textures and ﬂow structures clearly indicate that the frictional melt formed before the specimens broke into pieces.

12.4 Powder X-ray Diﬀraction Analysis of Run Products 12.4.1 Diﬀraction Patterns of Run Products The run products of gabbro and granite samples are commonly subjected to powder X-ray diﬀraction analysis to quantitatively determine the composition of molten material and the major constituent minerals within the experimentally derived pseudotachylyte (mixture of glass and fragments) (Lin 1991; Lin and Shimamoto 1994, 1998; Tsutsumi 1999). Figure 12.13 shows representative X-ray diﬀraction spectra for four run products of ﬁne-grained granite and two specimens of gabbro. The spectra for the gabbro and granite host rocks are also shown. All spectra for the run products exhibit broad bands of 2θ values from 15 to 42◦ , similar to

Fig. 12.13. X-ray diﬀraction spectra of experimentally generated pseudotachylytes derived from (A) ﬁne-grained granite and (B) ﬁne-grained gabbro. The diﬀraction data in (a) are for host rocks and those in (b)–(e) are for experimentally generated pseudotachylyte. Ma: mica; Pl: plagioclase; Qz: quartz; Py: pyroxene. The measurement conditions were as follows: ﬁltered CuKa (1.54050 ˚ A) radiation, X-ray generator 3 kW, 40 kV, 20 mA, sampling width 0.02˚, scanning speed 3.0◦ /min., divergence slit 1.0◦ , scattering slit 1.0◦ , receiving slit 0.15 mm, 2θ position 5.0◦ . (After c Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

12.4 Powder X-ray Diﬀraction Analysis of Run Products

301

the spectra for glassy volcanic rocks such as obsidian and fault-generated glassy pseudotachylyte (see Chap. 5). This broad band is absent in the spectra for the gabbro and granite host rocks (Fig. 13.13A-a and 12.13B-a). The distinctive broad band in the run products demonstrates the presence of a certain amount of glass or non-crystalline material within the experimentally generated pseudotachylyte (Fig. 12.13A-b∼e, 12.13B-b and B-c). The diﬀraction peaks of pyroxene and feldspar crystals are present in both the host gabbro and its run products, and those of feldspar and quartz are present in the host granite and its run products, although the integrated intensities of all peaks are smaller than those of the host rocks (Fig. 12.13). The mica peaks in the 2θ position of 9–10◦, clearly observed in both the gabbro and granite host rocks, are not recognized in the spectra of the run products (Fig. 12.13A-b∼e, 12.13B-b and B-c). These spectra indicate that all of the mica crystals within molten material were preferentially melted, even though other minerals were only partially melted.

12.4.2 Quantitative Analysis Quantitative analyses of molten material in the run products are performed using the same method as that described in Chap. 5. Six standard samples made up of 0, 10, 20, 50, 75, and 100 wt% artiﬁcially fused glass are prepared by weighing and mixing powders of the glass and host rocks for the gabbro and granite samples. Figure 12.14 shows the obtained X-ray diﬀraction spectra of the standard samples. All of the spectra except for those for the gabbro and granite host samples show a broad band at 2θ values of 12–42◦. As described in Chap. 5, the broad bands represent the integrated intensity generated by glassy material; the band increases with increasing glass content (Fig. 12.14A). The peak intensities for crystalline material decrease with increasing glass content. The integrated intensities of glass (Ig) in the standard samples are measured using the triangle approximation method, as shown in Fig. 12.15B, and are plotted against the glass content (Wg) in Fig. 12.15. Although there are personal equation on the measurements of the broad band area and peak intensity of the crystal, a least squares analysis of the measurement results reveals a linear relation between the extent of the broad area and glass content (Fig. 12.15). This means that the integrated intensity of glass increased linearly with the increasing amount of glass content. The integrated intensities of the experimentally generated pseudotachylytes are also measured and plotted on the calibration curves using the method described in Chap. 5. Figure 12.15A and 12.15B shows the measurement results for granite and gabbro run products, respectively (Lin 1991; Lin and Shimamoto 1998). The glass contents (Wg) of the molten materials, as taken directly from the calibration curves, are 43.4–62.5 wt% and 44.5–52.3 wt% for the granite and gabbro run products, respectively. This means that approximately 50 wt% of the molten material is made up of fragments of the host

302

12 Experimentally Generated Pseudotachylyte

Fig. 12.14. X-ray diﬀraction spectra of standard samples for (A) ﬁne-grained granite and (B) ﬁne-grained gabbro. Data for (a) to (f) correspond to glass contents of 0, 10, 20, 50, 75, and 100 wt%, respectively. Ma: mica; Pl: plagioclase; Qz: quartz; Py: pyroxene. The measurement conditions are the same as that shown in c Fig. 12.13. (After Lin and Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

Fig. 12.15. Calibration curves for the content (Wg) of glassy matrix relative to the bulk specimen for (A) ﬁne-grained gabbro and (B) ﬁne-grained granite. Vertical axes show the integrated intensities in the X-ray diﬀraction spectra. (After Lin and c Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

12.4 Powder X-ray Diﬀraction Analysis of Run Products

303

rock; this is consistent with observations made using the optical microscope and electron microscope, as described above. We also analyzed crystals of the main host-rock minerals, such as quartz and feldspar, which remained in the melt material generated from the granite samples. The integrated intensities (Iq) of quartz peaks in the 2θ position of 26–27◦ (b–f in Fig. 12.14A), which are the strongest quartz peaks, were measured and plotted against the quartz content measured from the standard samples. The intensity of equivalent quartz peaks increased linearly with increasing amount of crystalline material. The calibration curve shown in Fig. 12.16A was obtained by using the total original quartz content of the host granite sample as 100 wt% and the total original quartz content of the 100% glass standard sample as 0 wt%. The integrated intensities of equivalent quartz peaks in the run products were then measured and plotted on the calibration curve. The quartz contents of the run products, as taken directly from the trend line, are 45.5, 71.9, 75.8, and 82.9 wt% of the original quartz content in the granite samples. This means that approximately 17–54 wt% of the quartz crystals in the host rocks were melted within the run products. In the same way, a calibration curve was also obtained for feldspar (using the plagioclase in the peak 2θ position of 27–28◦) from the spectra of standard samples shown in Fig. 12.14A. The integrated intensities of the plagioclase

Fig. 12.16. Calibration curves for the contents of (A) quartz crystals and (B) feldspar (albite) crystals in ﬁne-grained granite specimens. Vertical axes show the integrated intensities of albite peaks in the X-ray diﬀraction spectra. (After Lin and c Shimamoto 1998). 2007, with kind permission from Elsevier Science Ltd

304

12 Experimentally Generated Pseudotachylyte

peaks were measured and plotted upon the calibration curve (Fig. 12.16B), yielding plagioclase contents of 30.5, 35.0, 34.5, and 39.0 wt% of those within the host rocks.

12.5 Chemical Composition Data To understanding the chemical process of frictional melting, it is necessary to compare the chemical compositions of run products with those of the host rock. The bulk compositions of the host rock samples used in the frictional experiments are generally measured via XRD analysis, while the molten material within run products is analyzed using an electronic microprobe. The chemical compositions of the host gabbro and granite samples are shown in Table 12.2, while those of the molten run materials are shown in Tables 12.3–12.7. Although fragments of the host rock were avoided in undertaking the chemical analysis of the glassy matrix, ultraﬁne fragments could not be completely excluded because of their small size and compositional similarity between the matrix and fragments when viewed in SEM-EDX images. 12.5.1 Gabbro Samples The glassy matrix derived from medium-grained gabbro samples is heterogeneous in chemical composition, whereas that of the ﬁne-grained gabbro exhibits a more homogeneous composition (Fig. 12.17). The chemical compositions of the glassy matrices are somewhat lower in SiO2 content than those Table 12.2. Bulk compositions of rocks used in high-velocity frictional melting experiments analyzed by XRF Wt% SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 Total

01 52.51 0.19 18.44 5.97 0.15 6.94 13.25 2.38 0.18 0.01 100.02

02 49.85 2.85 16.56 10.11 0.15 4.90 8.20 3.13 2.76 1.58 100.09

03 72.75 0.24 14.07 2.14 0.06 0.37 1.40 3.87 5.01 0.10 100.91

04 71.93 0.30 15.21 2.39 0.06 0.48 2.37 3.50 3.16 0.09 99.49

05 71.60 0.49 15.10 2.94 0.07 0.64 2.16 3.05 4.00 0.15 100.20

01: medium-grained gabbro; 02: ﬁne-grained gabbro; 03: coarse-grained granite; 04: medium-grained granite; 05: ﬁne-grained granite. FeO*: total Fe calculated as FeO. (Data from Lin 1994, 1998)

12.5 Chemical Composition Data

305

Table 12.3. EDX-determined chemical compositions of glass matrices (Mg1–Mg9) within pseudotachylyte generated experimentally from medium-grained gabbro samples wt%

Mg1

Mg 2

Mg 3

Mg 4

Mg 5

Mg 6

Mg 7

Mg 8

Mg 9

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 Total

53.09 0.18 16.88 6.15 0.26 7.52 12.27 1.83 0.23 0.05 0.62 0.18 99.26

53.22 0.13 16.28 6.51 0.00 9.30 12.26 1.90 0.27 0.00 0.29 0.43 100.59

53.24 0.16 18.92 5.42 0.11 5.80 12.04 2.54 0.13 0.19 0.00 0.32 98.87

53.94 0.00 24.24 2.46 0.04 3.12 12.21 3.33 0.24 0.03 0.00 0.13 99.72

52.17 0.36 1.20 8.24 0.16 13.60 20.68 0.00 0.11 0.16 0.07 0.19 96.94

52.83 0.00 25.25 2.56 0.04 3.02 12.20 3.43 0.24 0.04 0.00 0.12 99.94

52.36 0.25 14.97 7.94 0.06 8.32 11.80 1.42 0.26 0.16 0.35 0.27 98.16

52.65 0.50 1.11 8.53 0.45 14.42 21.10 0.16 0.00 0.00 0.36 0.21 99.49

52.38 0.35 1.36 9.11 0.16 14.08 20.60 80.01 0.02 0.23 0.23 0.29 98.82

(Data from Lin 1994, 1998)

Table 12.4. EDX-determined chemical compositions of glass matrices (fg1–fg9) within pseudotachylyte generated experimentally from ﬁne-grained gabbro wt%

fg1

fg2

fg3

fg4

fg5

fg6

fg7

fg8

fg9

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 SeO Total

47.53 3.54 16.33 9.88 0.16 6.32 6.98 2.64 3.78 0.14 0.08 1.77 0.07 99.22

49.18 3.02 16.86 8.57 0.20 5.18 5.88 2.63 5.14 0.08 0.00 1.43 0.70 98.87

46.91 3.56 16.23 10.11 0.15 5.41 7.46 2.66 2.92 0.00 0.08 1.60 0.26 97.35

47.78 3.28 16.88 9.60 0.22 5.95 7.25 3.02 2.98 0.00 0.26 1.80 0.40 99.42

46.83 2.99 16.40 10.43 0.21 5.74 7.09 3.07 3.16 0.00 0.00 1.59 0.17 97.68

47.27 3.35 16.97 9.96 0.24 5.60 7.37 2.70 3.10 0.00 0.15 1.55 0.20 98.46

47.17 3.23 15.75 10.31 0.11 6.27 7.48 2.80 3.48 0.05 0.07 1.58 0.39 98.69

46.32 3.45 16.40 10.04 0.00 6.28 6.93 2.76 3.54 0.14 0.48 1.45 0.41 98.20

46.62 3.62 16.35 10.28 0.09 6.22 6.92 2.48 3.59 0.06 0.01 1.60 0.34 98.18

(Data from Lin 1994, 1998)

306

12 Experimentally Generated Pseudotachylyte

Table 12.5. EDX-determined chemical compositions of glass matrices (Cg1– Cg9) within pseudotachylyte generated experimentally from coarse-grained granite samples wt%

Cg1

Cg 2

Cg 3

Cg 4

Cg 5

Cg 6

Cg 7

Cg 8

Cg9

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 SeO Total

65.66 0.09 18.75 0.94 0.00 0.00 1.02 5.81 5.00 0.00 0.00 0.35 0.60 98.12

65.62 0.20 19.12 2.41 0.00 0.30 1.02 4.02 6.83 0.00 0.10 0.68 0.20 100.50

66.24 0.31 17.26 1.73 0.00 0.04 0.85 4.09 6.29 0.07 0.08 9.11 0.00 97.07

64.18 0.09 18.49 2.18 0.04 0.30 1.02 3.98 6.24 0.09 0.00 0.36 0.19 97.36

70.27 2.99 16.56 0.71 0.05 0.18 0.71 4.29 6.20 0.09 0.04 0.54 0.52 100.15

67.40 0.26 17.86 1.01 0.05 0.25 0.93 4.22 7.15 0.05 0.06 0.37 0.62 100.15

68.36 0.32 17.23 1.29 0.02 0.21 0.90 4.38 5.59 0.06 0.03 0.48 0.53 99.50

47.78 11.08 13.67 7.82 1.49 0.38 1.74 2.79 4.67 0.03 0.04 0.68 0.50 92.50

46.72 11.31 13.26 8.23 1.56 0.22 1.73 2.73 4.47 0.04 0.00 1.66 0.52 91.46

(Data from Lin 1994, 1998).

Table 12.6. EDX-determined chemical compositions of glass matrices (Mg01– Mg09) within pseudotachylyte generated experimentally from medium-grained granite wt%

Mg01

Mg02

Mg03

Mg04

Mg05

Mg06

Mg07

Mg08

Mg09

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 SeO Total

61.12 0.51 20.52 3.72 0.18 0.00 1.02 3.71 4.36 0.00 0.00 0.34 0.00 98.19

64.53 0.17 20.37 1.52 0.02 0.30 1.02 5.00 3.67 0.01 0.00 0.14 0.33 99.65

65.51 0.14 20.68 1.74 0.08 0.04 0.85 4.98 3.06 0.0 0.10 0.28 0.07 100.83

61.17 0.42 19.78 3.80 0.01 0.30 1.02 3.48 4.51 0.00 0.00 0.28 0.10 97.90

63.12 0.48 20.54 3.29 0.00 0.18 0.71 3.92 3.40 0.05 0.03 0.29 0.00 98.26

63.10 0.35 21.39 3.65 0.01 0.25 0.93 4.54 3.26 0.03 0.05 0.26 0.00 100.91

60.81 0.49 20.58 3.76 0.08 0.21 0.90 3.56 5.14 0.00 0.06 0.37 0.11 98.63

67.31 0.29 17.15 2.65 0.10 0.38 1.74 3.59 3.40 0.06 0.06 0.49 0.59 99.02

41.13 1.86 20.21 18.45 0.74 8.70 2.20 1.58 3.16 0.03 0.00 0.15 0.17 98.37

(Data from Lin 1994, 1998)

12.5 Chemical Composition Data

307

Table 12.7. EDX-determined chemical compositions of glass matrices (Fg01-Fg09) within pseudotachylyte generated experimentally from ﬁne-grained granite wt%

Fg01

Fg02

Fg03

Fg04

Fg05

Fg06

Fg07

Fg08

Fg09

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 SeO Total

65.70 0.41 17.48 2.48 0.04 0.44 2.56 4.18 3.64 0.00 0.01 0.05 0.04 97.03

61.58 0.45 19.98 3.56 0.05 0.79 2.85 3.88 4.58 0.01 0.01 0.06 0.06 97.86

63.33 0.47 17.87 3.85 0.00 0.69 2.57 3.70 4.40 0.01 0.00 0.05 0.07 97.21

62.43 0.37 19.32 3.20 0.05 0.74 2.56 4.00 4.93 0.01 0.00 0.04 0.04 97.69

61.26 0.48 19.82 3.38 0.18 0.60 2.82 3.86 4.86 0.01 0.00 0.04 0.06 97.36

61.73 0.52 19.97 3.61 0.08 0.57 2.91 3.81 4.40 0.00 0.00 0.07 0.04 97.71

64.02 0.25 19.93 1.95 0.01 0.17 3.21 4.76 3.43 0.00 0.00 0.01 0.02 97.76

64.86 0.52 18.61 3.59 0.03 0.72 2.42 3.26 4.46 0.00 0.02 0.06 0.02 98.57

50.07 1.04 19.61 11.98 0.48 1.11 3.87 3.40 2.76 0.00 0.01 0.05 0.03 94.41

(Data from Lin 1994, 1998)

of the host rock for ﬁne-grained and medium-grained gabbro samples, and are locally similar to those of plagioclase (Tables 12.3 and 12.4; Fig. 12.17) and pyroxene found within ﬁne-grained and medium-grained gabbro samples. This indicates that some of the melts that formed via the fusion of plagioclase within medium-grained gabbro samples were not contaminated by other melts of contrasting chemical composition prior to being quenched into glass. 12.5.2 Granite Samples The glassy matrices of the three types of granite samples all have SiO2 contents that are lower (by ∼2–30 wt%) than that of the host granite (Tables 12.5–12.7; Figs. 12.18 and12.19). The matrix compositions of the coarse-grained granite are mainly concentrated in two areas of the plot shown in Fig. 12.18: one is similar to the composition of pyroxene within the host rocks (Table 12.5, columns Cg8–Cg9), and the other is similar to that of feldspar composition (Table 12.5, columns Cg1–Cg7). No pure silica matrix was found in the run samples. The data for the oxides SiO2 , Al2 O3 , FeO, MgO, CaO, Na2 O, and K2 O also clearly demonstrate that both the original pyroxene and feldspar melted but remained largely unmixed (Figs. 12.20 and 12.21). There is also a relatively large variation in the chemical compositions of the matrices of run products for ﬁne- and medium-grained granite samples (Tables 12.6 and 12.7; Fig. 12.19). The Al2 O3 content is similar to that of plagioclase within the host granite. Generally, the FeO and MgO contents are

308

12 Experimentally Generated Pseudotachylyte

Fig. 12.17. SiO2 –Al2 O3 –“All other components” diagram showing compositional variation in the glassy matrix generated in ﬁne-grained gabbro, medium-grained gabbro, and host gabbro. Pl and Py: plagioclase and pyroxene included in the host gabbro (After Lin and Shimamoto 1998)

approximately 2.0 wt% higher than those in the bulk composition of the host granite samples. The CaO, Na2 O, and K2 O contents are also a little higher than those of the host rocks (Figs. 12.20 and 12.21). 12.5.3 Albitite–Quartz and Anorthosite–Anorthosite Pairs The chemical compositions of the albitite, quartzite, and anorthosite samples, together with their normative mineral compositions, are shown at the bottom of Table 12.8. In a phase diagram that shows eutectic melting of the NaAlSi3 O8 –SiO2 system (Fig. 12.22), the silica contents of the quartzite and albitite are shown by the lines AB and CD, respectively. In Fig. 12.23, the anorthite content of the anorthosite is shown by the line HC, which represents the solid solution of plagioclase.

12.5 Chemical Composition Data

309

Fig. 12.18. SiO2 –Al2 O3 –“All other components” diagram showing compositional variation in the glassy matrix generated in coarse-grained granite and the host granite used in the experiments

Fig. 12.19. SiO2 –Al2 O3 –“All other components” diagram showing compositional variation in the glassy matrix generated in ﬁne-grained and medium-grained granc ite. (After Lin and Shimamoto, 1998). 2007, with kind permission from Elsevier Science Ltd

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12 Experimentally Generated Pseudotachylyte

Fig. 12.20. Variations in oxide concentrations in the glass matrix within experimentally generated pseudotachylyte derived from ﬁne-grained granite (Modiﬁed from Lin 1991, Lin and Shimamoto 1994)

12.5 Chemical Composition Data

311

Fig. 12.21. Variations in oxide concentrations in the glass matrix within experimentally generated pseudotachylyte derived from coarse-grained granite. (After Lin 1991, Lin and Shimamoto 1994)

312

12 Experimentally Generated Pseudotachylyte

Table 12.8. Bulk chemical compositions of albitite from Itoigawa, Niigata Prefecture, Japan, greenish quartzite from Brazil, and anorthosite from the Outer Hebrides, Scotland, as used in high-velocity frictional melting experiments wt%

Albitite

Quartzite

Anorthosite

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O P2 O5 Total

81.68 0.02 11.93 0.43 0.01 0.18 0.27 6.82 0.04 0.05 101.43

96.87 0.09 1.79 2.72 0.26 0.59 0.28 0.03 0.20 0.09 102.92

51.52 0.16 29.71 1.26 0.02 0.63 14.35 3.12 0.20 0.02 100.99

Normative mineral compositions: Albite 56.9 0.3 Anorthite 1.4 0.8 Quartz 40.2 89.6 Others 1.5 9.3

25.7 67.9 0.5 5.9

(Data from Lin 1991; Shimamoto and Lin 1994)

Fig. 12.22. NaAlSi3 O8 –SiO2 binary diagram of the glass matrix within experimentally generated pseudotachylyte derived from albitite–quartzite pairs. The numbers of (25) and (2) represent the number of measured data. The total number of analyzed data is 27, of which 25 are composed of albite, the same as that of the host albitite (Modiﬁed from Lin 1991; Shimamoto and Lin 1994)

12.5 Chemical Composition Data

313

Fig. 12.23. NaAlSi3 O8 –CaAlSi2 O8 solid system diagram of the glass matrix within experimentally generated pseudotachylyte derived from anorthosite pairs (After Lin 1991; Shimamoto and Lin 1994)

The chemical compositions of molten material from two sets of specimens are shown in Tables 12.9 and 12.10 and Figs. 12.22 and 12.23. Although fused glass that contained large clasts was avoided in performing the chemical analyses, analyzed portions might have contained very small fragments. Table 12.9 lists 9 representative results from 27 chemical analyses of the frictional melt from the albitite–quartzite pair. Except for two cases (columns 8 and 9), the silica content of the molten material is less than several percent. This is clearly shown by the gray area marked by the number (25), which indicates the number of data points, near the left-hand side of Fig. 12.22. Two data points located near the eutectic point represent the compositions for columns 8 and 9 (Table 12.9; Fig. 12.22). Molten material derived from the anorthosite specimen has an An content that is nearly identical to the original composition of the anorthosite (Table 12.10; Fig. 12.23).

314

12 Experimentally Generated Pseudotachylyte

Table 12.9. EDX-determined chemical compositions of glass matrices (Ab1–Ab9) within pseudotachylyte generated experimentally from albitite and quartzite samples. The 9 data sets listed here are representative examples from a total of 27 analyses wt%

Ab1

Ab2

Ab3

Ab4

Ab5

Ab6

Ab7

Ab8

Ab9

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 SeO Total

68.12 0.03 20.29 0.10 0.00 0.00 0.58 10.51 0.00 0.01 0.00 0.38 0.72 100.74

67.91 0.00 19.34 0.06 0.00 0.02 0.03 11.16 0.00 0.00 0.05 0.24 0.75 99.56

67.74 0.00 19.40 0.06 0.01 0.00 0.08 10.92 0.02 0.00 0.07 0.42 0.64 99.36

67.98 0.00 20.39 0.00 0.07 0.00 0.39 11.00 0.00 0.01 0.00 0.45 0.53 100.82

67.92 0.03 19.82 0.04 0.00 0.00 0.24 10.51 0.00 0.03 0.00 0.50 0.56 99.65

67.98 0.05 20.05 0.04 0.00 0.02 0.45 10.68 0.00 0.05 0.02 0.53 0.89 100.76

67.63 0.00 19.96 0.24 0.12 0.02 0.61 10.29 0.11 0.01 0.02 0.25 0.52 99.78

74.23 0.00 14.78 0.30 0.00 0.00 0.32 7.04 0.09 0.00 0.00 0.16 0.16 97.08

76.51 0.10 12.76 0.04 0.00 0.07 0.26 8.29 0.15 0.19 0.22 0.57 0.57 99.73

90.7 6.7 2.6 93/7

91.3 6.1 2.6 94/6

88.2 6.6 5.2 93/7

62.0 33.3 4.7 65/35

63.2 30.4 6.4 68/32

Normative mineral compositions: Albite 90.0 95.8 94.1 93.3 Quartz 6.7 3.1 4.4 4.2 Others 3.3 1.1 1.5 2.5 Ab/Qz 93/7 97/3 96/4 96/4

(Data from Lin 1991; Shimamoto and Lin 1994) Table 12.10. EDX-determined chemical compositions of glass matrices (An1–An9) within pseudotachylyte generated experimentally from anorthosite samples wt%

An1

An2

An3

An4

An5

SiO2 TiO2 Al2 O3 FeO∗ MnO MgO CaO Na2 O K2 O Cr2 O3 NiO P2 O5 SeO Total

48.79 0.51 29.25 2.82 0.25 1.93 13.41 2.43 0.24 0.03 0.28 0.00 0.31 100.25

49.75 0.34 29.21 2.66 0.00 1.55 13.04 2.85 0.31 0.04 0.00 0.21 0.00 99.96

48.99 51.57 49.72 49.73 49.93 66.88 0.63 0.00 0.33 0.31 0.37 0.01 27.81 31.67 30.60 30.27 28.75 20.11 3.39 0.11 1.78 1.99 2.61 0.00 0.14 0.00 0.04 0.07 0.00 0.00 2.40 0.00 0.97 0.88 1.44 0.52 12.54 10.42 13.28 13.36 13.48 0.00 3.13 3.50 2.97 2.64 3.31 11.14 0.42 0.01 0.15 0.20 0.39 0.02 0.25 0.00 0.00 0.09 0.07 0.00 0.00 1.74 0.21 0.29 0.09 0.00 0.37 0.05 0.13 0.31 0.09 0.57 0.05 1.76 0.71 0.00 0.17 0.91 100.12 100.83 100.89 100.14 100.70 100.16

Plagioclase composition: An/Ab 76/24 73/27 70/30 64/36 72/28 (Data from Lin 1991; Shimamoto and Lin 1994)

An6

75/25

An7

70/30

An8

0/100

An9 49.24 0.04 30.75 1.55 0.00 0.77 13.61 2.78 0.18 0.14 0.14 0.03 0.42 99.65 74/26

12.6 Discussion

315

12.6 Discussion 12.6.1 Vein Geometry Most melt-origin pseudotachylytes are strikingly similar in terms of a dense and aphanitic appearance, their occurrence as irregular injection veins intruded into country rocks as both simple and complex networks, and vein widths in the order of several millimeters to several centimeters. Sibson (1975) classiﬁed fault-generated pseudotachylyte veins into two fundamental types: i) fault veins, which occur along markedly planar shear fractures on which the pseudotachylyte was generated, and ii) injection veins that were intruded into the country rocks and commonly appear as dilational veins that record no lateral oﬀset (see Chap. 4). The diﬀerent styles of fault veins of variable thicknesses and injection veins that form complex vein networks and pinch-andswell structures have been successfully reproduced in high-velocity frictional melting experiments (Fig. 12.8; e.g., Lin 1991, 1992; Lin and Shimamoto 1998; Tsutsumi 1999). Injection veins of crushing-origin pseudotachylyte and fault gouge (see Chap. 10) are similar to those of melt-origin pseudotachylyte in geometric shape. Field occurrences, powder X-ray diﬀraction patterns, chemical composition data, and size–distribution patterns of fragments reveal that these injection veins are derived from the same source material as the fault veins; i.e., the granitic country rock (Lin 1996, 1997a). On this basis, it is suggested that injection of the veins occurred via the rapid ﬂuidization of ﬁne-grained fragments that were generated in the shear zone during seismic faulting. It has also been demonstrated that experiment-derived pseudotachylyte injection veins contain approximately 50 wt% fragments mixed with melt material. These experimental results support the proposal that mixtures of fragments and melt were injected into coseismic voids by the rapid intrusive-like spraying of a gas–solid–liquid system (see Chap. 10). 12.6.2 Melting Textures The evidence obtained from natural pseudotachylytes (e.g., Sibson 1975; Lin 1994a, b) and experimentally generated pseudotachylytes (e.g., Spray 1987, 1988, 1992, 1993; Killick 1990; Lin 1991) demonstrates beyond doubt that melt-origin pseudotachylytes are able to form by frictional heating at shallow depths within the crust. The vein geometry and structures developed in the experiments clearly reveal that frictional melting can occur under normal stresses as small as 1.0–1.5 MPa and high slip rates of 1.0–2.0 m/s, corresponding to the conditions of seismic faulting at shallow depths within the crust (Lin 1991; Lin and Shimamoto 1998). The vesicles and angular to subangular and rounded fragments that have been reported from natural pseudotachylyte are also observed in experimentally generated pseudotachylyte. The degree of rounding of the fragments

316

12 Experimentally Generated Pseudotachylyte

within pseudotachylyte veins is a possible indicator of melting (see Chap. 7). The degree of rounding of quartz and plagioclase fragments from the wellknown melt-origin pseudotachylytes described by Sibson (1975), Toyoshima (1990), and Lin (1991, 1994a) have been measured by Lin (1997d, 1999b), as well as the degree of rounding within crushing-origin pseudotachylyte and cataclastic rocks. Fragments observed within experimentally generated pseudotachylytes also show a high degree of rounding (Rd >0.4; Fig. 12.9). This shows that the rounded fragments formed by frictional melting during the frictional melting experiment. Fault veins contain a higher proportion of fragments than injection veins (Fig. 12.9). This may provide a criterion for distinguishing fault veins from injection veins in natural pseudotachylyte (Magloughlin and Spray 1992). Microlitic structures found in natural pseudotachylyte (e.g., Lin 1994b) and experimentally generated pseudotachylyte (Spray 1988) have yet to be simulated in the experiment carried out by Lin and Shimamoto (1998). It may be that the cooling rate was too rapid to enable the formation of microlites at low normal stress. 12.6.3 Non-equilibrium Melting Processes The frictional melt temperatures attained during pseudotachylyte formation are commonly estimated by assuming conditions of chemical equilibrium during frictional melting. If frictional melting occurs in chemical equilibrium, minerals will begin to melt at their established melting points. As shown in Fig. 12.22, if frictional melting of a albitite–quartzite pair occurs under conditions of chemical equilibrium, the composition of the initial melt that forms at the lowest temperature must be close to that for the eutectic point (E). Even if only albite melts in the experiments, the melting is not due to albite-onalbite frictional melting because albite is not sliding upon albite in albitite– quartzite experiments. The composition of the initial melt of a solid solution under equilibrium melting is diﬀerent from the composition of the solid solution (plagioclase); this provides another test of whether melting occurs under conditions of chemical equilibrium. In albitite–quartzite experiments, albite is always in sliding contact with quartz; consequently, it is inevitable that melt forms with the eutectic composition given equilibrium melting. Thus, the compositions of the molten materials in Fig. 12.22 and Table 12.9 clearly indicate that frictional melting did not occur under conditions of chemical equilibrium. Among 27 chemical analyses, only two have a composition that is close to the eutectic point, while another sample contains more silica than that at the eutectic point. The reason for these two chemical compositions is not clear, but it could be due to the presence of minute fragments of quartz mixed with the molten material or the local formation of quartz melt associated with quartz-on-quartz sliding across the fault. Despite these uncertainties, the majority of the results indicate that albite, which has a lower melting temperature than quartz, melted during frictional melting that

12.6 Discussion

317

occurred in the albitite–quartzite experiment. This ﬁnding is consistent with the concept of selective melting (Lin 1991; Lin and Shimamoto 1998). Alternatively, if plagioclase within the anorthosite sample melts under equilibrium conditions, the chemical composition of the initial melt must be given by point B in Fig. 12.23; however, the composition of the molten anorthosite material is similar to the original composition of the anorthosite sample (Fig. 12.23; Table 12.10), indicating that frictional melting did not take place under conditions of chemical equilibrium. The case shown in column An8 in Table 12.10 has a composition that is close to that of pure albite. The reason for this composition is not clear, but albite might have formed locally in the anorthosite during a retrograde stage. In any case, a melt of albitic composition does not indicate equilibrium melting. Non-equilibrium melting is thus conﬁrmed for the two critical examples. This ﬁnding should be kept in mind when estimating paleotemperatures on the basis of the stability ﬁelds of minerals under the assumption of equilibrium conditions and homogeneous temperatures. The melting temperatures in the above experiments are 1118◦C for albite and approximately 1550◦ C for anorthosite (Figs. 12.22 and 12.23). These are thought to represent minimum temperature because overstepping of the melting temperature is possible if heat dissipates from the fault zone and the rate of heat absorption via melting due to latent heat is insuﬃcient to account for the heat produced by viscous shearing of the molten layer. The distribution of temperature within a fault zone is likely to be heterogeneous, as seismic fault shearing deformation is commonly concentrated within narrow zones of less than 3 mm in width (Lin et al. 2004) and because it takes some time before the temperature is homogenized. In the experiments involving ﬁne-grained gabbro, the chemical compositions of the matrices are heterogeneous but the SiO2 concentrations of glassy matrices are 5–10 wt% lower than those of the host rock. This observation can be explained by the fact that the original quartz crystals within the host rock were less susceptible to melting; instead, SiO2 -poor hydrous minerals were preferentially melted. Most of the analyzed glass matrices have contrasting chemical compositions to the bulk compositions of the host rocks, but they are similar to those of the original plagioclase and pyroxene contained within medium-grained gabbro samples (Fig. 12.17). This ﬁnding demonstrates that the plagioclase and pyroxene crystals melted without undergoing chemical reactions with other minerals. As most of the examples of pseudotachylyte reported to date are hosted within granitic rocks, frictional melting experiments involving granitic rock provide important information in terms of understanding the melting process of natural fault-related pseudotachylyte. These types of frictional melting experiments have also been performed using Westerly granite (Spray 1992, 1993), revealing that the SiO2 contents of the melts are approximately 5 wt% lower than those of the protolith (Spray 1993). It is not clear, however, whether quartz crystals were less susceptible to melting than feldspar, as the 5 wt%

318

12 Experimentally Generated Pseudotachylyte

discrepancy in SiO2 content may simply reﬂect the preferential melting of SiO2 -poor hydrous minerals. The glassy matrices have SiO2 concentrations that are ∼3–30 wt% less than those of the host granite (Figs. 12.18 and 12.19). Rapid friction-related heating induces the melting of plagioclase via the selective fusion of albite components (Johannes 1989). For rocks melted at atmospheric pressure, it has also been observed that biotite and amphibole are the ﬁrst materials to melt, with the residual materials being quartz and feldspar (Bossi`ere 1991). Analyses of the mineral contents of rock fragments within pseudotachylyte generated experimentally from a granite sample reveal that only 20–50% of the original quartz crystals and 60–70% of the original feldspar crystals were melted during the frictional melting experiments. These results indicate that quartz is the most resistant of the granite-forming minerals to melting, with biotite being the least resistant, and feldspar minerals having an intermediate resistance. Thus, the observed SiO2 depletion within pseudotachylyte glass is likely to reﬂect the selective melting of constituent minerals of host granitic rocks during frictional heating. The experiments described above successfully reproduced the chemical compositions of glass matrices reported by Lin (1994a). Many studies have reported that frictional melting is not an equilibrium phenomenon (Lin 1991; Lin et al. 1992; Spray 1993; Shimamoto and Lin 1994). Experimental results and chemical composition data for natural pseudotachylytes reveal that frictional melting mainly occurs via the preferential melting of low-melting-point minerals under conditions of chemical disequilibrium rather than by total or partial melting. Thus, predictions of the mechanical properties of faults during frictional melting or pseudotachylyte formation must be based on non-equilibrium analyses of the thermal evolution within a fault zone. Once a molten layer forms along a fault, viscous shearing of the melt is the primary heat source, and the fault strength is most likely controlled by the viscosity and thickness of the molten layer. Overstepping of the melting temperature would act to reduce the melt viscosity and the dynamic fault strength or shear strength; however, this reduction in the shear stress would in turn act to reduce the rate of heat production. Feedback mechanisms are therefore likely to operate among the rate of heat production, rate of melting, shear resistance of the melt, heat conduction away from the fault, and rate of melt loss from the fault zone. The dynamic properties of a seismic fault during frictional melting can only be understood once such highly nonlinear and non-equilibrium processes are fully considered and understood. 12.6.4 Melting Temperature The melt temperature that accompanies the formation of pseudotachylyte has previously been estimated using microlite geothermometry (e.g., Maddock 1983; Toyoshima 1990; Lin 1994b) and the chemical composition of the matrix (e.g., Wallace 1976; Lin 1991, 1994a). The estimated melt temperatures

12.6 Discussion

319

of natural and experimental pseudotachylytes range from 750 to 1400◦C, as shown in Table 8.7. These estimates are based on an assumption of equilibrium melting processes. As stated above, frictional melting generally occurs under non-equilibrium conditions; consequently, these estimated temperatures are likely to represent underestimates. Temperatures along simulated faults increase above 1000◦C within several seconds of the start of frictional experiments and develop a highly heterogeneous distribution (Tsutsumi and Shimamoto 1994). Powder X-ray diﬀraction analyses reveal that all of the micas in the analyzed granite samples melted along with some of quartz crystals (Lin 1991; Lin and Shimamoto 1998). This means that for high-velocity friction experiments of granite samples conducted under a low normal stress (approximately 1 MPa) over a short time interval (<30 seconds), the melting temperature is higher than the melting point of mica (650◦ C) and locally higher than that of quartz (1723◦ C; Fig. 8.6). Melt temperatures have also been directly measured in past experiments by setting a temperature sensor in the cylindrical specimen and placing it in contacting with the frictional surface (end face of the cylindrical specimen) (Tsutsumi and Shimamoto 1994). The results of this approach indicate that the frictional melting temperature was at least 1169◦ C in a granite experiment and 1145◦C in a gabbro experiment (Tsutsumi and Shimamoto 1994). The melting temperature estimated in frictional experiments involving gabbro is higher than 1100–1550◦C, which overlaps with the melting points of feldspar and pyroxene. 12.6.5 High-Velocity Slip Weakening High-velocity frictional experiments involving gabbro samples have revealed the mechanism of slip weakening during frictional melting (Tsutsumi and Shimamoto 1997a; Hirose and Shimamoto 2005). Steady-state friction increases slightly with increasing slip rate on the order of 100 mm/s; this indicates a velocity-strengthening mechanism at a slow slip-rate that decreases markedly at higher slip rates of 43–140 cm/s. This reveals velocity-weakening at higher slip rates, as with the seismic slip rate (Tsutsumi and Shimamoto 1997a). Hirose and Shimamoto (2005) reported two stages of slip weakening during high-velocity experiments: one after the frictional peak and the other following the second peak friction. The ﬁrst weakening is associated with ﬂash heating and incipient melting at an asperity junction, while the second weakening is due to the formation of a molten layer and its subsequent growth along the simulated fault. The two stages of weakening are separated by a pronounced strengthening regime in which the melt patches grow into a thin, continuous molten layer at the second peak friction. The frictional coeﬃcient μ gradually decreases from ∼0.9 toward a residual or steady-state friction at μ =∼0.6. These experimental results suggest that growth of the molten zone is the main mechanism of slip weakening during seismic frictional melting.

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Index

Abrasion, 1, 139, 157 Acicular, 106, 110, 113–114, 118, 120, 122, 135, 187, 190, 196, 204, 275–276 Active fault, 12, 17–18, 41–43, 48, 148, 155, 226, 230 Advection, 174 Albitite, 290–291, 293, 308, 312–314, 316–317 Alpine Fault, 42, 47–48, 54, 67, 142, 170, 172 Alteration, 7, 14, 75, 88, 90, 130, 136, 245, 254 Amorphous material, 90, 225, 240–341, 253–254 Amphibolite facies, 136, 178–179, 206, 219–221 Amygdule, 12, 14, 84–89, 97–99, 102, 169, 171, 175, 225, 254, 283 Amygdule structure, 12, 98 Andesitic melt, 99, 101, 103 Anorthite, 135, 219, 290, 308, 312 Anorthoclase, 112, 119, 122–123, 133, 135, 137 Anorthosite, 290–291, 293, 308, 312–314, 317 Arima-Takatsuki Tectonic Line, 42, 230 Aseismic deformation, 204, 210 Aseismic regime, 224 Assimilation, 65, 152 Bagnold eﬀect, 83 Biotite, 12, 33–36, 39–42, 80, 109, 111–112, 115, 118–121, 122–126,

128, 130, 133, 135–137, 151, 167, 170, 174, 176, 196, 199, 203–206, 211, 217, 219, 240–241, 245–246, 249–250, 254, 269, 275, 290–291, 298–299, 318 Bolivian Mw 8.3 earthquake, 179 Bow-tie, 111–112, 116 Box-skeletal, 109 Branching, 50, 109–110, 115, 296 Breccia, 9, 12, 19–20, 23, 25–27, 29–30, 32, 35–37, 44, 48, 140–141, 148, 150, 153, 156, 225–226, 230, 232–234, 243, 255, 258, 266 Brittle–plastic transition zone, 45, 178–179, 220, 224 Brittle regime, 18–19, 45–46, 188, 208, 220 Bubble, 84–85, 89, 256, 266, 298 Byerlee’s Law, 43, 46 Calcite, 88, 102, 225, 234–236, 238, 243–245, 249–250, 252, 256, 258, 260–261, 263 Carbon, 259–261 Carbonate material, 244, 252, 258–261, 263, 275–276 Carbonate vein, 253, 258–260 Cataclasite, 19, 25–27, 30, 32–36, 39–40, 42, 44–45, 48, 62, 148, 150, 153, 177–178, 181, 209, 211, 217, 225–227, 232, 240–241, 245 Cataclasite-related pseudotachylyte (C-Pt), 177–178, 185, 187–189, 204

342

Index

Cataclastic rock, 25–27 Cataclastic vein, 14, 72, 226, 245, 256 Cataclastic veinlet material, 256 C-foliation, 40 Chain-skeletal, 109 Chelungpu Fault Zone, 19 Chi-Chi Mw 7.6 (Taiwan) earthquake, 261 Chiufener-Shan landslide, 266, 269–273, 280 Chlorite, 43, 88, 102–103, 245–246 Chlorite vein, 136 Circular, 86, 89, 111–112, 115–116, 169, 277, 287–289, 293–294, 298 Clast, 19, 21, 30–38, 40–42, 139, 162, 179, 235, 243, 245, 258, 277, 313 Clastic-type pseudotachylyte, 225 Clinopyroxene, 127, 137, 187 Coesite, 197 Cohesion, 23, 25–26 Cohesive cataclasite, 19, 26–27, 30, 32, 39, 44 Cohesiveness, 23, 30 Collapse, 89, 98, 174–175, 234, 250, 272 Collisional orogenic belt, 47, 178–179, 197, 206–207, 216, 219 Collision zone, 47 Comminution, 150, 225, 242, 254 Compound spherulitic, 112, 120 Concordant vein, 60 Conglomerate, 9–10, 139–141 Continental crust, 40–42, 44–46, 218–220, 222–223 Contractional duplexe, 65 Contractional fault, 65 Convection, 174 Cooling joint, 57–58 Coseismic displacement, 215, 226, 263 Coseismic shear zone, 68 Coseismic slip, 2, 45, 222, 257 Coseismic surface rupture, 48, 50, 68, 70, 257, 269 Coulomb friction law, 43 Course-feathery, 109 Course-scoplitic, 109 Crack-ﬁll vein, 226, 234–237, 243–244, 245–250, 252, 258 Cross-shaped, 106, 110 Crotch, 109–110

Crushing-origin pseudotachylyte, 8, 13, 15, 32, 48, 148–149, 155–156, 225–230, 237–242, 245, 250, 253–255, 257, 315–316 Cryptocrystalline-type, 77 Crystalline material, 93, 95, 245, 301, 303 Crystal-plastic deformation, 13, 18, 21, 23–24, 33, 39, 40, 44–45, 183, 186–187, 190, 192, 214, 217, 218, 221, 223 C’-surface, 32–34, 37–38, 40–41, 206 C-surfaces, 32–34, 36, 40, 42, 206 Cylinder, 87, 288–289 Dabieshan collisional zone, 45 Dahezhen Shear Zone, 45, 179, 184, 197–199, 204, 208–217, 219–221 Dendritic group, 107, 109–112 Dendritic-skeletal, 109, 128 Dendritic-spherulite, 112, 115 Devitriﬁcation, 7, 12–13, 90, 137 Devolatilization, 98–99 Diﬀusion model, 167, 263 Dike-like rock, 5, 8 Dilatancy, 41, 261, 263 Dilatation, 237 Dilational fracture, 62 Discordant vein, 8, 60 Disintegration, 139, 158 Dissolution, 155, 255 Distortional strain energy, 222–223 Drill core, 17–19, 21, 28–29, 35, 37, 235–238, 243–246, 250–252 Dynamical recrystallization, 221, 223 Eclogite band, 197 Eclogite facies, 178, 197, 223 Elastic-frictional (EF), 44 Equilibrium, 172–173, 219, 258–259, 281, 283–284, 316–319 Equivalent slip rate, 289–290 Escape structure, 89 Eutectic point, 313, 316 Extensional reservoir, 65 Fan, 111–112, 116 Fault-fracture network, 225, 234–236, 258

Index Fault-fracture zone, 30, 48, 50–51, 63–64, 73, 226, 231 Fault gouge, 12, 19–20, 23, 30–31, 35–39, 42–44, 48, 128, 130–131, 135, 148–150, 153, 155–156, 225–226, 230–234, 238–240, 242, 244–246, 250, 254–257, 260, 266, 315 Fault model, 17, 40, 43–45, 223 Fault-parallel slip surface, 65 Fault plane, 2, 8, 13, 15, 19, 29–33, 37, 50–52, 60, 64, 66, 70, 72, 96–97, 157, 173, 175, 182, 199, 212, 226–233, 254–257, 263, 265, 283 Fault-related pseudotachylyte, 1–2, 8, 14–15, 47–48, 78, 91, 97, 105, 177, 179, 280, 283, 291, 317 Fault-related rocks, 6, 8, 17–18 Fault rocks, 3, 12–13, 15, 17, 18–40, 43–44, 98, 135, 139, 143, 148–149, 151–152, 220, 223, 225, 265, 283, 286 Fault vein, 60–65, 68, 70–73, 140–141, 199–200, 226–228, 230–232, 242, 245–246, 255, 296, 315–316 Fault zone, 1–3, 7–9, 12–13, 15, 17–19, 21, 23, 25–27, 29–30, 32, 36, 40–48, 50–51, 64–65, 68, 71–72, 78, 97–98, 105, 122, 135–136, 139–140, 149–150, 155, 157, 159–160, 169, 172, 174–175, 177, 181–182, 213, 215–218, 222–224, 225–232, 234–235, 239–240, 244–245, 254–258, 260–263, 284, 287, 291, 317–318 Feldspar, 9, 32–33, 36, 39–41, 45–46, 78, 80, 92–93, 95–96, 102, 106, 109, 111, 115, 118–123, 126, 128, 136–137, 147, 151, 153, 155, 166–167, 171, 187, 190, 194, 196, 199, 204–206, 209, 211, 216–220, 222–223, 237, 245, 275–277, 281, 291, 298–299, 301, 303, 307, 317–319 Fine-feathery, 109–110, 115 Fine-grained material, 12, 15, 23, 38, 48, 60, 62, 64–65, 75, 79, 84, 88, 195, 204, 209, 234–237, 243–244, 252, 254–256, 258

343

Fine-grained matrix, 6, 9, 13, 28, 30, 33, 35, 41, 68, 75, 78, 98, 140, 145, 148, 152, 159, 163–164, 167–168, 187, 191, 202, 204, 208, 211, 225, 240–241, 272 Fine-scoplitic, 109 Fir-like, 110, 115 Flinty crush-rocks, 6, 9 Flow fold, 82 Flow streak, 78, 81–87, 106, 275, 277 Flow structure, 9, 14–15, 38–39, 78, 80–86, 90, 97, 135–136, 141, 151, 188, 190, 196, 204, 214, 225, 266, 269, 274–275, 277, 295, 300 Flow velocity, 82–84 Fluid diﬀusion, 41, 261, 263 Fluidization, 12, 15, 254–256, 315 Foliated cataclastic rock, 19, 21, 27, 31–40 Foliation, 19–24, 31, 33–35, 37–42, 44, 46, 57, 60, 85, 142, 179, 184, 186–192, 194–196, 198–201, 204, 207–210, 214–217, 219, 233, 235, 238, 242 Fossil earthquake, 1, 225, 234, 266 Fractal, 145, 148, 150 Fractal dimension, 145, 148, 150–151 Fracture-ﬁll vein, 243, 250 Fragment, 62, 72, 81–83, 109, 112–113, 115–116, 120, 126, 139–140, 145–147, 149–150, 152–155, 157, 167, 202–203, 217, 275, 297 Fragmentation, 65, 145, 254 Fregon subdomain, 179, 181 Frictional coeﬃcient, 41, 43, 285, 289, 319 Frictional fusion, 105, 170, 175, 178, 265–266, 269 Frictional melt, 2, 13, 48, 60, 62, 64–66, 83–84, 86, 89, 92, 97–99, 105, 122, 127, 130–131, 145, 157, 159, 165, 170–171, 173, 175, 281, 283, 286, 289, 292–295, 300, 313, 316 Frictional melting, 2, 7–8, 12–15, 23, 26, 48, 92, 97–98, 105, 130–131, 148, 157–159, 170–177, 179, 212, 222, 254, 257, 265, 280, 283–287, 289–295, 298, 304, 312, 315–319 Frictional milling, 294

344

Index

Frictional regime (FR), 19, 44–45 Fuyun Fault, 15, 47–48, 50–52, 55, 57, 76–86, 91–92, 105, 109, 112, 122, 126, 128, 130, 135, 160, 169–172, 269 Fuyun M, 8.0 earthquake, 50 Gabbro, 172, 175, 287–288, 290, 292–295, 298, 300–302, 304–308, 317, 319 Garnet, 39, 80, 111, 119, 121, 126, 128, 131, 133, 137, 187, 194, 196–197, 203–205, 219 Gasiﬁcation, 97–98 Gas pressurization, 174–175 Gas–solid–liquid system, 255–256, 315 Gas–solid system, 83, 255–256 Generation zone, 51, 60, 62, 64–68, 70, 73, 100, 139–141, 147, 176, 181–182, 214, 295–297 Geothermal gradient, 40, 46, 136, 178, 212, 218–221, 258 Geothermometer, 171 Glass, 12–15, 48, 72, 75, 78–81, 90–99, 103, 106, 113, 136–137, 151, 162, 164–167, 169, 171–172, 175, 225, 237, 254, 265–266, 269, 274–280, 283, 300–303, 305–307, 310–314, 317–318 Glass material, 48, 90, 92, 97, 254, 275 Glass-type pseudotachylyte, 79–80, 93, 96, 97, 167 Glassy matrix, 7, 57, 72, 75, 78, 84, 91, 93, 96, 113, 136, 141, 159, 171, 277, 302, 304, 308–309 Globular, 111–112 Gole Larghe Fault Zone, 71–72 Granite, 9–10, 24, 28, 31–33, 35, 37–40, 93, 122, 126, 135–136, 141, 150, 163–164, 166, 179, 227–228, 237, 243–245, 269, 275, 287, 289–291, 293, 298–304, 306–311, 317–319 Granular, 106, 110, 128, 277 Granulite facies, 48, 178–179, 181, 185–186, 220, 223 Granulite-related pseudotachylyte (G-Pt), 178–179, 181, 183, 186, 189, 192–197, 199–203, 205, 219–220, 223

Greenschist facies, 46, 178, 212, 221 Grossular, 111, 121, 128, 130–131, 133, 137 Groundwater level, 41, 260–261, 263 Halo, 81, 89 Hardness high-velocity frictional experiment, 48, 97–98, 105, 172, 174–175, 257, 280, 284–291, 295, 319 Heterogeneous, 72, 75, 98–99, 159–160, 170, 210, 245, 284, 293, 299, 304, 317, 319 Hiddaka Metamorphic Zone, 172 High-velocity frictional melting, 15, 130–131, 172, 175–176, 212, 283–284, 289–290, 292–293, 298, 304, 312, 315 Holloysite, 43 Hornblende, 80, 109, 111–112, 114–115, 119, 121–122, 126–128, 130, 133, 135–137, 176, 199, 204, 206, 217–219, 249, 290, 298 Host rock, 6–8, 11, 13, 15, 23, 30, 62, 72, 93, 96, 112, 122–127, 137, 150, 153, 159–171, 175, 179, 181–184, 186, 190–192, 196, 204, 214, 225, 233, 235, 243, 245–246, 250, 255, 258, 260, 266, 269, 272, 278–279, 281, 283, 296, 298, 300–301, 303–304, 307–308, 317 Hyalomylonite, 6, 265 Hydration, 98, 174–175 Hydraulic friction, 286 Hydraulic pressure, 286–287 Hydrofracturing, 41 Hydrothermal ﬂuid, 254, 263 Igneous dike, 57 Iida-Matsukawa Fault, 8, 148–150, 152–153, 156, 226–227, 229–232, 239–242, 245–246, 254 Illite, 43 Ilmenite, 111, 119, 128, 130, 132, 196, 203, 205 Incohesive cataclastic rocks, 19, 23 India Mw 7.6 earthquake, 222 Inﬁltration of surface water, 257, 260–261, 263

Index Injection gouge vein, 245 Injection mylonite, 6 Injection vein, 9, 15, 29, 51, 60–64, 67, 73, 81, 99, 134, 141, 172, 175, 182, 187, 189, 195, 199, 204, 207, 211, 225–226, 228–232, 242, 245–246, 250, 254–256, 277, 293, 295–297, 315–316 Interlayered fault gouge, 232, 234, 238, 246 Interlayered pseudotachylyte, 233, 238–240, 246 Intracontinental earthquake, 41, 212, 223 Intracontinental fault zone, 41, 47 Isotherm, 157 Isotopic data, 258 Itoigawa–Shizuoka Tectonic Line Active Fault Zone, 226–229 Ivrea Verbano Fault Zone, 169 Kaolinite, 43 Kobe Mw 7.2 earthquake, 17, 226 K¨ ofels landslide, 265–266 Kunlun Mw 7.8 earthquake, 70, 222, 257 Landslide, 96, 265–267, 269, 271–273, 277, 280–281 Landslide-generated pseudotachylyte, 265–266, 280 Landslide-origin molten material, 265 Langtang Himayala landslide, 102, 143, 266–269, 271, 280 Lath-like, 106, 110, 114, 122, 187, 190, 196, 204, 276–277 Laumontite, 245, 249–250 Layered fault gouge, 232–234, 238, 246 Lewisian gneiss, 10 Liquid pressure, 99–100 Lithic clast, 139 Lithic fragment, 139 Lithic porphyroclast, 139 Lithoclast, 139 Lithology, 43, 140–141, 181 Lithostatic depth, 42 Lithostatic pressure, 98–100, 103, 280 Lubricant, 2, 98 Macpherson model, 103

345

Magnetic susceptibility, 82, 129, 134, 135 Magnetite, 88, 102, 111, 128–131, 205, 291 Magnus eﬀect, 83 Marine-origin ﬂuid, 260 Matrix, 6–7, 9–10, 12–14, 24–25, 28, 30, 33, 35, 39, 41, 57, 65, 68, 72, 75, 78–84, 86, 90–91, 93–94, 96–98, 106, 113, 118–119, 128, 136, 140–141, 145, 148, 151–152, 159–160, 162–171, 175, 187–188, 191, 195–196, 202, 204–206, 208–211, 214, 225, 238, 240–241, 254, 265, 272, 274–277, 280, 283, 300, 302, 304, 307–313, 318 Melt, 1–2, 7–8, 13–15, 23, 48, 60, 62–66, 68, 72–73, 78, 82–84, 86, 88–90, 92, 95–103, 105–106, 122, 127–131, 135–137, 141, 143–145, 151–152, 155–157, 159–160, 162–165, 167, 169–176, 179, 181, 187, 204, 211–215, 225–227, 240, 253–256, 265, 269, 272, 275, 277, 281, 283–284, 286, 288–289, 292–295, 298, 300, 303, 313, 315–319 Melting point, 153, 171, 173–174, 275, 281, 316, 318–319 Melt-origin pseudotachylyte, 7–8, 48, 72, 78, 83–84, 90, 98, 102, 105, 128–130, 141, 143, 151–152, 155–156, 159–160, 164, 171, 204, 213–214, 225–227, 240, 253, 255–256, 275, 283–284, 315–316 Metamorphism, 7, 14, 57, 75, 90, 95, 130, 136, 186, 197, 220 Meteoric water, 258–263 Meteorite impact, 1, 5–7, 9–10, 140–141 Mica ﬁsh, 32–33, 35–36 Microcrack, 28, 206, 256, 258, 291 Microcrystalline-type, 81–82 Microfault, 71, 191 Microlite, 9, 12, 14–15, 57, 75, 78, 80–81, 83–84, 90, 92, 95, 97, 105–137, 141, 144, 151, 154, 167, 171, 187, 190, 193, 195–198, 202–205, 211–212, 214, 217, 225, 254, 266, 275–276, 280, 283–284, 316, 318

346

Index

Microlitic mylonite, 6 Microlitic-type, 83 Microshear, 32–34, 36–37, 40–42 Mixed-type, 84, 91, 111 Moine Thrust, 10 Molten material, 7–8, 96, 129, 147, 170, 256, 265–266, 272, 284, 294, 298–301, 304, 313, 316 Molten zone, 283–284, 319 Montmorillonite, 43, 245–246 Mulga Park subdomain, 179 Mylonite, 10, 19, 23–24, 26, 33, 35–37, 39, 45–46, 60, 136, 178–179, 181–184, 186–187, 190, 192–194, 196, 198–200, 204, 206–207, 209–211, 213–217, 219–221 Mylonite-related pseudotachylyte (M-Pt), 45, 177–179, 181–184, 186–201, 204–212, 213–223 Mylonitic foliation, 21, 57, 85, 179, 184, 189–190, 194–195, 199–200, 204, 207, 214 Mylonitic rocks, 10, 19–25, 32, 40, 44–46, 48, 82, 135, 151, 177, 190, 209, 219–220, 223 Mylonitization, 5, 12–13, 39–40, 46, 183, 186, 190, 200, 204, 208–210, 218–220, 223, 265 Newtonian ﬂuid, 83 Newtonian suspension, 84, 87 Nojima Fault, 17, 28–29, 33, 37–39, 47, 130, 153, 156, 226, 232–240, 243–246, 250–253, 256–258, 260–261, 263 Non-crystalline, 90–91, 93, 96–97, 241, 245, 254, 275, 301 Non-equilibrium, 122, 127–128, 281, 284, 316–318, 319 Non-foliated cataclastic rock, 28–31 Normal stress, 41–43, 173–175, 212, 281, 287, 289–290, 292, 315–316, 319 North Anatolian Fault Zone, 258, 260 Osumi Shear Zone, 156 Outer Hebrides Thrust, 9, 13, 63, 65–66, 71, 86, 90, 93, 140, 142–143, 172, 177, 221

Overgrowth, 111–112, 115, 118, 136 Paired generation zone, 64–66, 68 Peru Mw 8.2 earthquake, 222 Pine-like, 109–110, 115 Plagioclase, 80, 90–92, 109, 111–112, 115, 117–120, 122, 124, 133, 135–137, 167, 187, 193–194, 196, 200, 202–205, 209, 211–212, 216–219, 246, 250, 279, 290–291, 297, 300, 302–304, 307–308, 314, 316–318 Plastic-dominated regime, 177, 179, 188, 215, 217–221, 223 Plastic ﬂow regime, 18, 46, 193, 218, 223–224 Plastic instability, 177, 221–223 Plastic-instability model, 221 Plumose, 109–110, 115 Pore ﬂuid pressure, 41–42, 97, 175, 287 Porphyroclast, 24, 32, 46, 102, 139, 206, 208–210, 216–217 Pressurization, 62, 98–99, 174–175 Primary cohesion, 23, 25–26 Propagation model, 221 Protocataclasite Protomylonite, 19, 24–25 Pseudotachylyte Pyroxene, 80, 111, 119, 121, 126, 127–129, 133, 136–137, 187, 194, 202, 290–291, 297–298, 300–302, 307–308, 317, 319 Qinling-Dabie Shan collisional orogenic belt, 178–179, 197, 206–207, 216, 219 Quartz, 32–33, 36, 40–41, 45–46, 78–83, 88–96, 102–103, 106, 109, 111–113, 115–116, 118, 120–121, 136, 151, 153, 155, 165, 167, 171–172, 176, 187, 190, 194, 196, 198–200, 202, 204, 206, 209, 211, 214, 216–217, 219–221, 223, 237, 245–246, 249–250, 275, 277, 279, 281, 289–291, 293, 299–303, 308, 312, 314, 316–319 Quartz-feathery, 109–110 Quartzite, 290–291, 293, 308, 312–314, 316–317

Index Quartzo-feldspathic continental crust, 44–46 Quartz vein, 291 Quasi-conglomerate, 140–141 Quasi-plastic (QP), 44–45 Quenching, 13–14, 97, 105, 127, 136–137, 169, 187, 204 Recrystallization, 22–24, 46, 88, 139, 214, 217, 221, 223 Redbank Shear Zone, 177 Reservoir zone, 60 Rheology, 41, 43, 286 Riedel shear, 32, 65, 68 Rock-forming mineral, 40–41, 46, 96, 165, 167, 171–172, 174, 214, 245, 250, 275, 298 Rotary-shear high-velocity friction device, 287 Roundness, 153–155, 157–158, 209 Saint-Barthkemy Massif, 177–178 San Andreas Fault, 10, 17, 40–42, 148, 150 Sandwich structure, 232–233 San Gabriel Fault Zone, 150 Sanidine, 84, 111, 113, 119, 122, 133, 135, 137 Sanidinite facies, 135 Santa Rose mylonite shear zone, 8, 15, 129, 133–134 S-C cataclasite, 32, 40–43, 45–46 S-C fabrics, 24, 32–34, 39, 42–44, 207, 213 S-C mylonite, 36, 45–46 Seawater, 258–263 Seismic faulting, 1, 12, 15, 17, 19, 48, 51, 97–98, 105, 130–131, 139, 173–176, 178, 192–193, 204, 214–215, 217, 219–220, 223, 225–226, 229, 234, 242, 245, 255–258, 261–263, 265, 283–284, 292, 315 Seismic pumping model, 263 Seismic slip, 13, 15, 18, 45–46, 72, 157, 177, 209, 212, 218, 220, 222–223, 225–226, 254–257, 283, 319 Seismite, 8 Seismogenic fault zone, 40–43, 159, 223

347

Seismogenic zone, 41, 45–46, 178, 212, 218–219, 221–223, 283 Self-similar, 145, 150–151 Serpentinite, 43 S-foliation, 33, 40–42 Sheaf, 111–112, 115, 117, 136 Sheaf aggregate, 111–112, 115 Shear band, 21, 24, 32–35, 37, 40–42, 206, 209, 234 Shear resistance (τ ), 41, 97–98, 218, 222–223, 318 Shear strain energy, 221, 223 Shear strength, 41, 46, 97–98, 176, 318 Shear stress, 9, 98, 174–175, 219, 283, 286, 289–290, 318 Shear zone, 18–19, 22–23, 32, 38–39, 44–45, 48, 64–65, 67–69, 82, 99, 156, 172, 177, 179, 181–184, 186, 188, 192, 197–199, 214–215, 217, 219–223, 258, 292, 295–296, 298, 315 Sheath-fold-like structure, 82, 86 Siderite, 245–246, 249–250, 260 Sidewall ripout, 64–65, 68 Similar-type fold, 82–83, 85–86 Simple group, 106–109 Single generation zone, 64–65, 67–68 Skeletal group, 109–110 Slab duplex, 65 Slip weakening, 8, 319 Solid-gas-ﬂuid system, 15 Solid-gas system, 12 Solid solution, 290–291, 308, 316 Spherulite, 9, 12, 14, 108, 112, 115–116, 118–119, 121–122, 167, 169, 187 Spherulitic, 83, 102, 106, 111–120, 126, 128, 131, 135–137, 141, 144, 154, 187, 190, 193, 196, 198, 203 Spherulitic group, 112–118 Spider-like, 84, 106, 110, 113, 135–137 Spinel, 111, 119, 128, 132 S-surface, 32–34, 36, 40, 206, 213 Stellate, 111–112, 115, 118 Stress relaxation, 221 Stretching lineation, 197–199 Subduction zone, 47, 176 Suction pumping model, 262–263 Surface water, 30, 252, 257–258, 260, 261, 263

348

Index

Tabular-skeletal, 109 Tachylyte, 5 Taiwan Mw 7.6 earthquake, 266 Tectonic-generated pseudotachylyte, 1 Tectonic-related pseudotachylyte, 47 Tensile strength, 99–100 Tensional fracture, 60, 65, 236, 258 Thermal ﬂuctuation, 221 Thermal fracturing, 289, 291, 293 Thermal pressurization, 62, 98–99, 174–175 Three-layer fault model, 45 Transition regime, 45 Transportation, 83, 152 Trap-shotten, 10 Trap-shotten gneiss, 6 Trichitic, 106, 113, 118, 120 Tsaoling landslide, 271–272 Two-layer fault model, 44–45 Ultracataclasite, 323 Ultrahigh-pressure (UHP), 178–179, 216–217, 219 Ultraman-daina, 1–2 Ultramylonite, 19, 24–25, 27, 45, 177, 181–184, 188, 190, 192, 194–195, 199, 207, 219–220 Ultramylonite-associated pseudotachylyte (Um-Pt), 177, 179, 181–184, 186, 188–192, 195, 218–220, 223 U-shaped ﬂow streak, 84

Veinlet, 7–9, 12, 15, 29, 60, 194, 234, 244, 256, 258 Veinlet cataclastic rock, 12, 225–226, 237, 244, 253, 257 Velocity-weakening, 319 Vermiculite, 43 Vesicle, 12–15, 84–89, 97–99, 101–103, 169, 171, 175, 225, 254, 269–270, 275, 277, 283, 294, 297–298, 315 Vesiculation, 103 Viscosity, 62, 173–175, 284, 318 Vitreous luster, 79–80, 227 Volcanic glass, 12, 91, 96, 137, 274, 277 Vredefort Dome structure, 5 Wall eﬀect, 83, 141 Woodroﬀe Thrust, 45, 48, 50, 53–54, 56–60, 62–63, 65–66, 68, 70, 90, 93, 156, 162, 172, 177–179, 181–183, 185–203, 205, 208, 214–215, 219–221, 223 Xenoclast, 139 Xenolith, 139 X-ray diﬀraction pattern, 78–80, 90–93, 96, 279, 315 X-ray ﬂuorescence, 159 X–Z section, 19, 31, 33–34, 210 Yangtze continental block, 216 Zoning structure, 102, 116, 169