Active global seismology : neotectonics and earthquake potential of the eastern Mediterranean region 1118944992, 9781118944998, 978-1-118-94498-1

Citation preview

Geophysical Monograph Series

Geophysical Monograph Series 175 A Continental Plate Boundary: Tectonics at South Island, New Zealand  David Okaya, Tim Stem, and Fred Davey (Eds.) 176 Exploring Venus as a Terrestrial Planet  Larry W. Esposito, Ellen R. Stofan, and Thomas E. Cravens (Eds.) 177 Ocean Modeling in an Eddying Regime  Matthew Hecht and Hiroyasu Hasumi (Eds.) 178 Magma to Microbe: Modeling Hydrothermal Processes at Oceanic Spreading Centers  Robert P. Lowell, Jeffrey S. Seewald, Anna Metaxas, and Michael R. Perfit (Eds.) 179 Active Tectonics and Seismic Potential of Alaska  Jeffrey T. Freymueller, Peter J. Haeussler, Robert L. Wesson, and Göran Ekström (Eds.) 180 Arctic Sea Ice Decline: Observations, Projections, Mechanisms, and Implications  Eric T. DeWeaver, Cecilia M. Bitz, and L.-Bruno Tremblay (Eds.) 181 Midlatitude Ionospheric Dynamics and Disturbances  Paul M. Kintner, Jr., Anthea J. Coster, Tim Fuller-Rowell, Anthony J. Mannucci, Michael Mendillo, and Roderick Heelis (Eds.) 182 The Stromboli Volcano: An Integrated Study of the 2002–2003 Eruption  Sonia Calvari, Salvatore Inguaggiato, Giuseppe Puglisi, Maurizio Ripepe, and Mauro Rosi (Eds.) 183 Carbon Sequestration and Its Role in the Global Carbon Cycle  Brian J. McPherson and Eric T. Sundquist (Eds.) 184 Carbon Cycling in Northern Peatlands  Andrew J. Baird, Lisa R. Belyea, Xavier Comas, A. S. Reeve, and Lee D. Slater (Eds.) 185 Indian Ocean Biogeochemical Processes and Ecological Variability  Jerry D. Wiggert, Raleigh R. Hood, S. Wajih A. Naqvi, Kenneth H. Brink, and Sharon L. Smith (Eds.) 186 Amazonia and Global Change  Michael Keller, Mercedes Bustamante, John Gash, and Pedro Silva Dias (Eds.) 187 Surface Ocean–Lower Atmosphere Processes  Corinne Le Quèrè and Eric S. Saltzman (Eds.) 188 Diversity of Hydrothermal Systems on Slow Spreading Ocean Ridges  Peter A. Rona, Colin W. Devey, Jérôme Dyment, and Bramley J. Murton (Eds.) 189 Climate Dynamics: Why Does Climate Vary?  De-Zheng Sun and Frank Bryan (Eds.) 190 The Stratosphere: Dynamics, Transport, and Chemistry  L. M. Polvani, A. H. Sobel, and D. W. Waugh (Eds.) 191 Rainfall: State of the Science  Firat Y. Testik and Mekonnen Gebremichael (Eds.) 192 Antarctic Subglacial Aquatic Environments  Martin J. Siegert, Mahlon C. Kennicut II, and Robert A. Bindschadler 193 Abrupt Climate Change: Mechanisms, Patterns, and Impacts  Harunur Rashid, Leonid Polyak, and Ellen  Mosley-Thompson (Eds.) 194 Stream Restoration in Dynamic Fluvial Systems: Scientific Approaches, Analyses, and Tools  Andrew Simon, Sean J. Bennett, and Janine M. Castro (Eds.) 195 Monitoring and Modeling the Deepwater Horizon Oil Spill: A Record-Breaking Enterprise  Yonggang Liu, Amy MacFadyen, Zhen-Gang Ji, and Robert H. Weisberg (Eds.) 196 Extreme Events and Natural Hazards: The Complexity Perspective  A. Surjalal Sharma, Armin Bunde, Vijay P. Dimri, and Daniel N. Baker (Eds.) 197 Auroral Phenomenology and Magnetospheric Processes: Earth and Other Planets  Andreas Keiling, Eric Donovan, Fran Bagenal, and Tomas Karlsson (Eds.) 198 Climates, Landscapes, and Civilizations  Liviu Giosan, Dorian Q. Fuller, Kathleen Nicoll, Rowan K. Flad, and Peter D. Clift (Eds.)

199 Dynamics of the Earth’s Radiation Belts and Inner Magnetosphere  Danny Summers, Ian R. Mann, Daniel N. Baker, and Michael Schulz (Eds.) 200 Lagrangian Modeling of the Atmosphere  John Lin (Ed.) 201 Modeling the Ionosphere-Thermosphere  Jospeh D. Huba, Robert W. Schunk, and George V Khazanov (Eds.) 202 The Mediterranean Sea: Temporal Variability and Spatial Patterns  Gian Luca Eusebi Borzelli, Miroslav Gacic, Piero Lionello, and Paola Malanotte-Rizzoli (Eds.) 203 Future Earth - Advancing Civic Understanding of the Anthropocene  Diana Dalbotten, Gillian Roehrig, and Patrick Hamilton (Eds.) 204 The Galápagos: A Natural Laboratory for the Earth Sciences  Karen S. Harpp, Eric Mittelstaedt, Noémi d’Ozouville, and David W. Graham (Eds.) 205 Modeling Atmospheric and Oceanic Flows: Insightsfrom Laboratory Experiments and Numerical Simulations  Thomas von Larcher and Paul D. Williams (Eds.) 206 Remote Sensing of the Terrestrial Water Cycle  Venkat Lakshmi (Eds.) 207 Magnetotails in the Solar System  Andreas Keiling, Caitríona Jackman, and Peter Delamere (Eds.) 208 Hawaiian Volcanoes: From Source to Surface  Rebecca Carey, Valerie Cayol, Michael Poland, and Dominique Weis (Eds.) 209 Sea Ice: Physics, Mechanics, and Remote Sensing  Mohammed Shokr and Nirmal Sinha (Eds.) 210 Fluid Dynamics in Complex Fractured-Porous Systems  Boris Faybishenko, Sally M. Benson, and John E. Gale (Eds.) 211 Subduction Dynamics: From Mantle Flow to Mega Disasters  Gabriele Morra, David A. Yuen, Scott King, Sang Mook Lee, and Seth Stein (Eds.) 212 The Early Earth: Accretion and Differentiation  James Badro and Michael Walter (Eds.) 213 Global Vegetation Dynamics: Concepts and Applications in the MC1 Model  Dominique Bachelet and David Turner (Eds.) 214 Extreme Events: Observations, Modeling and Economics  Mario Chavez, Michael Ghil, and Jaime Urrutia-Fucugauchi (Eds.) 215 Auroral Dynamics and Space Weather  Yongliang Zhang and Larry Paxton (Eds.) 216 Low‐Frequency Waves in Space Plasmas  Andreas Keiling, Dong‐Hun Lee, and Valery Nakariakov (Eds.) 217 Deep Earth: Physics and Chemistry of the Lower Mantle and Core  Hidenori Terasaki and Rebecca A. Fischer (Eds.) 218 Integrated Imaging of the Earth: Theory and Applications  Max Moorkamp, Peter G. Lelievre, Niklas Linde, and Amir Khan (Eds.) 219 Plate Boundaries and Natural Hazards  Joao Duarte and Wouter Schellart (Eds.) 220 Ionospheric Space Weather: Longitude and Hemispheric Dependences and Lower Atmosphere Forcing  Timothy FullerRowell, Endawoke Yizengaw, Patricia H. Doherty, and Sunanda Basu (Eds.) 221 Terrestrial Water Cycle and Climate Change: Natural and Human-Induced Impacts  Qiuhong Tang and Taikan Oki (Eds.) 222 Magnetosphere-Ionosphere Coupling in the Solar System  Charles R. Chappell, Robert W. Schunk, Peter M. Banks, James L. Burch, and Richard M. Thorne (Eds.) 223 Natural Hazard Uncertainty Assessment: Modeling and Decision Support  Karin Riley, Peter Webley, and Matthew Thompson (Eds.) 224 Hydrodynamics of Time-Periodic Groundwater Flow: Diffusion Waves in Porous Media  Joe S. Depner and Todd C. Rasmussen (Eds.)

Geophysical Monograph 225

Active Global Seismology Neotectonics and Earthquake Potential of the Eastern Mediterranean Region ̇ Ibrahim Çemen Yücel Yılmaz

This Work is a copublication of the American Geophysical Union and John Wiley & Sons, Inc.

This Work is a copublication of the American Geophysical Union and John Wiley & Sons, Inc.

Published under the aegis of the AGU Publications Committee Brooks Hanson, Director of Publications Robert van der Hilst, Chair, Publications Committee © 2017 by the American Geophysical Union, 2000 Florida Avenue, N.W., Washington, D.C. 20009 For details about the American Geophysical Union, see www.agu.org. Published by John Wiley & Sons, Inc., Hoboken, New Jersey Published simultaneously in Canada No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording, scanning, or otherwise, except as permitted under Section 107 or 108 of the 1976 United States Copyright Act, without either the prior written permission of the Publisher, or authorization through payment of the appropriate per‐copy fee to the Copyright Clearance Center, Inc., 222 Rosewood Drive, Danvers, MA 01923, (978) 750‐8400, fax (978) 750‐4470, or on the web at www.copyright.com. Requests to the Publisher for permission should be addressed to the Permissions Department, John Wiley & Sons, Inc., 111 River Street, Hoboken, NJ 07030, (201) 748‐6011, fax (201) 748‐6008, or online at http://www.wiley.com/go/permissions. Limit of Liability/Disclaimer of Warranty: While the publisher and author have used their best efforts in preparing this book, they make no representations or warranties with respect to the accuracy or completeness of the contents of this book and specifically disclaim any implied warranties of merchantability or fitness for a particular purpose. No warranty may be created or extended by sales representatives or written sales materials. The advice and strategies contained herein may not be suitable for your situation. You should consult with a professional where appropriate. Neither the publisher nor author shall be liable for any loss of profit or any other commercial damages, including but not limited to special, incidental, consequential, or other damages. For general information on our other products and services or for technical support, please contact our Customer Care Department within the United States at (800) 762‐2974, outside the United States at (317) 572‐3993 or fax (317) 572‐4002. Wiley also publishes its books in a variety of electronic formats. Some content that appears in print may not be available in electronic formats. For more information about Wiley products, visit our web site at www.wiley.com. Library of Congress Cataloging‐in‐Publication Data is available. ISBN: 978-1-118-94498-1 Cover image: The image presents the water‐level response in an aquifer due to periodic excitation of four groundwater wells at different amplitudes and frequencies. Printed in the United States of America

10 9 8 7 6 5 4 3 2 1

CONTENTS Contributors���������������������������������������������������������������������������������������������������������������������������������������������������������vii Preface������������������������������������������������������������������������������������������������������������������������������������������������������������������ix 1 Neotectonics and Earthquake Potential of the Eastern Mediterranean Region: Introduction ̇ Ibrahim Çemen and Yücel Yılmaz.....................................................................................................................1 Part I: Morphotectonic Characteristics of Neotectonics in Anatolia and Its Surroundings

9

2 Morphotectonic Development of Anatolia and the Surrounding Regions Yücel Yılmaz..................................................................................................................................................11 3 Diversion of River Courses Across Major Strike‐Slip Faults and Keirogens A. M. Celâl Şengör.........................................................................................................................................93 Part II: Neotectonics of the Aegean-Western Anatolian Region

103

4 Effect of Slab‐Tear on Crustal Structure in Southwestern Anatolia: Insight From Gravity Data Modeling ̇ Rezene Mahatsente, Süleyman Alemdar, and Ibrahim Çemen.......................................................................105 5 Geodynamical Models for Continental Delamination and Ocean Lithosphere Peel Away in an Orogenic Setting Oğuz H. Göğüş, Russell N. Pysklywec, and Claudio Faccenna.....................................................................121 6 Major Problems of Western Anatolian Geology Yücel Yilmaz................................................................................................................................................141 7 The Çataldağ Plutonic Complex in Western Anatolia: Roles of Different Granites on the Crustal Buildup in Connection With the Core Complex Development Ömer Kamacı, Alp Ünal, Şafak Altunkaynak, Stoyan Georgiev, and Zeki M. Billor.........................................189 Part III: Seismotectonics in the Eastern Mediterranean Region

223

8 Fault Structures in Marmara Sea (Turkey) and Their Connection to Earthquake Generation Processes Mustafa Aktar..............................................................................................................................................225 9 North Aegean Active Fault Pattern and the 24 May 2014, Mw 6.9 Earthquake Sotiris Sboras, Alex Chatzipetros, and Spyros B. Pavlides..............................................................................239 10 Seismic Intensity Maps for the Eastern Part of the North Anatolian Fault Zone (Turkey) Based on Recorded and Simulated Ground‐Motion Data Aysegul Askan, Shaghayegh Karimzadeh, and Mustafa Bilal..........................................................................273 Index������������������������������������������������������������������������������������������������������������������������������������������������������������������289

v

CONTRIBUTORS Mustafa Aktar Kandilli Observatory and Earthquake Research Institute Bogazici University ̇ Kandilli‐Istanbul, Turkey

Oğuz H. Göğüş Eurasia Institute of Earth Sciences ̇ Istanbul Technical University (ITU) ̇ Istanbul, Turkey

Süleyman Alemdar Department of Geological Sciences The University of Alabama Tuscaloosa, Alabama, USA

Ömer Kamacı Department of Geological Engineering ̇ Istanbul Technical University ̇ Istanbul, Turkey

Şafak Altunkaynak Department of Geological Engineering ̇ Istanbul Technical University Iṡ tanbul, Turkey

Shaghayegh Karimzadeh Civil Engineering Department Middle East Technical University Ankara, Turkey Rezene Mahatsente Department of Geological Sciences The University of Alabama Tuscaloosa, Alabama, USA

Aysegul Askan Civil Engineering Department Middle East Technical University Ankara, Turkey

Spyros B. Pavlides Department of Geology Aristotle University of Thessaloniki Thessaloniki, Greece

Mustafa Bilal Civil Engineering Department Middle East Technical University Ankara, Turkey

Russell N. Pysklywec Department of Earth Sciences University of Toronto Toronto, Ontario, Canada

Zeki M. Billor Department of Geology and Geography Auburn University Auburn, Alabama, USA

Sotiris Sboras Institute of Geodynamics, National Observatory of Athens Greece

̇ Ibrahim Çemen Department of Geological Sciences The University of Alabama Tuscaloosa, Alabama, USA

A. M. Celâl Şengör ̇ ITÜ Maden Fakültesi, Jeoloji Bölümü ve Avrasya Yerbilimleri Enstitüsü Ayazağa ̇ Istanbul, Turkey

Alex Chatzipetros Department of Geology Aristotle University of Thessaloniki Thessaloniki, Greece

Alp Ünal Department of Geological Engineering ̇ Istanbul Technical University ̇ Istanbul, Turkey

Claudio Faccenna Laboratory of Experimental Tectonics Università Roma TRE Roma, Italy

Yücel Yılmaz Department of Geological Engineering ̇ Istanbul Technical University ̇ Istanbul, Turkey

Stoyan Georgiev Department of Geochemistry and Petrology Geological Institute Bulgarian Academy of Sciences Sofia, Bulgaria vii

PREFACE Within the last two decades or so, earth scientists realized that the ultimate goal of global seismology is to invent a method to forecast earthquakes similar to meteorologists forecasting the weather. Presently, almost all earthquake scientists believe that this goal is still out of reach. However, they are trying to develop algorithms and computer models, using new satellite data and other resources, to predict when an earthquake will occur. Many earthquake scientists hope that earthquake forecast will be possible within the 21st century. Until then, we need to investigate earthquake potential of the regions where large historical earthquakes have occurred and future ones are expected. These investigations involve interdisciplinary research. The eastern Mediterranean is one of the most seismically active continental regions of the world. It includes three of the most active strike‐slip faults: the North Anatolian, East Anatolian, and Dead Sea fault zones. Many devastating historical earthquakes have occurred along these fault zones and along other active tectonic features in the region, such as the Hellenic and Cyprus subduction zones, extensional faults in the Aegean region, and the Zagros suture zone. Considering that the eastern Mediterranean has experienced many devastating earthquakes throughout history, including numerous ones of Mw > 7.0 during the 20th century, it is safe to assume that the region can expect many large earthquakes during the 21st century and beyond. Consequently, there is an urgent societal need to understand the neotectonics and earthquake potential of the region.

This AGU‐Wiley book is an outgrowth of a research symposium titled “Neotectonics and Earthquake Potential of the Eastern Mediterranean Region” at the 2013 AGU fall meeting. The symposium, organized by the editors of this book, was well attended by researchers from all over the world and provided a formal discussion on s­ everal important issues related to neotectonics and earthquake potential of the region. The book contains 10 chapters organized under three thematic groups: (1) morphotectonic characteristics of neotectonics in Anatolia and its surroundings, (2) neotectonics of the Aegean‐western Anatolian region, and (3) seismotectonics in the eastern Mediterranean region. We hope this book will provide basic knowledge for the development of new earthquake research projects in the region and elsewhere in the world. We thank all of our contributors and reviewers for their excellent and timely work that helped realize this book. We are in debt to Rituparna Bose and Mary Grace Hammond for their efforts on behalf of this volume, which would not have been realized without their constant push on our editorial duties. Most important, we thank our families for not only their support during the long hours of editing this book but also for supporting us constantly through the years while we were conducting research to contribute to the science of geology. ̇ Ibrahim Çemen Yücel Yılmaz

ix

1 Neotectonics and Earthquake Potential of the Eastern Mediterranean Region: Introduction Iḃ rahim Çemen1 and Yücel Yılmaz2

1.1. INTRODUCTION

USA, which included 8 oral and 12 poster presentations. This book is a collection of the research that was pre­ sented at the meeting. The eastern Mediterranean region is one of the most dynamically complex and seismically active neotectonic settings on Earth (Figure 1.1). It includes the following major geographic divisions: the Aegean Sea region, the  Anatolian Peninsula, and the northern part of the Arabian Peninsula. Each of these geographic domains corresponds to a distinctly different and composite ­tectonic entity. The Anatolian Peninsula is part of the Alpine‐Himalayan orogenic belt. Along its northern and southern edges lie approximately E‐W trending mountain ranges known as the Pontides (the northern range) and the Taurides (the southern range). In Anatolia, the orogeny started in the north, migrated progressively to the south, and ended up in the Bitlis‐Zagros orogenic belt. Following the latest phase of the collision along the Bitlis‐Zagros suture, the Arabian Plate continued moving northward and gener­ ated a north‐directed contraction (Figure  1.1). Conse­ quently, the East Anatolian crust and lithosphere have been thickened, and the region was elevated to form the East Anatolian‐Iranian high plateau. This shortening gave way to the formation of the North Anatolian and East Anatolian fault zones (Figure 1.1). The initiation of the two fault zones is generally considered as the beginning of neotectonics in Anatolia and surrounding regions. Neotectonics of the eastern Mediterranean region is dominated by the African Plate subduction along the Hellenic and Cyprus trenches, collision between the Anatolian and Eurasian plates, and westward extrusion

Neotectonics is a subdiscipline of tectonics and involves the study of recent motions and deformation of the Earth’s crust. These recent motions, particularly those produced by earthquakes, can provide insights on the physics of earthquake recurrence, the growth of mountains, oro­ genic movements, and the seismic hazard. This volume focuses on neotectonics of the eastern Mediterranean region (Figure  1.1), which has experienced many major devastating earthquakes throughout its recorded history. A major devastating earthquake in the region occurred at 3:02 a.m. on 17 August 1999 in Izmit, Turkey (Mw = 7.4), lasted for 37 sec, killed around 17,000, injured 44,000 people, and left approximately half a million people homeless. Economic loss due to this earthquake is estimated at around $20 billion. Since the Izmit earthquake, several North American, European, and Turkish research groups have been studying the neotectonics and earthquake potential of the eastern Mediterranean region by using different geological and geophysical methods, including GPS studies, geodesy, and passive source seismology. Some results from these studies were presented in major North American and European geological meetings and published in major earth science journals. However, the first comprehensive collection of research case studies of this region was convened by the editors of this book at the 2013 AGU fall meeting in San Francisco, California, Department of Geological Sciences, The University of Alabama, Tuscaloosa, Alabama, USA 2  ̇ Department of Geological Engineering, Istanbul Technical ̇ University, Istanbul, Turkey 1 

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 1

2  ACTIVE GLOBAL SEISMOLOGY

Figure  1.1  Digital elevation map of the eastern Mediterranean region showing major neotectonics structural features, volcanic centers (red triangles), and epicenters of the earthquakes (M > 5.0) since 1950. A = Ankara; EAFZ = East Anatolian fault zone; EF = Ecemis fault; I = Istanbul; MM = Menderes Massif; NAFZ = North Anatolian fault zone; T = Thessaloniki; TGF = Tuz Golu fault.

of the Anatolian Plate along the north and east Anatolian fault zones [Sengor and Yilmaz, 1981; Sengor et al., 1985; Robertson and Dixon, 1984; Çemen et  al., 1999, 2006; Aksu et al., 2005]. The convergent zones are characterized by deep earthquakes along the Hellenic and western seg­ ment of the Cyprus arcs [Di Luccio and Pasyanos, 2007], volcanism [Pe‐Piper and Piper, 2006; 2007; Altunkaynak and Dilek, 2006; Prelević et al., 2012; Jolivet et al., 2013], large‐scale continental extension [Faccenna et al., 2003; Çemen, 2010; and Ersoy et  al., 2014], uplift [Schildgen et  al., 2012; 2014], trench retreat, slab tear, and slab detachment [Faccenna et  al., 2006; Biryol et  al., 2011; Hall et al., 2014]. The extension and uplift are related to the southwest retreating Hellenic trench and westward movement of the Anatolian Plate [Çemen et al., 2006 and 2014; Reilinger et  al., 2010; Cosentino et  al., 2012; Schildgen et al., 2014]. This book contains nine chapters covering a wide range of contributions to the neotectonics and earthquake

potential of the eastern Mediterranean region. The chapters cover an extensive and overlapping tectonic mosaic of new data that contribute significantly to our understanding of the crustal and lithospheric behavior manifested by tectonic, seismotectonic, and morphotectonic elements in the region. The chapters are organized under the following the­ matic groups. 1.1.1. Part I: Morphotectonic Characteristics of Neotectonics in Anatolia and Its Surroundings Two chapters are in this section of the book. 1.1.1.1. Chapter 1. Morphotectonic Development of Anatolia and the Surrounding Regions by Yücel Yılmaz This chapter may be regarded as a tectonic backbone of the book in the sense that it covers tectonic framework of Anatolia and its surroundings. Several local and

Neotectonics and Earthquake Potential of the Eastern Mediterranean Region: Introduction  3

regional morphological studies were conducted on differ­ ent parts of Anatolia. However, this chapter is the first attempt to encompass morphological treatment of the whole Anatolian peninsula to evaluate interactions of morphotectonically different regions and major tectonic elements, and along this direction it provides a platform for similar future studies. The chapter’s major points are summarized as follows: Anatolia is being deformed presently under an ongoing severe post‐late orogenic tectonic regime, which is expressed by the GPS data; frequent earthquakes that occur in a vast terrain from the east to the west; and rug­ ged, irregular, and tectonically controlled morphology. In order to understand the tectonics of Anatolia, structural analyses of the tectonically different regions and the earthquakes are studied extensively using a variety of methods and techniques. but the morphotectonic features that are also equally important are commonly ignored. Therefore, this chapter is complementary to most of the structurally and tectonically oriented regional and local treatments. Anatolia and the surrounding regions contain a num­ ber of morphotectonic subdivisions including the East Anatolian‐Iranian high plateau. The other subdivision are the peripheral mountain ranges (the Pontides and Taurides), the central Anatolian plateau, and the west­ ern Anatolian extensional region. They have all essen­ tially formed during the Neotectonics period. Therefore, a critical period in the geologic‐tectonic and particularly morphotectonic history of the region corresponds to a change from the Paleotectonic to Neotectonic periods. This chapter first addresses the timing and cause of the transition between these periods for each tectonic subdi­ vision of the region, and then discusses at length mor­ photectonic character and characteristic features of each of the morphotectonic subdivisions, starting from the northwestern edge of the Arabian Peninsula around the Bitlis‐Zagros suture mountains because this belt is the latest product of the Anatolian Orogen that formed as a result of the collision between the Arabian and Anatolian plates. The northward advance of the Arabian Plate con­ tinuing after the collision generated a north‐directed severe contraction to push Anatolia northward. The N‐S contraction initially deformed eastern Turkey. East and central Anatolia began to rise together. In this, slab break off of the northerly subducting plate and ­lithospheric delamination played a significant role. The contraction then formed the North and East Anatolian transform faults. These faults border the Anatolian Plate, which began escaping to the west. Major morpho­ tectonic features, the peripheral mountains (the Pontides and Taurides) and the western Anatolian extensional region, have evolved together with the transform faults,

which played an important role in transfer of the stress in the region. Compared to the east, the western of Anatolia has fol­ lowed a different path of morphotectonic development. The region was a high land during the Early Miocene period while eastern Anatolia was under a shallow sea. The environments began to reverse from Late Miocene onward. Southerly retreat of the subducting eastern Mediterranean oceanic slab has generated N‐S extension in western Anatolia, Turkey. As a result, the present mor­ phology began to develop. The E‐W trending horst and graben structures that dominate the landscape today began to form during later periods of extension in the Quaternary. 1.1.1.2.  Chapter  2. Diversion of  River Courses Across Major Strike‐Slip Faults and  Keirogens by A. M. C. Şengör The second chapter describes a pioneering morpho­ tectonic study. It explores some theoretical possibilities of river bends along strike‐slip fault zones using the example of the North Anatolian fault that forms a family of faults, which constitute the North Anatolian keirogen [Şengör et  al. 2005]. The author, documenting prelimi­ nary results of his research, draws our attention to the following points. Studies along the North Anatolian fault and other major active strike‐slip faults and keirogens in the world have revealed complications in river offsets that cannot simply be explained by preexisting slope con­ ditions and capture events. These seem to result from the presence of numerous lesser strike‐slip faults parallel with the main displacement zone of a large strike‐slip fault, and from the structure and topographic evolution of synthetic and antithetic pull‐apart basins. Some cuspate pull‐apart‐ basin‐bounding normal faults may give the mistaken impression of a river bending into a strike‐slip fault because of numerous parallel faults. Other complications result from the presence of structures that predate the formation of a through‐going, main strike‐slip fault. All strike-slip faults consist of surfaces of slip anasto­ mosing along the strike of a fault zone, when a fault zone is narrow as a line. However, when its width exceeds a few kilometers, the motion of individual lozenges or phacoids surrounded by the anastomosing branches visi­ bly influence the topography creating whaleback ridges, that in places may function as shutter ridges at the mouths of valleys consequent to the drainage of the main fault valley, sag ponds, and pull‐apart basins that can be of various sizes and aspect ratios, and push‐up ridges that may be simple folds or thrust blocks. Whatever basins form along a strike‐slip fault zone, their floors may assume various slopes, both in direction and amount, depending on the geometry of the down‐dropping fault(s).

4  ACTIVE GLOBAL SEISMOLOGY

1.1.2. Part II: Neotectonics of the Aegean‐Western Anatolian Region This section contains four chapters. 1.1.2.1.  Chapter  4. Effect of  Slab‐Tear on  Crustal Structures in  Southwestern Anatolia: Insight from  Gravity Data Modeling by R. Mahatsente, S. Alemdar, and I.̇ Çemen This chapter examines the effects of the asthenospheric window on major crustal structures in western Turkey and the upper mantle using gravity data modeling. The authors use a combination of terrestrial and satellite gravity data. Their gravity model is also constrained by the results of recent receiver function and seismic tomog­ raphy studies [e.g., Biryol, 2011]. The gravity model in this chapter suggests that depth to the top of the asthenospheric material (i.e., the crust of the Earth), ranges from 24 to 29 km below the Menderes Massif. The location of this thinned crust coincides with high heat flow of magmatic centers in the Menderes Massif complex. The asthenospheric material, as deduced from its density value and dimensions, is most probably deep in origin (asthenospheric and lithospheric mantle origin) and may be related to the low‐velocity astheno­ spheric material in the upper mantle imaged by seismic tomography. The absence of no deep earthquakes in the asthenospheric window area is also a line of evidence for the presence of the low‐velocity zone. This indicates that the subducting African slab has experienced major slab‐tear beneath southwestern Anatolia, and the gap in the slab may be a channel through which asthenospheric material is rising up to the uppermost mantle [e.g., Chang et al., 2010; Biryol et al., 2011; Salaün et al., 2012]. The crustal thinning in the Menderes Massif area is partly attributed to the hot asthenospheric material in the upper mantle and extensional tectonics related to the  southwest retreating Hellenic trench and westward ­movement of the Anatolian Plate. The authors suggest that hot asthenospheric material in the upper mantle may have induced thermal erosion in the overlying crystalline basement and the lower crust. They use the slow average shear wave velocity [e.g., Delph et al., 2015] of the crust in southwestern Anatolia as a line of evidence to indicate a thermally altered crust. Moreover, the presence of vol­ canic centers and high geothermal gradients in the Menderes Massif complex indicate the existence of asthe­ nospheric flow beneath southwestern Anatolia. The gravity model in this chapter suggests that crust thickens from southwestern Anatolia toward the Hellenides in western Greece and central Anatolia in Turkey, respec­ tively. The regions outside the asthenospheric window show, by far, the largest crustal thickness (30–42 km). This basically leads to the conclusion that the observed

crustal thinning in southwestern Anatolia may be partly attributed to thermal erosion induced by an upwelling hot asthenosphere and extensional tectonics related to the southwest retreating Hellenic trench and westward movement of the Anatolian Plate. 1.1.2.2. Chapter 5. Geodynamical Models for Litho­ spheric Delamination in an Orogenic Setting by O. Gögüs, R. Pysklywec, and C. Faccenna In this chapter, the authors use a synthesis of geologi­ cal, geophysical, and petrological data to infer that a por­ tion of the mantle part of the lithosphere may have been removed from beneath the crust in several orogenic regions. To quantify the response to delamination, they applied numerical and laboratory‐based analogue experi­ ments. Numerical model predictions show that the lith­ ospheric delamination is associated with broad surface uplift due to the thermal and isostatic effect driven by mantle upwelling. They claim that mantle lithosphere delamination can occur with slow plate convergence, where the slab peels off/rolls back similar to a retreating ocean slab subduction. The results suggest that continental delamination may be a natural progression from prior ocean plate subduction and illustrate also that the removal of mantle lithosphere does not necessarily require a significant density hetero­ geneity to initiate. Their experiments reveal that when the plate convergence is higher, the mantle lithosphere is less prone to delaminate from the crust. With higher plate convergence, the consumed mantle lithosphere can drape forward instead. The proplate crust separates from the mantle lithosphere only at the collision zone and is over­ thrusted/accreted on top of the retroplate. The numerical results may satisfy geological and geophysical observa­ tions for the East Anatolia plateau uplift that occurred since the last 13 Myrs. The delaminating slab may produce subsidence over the crust in response to the migration of the mantle lithosphere. The surface uplift may increase with higher plate convergence. Laboratory based experi­ ments show that slower plate convergence with retreating ocean lithosphere subduction can develop into delamina­ tion whereas for the experiments with higher plate con­ vergence, the crust above the consumed mantle lithosphere becomes accreted on the retro‐plate similar to flake tectonics. 1.1.2.3.  Chapter  6. Major Problems of  Western Anatolian Geology by Y. Yılmaz The western Anatolian and Aegean regions have long been known to represent a broad zone of N‐S extension stretching from Bulgaria in the north to the Hellenic arc in the south [McKenzie, 1972, 1978; Jackson and McKenzie, 1978; McKenzie and Yılmaz, 1991; Taymaz, 1996]. Under the close tectonic control of the extension, the western

Neotectonics and Earthquake Potential of the Eastern Mediterranean Region: Introduction  5

Anatolian region is characterized by a number of approxi­ mately E‐W trending, subparallel, normal fault zones, which border a swarm of grabens and the intervening horst blocks. As a consequence of this, there is an intense seismic activity. The author defines the aim and approach adopted in the chapter as follows: despite a pile of new data that has been collected during the last two decades, some major problems of western Anatolian geology still remain con­ troversial. Among these cause and timing of generation of the Menderes Massif and the magmatic associations, the N‐S trending grabens, and time of inception and con­ tinuity of the E‐W grabens are at the forefront. A number of different views have been proposed on each one of these subjects. Models proposed by different authors were commonly incompatible with one another. As a consequence of the nature of the problem, to establish a cross connection between the different events and to evaluate them in time‐space and regional geological ­ ­perspective is critical. In this chapter, main geological entities of western Anatolia are reviewed under separate headings, the ongo­ ing controversies around them are discussed first, and then some solutions are proposed. 1.1.2.4. Chapter 7. The Çataldağ Plutonic Complex in Western Anatolia: Roles of  Different Granites on the Crustal Build‐up in Connection with the Core Complex Development by O. Kamacı, A. Unal, S. Altunkaynak, M. Z. Billor, S. Georgiev, P. Marchev This chapter provides a detailed geological map of the Çataldağ area of western Anatolia, Turkey, accompanied by structural and geochemical data set to review origin of granites generated during the Neotectonics extensional setting. The metamorphic core complex in the Çataldağ area was exhumed in Early Miocene as a dome structure in the footwall of a low‐angle detachment surface. A number of micro‐ and mesoscale shear sense indicators display evidence that the rocks underwent ductile defor­ mation in the earlier stage of the exhumation, which was superimposed later by a semibrittle and brittle deformation. They indicate a top‐to‐north and top‐to‐NE sense of movement. The exhumation process was partly contem­ poraneous with the development of the major core complexes of the region (e.g., the Menderes and Kazdağ massifs) as a result of combined effects of thermal weak­ ening and rollback of the Aegean subducted slab during the Oligocene–Early Miocene. Closely associated with the development of the core complex, this study documents in detail, geology, structure, and age of the Çataldağ Plutonic Complex (CPC) as the main rock association within the footwall of the Çataldağ Detachment surface. CPC consists of two contrasting granitic bodies; an older granite‐gneiss‐migmatite complex (GGMC) and a

younger I‐type granodioritic body: Çataldağ granodior­ ite (CG). The former is a heterogeneous body consisting of migmatite, gneiss, and two‐mica granite, and represents a deep‐seated pluton. By contrast, the latter represents a discordant, shallow level intrusive body. New U‐Pb zir­ con (LA‐ICPMS) and monazite ages of GGMC yield magmatic ages of 33.8 and 30.1 Ma (Latest Eocene‐Early Oligocene). The 40Ar/39Ar muscovite, biotite, and K‐ feldspar from the GGMC yield the deformation age span 21.38 ± 0.05 Ma and 20.81 ± 0.04 Ma, which is also the age of the emplacement (20.84 ± 0.13 Ma and 21.6 ± 0.04 Ma) of ÇG. The age data when evaluated together with the contact relationships, internal petrological, and primary structural textural features indicate collectively that the two plutons were formed at different times, and were emplaced at different levels in the crust. 1.1.3. Part III: Seismotectonic in the Eastern Mediterranean Region This section includes three chapters dealing with the recent earthquakes in the eastern Mediterranean region. 1.1.3.1. Chapter 8. Fault Structures in Marmara Sea (Turkey) and Their Connection to Earthquake Generation Processes by M. Aktar This chapter investigates seismotectonics of the North Anatolian fault around the Marmara basin based on data  previously derived from bathymetry and seismic reflection profiles. The investigation concentrates on the high‐resolution seismological data collected in recent years to verify if earthquake occurrences are conformal with the structural elements. The chapter contains a short compilation of the structural elements in the Marmara basin and evaluates the high‐resolution seismological data. The author also analyzed sensitivity limits of the seismological data in detail and determined error bounds. In the major part of the chapter, the seismicity and inferred fault structures are analyzed in detail for the western high, central, and Kumburgaz basins in the Marmara basin. The chapter concludes that a single rectilinear fault plane is likely to stand as the single source for the majority of earthquakes occurring along the central axis of the Marmara Sea. A single fault plane hypoth­ esis is seen to be largely supported by the seismological observations. The western Marmara high is modeled as a pressure ridge. The central Marmara basin is con­ firmed to reflect a negative flower structure. No clear evidence is found for a major step over or pull‐apart structure. The chapter concludes also that resolution of seismological data is insufficient to study small‐scale secondary fractures such as Riedel structures along the single rectilinear fault.

6  ACTIVE GLOBAL SEISMOLOGY

1.1.3.2.  Chapter  9. The  North Aegean Active Fault Pattern and the 24 May 2014, Mw 6.9 Earthquake by S. Sboras, A. Chatzipetros, and S. Pavlides This chapter provides an excellent overview of the Aegean geodynamics with a particular emphasis on the active fault geometry of northern Greece and especially the North Aegean Trough (NAT). Findings of this study may be summarized as follows. The North Anatolia fault extends westward from the Sea of Marmara and the Gulf of Saros into the Aegean Sea. The fault strike changes from WSW‐ENE in the Gulf of Saros and Samothraki Island, to SW‐NE south of Chalkidiki Peninsula and reaches to the coast of the Greek mainland (Thessaly), where it terminates. The fault displays almost pure strike‐ slip character within Turkish territory, while it shows oblique‐slip to normal sense of movement in the North Aegean Sea (transtensional tectonics). The causative fault of the 24 May 2014 strong earthquake is a segment (45 km long and 12 km wide) of the NAT, part of the North Anatolian fault (NAF) system, located offshore between Samothraki and Lemnos islands. This interpre­ tation is supported by the earthquake epicenter, the aftershock distribution, and the seafloor morphology. In this chapter, an ENE‐WSW striking right lateral strike‐slip, SSE‐dipping fault plane has been modeled. The receiver faults have been modeled according to the Greek Database of Seismogenic Sources (GreDaSS) and include both Individual Seismogenic Sources (ISSs) and Composite Seismogenic Sources (CSSs). The static stress change after the 2014 mainshock on the nearby faults shows that only the immediately eastern segment of the “North NAT” CSS (CSS290), that is, the “Samothraki SE” (ISS‐ISS291) bears stress rise. This can explain the eastern aftershock cluster that lies along its fault plane. Static stress rises on the Samothraki SE ISS and triggering effect could be expected. Although this source was reacti­ vated during the 1975 earthquake, rapidly deforming crust in this region and the effects of other earthquakes since then, either strong or weak, left the triggering issue open to discussion. Moreover, it is not clear how the 2014 aftershock eastern cluster affects the stress state of the fault. The normal dip‐slip “North Samothraki” ISS (ISS288) is situated in the stress drop area, as well as the entire Samothraki Island (for faults of similar geometry and kinematics). The last fault that is affected by the 2014 rupture is the “South NAT” CSS (CSS800), which is almost entirely situated in the stress drop area, while a small part of it (toward its northeastern tip) is lying in an insignificant stress rise area. The 2014 earthquake fault plane rupture was not enough for the static stress change to reach more distant faults (“Saros Gulf”: ISS290, “Athos”: ISS282, “NAB segment A”: ISS810 and “NAB segment B”: ISS811). More impor­ tantly, the “Athos ” ISS, which is located at the western cluster of the aftershock sequence, is too far away from any

calculated stress change. Thus, static stress transfer cannot explain the nucleation of the western cluster. The “Lemnos” CSS (CSS825) was intentionally left out of the calculations, due to the presence of several similar faults on the northern part of Lemnos Island. However, the effects of stress changes can be inferred from receiver faults with similar properties that have been part of the calculations, such as the “NAB segment B” ISS and the “South NAT” CSS. Thus, for this kind of receiver fault, the entire island demonstrates stress drop and, hence, a probable earthquake delay. 1.1.3.3.  Chapter  10. Seismic Intensity Maps for the  Eastern Part of  North Anatolian Fault Zone Turkey Based on Recorded and Simulated Ground Motion Data by A. Askan, S. Karimzadeh, and M. Bilal This chapter provides synthetic intensity maps for a selected set of earthquake scenarios for the sparsely monitored and relatively unstudied eastern part of the North Anatolian fault zone (NAFZ). The maps are pro­ duced to evaluate connections between intensity and peak ground motion values. The study focuses on the eastern segments of the NAFZ around the Erzincan region where there are only sparse seismic networks. The city of Erzincan in eastern Turkey is located in the area where three active faults intersect: the North Anatolian, Northeast Anatolian, and Ovacik faults. The city center is in a pull‐apart basin underlain by soft sediments, which significantly amplify the ground motions. The seismicity in the region through ground motion simulations is used for potential earthquake scenarios of various magni­ tudes. The combination of the tectonic and geological settings of the region have led to destructive earthquakes such as the 27 December 1939 (Ms = 8.0) and the 13 March 1992 (Mw = 6.6) events resulting in extensive losses. In this chapter, first ground motion simulations for a set of hypothetical events as well as the 1992 Erzincan earthquake are performed. Second, local relationships between MMI (Modified Mercalli Intensity) and PGA (Peak Ground Acceleration) as well as PGV(Peak Ground Velocity) are utilized to obtain the correspond­ ing MMI values. The study presents the results in the form of synthetic intensity maps for the 1992 event and the earthquake sce­ narios. The maps are useful for the earthquake hazards reduction program in the region, especially within the area of the city of Erzincan where a devastating earth­ quake of Ms = 8.0 occurred in 1939. ACKNOWLEDGMENT We thank very much Ms. Rituparna Bose and Mary Grace Hoboken‐Hammond for their endless support and encouragement during the preparation of this volume. We are indeed in debt to them. This volume would not be

Neotectonics and Earthquake Potential of the Eastern Mediterranean Region: Introduction  7

realized without their relentless pursuit for perfection and constant push on us, as editors, our contributors, and reviewers. A number of reviewers provided their reviews for the chapters in this volume. These reviews definitely helped us to elevate the scientific standards of the chapters included in the volume. We would like to take this oppor­ tunity to thank our colleagues who shared their geological knowledge freely with us over the years. We learned tre­ mendously from these interactions. Last but not least, we would like to thank our families who supported us through­ out the completion of this volume. We would also like to thank our families for supporting our constant desire to contribute to the science of geology, and, in many respects, paying the price of not seeing us enough due to our con­ stant travels as geologists throughout the years. REFERENCES Aksu, A. E., J. Hall, and C. Yaltirak (2005), Miocene to recent tectonic evolution of the eastern Mediterranean: New pieces of the old Mediterranean puzzle, Marine Geol., 221, 1–13. Altunkaynak, S. A., and Y. Dilek (2006), Timing and nature of postcollisional volcanism in western Anatolia and geodynamic implications, Geological Society of America Special Paper. Biryol, C. B., S. L. Beck, G. Zandt, and A. A. Özacar (2011), Segmented African lithosphere beneath the Anatolian region inferred from teleseismic P‐wave tomography, Geophys. J. Int., 184(3), 1037–1057. Çemen, I.̇ (2010), Extensional tectonics in the basin and range, the Aegean, and western Anatolia: Introduction, 1–6, in Extensional Tectonics in the Basin and Range, the Aegean, and Western Anatolia, edited by I. Çemen, Tectonophysics, 488. Çemen, I.,̇ C. Goncuoglu, and K. Dirik (1999), Structural evolution of the Tuzgolu basin in central Anatolia, Turkey, J. Geol., 107, 693–706; doi: 10.1086/314379. Çemen, I.,̇ C. Helvaci, and Y. Ersoy (2014), Cenozoic exten­ sional tectonics in western and central Anatolia, Turkey: Introduction, Tectonophysics, 635, 80–99, 10.1016/j.tecto. 2014.09.004. Çemen, I.,̇ E. J. Catlos, O. Göğüs, and C. Özerdem (2006), Postcollisional extensional tectonics and exhumation of the Menderes massif in the western Anatolia extended terrane, Turkey, Geological Society of America Special Papers, 409, 353–379. Chang, S. J., S. van der Lee, M. P. Flanagan, H. Bedle, F. Marone, E. M. Matzel, and C. Schmid (2010), Joint inversion for three‐dimensional S velocity mantle structure along the Tethyan margin, J. Geophys. Res. Sol. Earth (1978–2012), 115(B8). Cosentino, D., T. F. Schildgen, P. Cipollari, C. Faranda, E. Gliozzi, N. Hudáčková, S. Lucifora, and M. R. Strecker (2012), Late Miocene surface uplift of the southern margin of the central Anatolian plateau, Central Taurides, Turkey, Geol. Soc. Am. Bull., 124, 133–145. Delph, J. R., C. B. Biryol, S. L. Beck, G. Zandt, and K. M. Ward (2015), Shear wave velocity structure of the Anatolian plate: Anomalously slow crust in southwestern Turkey, Geophys. J. Int., 202(1), 261–276.

Di Luccio, F., and M. E. Pasyanos (2007), Crustal and upper‐mantle structure in the eastern Mediterranean from the analysis of sur­ face wave dispersion curves, Geophys. J. Int., 169(3), 1139–1152. Ersoy, Y., I. Çemen, C. Helvaci, and Z. Billor (2014), Tectono‐ stratigraphy of the Neogene basins in western Turkey: Implications for tectonic evolution of the Aegean extended region, Tectonophysics (2014), 10.1016/j.tecto.2014.09.002. Faccenna, C., L. Jolivet, C. Piromallo, and A. Morelli (2003), Subduction and the depth of convection in the Mediter­ ranean mantle, J. Geophys. Res., 108 (B2), 2099, 10.1029/ 2001JB001690. Faccenna, C., O. Bellier, J. Martinod, C. Piromallo, and V. Regard (2006), Slab detachment beneath eastern Anatolia: A possible cause for the formation of the north Anatolian fault, Earth Planet. Sci. Lett., 242, 85–97. Hall, J., E. A. Aksu, I. Elitez, C. Yaltırak, and G. Çifçi (2014), The Fethiye‐Burdur fault zone: A component of upper plate extension of the subduction transform edge propagator fault linking Hellenic and Cyprus arcs, eastern Mediterranean, Tectonophysics, 635, p. 80–99. Jackson, J., and D. McKenzie (1988), The relationship between plate motions and seismic moment tensors, and the rates of active deformation in the Mediterranean and Middle East, Geophys. J., 93, 45–73. Jolivet, L., C. Faccenna, B. Huet, L. Labrousse, L. Le Pourhiet, O. Lacombe, E. Lecomte, E. Burov, Y. Denèle, J. P. Brun, M. Philippon, A. Paul, G. Salaün, H. Karabulut, C. Piromallo, P. Monié, F. Gueydan, A. I. Okay, R. Oberhänsli, A. Pourteau, R. Augier, L. Gadenne, O. Driussi (2013), Aegean tectonics: Strain localisation, slab tearing and trench retreat, Tectonophysics, 597, 1–33. McKenzie, D. (1972), Active tectonic of the Mediterranean region, Geophys. J. R. Astrol. Soc., 30, 109–185; doi: 10.1111/ j.1365‐246X.1972.tb02351.x. McKenzie, D. P.(1978), Some remarks on the development of the sedimentary basins, Earth Planet. Sci. Lett., 40, 25–32. Mc Kenzie, D., and Y. Yılmaz (1991), Deformation and volcan­ ism in western Turkey and the Aegean, Bull. Tech. Univ. Istanbul, Spec. Issue on Tectonics, 44, 345–373. Pe‐Piper, G., and D. J. Piper (2006), Unique features of the Cenozoic igneous rocks of Greece, Geological Society of America Special Papers, 409, 259–282. Pe‐Piper, G., and, D. J. Piper (2007), Neogene backarc volcan­ ism of the Aegean: new insights into the relationship between magmatism and tectonics, Geological Society of America Special Papers, 418, 17–31. Prelević, D., C. Akal, S. F. Foley, R. L. Romer, A. Stracke, and P. Van Den Bogaard (2012), Ultrapotassic mafic rocks as geo­ chemical proxies for post‐collisional dynamics of orogenic lithospheric mantle: The case of southwestern Anatolia, Turkey, J. Petrol., 53(5), 1019–1055. Reilinger, R., S. McClusky, D. Paradissis, S. Ergintav, and P. Vernant (2010), Geodetic constraints on the tectonic evolu­ tion of the Aegean region and strain accumulation along the Hellenic subduction zone, Tectonophysics, 488(1), 22–30. Robertson, A. H. F., and J. E. Dixon. (1984), Aspects of the geological evolution of the Eastern Mediterranean, 1–74, in The geological evolution of the eastern Mediterranean, edited by J. E. Dixon and A. H. F. Robertson, Spec. Publ. Geol. Soc., 17, London, Blackwell Scientific.

8  ACTIVE GLOBAL SEISMOLOGY Salaün, G., H. A. Pedersen, A. Paul, V. Farra, H. Karabulut, D. Hatzfeld, and C. Pequegnat (2012), High‐resolution surface wave tomography beneath the Aegean‐Anatolia region: Constraints on upper‐mantle structure, Geophys. J. Int., 190(1), 406–420. Schildgen, T. F., C. Yıldırım, D. Cosentino, and M. R. Strecker (2014), Linking slab break‐off, Hellenic trench retreat, and uplift of the central and eastern Anatolian plateaus, Earth‐ Science Rev., 128, 147–168. Schildgen, T. F., D. Cosentino, A. Caruso, R. Buchwaldt, C. Yıldırım, S. A. Bowring, B. Rojay, H. Echtler, and M. R. Strecker (2012), Surface expression of eastern Mediterranean slab dynamics: Neogene topographic and structural evolution of the southwest margin of the central Anatolian plateau, Turkey, Tectonics, 31(2).

Şengör, A. M. C., and Y. Yilmaz (1981), Tethyan evolution of Turkey: A plate tectonic approach, Tectonophysics, 75(3), 181–241. Şengör, A. M. C., N. Görür, and F. Saroglu (1985), Strike‐slip deformation, basin formation and sedimentation: Strike‐slip faulting and related basin formation in zones of tectonic escape, Turkey as a case study, Society of Economic Paleontologists and Mineralogists, Special Publication, 37, 227–264. Şengör, A. M. C., O. Tuysuz, C. Imren, M. Sakınc, H. Eyidogan, N. Gorur, X. Le Pichon, and C. Rangin (2005), The north Anatolian fault: A new look, Ann. Rev. Earth Planet. Sci., 33, 37–112; doi: 10.1146/annurev.earth.32.101802.120415 Taymaz, T. (1996), Wave travel‐time residuals from earthquakes and lateral inhomogeneity in the upper mantle beneath the Aegean and the Aegean trench near Crete, Geophys. J. Int., 127, 545–558.

Part I Morphotectonic Characteristics of Neotectonics in Anatolia and Its Surroundings

2 Morphotectonic Development of Anatolia and the Surrounding Regions Yücel Yılmaz

ABSTRACT In Anatolia, the late–post orogenic deformation is still continuing severely today. The north-south compression initially deformed eastern Turkey. East and central Anatolia began to rise together. Slab break off of the northerly subducting plate and lithospheric delamination played a significant role in this. Compression then formed the North and East Anatolian transform faults. They defined the Anatolian Plate, which began escaping to the west. Major morphotectonic features, the peripheral mountains, the Pontides and the Taurides, and the western Anatolian horsts and grabens, have formed from this time onward. The transform faults played roles in the transfer of the stress. Compared to the east, the west of Anatolia has followed a different path of morphotectonic development. The region was a highland during the Early Miocene period, while eastern Anatolia was under shallow sea. The environments began to change in opposite ways from Late Miocene onward. Southerly retreat of the subducting eastern Mediterranean oceanic slab has generated north-south extension on the upper plate. As a result, the present morphology began to develop. The east-west trending horst and graben structures that dominate the landscape today began to form during later periods of the extension in the Quaternary. 2.1. INTRODUCTION

and major tectonic elements. This is a first attempt in this direction to establish a platform for future studies. Anatolia is located in the middle of the Alpine‐Himalayan mountain ranges. The orogeny started north of the Anatolian orogenic belt (the Pontides) and migrated progressively to the south and ended up in the Bitlis‐Zagros orogenic belt (Fig. 2.1). The late‐post orogenic deformation is still continuing severely as expressed by GPS data, which display clearly that Anatolia is moving in an anticlockwise sense from the east to the west (Fig. 2.2). The Arabian Plate is moving in a north‐northwest direction. Rate of motion of the Anatolian Plate increases to the west from 15–20 mm/yr to 30–35 mm/yr [Reilinger et al., 2006, 1997, 2006, 2010; Kahle et al., 2000; Taymaz et al., 2007; and references therein]. This motion is controlled by the two major parameters: (1) the Arabian Plate steadily advancing northward and bulldozing eastern Anatolia, and (2) the Hellenic trench, along which the eastern Mediterranean Oceanic Lithosphere is being consumed

Anatolia is being deformed presently under an ongoing severe post‐late orogenic tectonic regime. The two lines of evidence manifest this: (1) frequent earthquakes that occur in a vast terrain from the east to the west, and (2) a rugged, irregular, and tectonically controlled morphology. In order to understand the tectonics of Anatolia, structural analyses of the tectonically different regions and the earthquakes are studied extensively using a variety of methods and techniques. Some local and regional morphological studies were also conducted on different parts of Anatolia. However, a morphological treatment of the whole of Anatolia has not yet been attempted to evaluate interactions of different morphotectonic regions ̇ Department of Geological Engineering, Istanbul Technical ̇ University, Istanbul, Turkey

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 11

12  ACTIVE GLOBAL SEISMOLOGY

Figure 2.1  Morphotectonic map of Anatolia [modified from Yolsal‐Çevikbilen et al., 2012]. Global topography data taken after USGS. Bathymetry data are derived from GEBCO/97–BODC, provided by GEBCO [1997] and Smith and Sandwell [1997a, b]. Summary sketch map of the faulting and bathymetry in the eastern Mediterranean region, compiled from our observations and those of Le Pichon and Angelier [1981], Taymaz [1990], Taymaz et al. [1990, 1991a, b], Şaroğlu et al. [1992], Papazachos et al. [1998], McClusky et al. [2000], and Tan and Taymaz [2006]. Big black arrows show relative motions of plates with respect to Eurasia [McClusky et al., 2003]. Bathymetry data are derived from GEBCO/97–BODC, provided by GEBCO [1997] and Smith and Sandwell [1997a, b]. Shaded relief map derived from the GTOPO‐30 global topography data taken after USGS. The yellow lines with red glow are peripheral mountains: P = Pontide mountain range; T = Tauride mountain range; IA = Isparta angle; MA = Misis‐Andırın wrench fault zone; ED = Erzincan depression; EACP = Eastern Anatolian compressional province; CATP = Central Anatolian transitional province; AWAEP = Aegean‐western Anatolian extensional province. The red ellipse and white arrow indicates orthogonal shortening area within the East Anatolian compressional province. The black and red arrows indicate directions of motions. NAF = North Anatolian fault; EAF = East Anatolian fault; DSF = Dead Sea fault; NEAF = northeast Anatolian fault; EPF = Ezinepazarı fault; PTF = Paphos transform fault; CTF = Cephalonia transform fault; PSF = Pampak‐Sevan fault; AS = Apsheron sill; GF = Garni fault; OF = Ovacık fault; MT = Muş thrust zone; TuF = Tutak fault; TF = Tabriz fault; KBF = Kavakbaşı fault; MRF = Main Recent fault; KF = Kağızman fault; IF = Iğdır fault; BF = Bozova fault; SaF = Salmas fault; SuF = Sürgü fault; G = Gökova graben; BMG = Büyük Menderes graben; Ge = Gediz graben; Si = Simav graben; BuF = Burdur fault; BGF = Beyşehir Gölü fault; TF = Tatarlı fault; SuF = Sultandağ fault; TGF = Tuz Gölü fault; EcF = Ecemiş fault; ErF = Erciyes fault; DF = Deliler fault; MF = Malatya fault; KFZ = Karataş‐Osmaniye fault zone (part of MA); K = Kaçkar Mountain; D = Düzce.

by subducting northward under the Aegean‐Anatolian Plate (Fig. 2.2). The continuing northward advance of the Arabian Plate and the resulting compression has been accommodated in eastern Anatolia by shortening deformation. As a consequence, the East Anatolian crust has been thickened and the region was elevated to form the East

Anatolian‐Iranian high plateau to the end of Miocene (Fig.  2.1). Further shortening generated two transform faults, namely the North Anatolian transform fault (NATF) and the East Anatolian transform fault (EATF). The two transform faults together define an independent tectonic entity known as the Anatolian Plate, which has been protruding away from the thickened and shortened

Morphotectonic Development of Anatolia and the Surrounding Regions  13

Figure 2.2  GPS observations of relative motion of Anatolia with respect to the Africa‐Arabia‐Eurasia plates [modified from McClusky et al., 2000; Reilinger et al., 2006]. The red arrow displays counterclockwise sense of motion of Anatolia. The yellow arrows display different circular motions of the northern (Pontides) and the southern (Taurides) peripheral mountain ranges of Anatolia. The white arrows are the major components of the stress field that effects the Pontides, and the light green arrows are the resultant.

eastern Anatolian region to the west possibly from late Pliocene‐Pleistocene onward (Figs.  2.1 and 2.2). In the Aegean region, such westward escape of the Anatolian Plate accelerated a north‐south extensional regime (Fig.  2.1). Combination of forces, the pull from the subduction zone and push from the convergent zone, is collectively causing the Turkish Plate to rotate southwestward (Fig. 2.2). The Anatolian orogenic belt is the product of the Tethyan oceanic realm. Initially the PaleoTethyan Ocean during the Late Paleozoic–Early Mesozoic period, and then the NeoTethyan Ocean from Early Mesozoic to the Tertiary Period are responsible from the development of this complex orogenic belt (see Ş engör [1979b]; Ş engör et al. [1980, 1982]; Ş engör and Yılmaz [1981]; Ş engör et al. [1985] for the PaleoTethyan and NeoTethyan evolution of the eastern Mediterranean regions). During the NeoTethyan period, some continental slivers rifted off of the northern edge of the Arabian Plate and new branches of the NeoTethyan oceans, marginal basins, related arcs, and back arc basins were developed, and then they were all destroyed during the demise of the oceanic environments. This began during the Late Mesozoic, and continued to the end of Tertiary. The collision of the NeoTethyan tectonic entities began in the north, and migrated to the south from Late Cretaceous–Early Eocene to Miocene [Ketin, 1966; Ş engör and Yılmaz, 1981; Yılmaz, 1993;

Yılmaz et  al., 1997]. Eventually, the continental slivers welded together to form the present amalgamated ­tectonic entities, the Anatolian orogenic belt. There is a critical period in the geologic‐tectonic and particularly morphotectonic history of Turkey that ­corresponds to a change from the PaleoTectonic to the NeoTectonic eras. Although the term NeoTectonic is widely used in geological literature, its application appears to be rather vague. Some use it as a time frame for referring to events that happened during or after Miocene or Neogene. Some others use it irrespective of any time‐space reference. Since we also use this division extensively in our narrative, we find it useful to state at the beginning of the chapter that we use it in the same sense as was first defined by Ş engör [1979a], and we will briefly explain the concept again in the following. During the late stage of the development of the Bitlis‐ Zagros orogenic belt, the suture mountain [Yılmaz, 1993] began elevating between the converging jaws of the Arabian Plate and the northerly located amalgamated tectonic entities during the Middle Miocene time [Yılmaz, 1993], and then the whole of eastern Anatolia and possibly central Anatolia began to rise as a coherent block, approximately to the end of the Late Miocene time. This rapid elevation may be the consequence of the break off of the subducting plate under East Anatolia (Piromallo

14  ACTIVE GLOBAL SEISMOLOGY

and Morelli, [2003]; see Chapter 3 for a discussion on this and related topics). As northward advance of the Arabian Plate continued after the collision, the compressive forces were initially accommodated along the orogenic belt, and then it began to deform the regions farther north more affectively. As a result of this, the east of Turkey (the region of eastern Anatolia and the Pontides) began to be severely deformed (the Neogene successions were tightly folded and faulted). The two transform faults, namely the northern Anatolian transform fault (NATF) and the East Anatolian transform fault (EATF) were formed during this period, when the shortening of the East Anatolian crust reached an excessive stage that could no longer be accommodated within the volume of eastern Turkey alone. From this time onward, the land bounded by the two transform faults defined an independent tectonic entity known as the Anatolian Plate, which began to move away from the point of convergence to the west. This event began possibly during the late Pliocene‐ Pleistocene time (see Chapter  3 for a discussion on the topic). This is the time of initiation of the NeoTectonic era. From this time onward, the morphotectonic nature of the eastern Mediterranean region and particularly Anatolia began to  change drastically. Three distinctly different tectonic zones were developed as a result of this event [Şengör, 1979a], (1) eastern Anatolian compressional province, (2) western Anatolian extensional province, and (3) the transitional zone of central Anatolia (the Ova region of Şengör [1979a]) that began to behave as a bridge between East and West Anatolia (Fig. 2.1). The eastern Mediterranean as a tectonic domain covers a vast region, which includes the following major geographic divisions. The eastern Mediterranean region (sensu stricto), northwestern part of the Arabian Peninsula, the Anatolian Peninsula, the Aegean Sea region. Each one of these geographic domains represents a distinctly different tectonic entity. The Arabian Plate represented initially a passive continental margin and remained as the passive continental margin till the present orogenic belt began to form during the Late Tertiary period (see Yılmaz [1993] for a detailed account of the development of the southeast Anatolian orogenic belt). The Anatolian Peninsula represents an orogenic belt from the northern edge to the southern edge. Along the edges lie approximately east‐west trending mountain ranges known as the Pontides and the Taurides in the north and south, respectively. They are peripheral mountain belts, which rise steeply like a wall from the surrounding seas, and separate the interior of Anatolia from the Black Sea and the Mediterranean Sea. During the initial stage, the uplift of eastern and central Turkey and the peripheral zones accompanied the elevation. Later the border zones were elevated to higher levels and formed the present peripheral mountains.

The Aegean region is an extensional tectonic province. Development of this tectonic entity is closely connected with the surrounding tectonic units, demise of the eastern Mediterranean Ocean, and the western escape of the Anatolian Plate (Fig. 2.1). The Black Sea is delimited in the south by the Pontides. Thus, development of the southern part of the Black Sea has been under the strict tectonic control of the Pontide Range. Each of these tectonic provinces may be divided into some subdivisions. Within the eastern Mediterranean Sea region (Fig.  2.3), there are remnant oceanic basins (e.g., the Herodotus and Levantine basins), young rifted marine basins (most of those are located in the south of the Tauride Range such as the Rhodes, Antalya, Cilicia‐ ̇ Adana, and Iskenderun basins), the trenches (e.g., the Hellenic, Pliny, Strabo, and Pytheus‐Cyprus trenches), seamounts (e.g., the Anaximander and Eratosthenes seamounts), and some ridges (e.g., the Mediterranean, Latacia, and Larnaca ridges). In the following paragraphs, outlines of the major tectonic, morphotectonic, and seismotectonic characteristics of the main tectonic provinces are summarized. 2.2. SOUTHEAST ANATOLIA (NORTHWESTERN PART OF THE ARABIAN PLATFORM) The northern periphery of the Arabian Plate curves around the southeast Anatolian orogenic belt, which delimits the Arabian platform in the north (Fig.  2.1). Approaching the mountain ranges, the Arabian platform turns to a foreland fold and thrust belt (Figs. 2.1, 2.4, 2.5, and 2.6). From the interior toward the orogenic belt, folding and tightening are gradual. Within the 30–100 km zone from the mountain range on the western part of the Arabian platform another gradual change is noticed. The folds and the closely associated faults began to display an en echelon pattern (Fig.  2.5). This is due to the replacement of the orthogonal shortening in the central part (apex of the curve) by an escape tectonic regime as expressed by the development of the oblique faults having strike‐slip and reverse‐slip components. The southerly overturned folds are truncated by the left‐stepping oblique faults [Yılmaz et al., 1985; Yılmaz and Yıldırım, 1996; Yiğitbaş and Yılmaz, 1997] (c in Fig. 2.5). The seismic data and the analyses of the frequently generated earthquakes also support motions along the oblique slip faults (Fig. 2.7 a, b). The northern edge of the Arabian platform in the southeast Anatolian region was a site of marine environment, which survived with three major interruptions from the Cambrian to the end of Middle Miocene (Fig. 2.4). Starting from the Late Miocene, a continental environment replaced the marine environment [Tuna, 1973; Yılmaz et al., 1988; Yılmaz, 1993; Yılmaz and Duran, 1997; Siyako et al., 2013].

Figure 2.3  Major tectonic units of the eastern Mediterranean region and their extension to the neighboring land areas. LvB = Levantine basin; TO = Troodos ophiolite; CB = Cyprus basin; LR = Larnaca ridge; GR = Gelendzik rise; KBB = Kiti Baer Bassit unit; LB = Larnaca basin; BB = Baer Bassit; DSFZ = Dead Sea fault zone; KD = Keldağ horst; KZO = Kızıldağ ophiolite; AD = Amik depression; Ama M = Amanos Mountains; ABF = western boundary fault of the Karasu depression (the fault that separates the Amanos Mountain from the Karasu depression). The sinusoidal lines refer to the general pattern of the river network on the Arabian platform.

Figure  2.4  Geology map of southeastern Anatolia [modified from Yılmaz, 1993]. H = Hazro high‐anticline; EM = Engizek metamorphic massif; DAF = East Anatolian transform fault; MTJ = Maraş triple junction; K = Kızıldağ ophiolite. The white and red rectangles refer to the locations of Figures 2.5 and 2.6, respectively.

16  ACTIVE GLOBAL SEISMOLOGY

Figure 2.5  Morphotectonic map of the western part of the Bitlis‐Zagros suture mountains; (a) and (b) in the red rectangles are two strike‐slip fault zones (the Deliçay fault and the Türkoğlu‐Haruniye fault that are detailed in Fig. 2.9) that cut the entire width of the Amanos Mountain (see also Fig. 2.8 for their locations); C = the northern boundary fault of the Maraş basin; MD = the Maraş triple junction (depression); K‐MF = Kyrenia‐Misis fault zone (MA in Fig. 2.1); SF = Sürgü fault; A, E, and GA are the locations of Afşin and Elbistan towns and the city of Gaziantep. Yellow arrows indicate relative motion of the fault‐bounded blocks. The white arrows display some prominent foreland folds in the fold and thrust belt of the Arabian platform. The current stress state of the region is transpressional field with southwest trending sigma 1, associated strain is strike‐slip with transpressional component.

Three major phases of deformation are differentiated in the northwestern part of the Arabian platform [Yılmaz, 1993; Yılmaz et al., 1993]. These are associated with the nappe emplacements from the north (Fig. 2.4), where the NeoTethyan Ocean was located during the Mesozoic and Cenozoic periods [Ş engör and Yılmaz, 1981]. During the first two phases, giant ophiolite nappes were obducted onto the Arabian platform (Fig.  2.4; Yılmaz [1985]; Yılmaz et al. [1993]). The last phase is connected with the development of the present orogenic belt (Fig. 2.4; Yılmaz [1993]; Yılmaz et al. [1993]). The first deformation phase began during the Turonian time and lasted until the end of Early Maastrichtian. Results of this deformation phase are clearly seen in the Amanos Mountains (Figs. 2.4). The second phase occurred in the Early‐Middle Eocene time, and its imprints were observed more clearly in the central Amanos Mountains (Yiğitbaş et al. [1992], Yılmaz et al. [1993]). The effects of the last phase are observed in the thick sedimentary succession covering an age range from the Late Eocene‐Oligocene, when the frontal parts of the southerly transported nappe package first hit the Arabian platform (Figs. 2.4, 2.5, 2.6, 2.7, and 2.8; Yılmaz

[1993]; Hüsing et al. [2009], Silja et al. [2009]), and then it survived to the end of Middle Miocene. The deformation became severe particularly to the end of Middle Miocene, when the allochthonous units were thrust onto the Arabian platform (Fig. 2.4; Yılmaz [1993]). Despite the collision, the northerly advance of the Arabian Plate has continued to the present. The convergence rate is established to be about 15 mm/yr reaching to 21 mm/yr, according to GPS measurements (Fig. 2.2; Reilinger et al. [1997, 2006, 2010]; Kahle et  al. [2000]; McClusky et  al. [2000]) (for recent kinematic data from the northwest part of the region, the reader is referred to Seyrek et al. [2014]). In light of the detailed geological studies [Yılmaz, 1984a; Yılmaz et al., 1985; Yılmaz, 1990b, 1993; Yılmaz et al., 1993a; Parlak et al., 2004; Rızaoğlu et al., 2009; and references therein], the evolution of the southeastern Anatolian orogen may simply be summarized as a progressive relative southward transport of the nappes toward the Arabian Plate during the Late Cretaceous to the Late Miocene period. This caused progressive accretion of different tectonic units into the nappes (Fig. 2.4). Within the nappe stack are the fragments of a small

Morphotectonic Development of Anatolia and the Surrounding Regions  17

Figure 2.6  Geology map of western part of the Southeast Anatolian orogenic belt (Bitlis‐Zagros suture mountains). EATF: East Anatolian transform fault; SF = Sürgü fault; CSGF = Çicekli‐Savrun‐Göksun fault; IZ = imbricated zone of the southeast Anatolian orogenic belt. Its western continuation is the escape zone that has escaped away from the collision. Asi G = Asi (River) graben; DSF = Dead Sea fault zone; MTJ = Maraş triple junction (note the pair of faults on both sides of the Maraş triple junction. The two faults in the west cut the entire width of the Amanos Mountains obliquely aligning along the extension of the major branches of the East Anatolian transform fault zone). The black rectangle indicates the location of Figure 2.11. The Sürgü fault is one of the old and yet still active major faults of the orogenic belt. It separates the nonmetamorphic Elbistan ophiolite (light green area) (an intact ophiolitic slice, above which an in situ, thick sequence was deposited from Late Cretaceous to the Eocene period, prior to the emplacement of the ophiolite on to the metamorphic massifs [Yılmaz et  al., 1985,1993a] and high‐grade metamorphic ophiolite assemblage (the Berit ­ophiolite and the associated units (including the Eocene successions) (dark green area) [Genç et al., 1993; Yılmaz et al., 1993b]. Across the fault, a kilometers‐thick column of rocks was exhumed during the Eocene time. Therefore, the Sürgü fault zone acted as a major normal fault (a detachment fault?). The fault reactivated under the transpressional tectonic regime as an oblique (strike‐slip dominant) fault zone during the NeoTectonic era. Therefore, the fault has changed its character from a normal fault to an oblique fault (possibly as an inverted fault). The Çicekli‐Savrun‐Göksun (ÇSGF) and the Sarız faults are two major faults that were formed under transpressional tectonic regime. Between the two faults, a metamorphosed slice of the Taurides (consisting mainly of marble and recrystallized limestone ­succession) wedged tectonically into the southern region. The southward advance of the tectonic wedge along the Misis‐Andırın fault zone into the Adana plain is manifested by the progressive southerly migration of the Plio‐ Quaternary coarse fluvial clastic wedges [Yılmaz and Gürer, 1996]. ÇSGF presently connects the Sürgü fault and the Misis‐Andırın fault zone. The Sarız fault is also a major fault zone of the region. Across the fault the nonmetamorphic Taurus units are in direct contact with the high‐grade metamorphic rocks of the Taurus Range. These two fault zones were formed during the PaleoTectonic era and have continued as major faults during the NeoTectonic era. They have played major roles in the development of the present morphology of the region. FW represents the flysch wedge that forms when the escape tectonic regime began.

oceanic basin, a back arc basin, and an island arc [Yılmaz, 1993; Yılmaz et  al., 1993; Parlak et  al., 2004]. The nappe stack has finally collided with and welded onto the Arabian plate, when the separating ocean floor

was totally obliterated by the Late Eocene time [Yılmaz, 1993; Yılmaz et al., 1993; Hüsing et al., 2009]. In the follo­ wing period, a remnant sea, which remained between the colliding continents, began retreating westward

18  ACTIVE GLOBAL SEISMOLOGY

Figure 2.7  Maps showing (a) the regional stress distribution in northwestern part of the Arabian Plate and the Bitlis‐Zagros orogenic belt; (b) major faults (see Fig. 2.6 for the names of the faults and seismicity of northwestern edge of the Arabian platform, and the earthquake fault plane solution data [modified from Eyidoğan, 1983]; EATF = East Anatolian transform fault; DSF = Dead Sea fault.

(Continued)

Morphotectonic Development of Anatolia and the Surrounding Regions  19

toward the Mediterranean [Yılmaz, 1985, 1988, and 1993] or eastward toward the Gulf [Elmas and Yılmaz, 2007]. The evidence for this are linear flysch units, accumulated in the remnant basin, which show westward regression from the Middle Miocene to the beginning of the Late Miocene (see FW in Fig.  2.6) [Yılmaz, 1993; Özdoğan et al., 2011; Siyako et al., 2013]. The continuing northward advance of the Arabian Plate squeezed the nappe pile, and formed an imbricated zone between the nappes and the Arabian plate [Yıldırım and Yılmaz, 1991; Yılmaz, 1993; Yılmaz and Yıldırım, 1996] (Figs. 2.4 and IZ in Fig. 2.6). As the convergence continued, the southeastern Anatolian mountains began to rise. This occurred from the Late Miocene onward [Yılmaz, 1993; Yılmaz et al., 1988]. The southeast Anatolian orogenic belt may resemble the Cordilleran‐type accretionary complex. As in the case of the Circum Pacific orogenic belt, the disrupted ophiolite, mélanges, and deep marine sedimentary rocks were buried along the subduction zone and recrystallized under HP and HT conditions at a depth of 15–70 km and formed metamorphic tectonic units as exemplified from the Engizek, Pötürge, Malatya, and Bitlis, metamorphic massifs of the southeastern Anatolian orogenic belt (Fig. 2.4; Yılmaz et al. [1992, 1993]; Yılmaz and Yıldırım [1996]; Yılmaz [2011]). As a result of the successive deformation phases, the northern part of the Arabian platform successions ­suffered superimposed deformations. The late phase of deformations produced the present fold and thrust belt (Figs.  2.4, 2.5). Approaching the orogenic belt, the old basement of the Arabian Plate is seen to have been incorporated within the shortening zone of the fold and thrust belt. Basement involved thrusts are observed on the surface only in a few places, for example, the Hazro anticline‐ thrust structure, which is located close to the apex of the curve in the northeastern Diyarbakır area (h in Fig. 2.4). There, a thick‐skin deformation occurred and possibly propagated inward as the orogenic deformation evolved. This zone is narrow and does not extent much farther

south (not more than 100 km). Rarity of midcrustal seismic activity supports this view. Away from the orogenic belt, involvement of the basement is difficult to ascertain from the surface. However, some morphological features may help to assess its possible role, such as sudden vertical steps in front of the southerly inclined hills (anticlines). Away from the orogenic belt, the thick‐skin deformation is replaced by thin‐skin deformation. The seismic data (unpublished Turkish Petroleum Co. [TPAO] seismic data along and across the fold and thrust belt) show that the folds observed on the surface commonly do not penetrate to depths. Above the resistant Paleozoic basement rocks, the Mesozoic and Cenozoic carbonate successions were commonly decoupled from the base and moved independently as the Jura‐type (decollement) detachment folds. Thrusting appears to have followed the decollement. Some thrusts branching from the decollement level display progressive thrust‐fold propagation pattern (unpublished TPAO seismic data) as expressed by the fold propagation folds and fault‐bend faults. The first morphological clue to differentiate the “detachment‐fault‐ related” fold style from the “fault‐bent fault type” fold style is that the latter displays unexpectedly high topography and asymmetrical geometry in the morphology. One of the important structures of the western part of the Arabian platform is the Dead Sea transform fault zone (DSF; Figs.  2.1 and 2.3) [Girdler, 1990; Garfunkel, 1997; Garfunkel et al., 2014; and references therein]. Presently, it is a seismically active fault in southeastern Anatolia (Akyüz et al., 2006; Seyrek et al., 2014]. It is a plate boundary that defines the eastern boundary of the Arabian Plate separating it from the African Plate (Fig. 2.1). Along this boundary, the Arabian Plate moves to the north with a faster rate than the African Plate [DeMets et al., 1994; Masson et al., 2010]. It runs approximately in a north‐south direction for a distance more than a 1000 km from Gulf of Akaba in the south to the Maraş depression in the north (Figs 2.1, 2.6, and 2.8). It is a sinistral, strike‐slip fault displaying en echelon geometry. Left stepping along the fault generated a

Figure  2.8  (Continued) Morphotectonic map showing the major structural features of the northern Amanos Mountain (see Fig. 2.5 for location of the photo). See in white shaded letters: DF = Deliçay fault; THF = Türkoğlu‐ Haruniye fault. Two strike‐slip fault zones that run obliquely cutting the entire width of the Amanos Mountain Range are the Türkoğlu‐Haruniye fault (THFZ) and the Deliçay fault (DFZ). The white square shows the area where the two fault zones converge (see Fig 2.9 for the structural maps of the faults). Note the major branches of EATF are nearly connected with the Deliçay and the Türkoğlu‐Haruniye faults in the Karasu graben depression. Note also the easterly offset of the edge of the Karasu graben between the town of Türkoğlu (TR) and the village of Kılılı (KL). AD = The Ahır Dağı Anticline (horst), ADF = The Ahır Dağı fault zone, the white thick arrow indicates the principal stress direction, under which the oblique faults (the concave red lines in front of the Ahır Dağı anticline) were generated; DSF = Northern extension of Dead Sea fault that defines the western edge of the Karasu graben (KG); EBFKB = Eastern boundary fault zone of the Karasu graben. Note that along this side of the graben there are a number of small faults that collectively define the eastern boundary of the graben. MTJ = Maraş triple junction; ND, BÇ, Dc, BY, TR, KL, M, BE, PZ, CC, and GB are the locations of the settlements (pale brown shaded letters): Nurdağ, Bahçe, Düziçi (ancient Haruniye), Beyoğlu, Türkoğlu, Kilili, KahramanMaraş, Bahçelievler, Pazarcık, Çağlayan Cerit, Gölbaşı.

20  ACTIVE GLOBAL SEISMOLOGY

Figure  2.9  (a) Morphotectonic index map showing location of the Türkoğlu‐Haruniye fault zone (THF) and Deliçay fault zone (DF) within the major fault systems of the northern edge (SE Anatolian Region) of the Arabian platform. Abbreviations are the same as in Figure 2.8. (b and c) Structural maps of (b) the Deliçay fault zone and (c)Türkoğlu‐Haruniye fault zone. The red squares in (b) and (c) refer to the area where the two fault zones intersect. They correspond also to the yellow square in Figure 2.8. DFs in (c) refer to the branches of the Deliçay fault. Note the map pattern of the faults, which is characteristic for a strike‐slip fault zone. Segments of the faults are nonparallel to each other and to the overall applied shear strain. They collectively form a set of subparallel linear ridges along the trend. The ridges bound long and narrow valleys. The faults are commonly steep as the map pattern manifests. Across the faults, commonly members of an ophiolite are in contact with Mesozoic limestone successions. The fault zones are about 5 km wide. They become wider between Hodu Sazlığı and Çakıroğlu. Around the Hodu Sazlığı, good exposures of the fault are observed around Kıymatepe, Süttepe, Kocadaz Tepe hills, and the Bugansak Dere Creek. One of the major branches of the Türkoğlu‐Haruniye fault zone may be observed in the area to the north of Imalı village, where the fault scarp is subvertical, and the stretching lineations indicate left‐lateral displacement. A number of pressure ridges are recognized in the area between the north of the Imalı ̇ village and Ispirler village, where reverse‐slip component is more prominent. The structural maps of the Deliçay and Türkoğlu‐Haruniye fault zones were simplified after Yılmaz [1984, vol. 3, Figs. C22 and C29). For a detailed description of geology and morphology of the fault zones, see Yılmaz [1984, vol. 3, 430–439].

number of pull‐apart basins as exemplified from the Gharb, Hula, and Dead Sea rift basins. They are young basins. Their fill is Neogene and Quaternary in age. There is a regional uplift formed in association with the DSF, reaching up to 1 km height in the central region but dying out to the north. The age of the DSF is the subject of a

long controversy. Dates proposed on the initiation vary from Early‐Middle Miocene [Garfunkel, 1981; Garfunkel et al., 1996] to Late Miocene [Lyberis, 1988]. Along the northwestern edge of the Arabian Peninsula lies the Amanos Mountains as a north‐northeast–south‐ southwest trending range (Figs. 2.3, 2.4, 2.6, 2.9, and 2.10).

Morphotectonic Development of Anatolia and the Surrounding Regions  21

Figure 2.10  (a) Simplified geology map of the Amanos Mountains [modified from Yılmaz, 1984], which shows its main tectonic subdivisions. SA, CA, and NA correspond to the southern, central and northern Amanos regions, respectively. They are separated by strike‐slip faults, which cut the entire width of the mountain range obliquely and divide it into distinctly different tectostratigraphic zones [Yılmaz, 1984]. The oblique faults were generated probably during the Oligocene period. As a result of this, major structures of the mountain range were displaced by sinistral faults. Some of the oblique faults were reactivated during the NeoTectonic era. EATF = the East Anatolian transform fault zone; AD = Amik depression; MD = Musadağ; SF = Samandağ fault; KD = Keldağ horst‐ dome; DSF = Dead Sea fault. The black arrows show the location of (b) and (c) on the map. The red elliptical area shows the approximate coverage of the region occupied by (b). (b) The block diagram of the central Amanos Mountain (the red ellipse in [a]), which shows the ophiolite nappes that emplaced on the Arabian platform during the Middle‐Late (?) Eocene time. The nappe tectonically overlies the autochthonous succession that includes the Lower‐Middle Eocene sedimentary rocks at the top of the sequence (see the columnar Eşmişek section in [B]). The nappe is unconformably overlain by sediments of the Upper Eocene age. On the ophiolite slices a continuous, epi‐ophiolitic, deep‐sea sedimentary succession (mainly the pelagic limestone and red radiolarite) of Upper Cretaceous– Middle Eocene age span were also transported to the region (see the Değirmendere, Yılanlı, and Moralı columnar sections in [B]). (c) A view of the Karasu graben fault zone showing strike‐slip character of the graben boundary fault, manifested by the pressure ridges (the lenticular long and narrow, subparallel hills running along the graben boundary. An offset stream (green line) reflects left lateral displacements among the fault branches.

In the southern part, the mountain range slightly curves outward (Fig. 2.4). The curved area corresponds to the Amik structural depression (Figs. 2.3, 2.6, and 2.11a). Geology of the Amanos Mountains is critically important to understand evolution of the northwestern part of

the Arabian platform and the orogenic belt because this range provides critical evidence for (1) identifying more clearly the roles of each one of the nappe emplacement phases, because in this region the Miocene nappes did not extend far south to obscure the previous events

22  ACTIVE GLOBAL SEISMOLOGY

Figure 2.11  (a) Geology map of the southern Amanos Mountains [modified from Yılmaz, 1984]. AmaF = western boundary fault of the Karasu graben depression (Karasu graben fault); Amik DP = Amik depression; AF = Antakya fault (Asi graben fault); AG = Asi graben; SF = Samandağ fault; MH = Musa Dağ horst; DSFZ = Dead Sea fault zone; the black arrows correspond to the cross‐section directions that are displayed in Figure  2.11b (D), Figure 2.11c (E), Figure 2.13a (C), Figure 2.13c (A), and in Figure 2.13d (B). (The rectangle refers to the location of Fig. 2.11.) (b) A Google image displaying 3D morphology of the Asi graben and the adjacent horst. The view is from north‐northeast to south‐southwest. The red lines are the graben bounding, parallel (partly en echelon) faults. Note that the fault scarps are steep (pink arrows), and the upthrown blocks are back‐tilted to SE (white arrows). (c) Geological cross section based on the geology map displayed in (a), showing a more than 3.5 km thick rock sequence in the southern flank of the Keldağ dome (horst). The sequence covers an age span from Jurassic to Early‐Middle Miocene. From the northern flank of the anticline, the thick Mesozoic carbonate sequence was tectonically removed as a result of the tectonic uplift (exhumation; brittle thinning) during the Late Miocene. As a consequence of this event the Upper Miocene–Lower Pliocene sediments were deposited directly (unconformably) above the different units of the underlying succession (see also Fig.  2.12c). In the cross section of Figure  2.12c, the formation names of Besni (the uppermost Cretaceous), Belveren (Paleocene), and Midyat (Eocene) are regionally known rock units of southeast Anatolia.

Morphotectonic Development of Anatolia and the Surrounding Regions  23

(Fig. 2.4); (2) making correlations between major tectonic zones and units of the orogenic belt and the southeastern Anatolian units. Three distinctly different tectonic regions may be readily distinguished in the Amanos Mountains as the southern, central, and northern regions separated from one another by major fault zones (Fig.  2.10a). The three blocks, separated by the major faults, have their own distinctive lithological and tectonic character, for example., the Paleozoic basement rock in the northern region is weakly metamorphosed. In the southern Amanos, mainly the late Cretaceous nappes and related deformations are observed (Figs.  2.11c and 2.12c). In the central and the northern Amanos regions, the Eocene (Fig.  2.10a and b) and Miocene nappes (Fig. 2.4) and related deformations are distinguished [Yılmaz, 1984a]. Two young (post‐Miocene), north‐south trending structural depressions bound the Amanos Mountain range. These are the Karasu depression and the Iskenderun basin in the east and the west, respectively (Figs. 2.6, 2.10a, and 2.11a). The Karasu basin is located along the northward extension of DSF (Figs.  2.1, 2.3, and 2.9a). On the time of the extension of DSF to the Maraş depression, the views are diverse varying from the Early Miocene [Garfunkel et  al., 1996] and Late Miocene [Lyberis, 1988] to the Late Pliocene–Pleistocene [Karabacak and Altunel, 2013; Seyrek et al., 2014]. The offset along the DSF has also been calculated to vary from about 100 km in the south [Garfunkel, 1981] to 60 km in the north [Westaway, 2004]. Based on different sets of data (morphological, structural, seismological, and GPS) different slip rates have been estimated along the DSF, which varies from 3 mm to 10 mm/a. It is suggested however that the Syrian arc (Palmyride fold belt in Fig. 2.1), a northerly arching fault‐bend fold, accommodates some of the slip [Eyal, 2011; Seyrek et  al., 2014]. Therefore, in the regions located farther north with respect to the Syrian arc, particularly in the Karasu graben, the slip rate decreases significantly as a result of the slip partitioning, to the low values of 1–2 mm/y [Reilinger et  al., 2006; Masson et  al., 2010; Alchalbi et  al., 2010; Mahmoud et al., 2013]. Morphotectonic development of the Karasu depression and its immediate surroundings is complex. Problems, such as where in the north the Karasu depression terminates (was it a normal fault‐bounded graben or does it represent the northerly continuation of the Dead Sea fault zone), and how it is genetically connected with the adjacent Amik and Maraş depressions (Figs. 2.1, 2.5, and 2.6), were tackled at length in a number of studies [Ş engör, 1979a; Karabacak and Altunel, 2013; Perinçek and Çemen, 1990; Rojay et al., 2001; Yurtmen et al., 2002; Tatar et  al., 2004; Akyüz et  al., 2006; Westaway et  al.,

2006; Bolton and Robertson, 2007, 2008; Över et al., 2004; Mahmoud et  al., 2013; Masson et  al., 2010; Karabacak et  al., 2010; Karabacak et  al., 2010; Karabacak and Altunel, 2013; Duman and Emre, 2013; Seyrek et  al., 2014], and different kinematic solutions were proposed. The diversity of views is due possibly to the critical location of the region, where three major active tectonic elements converge. These are the EATF, DSF, and tectonic elements of the eastern Mediterranean region (Figs. 2.1 and 2.3). In the following paragraphs, the major morphotectonic ­features of the region and their possible role in the development of the Maraş depression will be briefly reviewed. Before this, however, major tectonic belts and the structures that are encountered in the eastern Mediterranean region will be briefly outlined to compare and correlate them with the surrounding land areas in the north and the east (Fig. 2.3) to form a base for the following discussion. Troodos Mountain (the Troodos ophiolite, TO, in Fig.  2.3) in Cyprus extends through the Latacia subsea ridge to the Kızıldağ ophiolite in the Amanos Mountains (Fig.  2.3) [Yılmaz et  al., 1985; Yılmaz and Gürer, 1996; Hall et al., 2005; Bolton and Robertson, 2007; Tarı et al., 2014]. The Cyprus trench, located in the south of the island extends eastward to the southern boundary fault of the Latacia ridge [Woodside et al., 2002], and continues farther east, along a corridor as a fault zone separating the Musadağ and the Keldağ in the Hatay region (Figs.  2.10a and 2.12a). The Kyrenia range in northern Cyprus is cut and displaced from the Misis‐Andırın system of southern Turkey (Fig. 2.3) by a north‐northeast trending transpressional steep fault zone (Figs. 2.3 and 2.13c) [McKay and Robertson, 2013]. It forms the eastern subsea boundary of the Kyrenia‐Misis ridge, which in turn divides the eastern Mediterranean into two subbasins, the Iskenderun basin and the Adana‐Cilicia basin in the east and west, respectively (Fig. 2.3). Farther north, it extends to the Misis (Misis‐Andırın) fault zone (Figs. 2.3, 2.5, and 2.6). The land area has also suffered from a complex deformation [Yılmaz and Gürer, 1996; Gürsoy et al., 2003; Robertson et al., 2004] as displayed in Figure 2.5. Seismic activity along these fault zones is frequent and severe (Fig. 2.7b). According to Duman and Emre [2013], about one‐third of the relative motion between the Turkish and Arabian plates is accommodated on the Misis‐Kyrenia fault zone. Thus, the Misis and the Sarız faults farther north may be evaluated together with the DSF as the transform‐edge propagator accommodating the northerly advance of the Arabian Plate (Fig.  2.1). A  considerable amount of the motion is undoubtedly accommodated along the Escape Zone (located to the north of Maraş) (Figs. 2.6 and 2.7). The motion that is accommodated within, along and around the Amanos

24  ACTIVE GLOBAL SEISMOLOGY

Figure 2.12  (a) A northerly view of the 3D topography of the Asi River valley and the surroundings. Red lines are the faults. Asymmetry of the Asi graben and the adjacent horsts is evident. The northwest‐ and west‐facing slopes are steep. Southeast‐ and east‐facing slopes are smooth as displayed by the black arrows. The concave fault plane (pink areas) reveals listric nature of the normal (step) faults. The en echelon map pattern of the graben bounding faults of the Asi graben indicates that strike‐slip components of the faults also have played a significant role in the displacement, along which the Keldağ horst (KDH) has slid into the sea region and cut abruptly the southward continuation of the Samandağ fault. The listric character of normal faults is more pronounced in the fault zone that bounds the Kızıldağ horst (note the concave map pattern of the faults facing the Arsuz plain on the Gulf of ̇ Iskenderun). The Samandağ fault has reactivated later as indicated by the morphological data; note the elevated sea terraces (blue horizontal planes) in the Samandağ area (S), and the steep, subvertical cliff (yellow arrow) along the Cevlik (C) coastline. T1 and T2 represent hanging river terraces lying at different heights (about 80 m on the T1 and 200 m on the T2) due to the uplifting of the fault block by the Samandağ fault. Accordingly, incision along the Asi graben carved a canyonlike deep and narrow river valley near the sea coast (see Erol [1963] for a detailed account of the morphological development of the Asi River valley). Sea invasions into the Asi graben depression occurred during the Plio‐Quaternary period as revealed by marine sands interfingering with the fluvial conglomerate and sandstone. The Dead Sea fault is represented by a number of subparallel faults (continuous red lines). Each one of the branches may be identified clearly by its imprint in the morphology as exemplified from the linear ridges (pressure ridges), long and narrow valleys, and so on. Collectively, the faults form a more than 10 km wide zone, and an anastomosing map pattern. Three major faults–Dead Sea fault (DSFZ), Karasu graben fault (ABF), and Antakya fault (the Asi graben fault)–converge in an area where their interactions have generated development of the Amik depression (Amik D). It is a complex structure as suggested by the seismic data [TPAO, unpublished data]. The long and continuous yellow lines are the latitude and longitude. (b) Google image showing 3D topography of the Keldağ dome. The broken white lines are the contour lines displaying attitude of the bedding planes that define the dome. The brown lines are the listric normal faults. Note the concave fault planes; the fault along the coastline is steep (note the brown colored fault plane and subvertical cliff). The sea terraces (blue‐colored planes near sea level) and an erosional terrace (blue plane at about 500 m height) reveal that the uplift has occurred recently. The white line displays topographic profile across the dome. DZ = Denizgören Village; YV = Yeditepe Village. Yellow line is the road. (c) Block diagram showing geology and structure of the Keldağ dome. The double red line indicates the unconformity surface, above which Upper Miocene–Lower Pliocene sediments of low‐energy environment consisting of a limestone‐marl‐siltstone alternation overlie the thick pile of succession. Tectonic significance of this event is discussed in the text and in the caption of Figure 2.11c. The Asi graben depression represents an advance stage of the northwest‐southeast directed extension and the consequent localization of the areas of subsidence that occur during the Plio‐Quaternary period as revealed by the young fill of the graben; only the Plio‐Quaternary sediments are confined to the graben depression.

Morphotectonic Development of Anatolia and the Surrounding Regions  25

Figure  2.13  Schematic cross sections from the southern Amanos Mountains and the adjacent regions. See Figure 2.11a for the approximate directions of the cross sections. (a) The cross section in (a) covers the entire southern Amanos Mountains and the surrounding areas from the Mediterranean Sea in the north to the Baer Bassit in the south. Across the section, major horsts are observed, which display that under the ongoing northwest‐southeast directed regional extensional regime and the consequent thinning, a thick rock sequence was tectonically removed. As a result of this, progressively deeper parts of the sequence were exposed toward the south: Tr = Triassic; JR = Jurassic; and UCT = Upper Cretaceous. Approximate lateral distance along the section is about 80 km. (b) Northerly view from the Keldağ area. A set of normal listric faults (broken thin red lines and the blue arrows) is observed as westerly steps in the morphology. The brown arrows display back‐tilting of the downthrown blocks. Approximate lateral distance along the section is about 2 km. (c) An approximately west‐northwest–east‐southeast schematic cross section (section [a] in Fig. 2.11a) from the Kızıldağ Mountain to the Kyrenia subsea ridge in the Mediterranean Sea, based on the seismic sections (unpublished TPAO data). The section displays the extensionally induced subsidence. The listric normal faults that collectively form a rollover anticline (double red line) and tectonic uplift of the Kızıldağ horst, where the Kızıldağ ophiolite is located (green‐colored rock unit). The straight red arrow, which defines the eastern boundary of the Kyrenia ridge indicates transpressive character of the fault that bound the Kyrenia ridge. An approximate lateral distance along the cross section is about 40 km. (d) An east‐west schematic cross section from western edge of the Kızıldağ horst to the Mediterranean Sea across the Arsuz plain (section [b] in Fig. 2.11). The yellow rock unit is the Miocene sediments. Note the rollover anticlines in the sequence that formed due to rotational motions along the downward flattening fault planes during the development of listric normal faults. The gray rock units are Plio‐Quaternary clastic sediments. The listric character of the normal fault, and back‐tilting of the downthrown blocks, are clearly identified in the field; different elevated terraces and dip angles of the back‐tilted terraces were differentiated and mapped [Yılmaz, 1984]. The inset shows the section direction in the Arsuz plain. An approximate distance along the section is about 7 km.

Mountains including the Amik and Karasu depressions, is less compared to the other two zones. The Latacia‐Antakya fault zone extending as the southern boundary fault of the Latacia ridge stretches along the narrow valley of the Asi River to the Amik Lake depression (Figs. 2.3 and 2.11), where it intersects with the Dead Sea fault (Figs.  2.1, 2.3, and 2.10). The development of the Amik depression was regarded as a

triple junction [Perinçek and Çemen, 1990; Över et  al., 2004], as a superimposed basin [Boulton, 2013], or as a pull‐apart basin located at a leftward step in a left‐lateral transtensional fault zone [Yılmaz, 1984; Seyrek et al., 2014]. The structural pattern does not support a triple junction model. The structural depression does not seem to be continuously and constantly enlarging in the favor of any of the major fault zones. A number of seismic sections

26  ACTIVE GLOBAL SEISMOLOGY

along and across the Amik depression (mostly unpublished TPAO data) indicate collectively that the tectonics of the depression is much more complex than the published data display. Therefore, it is not easy to classify this structure that formed in a region where the three major fault zones emerge and interact. The structural depression along the Asi River valley is known as the Asi River graben or Antakya graben (Figs. 2.11a and 2.12a), which is an active fault zone. One of the pioneering and detailed morphological studies on the Asi River valley and also on the river and the sea terraces was carried out by Erol [1991]. The graben is asymmetrical. The eastern side is steep and represented by active oblique faults with dominant dip‐slip component, behaving more like normal faults as suggested by Bolton [2007]. Although Seyrek et al. [2014] estimate rates of the left‐lateral slip and extension across the depression is of ~4–5 mm a‐1 and ~3–4 mm a‐1, respectively, the normal fault character is morphologically and geologically more distinct and dominant (Figs. 2.11b, 2.12a and b,). However, the eastern (upthrown) block was apparently displaced sideways and has slid into the sea region (Figs. 2.11a and 2.12a and b). There is no morphologically distinct major fault along the western flank of the Asi River graben, where the slope is smooth, the slope angle is moderate and the drainage network is uninterrupted from the top of the mountain to the Asi River valley (the yellow arrows in Fig. 2.12a). Along the eastern slope, on the other hand, a number of steep steps formed as a result of the dip‐slip displacement of the parallel faults (Figs. 2.11b, 2.12a and b). The fault, which cuts and displaces the northwestern part of the Keldağ anticline (and horst), is known as the Denizgören fault, named after Denizgören Village (Fig.  2.12b). Across the fault, the horst rises steeply above 1500 m heights. Some other high‐angle normal‐step faults also cut the flanks of the anticline (Figs. 2.11a, b, c, and 2.12a, b, c). As a result of the uplifting by the dip‐slip dominant faults, a number of hanging river terraces along the Asi River valley and sea terraces along the Mediterranean coastal area [Pirazzoli et  al., 1991; Koral et  al., 2001; Blackwell et al., 2011] are observed on the flanks of the Musadağ and the Keldağ highs (Figs. 2.12a and b). Koral et  al. [2001] estimated about a 0.5–1.5 mm/y average uplift rate. The hanging river terraces display shallow valley profiles at the higher elevations (T1 in Fig. 2.12a). They become progressively narrower downward to form a canyonlike, narrow V‐shaped profile at the bottom of the valley (T2 in Fig. 2.12a). The high‐level hanging river terraces may be traced northeastward to the broad and shallow river valley of the Asi River around the city of Antakya, where the altitude is about 80 m. But the same terrace stands at an elevation above 200 m near the sea. This observation alone reveals that across the Samandağ

fault (SF in Figs. 2.11a and 2.12a; the fault that strikes perpendicular to the Antakya graben), the northeastern block was elevated and gently tilted to northeast. Based on the elevated terraces, Erol [1963] suggested an uplift rate of ~0.2 mm a‐1. The Musadağ (MDH in Fig.  2.12a) has also been uplifted as a horst block between the graben‐boundary faults of the Antakya graben and the Samandağ fault. The latter fault delimits abruptly the seaward extension of the Musadağ horst and terminates the Antakya graben (Figs.  2.11a and 2.12a). Above the upthrown block, terraces (some of which are marine terraces and are represented by beach sands [Blackwell et  al., 2011]) are observed at high elevations up to about 180 m [Tarı et al., 2014]. Morphological identities of the uplifted terraces may be traced to heights over 300 m. A number of remnant erosional terraces may also be identified on the northern slope of the Keldağ horst (also anticline) [Yılmaz, 1993] (Fig. 2.12b) at heights above 400 m. The Samandağ fault is abruptly terminated by the Denizgören fault (the Asi graben faults). Yet it is still an active fault as the distinct fault morphology indicates. The fault block rises steeply to above 500 m heights across the fault. The slope along the fault is commonly subvertical, and drainage between the sea and the top of the plateau is yet in an incipient stage. The Samandağ fault may be identified as a member of the group of spacely developed faults that cut the entire width of the Amanos Mountains (Fig. 2.10; Yılmaz [1984]). It was later cut and displaced by the approximately north‐ south striking Asi graben fault. Undoubtedly, the Samandağ fault was reactivated later, since it is presently a very distinct structural feature in the morphology of the region. Based on ESR data, the rise of the terraces was recently documented in detail [Blackwell et al., 2012; Tarı et al., 2014], which reveals that the elevation of the sea terraces occurred within the last 400 ka. In fact, the Antakya depression itself is also a young morphological feature. It is filled with marine sediments of Pliocene age (Fig. 2.11). Quaternary fluvial conglomerates of the Asi River overlie them. Morphological and geological features of the Antakya graben are described and discussed in detail by Erol [1963], Boulton and Roberts [2008], Tüysüz et  al. [2013], and Tarı et al. [2014]. The Karasu graben (Figs. 2.5, 2.6, 2.8, 2.9a, and 2.10a) is also an asymmetrical depression. The western boundary is steep against the Amanos Mountains. Despite a number of en echelon faults lying obliquely to the axis of the depression, there is not one single major fault along the eastern edge of the depression (Fig. 2.8). A broadly north‐south striking fault zone consisting of a number of subparallel faults represents the western boundary of the Karasu graben. Although extending along the Dead Sea transform fault, the present eastern boundary of the Amanos horst is evaluated in some

Morphotectonic Development of Anatolia and the Surrounding Regions  27

­ revious studies [e.g., Arpat and Şaroğlu, 1975; Boulton p and Robertson, 2008] as a set of normal faults, which has elevated the Amanos Mountains and formed the Karasu graben, and the major strike‐slip fault runs along the graben axis. However, there is a significant strike‐slip ­displacement along the western boundary fault zone of the Karasu graben, which is detected clearly in the morphology (Fig. 2.10c). The offset streams are common, a bunch of linear ridges trending parallel to the Amanos Mountains are frequent, and they display an anastomosing pattern typical for a strike‐slip fault zone (Fig. 2.10c). However, presence of the lateral displacement does not rule out the equally important role of the transtensional dip‐slip displacement as previously suggested [Bolton and Robertson, 2008]. Total lateral displacement across the fault zone is difficult to establish due to the absence of reference points on the eastern side of the Karasu graben. To the south, the transtensional regime is gradually replaced by increasing amount of the northwest‐southeast extension (Fig. 2.7a). This is observed particularly in the southern part of the Amanos Mountains. Going away from the orogenic zone to the southwest, a significant extensional component is determined clearly in the region, located in and around the Iskenderun Bay [Aksu et  al., 2005; Hall et  al., 2005] and the adjacent areas in the Amanos Mountains. The cross sections across the southern Amanos Range display a number of fault zones with a major northwest dipping dip‐slip component (Fig. 2.13). Farther west the Iskenderun basin was formed as an asymmetrical graben (Fig. 2.13c) under the northwest‐ southeast or east‐west extensional regime. Similarly, the Kızıldağ and Keldağ horsts and the Iskenderun and the Asi River grabens are also asymmetrical (Fig. 2.13). The western sides of the horsts are steep (Figs.  2.11b, 2.12, and 2.13) and are defined by a number of listric normal faults (see the concave map pattern of the major fault zones in Figs. 2.11b, 2.12a, and 2.13), where the downthrown blocks are back tilted (Figs.  2.11b, 2.12a, and 2.13a, b, c, d). The field and seismic data (unpublished numerous seismic sections of TPAO along and across the major depressions and the surrounding regions) reveal that the fault planes are flattened downward (Figs. 2.13a, d). Consequently, to the southeast and east, the deeper and older units in the successions have been progressively elevated in the upthrown blocks, for example, the Triassic rocks in the Baer‐Bassit, the Jurassic rocks in the Keldag,̆ and the Upper Cretaceous rocks in the Kızıldağ (Fig. 2.13a). The data collectively suggest that the listric normal faults that flattened downward possibly converge toward a major fault zone, which is accommodating the bulk of the extension. The terraces developed on the downthrown blocks were also back tilted. The tilt angles and the thickness of the sediments deposited on

the downthrown blocks get younger toward the horst axes (Fig. 2.13d; Yılmaz [1984]). A major phase of the extension in the southern Amanos and surrounding region began during the Late Miocene and continued uninterruptedly to the present. Data to support this view may be given from the Keldağ region, where a thick sequence is observed on the Keldağ horst and the surrounding region. This succession that is exposed, covers an age span from Jurassic to Middle Miocene (Figs.  2.11c and 2.12c). A more than 3.5 km thick sequence was totally removed mainly by tectonic erosion during the Late Miocene period. As a result of this, the Upper Miocene–Lower Pliocene sediments of low‐energy marine environment, represented by siltstone, marl, and fine‐grained sandstone alternation, deposited directly on the Mesozoic carbonate sequence above a regional unconformity (Figs.  2.11a, c and 2.12c). The Upper Miocene–Lower Pliocene sediments form the base of the younger depressions. This vast area of shallow subsidence was localized later to form the present narrow graben depressions (e.g., the Asi graben), during the Quaternary. Consequently, a sea incursion occurred into this deep and narrow graben depression. The elevation of the adjacent horst blocks is still continuing today as revealed by the following morphological observations: (1) the hanging valleys are common between the fault bounded blocks, (2) the stream valleys on the western, fault‐bound side of the horsts display well‐developed champagne glass morphology (particularly those on the western slope of the Kızıldağ horst facing ̇ the Iskenderun basin), (3) headword erosion along the streams is yet in an incipient stage, (4) the deep carving has not yet reached to the top of the horst block. Supporting the morphological observation Seyrek et al. [2008] based on the age data from basalt lavas from the Ceyhan Gorge demonstrated that the elevation rate is between 0.25 and 0.4 mm a − 1. They state also that the entire ~2300 m height of present‐day relief has developed since the Mid‐Pliocene. The uplifting of the Amanos range cannot be attributed solely to the isostatic response generated as a consequence of the erosion as suggested by Seyrek et al. [2008]. On this, the northwest extension (of about 4 mm a − 1) across the Haruniye (Düziçi; Fig.  2.8) and Asi graben faults as suggested by Seyrek et al. [2014] plays a major role. The East Anatolian transform fault extends to the northern part of the Amanos Mountains around the Maraş depression (Figs. 2.8). Although westward continuation of the branches of EATF to the Amanos Mountains is hidden under the present basin fill, they extend along the two strike‐slip fault zones in the Amanos Mountains (Figs.  2.8 and 2.9a). They are the Deliçay fault zone and the Türkoğlu‐Haruniye fault zone, both of which are morphologically distinct major faults (Figs. 2.8

28  ACTIVE GLOBAL SEISMOLOGY

and 2.9). They strike obliquely and cut the whole width of the Amanos Mountain range (the faults a and b in Figs. 2.5, 2.8, and 2.9). The faults, although young and morphologically prominent, may be comparatively older than the north‐south striking faults that bound the surrounding Karasu graben and the Düziçi plains, and delimit the mountains along both edges (Figs.  2.5, 2.8, and 2.9a). The Deliçay and the Türkoğlu‐Haruniye faults may be regarded as the southwesterly extension of EATF, and their continuation was disrupted during the development of the Karasu depression (Fig. 2.8). Despite this, in the areas where the two fault zones are intersected (e.g., the depression between the villages of Türkoğlu and Kılılı, a village located about 2 km to the north of Türkoğlu Town; Fig.  2.8), the present morphology has been developed under the influence and interaction of both fault zones. The east‐west trending fault plane, which delimits the small pocket of depression, where Türkoğlu Town is placed, displaces the north‐south Karasu depression. As a result of the displacement, the Türkoğlu‐Haruniye fault extends a few tens of meters eastward into the depression (there the EATF and Türkoğlu‐Haruniye faults are very close to one another; see Figs.  2.8 and 2.9a). This may be evaluated that the Türkoğlu‐Haruniye fault has been reactivated after the development of the Karasu depression, and therefore may be considered also as an active fault zone in the sense defined by the active fault map of Turkey [Ş aroğlu et al., 1992; Emre et al., 2012‐2013]. Similarly, the Deliçay fault may also be traced within the Maraş depression. There the fault may be traced to a point that is very close to one of the major branches of EAFT (Figs. 2.8 and 2.9a). The Deliçay and Türkoğlu‐Haruniye fault zones consist of left‐lateral strike‐slip fault branches (Fig. 2.9b and c). The fault planes are commonly steep. The fault zone runs along the northeast trending two main valleys. To the west, they stretch to the Sabunsuyu valley located to the north of Haruniye (Düziçi). Farther west in the Haruniye plain beyond the Amanos Mountains, they are cut by a steeply dipping, north‐south‐striking normal fault, which separates the northern Amanos Mountains from the flatland. Extending to the north, the normal fault may be observed in the Ceyhan River gorge (Figs. 2.5, 2.8, and 2.9a). The strike‐slip fault zones, on the other hand, are buried under the alluvial deposits, and their possible continuation farther west is difficult to trace in the field. Within the Amanos Mountains, branches of the strike‐ slip fault zone form a number of step‐overs and an anastomosing map pattern (Fig. 2.9b and c). In the fault zone, the lenticular, fault‐bounded blocks are common, in which different members of an ophiolitic suite and different rock units of the Mesozoic limestone sequence are brought together across steeply dipping faults, which collectively

form the present ragged fault morphology. The faulted blocks represent the structures that formed under a sinistral shear [see Yılmaz, 1984, vol. 3, for the maps and detailed descriptions of the fault zones). The Maraş (Kahraman‐Maraş) depression is an asymmetrical structural basin located among the branches of a number of faults (Figs. 2.5, 2.6, 2.8, and 2.9a). There is not a major fault along its southern edge against the central part of the Amanos Mountains. The northern boundary of the Maraş depression, on the other hand, is a fault zone located in the southern front of the Ahır Dağı Mountains (Figs. 2.5, 2.6, 2.8, and 2.9a), which rises steeply from the Maraş plain of about 600 m high to heights of over 2300 m. The Ahır Dağı is a southerly inclined and partly overturned asymmetrical, westerly plunging anticline (and a horst). The revers slip component of the oblique faults, which presently define the southern boundary of the Ahır Dağı anticline, appears to be solely responsible from the initial stage of the development of the Maraş basin. The Ahır Dağı was thrust to the south in a late stage during the formation of the fold and thrust belt. The region lying in front of the thrust was depressed under the heavy load of the overlying block to form the initial depression of the Maraş basin as a foredeep. This was developed possibly during the Middle to Late Miocene period. The southerly overturned flank of the anticline was later sliced. Consequently, fault‐bounded, semi‐independent blocks were formed. They began to move westerly along the branches of an oblique and narrow (2500 m) fault zone during progression of the deformation (Figs.  2.8 and 2.9a). The fault zone consists of left‐stepping en echelon faults (Figs. 2.8 and 2.9a). Within the zone, the faults are steep, dipping 50–70° to the north. When EATF and the Dead Sea fault intersected, which occurred possibly during the Late Pliocene–Pleistocene time, the Maraş depression (Figs. 2.5–2.8 and 2.9a) began to evolve as a triple junction [Ş engör, 1979b]. In summary, the Maraş basin has passed through multiple deformational phases. The initial phase started during the formation of the southeastern Anatolian orogen, and its development is still continuing today under the late‐post orogenic deformation. Starting from the establishment of the continental envi­ ronment in the Late Miocene, a newly developed river network was formed on the Arabian continent (Figs. 2.3 and 2.14) and the river beds were later elevated and/or displaced during the progression of the deformation [Westaway et al., 2006; Demir et al., 2008, 2012]. In the areas close to the mountain range, major structural features of the fold and thrust belt have controlled the topography (Figs.  2.14). The lowlands intervened the east‐west trending hills. They correspond commonly to the anticlines and synclines, respectively (Figs. 2.4 and 2.5).

Morphotectonic Development of Anatolia and the Surrounding Regions  29

thickness of up to 250 m near the center. Activities of the volcanic center began about 10–11 m years ago and continued interruptedly to the Quaternary [Ş aroğlu and Emre, 1987; Keskin et al., 2012]. 2.3. THE EAST ANATOLIAN HIGH PLATEAU

Figure  2.14  Morphology map of eastern Anatolia. Red lines with yellow glows are the crest lines of the peripheral mountains (Pontides and Bitlis‐Zagros suture mountains) corresponding also to the main water divides. The red line represents the secondary water divides in the central part of the region. The white arrows display a major drainage network fanning away from a central high, the yellow circle with red glow.

To the north, their wavelengths diminish and the amplitudes increase (Fig.  2.14). They control the present drainage pattern (Figs. 2.5 and 2.14). Streams flow in the synclines. The valleys trend subparallel to the axes of the folds. Even the valleys of the Euphrates and Tigris, the two major rivers of the region, display a partly similar pattern running along a path, trending perpendicular to the regional dip. Thus, the drainage system displays a trellis pattern (Fig. 2.14). The drainage pattern of the streams is sinuous (Figs. 2.3 and 2.14). They turn around the dipping related periclinal terminations of the anticline axes. Thus, the newly formed synclines and anticlines are the depocenters and erosional lands, respectively. Therefore, sediment types and their facial distributions help to decipher how the topography evolved on the Arabian continent from the late Miocene to the present. In harmony with the structural displacements, due to combinations of the folding and thrusting accompanied by the erosion, the vertical stratigraphic offset varies greatly in the southeastern Anatolian plain from 1 km up to 6 km. However, vertical topographic offset is commonly less than 500 m. The only exception to this general morphological pattern is the Karacadağ basalt plateau (Fig. 2.4), where the drainage network is yet in an incipient stage. It is a shield volcano covering a large region of more than 10,000 km2. The basalt lavas poured from the north‐south extensional opening(s) (impactogene of Ş engör [1979a]; Yılmaz [1981b]). The lava pile reaches a

Eastern Anatolia is a high plateau with an average elevation of 2000 m. It is a relatively flat land, above which rise a number of conical individual peaks and northeast, east‐northease, and west‐northwest–northwest trending parallel hills and depressions (Fig. 2.14). The individual high peaks correspond to volcanic cones, which are observed in a vast region covering nearly the entire eastern Anatolian plateau (Figs.  2.14; Yılmaz et  al. [1987]; Yılmaz [1990]). The depressions and the intervening hills correspond commonly to the anticlines or horsts and synclines or grabens, respectively [Ş arogl̆ u et al., 1980; Ş aroğlu and Güner, 1981; Ş aroğlu, 1985; Ş aroğlu and Yılmaz, 1984, 1986, 1987], suggesting they are young structures and initial morphology has not yet been reversed. The time elapsed since their development is short [Ş aroğlu and Yılmaz, 1987]. This view is supported by field data obtained from regional geological mapping in the entire eastern Anatolia area (the unpublished maps of Turkish Petroleum Association), which show convincingly that only the post‐Pliocene sediments are confined to the present synclinal depressions (Figs.  2.15 and 2.16; Ş aroğlu and Yılmaz [1987]). The Miocene and Pliocene sediments on the other hand are regionally distributed and retain their major characteristics in the entire region irrespective of the present highs and lows. Geological mapping, supported by seismic data, reveals also that East Anatolia is underlain by an accretionary complex (Figs.  2.15 and 2.16) as initially displayed by Ş engör and Yılmaz [1981]. In the areas where the cover rocks were removed, always an ophiolitic mélange assemblage crops out at the base [Ş engör et  al., 2008]. The mélange was piled up during the demise of the NeoTethyan Ocean, which was located between the Pontide arc in the north and the continental slivers drifted away from the Arabian Plate in the south [Ş engör and Yılmaz, 1981; Yılmaz, 1993]. Between the bordering continents, the accretionary prism behaved like a wide and thick cushion and did not allow the bordering plates to get in direct touch with each other (the Turkic type orogen of Ş engör et al. [2008]). Above the ophiolitic basement, a sedimentary sequence was discontinuously deposited from Late Cretaceous until the Middle Miocene mostly in a marine environment. From the Late Miocene onward, a continental environment was established, and has remained until the present (Figs. 2.15 and 2.16). During the Oligocene most of East Anatolia was above sea level for a brief period (see the continental red beds of

Figure 2.15  Geological schematic cross sections illustrating consecutive stages (phases I to IV) of tectonic development of the eastern Anatolia. Phase I illustrates the rock sequence and tectonic position of the region during the Late Miocene–Early Pliocene. The region is underlain by an accretionary mélange prism (A), overlain by a locally developed succession that consists mainly of coarse clastic deposits of Oligocene age (B), Lower Miocene shallow marine limestone (C) (known as the Adilcevaz limestone), Upper Miocene lacustrine white limestone (D), and continental deposits of Pliocene age (E). The region has suffered increasing shortening deformations as expressed in the diagrams as open and tight folding (phase II), folding coupled with thrusting (commonly reverse faults; phase III), and folding and thrusting from the opposite sides, which forms ramp basins (phase IV). Starting Pleistocene onward, locally developed commonly fault‐bounded small depressions were formed. They were filled mainly with fluvial sediments (gray units in phases III and IV). As a result of interactions of the structures that are outlined above, a variety of structural forms were formed.

Figure  2.16  Geological north‐south cross section (modified from the section provided by Ö. Şahintürk; TPAO map) across the Erzurum‐Aşkale basin (see Fig. 2.22b for the location). Note that an ophiolitic mélange association of Cretaceous age forms the basement.

Morphotectonic Development of Anatolia and the Surrounding Regions  31

Figure 2.17  A photo showing gradual transition from the shallow marine limestone of the Early‐Middle Miocene age to the lacustrine limestone of Late Miocene age. The evaporate bed, in the middle, is a transitional unit. The photo is from the north of the Aş kale Town (see Fig. 2.22b for the location).

the Oligocene age (b in Figs.  2.15 and 2.16). A rather irregular topography was developed during this period as indicated by the regional unconformity surface below the Lower Miocene marine sediments. Locally developed coarse‐grained conglomerates of Oligocene age were deposited in narrow depressions. An exception to this general patterns is the northern region of East Anatolia, where a sea connection with the Back Sea [personal communication with F. Ş aroğlu and Ö. Ş ahintürk] and the Caspian Sea still remained during the Oligocene–Early Miocene period [Yılmaz, 1997b]. The evidence for this is (1) there is a gradual transition from the continental sediments of East Anatolia to the shallow marine sediments of the northern and northeastern areas of the East Anatolian region and the Caspian region [Harris, 1995]; (2) these regions share the same facies of the Oligo‐ Miocene sediments internationally known as the Maykop formation [TPAO reports and Harris, 1995]. During the Early Miocene, a new transgression invaded East Anatolia, and the previous irregular topography was smoothed as indicated by the fining upward profile in the succession. At the top of the Lower Miocene sequence, there is a shallow marine limestone unit known as the Adilcevaz Limestone [Ş engör and Yılmaz, 1981] (c in Figs. 2.15 and 2.16). It covered nearly the entire region, and its outcrops are particularly abundant in the southern and central parts of East Anatolia. From the Middle Miocene to the Late Miocene, a gradual change is observed. The shallow sea sediments are replaced nonconformably by lacustrine white limestone (d in Figs.  2.15, 2.16, and 2.17). The transition that is observed in the entire region suggests that interconnected lakes replaced the sea and covered a vast region in East Anatolia (see geology maps of the Mineral Research and Exploration General Directorate of Turkey [MTA] on the scale of 1/500,000).

This event may also be interpreted that the whole eastern Anatolia possibly together with central Anatolia (see the discussion in Chapter 4) began to rise as a coherent block (en mass) (phase I in Fig. 2.15) [Ş engör, 1979; Ş engör and Kidd, 1979; Dewey et al., 1986; Ş aroğlu and Yılmaz, 1987]. This indicates further that the tectonically controlled present local continental depressions and the adjacent peripheral highs had not yet formed during that period. The continuous sediment deposition, and the regionwide continuation of the units, changed drastically after the Late Miocene [Ş aroğlu et  al., 1980; Ş aroğlu, 1985; Ş aroğlu and Yılmaz, 1987]. Different sediment packages were formed within separate and independent depressions (Phase II in Figs. 2.15 and 2.16). Immediately after the disappearance of the inter­ connected lakes, the region suffered a severe phase of denudation. It formed a regionwide flat‐lying erosional surface above the lacustrine limestone and interbedded volcanic rocks of the Upper Miocene age (Fig.  2.18). Remnants of this surface are presently recognized at different heights, and may be used as a key horizon. Any major morphological feature that has developed above or below this marker postdates it, and thus is younger than the Late Miocene. Eastern Anatolia displays most of the geological ­features of a collisional orogenic belt. The continental crust is thick (>38 km) reaching up to 50 km in thickness. As a result of a continuous north‐south compression that followed the collision, a complex pattern of structures was developed (Fig. 2.19), as discussed and demonstrated by Şengör [1979a], Şengör and Kidd [1979], and Şarogl̆ u and Yılmaz [1986,1987]. Types and trends of the structures may be summarized as follows: (1) a conjugated pair of strike‐ slip faults striking northeast and northwest (the arrows a and b in Fig.  2.19); (2) approximately east‐west striking

32  ACTIVE GLOBAL SEISMOLOGY

Figure 2.18  Photo showing erosional surface (broken lines) developed above Plio-Quaternary basaltlavas of the Karst Plateau. Deepand narrow Arpacay river valleyformed carving the basalt and alternating pyroclastic succession. The yellowlines represent the elevated terraces, formed in association with the rise of the eastern Anatolian Plateau.

revers faults and folds (white lines with black glow in Fig.  2.19); (3) north‐south trending extensional openings (d in Fig. 2.19). Some of the volcanoes are located above the extensional openings (see the location of the Nemrut Volcano, NV, in Fig. 2.19) [Sa̧ rogl̆ u and Yılmaz, 1984, 1986, 1987; Yılmaz, 1990; Yılmaz et al., 1987, 1998]. The cover rocks of eastern Anatolia are commonly folded. The folds may locally be tight and even overturned (Fig. 2.20). The major faults strike northeast (e.g., NEAFZ; the northeast Anatolian fault zone, EAF; Erzincan‐Aşkale fault zone, OF; Oltu fault zone in Fig. 2.19) and northwest (e.g., TF, the Tutak fault zone, and DF, Doğu Bayazıt fault zone in Fig. 2.19). The faults of different strikes together form a conjugated pair in a regional scale (see Fig.  2.19). The northwest striking faults are more prominent and longer. The exceptions to this are the Oltu and the Kağızman‐Digor fault zones, which run parallel to the northern edge of the plateau. They are members of the northeasterly striking large fault system (the Erzurum‐Tbilisi fault zone of Westaway [1994]) that stretches to the Lesser Caucasus, where the Kasbeg fault is also a distinct member of this fault belt.

This fault zone was formed under the ongoing stress field forcing the plateau to move in northeastern direction as a consequence of the slow escape (see GPS vectors in Fig. 2.21). Most of the faults of East Anatolia are active, and responsible from the frequent earthquakes as exemplified from the recent (23  October 2011 and 24 June 2015) Van earthquakes [Elliot et al., 2013]. A majority of the present continental basins of eastern Anatolia may be viewed as ramp basins as exemplified from the Muş, Murat, and Bingöl basins in the south and the Pasinler, Tercan‐Aşkale basins in the north (Fig. 2.22). The eastern Anatolian plateau itself may also be regarded as a giant ramp basin between the peripheral mountain ranges, because revers fault zones bound the East Anatolian high plateau along both edges (Fig.  2.22a). Sediment fills of the fault‐bounded linear basins are younger than Pliocene (Fig. 2.15), and consist essentially of fluvial conglomerate and sandstone. Locally they are interfingering with lake deposits [Şaroğlu, 1985; Ş aroğlu and Yılmaz, 1987]. These fault‐controlled young depressions are aligned in two separate strips (Figs. 2.14, 2.22a and b). Both of the linear depressions are divided into

Morphotectonic Development of Anatolia and the Surrounding Regions  33

Figure 2.19  Morphotectonic map of eastern Turkey displaying major structural elements and the consequent morpholical features. Data derived from active tectonic map of MTA [2012], Şaroğlu [1985], Şaroğlu et al. [1980, 1992], Şaroğlu and Yılmaz [1984, 1986, 1987, 1991], Şengör and Kidd [1979],Yılmaz et al. [1998], Bozkurt [2001], Şengör et al. [2008], Özeren and Holt [2010]. Cities (the blue letters): El = Elazığ; M = Malatya; Erz = Erzincan; K = Kars; Ar = Ardahan; Ag = Ağrı; V = Van; LV = Lake Van. Faults (red lines; white letters with red glow) BZSMFTZ = Frontal thrust zone of the Bitlis‐Zagros suture mountains; FFTB; The foreland Fold and Thrust Belt; NATF = North Anatolian transform fault; EATF = East Anatolian transform fault; NEAFZ = Northeast Anatolian fault zone; EAF = Erzincan‐Aşkale fault zone; OF = Olur fault zone; KDF = Karayazı‐Digor fault zone; DF = Doğu Beyazıt fault zone; TF = Tutak fault zone. Volcanoes (green triangles; white letters with green glow) NV = Nemrut; SV = Süphan; EV = Etrüsk; MV = Meydan; AğV = Ağrı (Ararat); TV = Tendürek; SlV = Solhan; KBP = Karacadağ (basalt plateau); AlV  =  Alagöz (Aragats). ERZD = Erzincan depression; red lines = strike‐slip faults. Red arrows indicate two different trends of the strike‐slip fault zones: (1) the arrow with letter A indicates the region of northeast‐trending, left‐lateral, strike‐slip faults; (2) the arrow with letter B indicates the northwest‐trending right‐lateral, strike‐slip faults. The white arrow with red glow pointing to the black letter C refers to the area in the immediate east of Karlıova junction is the region of orthogonal shortening, where mainly east‐west striking reverse faults and folds dominate. White curvilinear lines = reverse faults or oblique faults with reverse‐slip components. The black letter D and the accompanying black arrow = area of north‐ south tensional openning under the north‐south compressional forces, where the Nemrut volcano is located.

smaller, local basins, bounded by complex oblique faults (displaying reverse‐slip and strike‐slip components). Therefore, the basins are commonly delimited by straight boundaries, which give the basins distinct geometrical patterns (Fig. 2.22b). The basement rocks of the peripheral mountains were thrust above the sedimentary and volcanic rocks of the depressions (Figs.  2.22 and 2.23). As a result of this, the crust in these areas is thicker than the surrounding regions reaching up to 50 km (48 km beneath the Bitlis‐Pötürge massifs in the south and 50 km underneath the Pontide basement in the north (Fig. 2.22). Average thickness of the crust along these zones is about 45 km [Barazangi et al., 2006; Ş engör et al., 2003, 2008].

The peripheral mountains, underlain by rigid old metamorphic basement rocks, were elevated at a higher rate than the relatively soft mélange material that forms the basement of the eastern Anatolian plateau. This is possibly due to the ophiolitic mélange, which partly absorbs the north‐south compression. In the central part of the eastern Anatolian plateau, the crust is thicker (see the circle defined by the broken lines in Fig. 2.14) with respect to the surroundings, because this region corresponds to the area of the intense orthogonal shortening (Figs. 2.14 and the area c pointed by two thick arrows in Fig. 2.19). As a consequence of the development of the giant ramp, the main flow direction of the rivers on the plateau is

34  ACTIVE GLOBAL SEISMOLOGY

Figure 2.20  Overturned folds observed in continental red beds of eastern Anatolia. The S vergent thrust cut the southern limp. Photos are from the Tuzluca area, southeast of Kars (see Fig. 2.22 for location). This phase of deformation corresponds to the structures that are developed in phase III of Figure 2.15.

approximately east‐west, being subparallel to the peripheral mountains (Fig. 2.14). It also implies that the rate of elevation of the peripheral mountains is greater than the rate of incision across the mountains; incision along the valleys of the major rivers does not keep pace up with the rate of elevation. Only one major river at the both sides of the plateau (Çoruh in the north and Fırat‐Euphrates in the south) has cut across the whole width of the mountains (Fig. 2.14). The north‐south shortening deformation is still continuing severely at the present time. Therefore, the whole

sequence is folded and faulted (Figs. 2.15, 2.19, and 2.20). The two transform faults, namely the North Anatolian transform fault (EATF) and the East Anatolian transform fault (EATF), were formed to accommodate the north‐south compressional deformation possibly during late Pliocene-Pleistocene when the crustal shortening had reached an  excessive stage. This may be attributed either to stress permutations between σ2 and σ3 or to triaxial strain conditions [Ş engör, 1979a, c; Ş engör and Kidd, 1979].

Morphotectonic Development of Anatolia and the Surrounding Regions  35

Figure 2.21  The GPS vectors from the east of Anatolia [modified after Şengör et al., 2008]: The purple heavy line along the Pontides is the major crest line. The blue arrows define the main plate motions. The black arrows represent major stress direction affecting the Pontides, which shows progressive westward decrease. The ellipse defines the orthogonally compressed region from which the stress vectors diverge.

Evidence for the time of the development of the Anatolian Plate and the intersection of the transform faults that bound the plate may be derived from the Karlıova Junction basin. It is a wedge‐shaped depression (Figs.  2.19 and brown triangular area in 2.22b) evolving between the two transforms faults. The oldest sediments confined to the basin start from the late Pliocene(?)-Pleistocene [Şarogl̆ u and Yılmaz, 1991]. They were derived from the fault‐ bounded elevations, and accumulated as lateral fan deposits. The region located to the immediate east of the Karlıova Junction, the north‐south compression, has deformed the rocks orthogonally (Figs.  2.14, 2.19, 2.21, and 2.22b) [Özeren and Holt, 2010, Karaogl̆ u et al., 2016]. Data for this may be listed as follows: (1) the revers faults and folds strike east‐west (see the area c in Fig. 2.19). An east‐west striking major revers fault that is located in the immediate east of the Karlıova Junction may be given as further proof. The northern block was thrust above the southern block. As a consequence, only the northern half of the caldera of the Solhan volcano (SlV in Fig. 2.19) is observed in the field. The southern half is apparently buried under the northern half. (2) A centrally located domelike morphological high was formed (Figs. 2.14, 2.19, and the central high in Fig. 2.22a), which may be resembled to a center of virgation (Figs.  2.14, 2.21, 2.22b). This high controls the

regional drainage network. The rivers fan away from this high dispersing to the east and the west (see the arrows in Fig. 2.14). The morphotectonic map pattern of the central part of eastern Anatolia may thus be compared to sheaves of wheat tied at the center. The peripheral mountains on both sides curve around this region forming a normal and a reverse V in the north and the south, respectively (Figs. 2.1, 2.14, and 2.19). This area may be viewed as the location of maximum indentation. Away from the high, toward the east and the west the faults are commonly oblique having a strike‐slip component (Fig. 2.19; Şarogl̆ u and Yılmaz [1987]). The morphological pattern of eastern Anatolia is in close harmony with the deformation pattern and the crustal motions as expressed by the GPS vectors (Fig. 2.21). A widespread volcanic activity began above the eastern Anatolian plateau starting from the Late Miocene onward, initiated possibly about 11 m yr ago in the northern areas, and migrated to the central areas about 5–6 m years ago. The volcanic activity was intensified in the central and southern areas about 3 m years ago [Yılmaz, 1990; Yılmaz et al., 1987, 1998; Pearce et al., 1990; Keskin, 2007; Keskin et al., 2012]. As a result, a thick volcanic blanket covered the eastern Anatolian plateau. More than a 2 km thick lava pile was measured and drilled in the Kars plateau in the northern part of the region (unpublished drilling data

36  ACTIVE GLOBAL SEISMOLOGY

Figure 2.22  (a) Schematic block diagram across the east of Turkey from the Black Sea to the Arabian platform. The elevated mountain ranges that border the east of Turkey are the two orogenic mountain ranges that are underlain by an old metamorphic basement. In contrast, an ophiolitic mélange association underlies the central high plateau, where the mélange material acted as a cushion and did not allow a direct contact of the bordering continents. The two red arrows represent the ongoing convergence following the collision. The thin black arrows indicate north and south vergent thrusting of the peripheral mountains as a consequence of the severe shortening deformation. The dark green belt under the Black Sea represents old, rigid remnant oceanic lithosphere, which resists the northward advance of Anatolia and thus generates the north‐south severe shortening deformation. The V‐ shaped areas of subsidence aligning along the mountain ranges represent the two peripheral troughs that are bordered by oblique faults, which commonly display a reverse‐slip component. A central high is also formed between the reverse faults. The northward subducting slab beneath the Arabian platform represents the demise of the NeoTethyan Ocean that once separated the Arabian Plate from Anatolia. The seismic images obtained by various analytical methods demonstrate that a remnant subducted oceanic lithosphere and its broken pieces are present underneath eastern Anatolia [Özaçar et al., 2008; Gans et al., 2009; Warren et al., 2013]. Hot asthenosphere is placed at about a 65 km depth. Therefore, the low‐velocity zone values are low, and extensive volcanic activity occurred starting from the Late Miocene to the present. The fast polarization direction trends appear to be in accordance with the direction of the plate motion [Sandvol et al., 2003]. (b) Morphotectonic map of eastern Anatolia. The thick red lines with yellow hue represent the peripheral mountains. The red triangles and rectangles represent small, fault‐bounded continental basins that are attached to one another along oblique slip faults. The small red arrow in the center represents a central high and also a secondary water divide in the central part of eastern Anatolia. The blue lines are the water divides of the Kopdağ and Kızıldağ ranges, which have been offset for more than 50 km by the NATF zone across the Erzincan depression. The brown triangular area at the intersection of the two transform faults is the Karlıova junction‐basin. BZSM = Bitlis‐Zagros suture mountains; FFTB = Foreland fold and thrust belt. The broken black lines correlate the morphotectonic features between (a) the cross section and (b) the map.

Morphotectonic Development of Anatolia and the Surrounding Regions  37

Figure 2.23  Photo showing northerly view of a U‐shaped, slightly overturned syncline in the Oltu depression north of Erzurum. The red arrows define the reverse faults that placed the basement above the Neogene continental red beds. The yellow lines indicate the ramp basin character of the structural depression. The photo reveals the initial morphology, the syncline, and the bordering anticlines (the valley and the hills) have survived to the present.

of TPAO). A number of well‐known volcanic centers, for example Nemrut, Süphan, Tendürek, and Ağri (Ararat), were formed during the Quaternary (see Fig. 2.19 for the location of the volcanoes; Yılmaz et al. [1998]). The volcanoes are preferentially aligned along either north‐south trending extensional opening (e.g., Nemrut volcano) as response to the north‐south compression or use the extensional openings associated with step‐overs formed along the two major strike‐slip fault zones (e.g., the Agr̆ ı volcano nested in a northeast trending sinistral pull‐apart basin [Yılmaz et al., 1998] or at the intersection of the two contrasting sets of fault zones as exemplified by the Süphan volcano [Fig. 2.19; Yılmaz et al., 1998]). According to a variety of geophysical data, there is either no mantle lithosphere or a very thin lithosphere that exists beneath East Anatolia [Zor et al., 2003; Türkelli et al., 2003; Sandvol et al., 2003; Angus et al., 2006; Özaçar et al., 2008, 2010; Bartol et al., 2012]. According to Barazangi et al. [2006], the crust floats on a partly molten asthenosphere. The uppermost mantle beneath the crust strongly attenuates Sn waves, and has very low Pn velocities (−7.6 km/Sn) [Barazangi et al., 2006; Gögü̆ ş and Pysklywec, 2008; Özaçar et  al., 2008, 2010]. This is thought to be related to slab steepening and break off of the northerly subducted slab, beneath the accretionary prism [Keskin, 2003; Facenna et  al., 2006; Lei and Zhao, 2007; Şengör et  al., 2008] as initially demonstrated by Piromallo and Morelli [2003] and Piromallo and Regard [2006]. It has been suggested also that

topographic uplift and the accompanying widespread volcanism may be closely related to slab break off and lithospheric delamination [Keskin, 2003; Keskin et  al., 2012; Gögü̆ ş and Pysklywec, 2008] as defined initially by Bird [1979]; absence of the mantle lithosphere allowed direct contact of the crust with hot asthenosphere, which caused melting and the consequent volcanism. The high topography is not isostatically supported by the thick crustal root. The crust beneath East Anatolia is not extremely thick [Gök et al., 2003, 2007; Tezel et al., 2013; Warren et al., 2013] as previously claimed [Ş engör and Kidd, 1979]. 2.4. THE CENTRAL ANATOLIAN PLATEAU Tectonically, central Anatolia is a zone of transition among the surrounding regions (see Figs. 2.1, 2.14, 2.24, and 2.25a). Morphotectonically, it is a bridge between the eastern Anatolian plateau and western Anatolia. Therefore, from the east to the west, central Anatolia shares many similar morphotectonic features that are outlined in the bordering regions and, therefore, only the differing major features will be summarized here. The major morphotectonic entity of central Anatolia is the Kırşehir‐Niğde Massif (Kırşehir continent of Ş engör and Yılmaz [1981]; Fig.  2.24). It occupies the largest part of the central Anatolian plateau. It is a 300 × 200 km large metamorphic massif displaying a triangular map ̆ e) Massif r­ epresents pattern. The Kırşehir (Kırşehir‐Nigd

38  ACTIVE GLOBAL SEISMOLOGY

Figure 2.24  Geology map that shows the major tectonic components of central Anatolia. KM = Niğde‐Kırşehir Massif; TM = Tokat Massif; IKM = Ilgaz‐Kargı Massif; DM = Devrekani Massif; ADSZ = Araç‐Daday shear zone; RBF = Remnant basin fill; IS = Intra Pontide suture; PMB = Peripheral molass basins; T = Taurus; CAF = Central Anatolian fault; EF = Ecemiş fault; TF = Tuzgölü fault; RTF = Reşadiye‐Tokat fault; ÇB = Çankırı basin.

a continental piece rifted from the Arabian Plate at the beginning of the Mesozoic [Ş engör and Yılmaz, 1981], and drifted away as an independent tectonic entity. This view is verified also by paleomagnetic data [Çinku et al., 2014, 2016]. Late stages of the emergence of the Kırşehir Massif occurred after the Late Eocene‐Oligocene during the late‐post tectonic deformation period that followed the collision between the Pontides, which represents an amalgamated tectonic entity in the north (Fig. 2.24) and the Tauride in the south [Yılmaz et al., 1997a, b; Jaffey and Robertson, 2001, 2005]. Some smaller faults, located in the immediate surrounding of the massif as peripheral faults, appear to have behaved possibly as the detachment faults during the exhumation of the Kırşehir Massif as a core complex [Whitney and Dilek, 1997; Whitney et  al., 2003]. As a consequence of the emergence, a number of tectonically controlled peripheral molasse basins (e.g., the Sıvas, Tuz Gölü, and Çankırı basins in Fig.  2.24) began to form around the Kırşehir Massif during the Late Eocene-Oligocene period [Yılmaz et  al., 1997b; Carter et  al., 1991]. At a later period, new faults were formed, particularly when the NeoTectonic regime started to deform central Anatolia. Some of these young faults cut, reactivated, and/or changed the character of some of the older faults [Koçyiğit, 1996; Bozkurt, 2001b; Umhoefer et al., 2014; Gökten et al., 2013; and references therein]. During this period, the Kırşehir Massif has been  forced to rotate internally and externally [Gökten et  al., 2013] in a counterclockwise sense (see the sense

of motion of the faults in Fig. 2.25a) [Tatar et al., 1996, 2000; Çinku et al., 2014 and 2016; and references therein]. This late period of deformation reorganized boundaries of the Niğde‐Kırşehir Massif and generated a number of new faults [Koçyiğit et al., 1996] (Figs. 2.24, 2.25a). They have defined some local young depressions. The term Ova regime (a Turkish word meaning a vast and flat lowland surrounded by hills) was applied to the basins of this nature that formed under the ongoing transtensional‐extensional deformation, which gave the central Anatolian plateau its distinct structural style [Ş engör, 1979a]. Some of these small basins were formed above the older ones and formed superimposed basins as exemplified from the Sıvas basin [Yılmaz et  al., 1997b; Carter et al., 1991], Tuz Gölü basin [Dirik and Erol, 2003; Koçyiğit et al., 1996; Fernández‐Blanco et al., 2013], and Kırşehir basin [Nairn et al., 2012, and references therein]. The basins have some common characteristics, which may be summarized from the evolution of the Tuz Gölü basin as follows: it started to develop long before the emergence of the Kırşehir Massif and the NeoTectonic era began (possibly during the Late Cretaceous, according to the unpublished data of TPAO). During the PaleoTectonic period, different Tectono‐stratigraphic units (tectonically controlled sediment packages separated by regional angular unconformities) were formed within the Tuzgölü basin [Görür et al., 1984, 1998; Çemen et al., 1999]. The basin survived during the NeoTectonic period, but the boundaries, and thus its shape, have been totally restructured

Morphotectonic Development of Anatolia and the Surrounding Regions  39

Figure 2.25  (a) Morphotectonic map of central Anatolia showing major faults and the consequent morphology [data compiled mainly from Şaroğlu et al., 1992; MTA active fault map of Turkey, 2012; Dirik and Göncüoğlu, 1996; Koçyiğit et al., 1996; Koçyiğit and Beyhan, 1998; Dirik et al., 2001; Bozkurt, 2001; Dirik and Erol, 2003; Altın, 2005; Gökten, 2013; Özsayın et al., 2013]. Heavy white lines with red glows are two transform fault zones: NATF = North Anatolian transform fault; EATF = East Anatolian transform fault. Half arrows show relative movement sense. Faults that diverge from NATF: LF = Ladik fault, SEF = Sungurlu‐Ezinepazarı fault; TrF = Turhal fault; ToF = Tokat fault; CAFZ = Central Anatolian fault zone; RF = Refahiye fault; MOF = Malatya‐Ovacık fault zone; YGF = Yakapınarı‐Göksun fault. Other major faults: TZF = Tuzgölü fault zone; EF = Eskişehir fault. Also: ALD = Aladağ Mountain; NEAFZ = Northeast Anatolian fault zone; HV = Hasandağ volcano and EV = Erciyes volcano (the two prominent volcanoes of central Anatolia); Cities (white letters): OR = Ordu; AK = Aksaray; T = Tokat; S = Sıvas; KY = Kayseri; M = Malatya; O = Osmaniye; ERZ = Erzincan. CTN = Central Taurus Mountains. (b) Morphotectonic map of the Isparta angle and the adjacent areas displaying major structural elements and the consequent morphological features. AG = Acıgöl graben; BG = Burdur graben; ÇBG = Çivril‐Baklan graben; AÇG = Afyon‐Çay graben; AŞG = Afyon‐Akşehir graben system; BG = Beyşehir graben; KG = Kovada graben; SG = Sandıklı graben. Curvilinear yellow lines are thrust or reverse faults. Red lines: Fethiye Burdur fault zone (FBF). They align along the trend of the Pliny and Strabo trenches on the land areas, and therefore are regarded as their extension on the land areas. White lines with strike‐slip displacement indicators are major strike‐slip faults. Lakes: Egl = Eğridir; BL = Burdur; BYL = Beyşehir. Cities: Ant = Antalya; F = Fethiye; S = Seydişehir [data derived from Şaroğlu et  al., 1992; Bingöl, 1980; Koçyiğit, 2000; Koçyiğit et al., 2000; Bozkurt, 2001; MTA map of active faults of Turkey, 2012].

[Bozkurt, 2001b; Özsayın et al., 2013; Fernández‐Blanco et al., 2013]. At the beginning phase, during the development of some transtensional faults, its original size was considerably reduced. Later, some of these faults were taken up by the major strike‐slip faults [Dirik, 2001] (for the main morphotectonic features and development of the Tuz Gölü basin, the reader is referred to Dirik and Erol [2000]). The Çankırı‐Çorum basin has also passed through a complex history of evolution similar to that of the Tuz Gölü basin. Initiation of this basin also probably began during the Late Cretaceous. But it definitely existed during the Eocene time. The basin thus has suffered consecutive deformational phases from the Eocene onward [Gökten

et  al., 2013]. A transpressional tectonic regime began deforming the basin since the beginning of the NeoTectonic regime [Lucifero et al., 2013]. The present landscape of central Anatolia (Fig. 2.25a) began to shape up during the NeoTectonic era [Schemmel et al., 2013]. In this, roles of the active faults are significant (Fig. 2.25a). Presently, some active fault zones define the ̆ e eastern and western boundaries of the Kırşehir‐Nigd Massif. These are the Sıvas‐Kayseri‐Ecemiş fault (known alternatively as central Anatolian fault, CAF, in Fig. 2.24, and CAFZ in Fig. 2.25a) [Koçyigĭ t and Beyhan, 1998], and the Tuz Gölü fault in the east and west, respectively [Dirik and Göncüogl̆ u, 1996]. The Tokat fault and Sungurlu‐ Ezinepazarı fault may also be regarded as other active

40  ACTIVE GLOBAL SEISMOLOGY

major faults of the region (Figs. 2.24 and 2.25a) [Koçyigĭ t et  al., 1996; Dirik and Göncüogl̆ u, 1996; Bozkurt, 2001b; and Gökten et  al., 2013]. Bozkurt [2001] describes many smaller active faults in the region (Fig.  2.25a). They are second‐order faults with respect to NATF and CAFz, but collectively, they divide central Anatolia into smaller subblocks (see Fig. 2.25a). These fault‐bounded discreet subblocks have rotated semi‐independently and have accommodated the bulk strain [Piper et al., 2006; Umhoefer et al., 2007; Iş̇ seven and Tüysüz, 2006, 2010; Özsayın and Dirik, 2011; Gürsoy et al., 2011; Lucifero et al., 2013]. The major faults of northern central Anatolia, the Malatya‐ Ovacık fault, the central Anatolian fault, the Tokat fault, and the Sungurlu‐Ezinepazarı fault, display a similar map pattern, splaying off (diverging) from the NATF zone, curving from northeast to southwest (Figs. 2.1 and 2.25a). Since the beginning of the NeoTectonic era, they have partly accommodated the westerly motions of the Anatolian Plate within the central Anatolian region [Bozkurt, 2001b; Koçyigĭ t and Beyhan, 1998; Jaffey and Robertson, 2001]. The Yakapınar‐Göksun fault (Fig. 2.25a), another splaying fault, which is located near the Erzincan region (ERZ in Fig. 2.25a), is claimed to have acted as the plate boundary fault of the Anatolian Plate before the development of the Karlıova Junction [Westaway and Arger, 1998]. The major faults splaying off of the NATF together with other smaller faults form a map pattern called fishbone structures [Şengör and Barka, 1992]. In the western part of the central Anatolian region, the  southerly curving major faults (e.g., the Sungurlu‐ Ezinepazarı fault; Fig.  2.25a) distribute transpressional forces along the curve [Gökten et al., 2013]. In the central Anatolian fault zone, CAFZ is more prominent and longer than the other faults. It stretches from the north of the Sıvas basin to the Mediterranean Sea (Figs.  2.1 and 2.25a). Koçyiğit and Beyhan [1998] differentiated 24 segments along CAFZ. They have all been active independently during its long history of development, for example, the northern and southern parts were active during the Eocene time [Yetiş, 1984a, b, Yetiş and Demirkol, 1984] and the Oligocene time, respectively [Jaffey and Robertson, 2001; Yılmaz et  al., 1997b]. These segments were reactivated during the NeoTectonic period after the Pliocene. They were attached to one another to form the present intracontinental, transcurrent fault. The central Anatolian fault is characterized by a number of pull‐apart basins that formed along the fault zone. They correspond commonly to step‐overs formed above sinistral double bend. The Erciyes volcano is located on an extensional opening of a similar nature (EV in Fig.  2.25a). Between Kayseri and Pozantı, the southern segment of CAFZ extends along a wide structural depression known as the Ecemiş corridor (the Ecemiş fault). This fault segment accommodates most of the ongoing deformation of

the southeastern part of central Anatolia. Within the fault zone, a left stepping is commonly observed, and about 60 km total offset is estimated [Jaffey and Robertson, 2001]. The curvature of the fault in the central part (see Fig. 2.25a) is regarded as a releasing bend. The Ecemiş fault has also generated a number of large transtensional, riftlike depressions [Dirik et al., 2001]. Major morphological features of CAFZ, such as the rifts and the linear ridges, were developed during the  NeoTectonic era [Şarogl̆ u et  al., 1992, 2001; Altın, 2003, 2005; Sarıkaya et al., 2015]. The young and active faults have obliterated most of the older morphological features. The Ecemiş fault zone forms the western boundary to the Aladağ Mountain (AlD in Fig. 2.25a; Tekeli et al. [1984]), the highest elevation of the region. It is a Late Quaternary glaciated mountain [Sarıkaya et al., 2015] as revealed by the following morphological data: the headword erosion forms deep and narrow canyons at the lower topographic levels. However, the narrow valleys die out before reaching the top of the plateau. The recently elevated north‐south striking small fault‐bounded blocks have forced the Ecemiş River valley to migrate steadily westward. The young lateral alluvial fans and cones that formed under the strict tectonic control of the active oblique faults (the strike‐slip and dip‐slip components) cascade backward, and thus determine the present elevation of the river terraces. The valley profile gets increasingly steep downward [Altın, 2003, 2005]. Recently, Sarıkaya et  al. [2015] determined precisely the slip rate along the Ecemiş fault based on analysis of airborne orthophotogrammetry and GNSS (Global Navigation Satellite System) surveys, which indicate 168 ± 2 m of left lateral and 31 ± 1 m vertical displacement. Accordingly, they obtained a geologic fault slip rate of 4.2 ± 1.9 mm a‐ 1 horizontally, and 0.8 ± 0.3 mm a‐ 1 vertically for the time frame between 104.2 ± 16.5 ka and 64.5 ± 5.6 ka. In addition to the major faults that are summarized above, there are other active important faults in the central Anatolian region. Among these are the Eskişehir fault (EF in Fig. 2.25a).It is located in the region of transition between the western Anatolian extensional province and the central Anatolian transtensional province (Fig. 2.1). The morphological data supported by GPS and seismic data reveal collectively that this is an active right‐lateral strike‐slip fault zone [Barka and Reilinger, 1997; Kahle et  al., 1998]. The GPS measurements across the fault yield 1.5 cm year–1 slip rates along the fault [Barka and Reilinger, 1997; Kahle et al., 1998]. The Eskişehir fault extends more than 200 km from the  Sultanhanı of Konya Province in the south (EF in Fig. 2.25a), through Eskişehir region in central Anatolia, to the Bursa region in the north (see Chapter  7). From south to north, its trend deviates from northwest to west‐ northwest. The fault consists of a number of fault segments.

Morphotectonic Development of Anatolia and the Surrounding Regions  41

Each one of the big segments is given a different name in the literature, such as the Cihanbeyli, Sultanhanı, Inönü, and Yeniceoba faults [Özsayın and Dirik, 2007]. The fault zone cut the Pleistocene‐Holocene sediments and controlled the deposition of the Quaternary alluvium. In the Marmara region, a prominent east‐west striking listric normal fault cuts and displaces the Eskişehir‐Bursa fault. The normal fault defines the steep northern slope of Uludağ Mountain, which rises steeply from the Bursa plain (about 200 m high) to 2500 m heights (see Chapter 7 for the master fault character of the Bursa fault). The apex region of the Isparta Angle is a transitional zone in the western Anatolian extensional province, the central Anatolian strike‐slip dominant tectonic regime, and the western Taurus (Fig. 2.25b). A group of northeast‐ southwest and northwest–southeast‐trending graben‐ horst structures dominates the region. Some of the best examples are the northwest–southeast‐trending Dinar, Beyşehir, Akşehir‐Afyon grabens, and northeast–southwest‐trending Burdur, Acıgöl, Sandıklı, and Çivril‐Baklan grabens (Fig. 2.25b). They are bounded by active, oblique‐ slip (transtensional) faults displaying a significant strike‐ slip component [Price and Scott, 1994, and references therein). However, this region is known to be a previous zone of compression, where the southeast‐vergent Lycian nappes (see Chapter  8) and the southwest‐vergent Akşehir‐Beyşehir nappes were developed to the end of Miocene (see the yellow curving lines in Fig.  2.25b) [de Graciansky, 1972; Ş engör and Yılmaz, 1981; Hayward, 1984a, b; Collins and Robertson, 2003]. Under the southwesterly directed transpressional tectonics, the western Taurus region was shortened and elevated. The Isparta Angle was formed as a consequence of this shortening deformation [de La Motte et al., 1995] (see the discussion on the possible causes of elevation of the western Taurus Mountains in Chapter 5). The geological and morphological evidence collectively reveal that the central Anatolian plateau has passed through different evolutionary stages compared to the neighboring Taurus Mountains since the Early‐Middle Miocene period [Fernández‐Blanco et al., 2012]. During the Early‐Middle Miocene time, the Taurus regions were mostly buried under a sea realm. The evidence for this is the presence of the marine limestone of the Lower‐ Middle Miocene age, observed mostly on the southern slopes of the Taurus up to heights of 1500 m (see the MTA Geology maps of Turkey on the scale of 1/100,000 and Fernández‐ Blanco et al. [2012]. The marine sediments of the Upper Miocene age are also locally identified in the central (e.g., Başyayla) and the northern edge of the Taurus Range [Schildgen et  al., 2014]. While the Taurus region was under the sea, central Anatolia was a land. Starting from the Oligocene‐Early Miocene time, continental units, mainly fluvial and lacustrine sedi-

ments, were deposited in central Anatolia [Fernández‐Blanco et al., 2013; Gökten et al., 2013; Lüdecke et al., 2013; and references therein]. Therefore, central Anatolia was standing at a higher elevation with respect to the Taurus Range during the entire Miocene period (the sea sediment of the Sıvas basin did not extend that far south to cover the central Anatolian plateau, and has a different origin and connection). Analyses of the Oligocene‐Early Miocene stable isotopes across central Turkey reveal also that there was no orographic barrier in the locations of the peripheral mountains at that period [Lüdecke et al., 2013; Schildgen et al., 2014]. The continental deposits of the central Anatolian ­plateau replaced the marine sediments without apparent angular unconformity indicating that the whole region was elevated en block, which is similar to the coeval event in eastern Anatolia. Therefore, it may be concluded that the two regions were parts of the same tectonic entity during that period. Schildgen et  al. [2012a, b, c] stated also that the uplifting of the eastern Anatolian region propagated westward through time. From the Late Miocene onward, the morphotectonic character of central Anatolia began to change with respect to the surrounding regions. The peripheral mountains on both sides began to rise with a higher rate of elevation to reach their present altitudes. The relative motions of central Anatolia with respect to the peripheral mountains may thus be compared to the motion of a seesaw. The isotope data also confirm that the uplifting of the central Anatolian plateau has occurred at a lower rate compared to the peripheral mountains [Schemmel et al., 2013]. According to Ş engül et al. [2013], it was only 1 km during the last 7 m yr period. The central and eastern parts of Turkey experienced two stages of uplifting: (1) wholesale uplift of central and eastern Anatolia together with the peripheral regions, (2) preferential (higher rate of) uplifting of the peripheral mountains. For the wholesale uplift, the following main mechanisms are proposed: (1) slab steepening and break off of the northerly subducting eastern Mediterranean oceanic lithosphere [Cosentino et  al., 2012]; and (2) the slab tear [Schildgen et  al., 2012a, b, c]. Removal of the lithospheric mantle and the associated vertical mantle flow beneath the central Anatolian plateau triggered the uplift [Ş engül et al., 2013]. The low‐velocity seismic zone beneath the crust is the indication of the hot / buoyant sublithospheric mantle underlying the plateau that is regarded as responsible from the epeirogenic elevation. It has been demonstrated recently that the wholesale uplift of the central Anatolian plateau began during the later period of the Late Miocene (>7) [Cosentino et  al., 2012; Aydar et al., 2013; and Schildgen et al., 2012a, 2014; Çiner et al., 2015]. The marine sediments along the southwestern margin of central Anatolia (the narrow belt between the

42  ACTIVE GLOBAL SEISMOLOGY

Tauride and central Anatolia) were uplifted about 1500 m [Schildgen et al., 2012] during this period. The surface uplift rate in the Mut basin located on the Tauride was calculated to be as an average at 0.24–0.25. But it was much higher (0.6–0.7 mm) since 1.6 m yr period [Cosentino et al., 2012, Schildgen et al., 2014]. Recently, Çiner et al. [2015] demonstrated also that the Quaternary uplift of the central Anatolian plateau is closely related to the uplift of the northern and southern plateau margins, and river ­terraces of the Kızılırmak indicate an average incision rate of 0.051 ± 0.01 mm/yr (51 ± 1 m/Ma) since ~1.9 Ma. 2.5. THE PERIPHERAL MOUNTAINS (PONTIDES AND TAURIDES) The peripheral mountains delimit the Anatolian Peninsula in the north and the south (Fig. 2.1). The northern range is known as the Pontides extending from the Lesser Caucasus to the Marmara region, where it terminates abruptly with an oblique fault zone (WBFP; see Chapter 7) displaying sinistral strike slip displacement [Yigĭ tbaş et al., 2004]. Along the southern side of Anatolia, there are two mountain ranges, the Taurus and the Bitlis‐Zagros suture mountains in the west and east, respectively (Fig.  2.1). They do not form one continuous mountain range as commonly thought. In the eastern areas, the Tauride Mountains align in a northeast trend, and terminate by the Erzincan depression (ED in Fig. 2.1), lying along the NATF zone (Fig. 2.14). The boundary between the Bitlis Range and the Taurus Range is a sharp division represented by a strike‐slip fault zone separating high‐grade metamorphic rocks of the Bitlis Range from the nonmetamorphic rocks of the Taurus carbonate platform (Figs. 2.1 and 2.6). The Pontides and Taurides as the peripheral mountains of the Anatolian Peninsula are geologically entirely different. The former was an active continental margin whiles the latter was a site of a passive continental margin during most of their geological histories. Their morphotectonic development, on the other hand, have similarities. They are both young mountains. They display east‐west zigzagging map patterns (Fig. 2.1), which is more pronounced in the southern range.In the following paragraphs, some important morphotectonic features of the peripheral mountains are outlined. The Pontides delimit the Black Sea in the south, and run parallel to its southern edge. Therefore, it is also known as the Black Sea Mountains. The Black Sea is a trapped remnant oceanic basin left after the consumption of the Tethyan oceans [Şengör et al., 1980, 1982a; Şengör and Yılmaz, 1981; Yılmaz et al., 1997]. The ocean floor character of the Black Sea lithosphere is in agreement with the seismic velocities [Yegerova et al., 2013]. The detailed geological account of the history of the Pontides is given by Yılmaz and Tüysüz [1984], Yılmaz et al. [1993b], and Yılmaz et al. [1997], which may simply be summarized as follows: the Pontides behaved

as an Andean type magmatic arc with respect to the surrounding oceans initially facing the PaleoTethys located in the north and the NeoTethys in the south. The demise of the oceans generated volcanic arcs along the Pontides [Şengör and Yılmaz, 1981; Yılmaz et al., 1997]. Therefore, the Pontides were formed as an arc, and remained as an arc to the very late  stages in the history of the NeoTethyan Ocean. The arc reversals occurred a number of times from the Late Paleozoic–Early Mesozoic to the Tertiary [Şengör, 1990; Yılmaz et al., 1997]. A late reversal phase generated the present Black Sea during the Cretaceous [Görür et al., 1983; Görür, 1988b; Yılmaz et al., 1997]. Present elevation of the Pontides is a young event occurring after the Late Miocene. Evidence for this may be summarized as follows: (1) there are Lower‐Middle Miocene marine sediments that are elevated on the shoulders of the Pontides. They are observed along the southern flank of the central Pontide around the Kelkit‐Reşadiye area and in the Artvin region of the eastern Pontide. Tunogl̆ u [1991] reports also from the Devrekani region of the western Pontide presence of Lower‐Middle Miocene marine sediments. (2) Poorly sorted and poorly lithified coarse clastic sediment, derived from the elevating mountains, and deposited in the adjacent lowlands is Pleistocene‐Quaternary in age. (3) Terrace deposits and sediment fills within the major valleys formed along and across the mountains are not older than Quaternary. The morphological features observed on the mountains (discussed next) and the geodetical measurements [Yıldırım et al., 2011, and references therein] also collectively support a rapid and recent elevation. Major morphological features of the Pontides are critically important to assess and evaluate their elevation history and the morphotectonic development. There is a close temporal association between the major tectonic events that affected the whole of eastern Turkey including the elevations of the Pontides, the eastern Anatolian high plateau, and the Bitlis Mountain Range, which are roughly synchronous. The Pontides began to rise to the present height after the Late Miocene, when the northward advance of the Arabian Plate began to squeeze the east of Anatolia (Fig. 2.21). The north‐south shortening deformation initially eliminated the sea realm from the region and then the wholesale uplift of eastern Turkey occurred. Following this event, the Pontides began to be more effectively squeezed between the old and resistant oceanic lithosphere underlying the Black Sea and eastern Anatolia (Fig. 2.22a). During this period, both peripheries of eastern Anatolia responded to the north‐south compression by thrusting to the north and to the south (Figs. 2.22a, b; 2.23). Finally, they began to rise and form the present peripheral mountain ranges. The thrusting of the Pontides above the post‐Miocene rocks in the offshore (unpublished seismic sections of TPAO) and onshore areas are clearly observed (for the north vergent thrusts of the central Pontides see Yıldırım et al. [2011],

Morphotectonic Development of Anatolia and the Surrounding Regions  43

Figure 2.26  (a) Geology map of the eastern Pontide (modified from the MTA 1/500,000 scale geology map of Turkey). The red arrows with yellow glow indicate relative motions of the fault‐bounded blocks as revealed by the relative offset of the granitic bodies. The ongoing crustal deformation of the east of Turkey generates the sinistral shear regime (red thick arrows). Straight white lines with black glow represent oblique faults with a dominant left‐lateral strike‐slip component. The white line along the coast displays zigzagging seashore, formed as a result of interactions of the two fault sets. (b) Photo from the eastern central Pontide showing the two sets of faults. The black arrows show strike‐slip faults and the resultant displacements. The red lines indicate normal faults and the associated vertical displacements. Interactions of the two different sets of faults formed a number of small bays, promontories, and zigzagging coastline.

and references therein). The south‐vergent thrusts may be observed along the southern flank of the Pontides, where the basement rocks and the ophiolites were thrust onto the Middle‐Upper Miocene sediments (Fig. 2.16). The northern flank of the Pontides facing the Black Sea is structurally defined by two different sets of faults (Fig. 2.26a, b): (1) oblique faults; they strike Northeast and northwest in the eastern and western parts Pontides respectively. (2) normal faults (Fig.  2.26a). The seismic data unravel continuation of these faults under the sea (unpublished TPAO offshore seismic data). The northeast‐striking oblique faults commonly display a significant reverse slip component; on the northern flank of the mountain, the

Pontides were thrust above the young (Post Miocene) sediments of the Black Sea. Its analogue to the east is the Achara thrust belt of Lesser Caucasus, which was thrust onto the young sediments of the Riouni basin [Banks et al., 1997]; (3) active normal faults, they accommodate the elevation of the mountain range along both slopes (Fig. 2.27a). Commonly, the oblique faults are longer in extent, but the normal faults are more numerous (Figs. 2.26 and 2.27). The latter faults commonly cut and displace the former faults. As a result of intersections of the oblique and the normal faults, zigzagging coastlines are generated, forming a number of little bays and promontories (Fig.  2.26a, b). Collectively, the two sets of faults divide the mountains into rectangular

44  ACTIVE GLOBAL SEISMOLOGY

Figure 2.27  (a) Schematic diagrams showing tectonic development of the Pontide Mountains. (b) The red arrows represent two opposite vergent thrust system that formed under the ongoing north‐south compressional tectonic regime, which is presently deforming all of eastern Turkey. The thick green arrow refers to the rise of the mountain as a consequence of the north‐south compression. The curvilinear red lines represent the normal faults enabling subsidence of the north and south flanks of the mountain range in response to the physical instability that formed as a result of continuing north‐south shortening and thickening, and the consequent elevation. The faults are normal listric in nature displaying a downward flattening pattern. The black arrows indicate relative motions of the fault blocks. (b) Photo from the Giresun region of the eastern Pontides showing perfectly planar steep slopes corresponding to traces (a) of the normal fault dividing the mountain as fault‐bounded blocks descending as fault steps. The red broken circle shows the location of the photo in (c). (c) Sudden steps in the morphology represent the faulted blocks bounded by the listric normal faults. The landward (backward) dipping of the top of the faulted blocks reveals the listric nature of the faults. (d) Photo from the central Pontides showing the traces of a steep listric normal fault (red arrow). The white lines in the horizon correspond to the step faults that formed sudden steps in the morphology The yellow arrows indicate the back‐tilted block resulting in a side valley (white arrow) running subparallel to the seacoast.

blocks (Fig. 2.27). Possibly because of this, some previous studies compared the Black Sea Mountain Range to the block faulted Harz‐type mountains rather than the Alpine‐ type mountains [Zankl, 1962; Westrum, 1962]. For the development of this complex fault pattern, the  following mechanism may be proposed. Under the ongoing north‐south compression, the Pontides were squeezed, shortened, thickened, and consequently elevated to above  3000 m heights (Fig.  2.28a). Therefore, they became physically unstable and the flanks of this elevated mass began to descend along step faults in the northern as well as the southern flanks (Figs. 2.27a, b, d). From the east to the west, height of the Pontides decreases steadily from above 2000 m (the Kaçkar Mountain; K in Figs.  2.1 and 2.26a) to 1000 m (Düzce region; d in Fig. 2.1). This appears to be connected with

the distance from eastern Anatolia, because the magnitude of the north‐south compressional component of the stress field decreases away from the east of Turkey (Fig. 2.2). Undoubtedly, NATF as northern boundary of the Anatolian Plate is mainly results from transfer and distribution of the compression (Figs. 2.2 and 2.21; Emre et al. [2009], Yildırım et al. [2013]). In fact, based on the InSAR analysis, Peyret et al. [2013] recently ­demonstrated the present‐day strain distribution across a segment of the central bend of the North Anatolian fault zone. A young mountain character of the Pontides is identified from its morphological features. Some of these may be listed as follows: (1) a perfectly planar, flat lying erosional surface is recognized at the top of the mountain range (Figs. 2.28a, b and 2.29a). This remnant erosional surface is traced all along the mountain range, and may well be

Morphotectonic Development of Anatolia and the Surrounding Regions  45

Figure 2.28  (a) Photo showing the flat top of the Pontide Range (yellow line and brown rectangle) representing the erosional surface elevated up to 3000 m heights on the mountain. The undulating lines are the hanging valleys revealing that away from the crest line, the carving increases steadily. However, the headword incision has not yet reached the top of the mountain to cut whole width of the mountain range. The green straight arrows represent abrupt changes of the slope angles. (b) A southeasterly aerial photo from the northern flank of the eastern Pontides between Trabzon and Giresun. The red arrows display angular slopes, steepening away from the top of the mountain. The curved green lines are the side valleys. The main curving drainage pattern that follows the side valleys developed as a result of the back‐tilting (yellow arrow) of the downthrown blocks.

Figure 2.29  (a) Views of valleys from the northern flank of the Pontide Mountains in the Giresun region. The slopes get progressively steeper approaching the Black Sea, and make angular connections with one another. Between them they form morphological angular unconformities revealing that the rise of the mountain is recent and episodic. The time elapsed since the elevation began is too short for erosional forces to obliterate the angular connections of the slope angles. (b) A view from the Çoruh River valley. Incision at the lower levels has formed a canyon. Within the valley, different sets of hanging terraces are observed at different heights. They were formed in response to the elevation of the mountain. Sediments on the terraces are fluvial deposits of Plio‐Quaternary age.

correlated with the marker erosional surface that is recognized on the eastern Anatolian high plateau. It stands at the 2000 m heights in eastern Anatolia but rose to more than 2000 m on the Pontides. The erosional surface above the eastern Anatolian plateau butts against steep slopes

of Pontides across the boundary faults, which apparently elevated the erosional surface to the top of the mountain. (2) At the top of the mountain, a set of hanging valleys is visible (Fig. 2.28a and 2.29a). Away from the main major divide, incision in the hanging valleys increases rapidly

46  ACTIVE GLOBAL SEISMOLOGY

along the southward and northward flowing streams (Figs. 2.28a, b and 2.29 a, b). However, the headword erosion from both flanks has not yet reached to the top of the mountain. The valleys with the exception of the Çoruh River gorge do not cut the entire width of the mountain to connect inland Anatolia with the Black Sea (Fig.  2.14). From the mountaintop to the low levels, the river valleys display increasingly steeper slopes and form deep canyons close to sea levels (Fig.  2.29 a, b). Accordingly, the slope angles also increase significantly to the lower levels. Changes of the slope angles are sharp and angular (Figs. 2.29a, b) indicating that (1) the time elapsed since the Pontides began to rise is too short for the erosional forces to smoothen and obliterate the angular connection; (2) the uplift possibly occurred episodically; and (3) within the valleys different sets of hanging river terraces are identified (Fig. 2.29b), and the fluvial deposits on the terraces are Quaternary in age. The slopes of the Pontides Mountains against the Black Sea are commonly steep (Figs.  2.27a, c, d), and form a number of clear steps in the morphology (Fig. 2.27). They are due to the normal faults, along which the slopes descend as a response to the uplifting of the mountain. On top of the fault, bounded blocks dip backward (back‐ tilted) (Figs.  2.27a–d) revealing the listric nature of the normal faults. Due to the back‐tilting, a set of side valleys running subparallel to the main trend of the mountain and the coastlines are formed (Figs.  2.27d and 2.28b). They stand against the major flow direction and divert the seaward flow of the streams. As a consequence, the rivers and streams follow an unexpectedly long path to reach the sea (Figs. 2.27d and 2.28b). Studies on rate and mechanism of the rise of the Pontides are few. Based on isotope ages from the granitic plutons (the red‐colored rock unit in Fig. 2.26a), Boztuğ et al. [2004] state that the rate of uplift was slow between 80.7 m yr and 62.4 m yr period (Late Cretaceous), was moderate between 57 m yr and 47 m yr (Eocene) period, and was rapid during the last 3.5 m yr period (since the Early Pliocene). The recent higher rapid rate of the rise is supported also by the morphotectonic data that are summarized above. Recently, Yıldırım et al. [2013] measured uplift rates of the marine terraces from the Sinop region of the central Pontides to be nonuniform, varying between 0.02 and 0.22 mm/yr. In the central part of the Pontides, on the other hand, they measured an average of 0.3 mm/yr uplift rate. Based on 10Be, 21Ne, and 36cl concentrations from gravel‐covered fluvial terraces and pediment surfaces along the trunk stream of the Gökırmak River, they yield model exposure ages ranging from 7 ± 1 ka to 346 ± 45 ka and average fluvial incision rates over the past ~350 ka of 0.28 ± 0.01 mm a‐1. Similarities between river incision rates and coastal uplift rates at the Black Sea coast

s­uggest that regional uplift is responsible for the river incision [Yıldırım et al., 2013]. Keskin et al. [2011] also found some scattered uplift rates from the marine terraces in the Trabzon region of the eastern Pontides. They vary between 0.07 and 0.017 mm/yr. Earlier Ertek and Aytaç [2001], studying the uplift rate from the entire Pontides coastal areas, reached a similar conclusion. The wide variation of the uplift rates obtained from the same areas appear to be connected with the main structural character of the Pontides, which may be summarized as an interaction of the two simultaneous events: (1) elevation of the mountain range, and (2) differential descent rate of each one of the fault‐bounded blocks. Since the two events occur partly coevally, they cause the widely scattering values varying from block to block. Therefore, in order to obtain more realistic rates of the elevation, a study needs to take into account (1) rate of elevation of the mountaintop and (2) the interaction of the two separate events. Yildırım et al. [2011] emphasized a possible role of the NATF as a driving force of the uplift (see Figs. 2.2 and 2.21). In the light of detailed geomorphological studies from the southern edge of the central Pontide near the central Anatolian plateau, they proposed that the strain accumulation along the bend of NATF was responsible from the elevation of the central Pontides range, which seems to be fairly plausible. The Taurus as the southern peripheral mountain range (Figs. 2.1 and 2.30) displays most of the morphotectonic characteristics that are similar to those described from the Pontides (for a treatise of the subject see Erol [1991]). Therefore, they will not be repeated here. The northern as well as the southern flanks of the Taurus Mountains are defined by steeply dipping normal faults (Fig.  2.31a). Recent and cyclic elevation of the mountain are expressed by the slopes and the valleys that were formed across the mountain (see Fig. 2.31b). For the young elevation of the Taurides, the following data may be listed: (1) all along the southern slope of the mountain range reefal limestone of Lower‐Middle Miocene crop out at various heights up to 2000 m (the MTA geological maps of Turkey on the scale of 1/100,000, and also Yıldız et al. [2003]; Cosentino et al. [2012]; and references therein); (2) in the Denizli area on the northern slope of the Babadag,̆ southwestern Anatolia, beach sediments of the Upper Miocene age crop out, and this unit passes laterally to the lacustrine limestone [Yılmaz et al., 1999, 2000; Gürer and Yılmaz, 2002]. The same lacustrine sediments, in turn, may be traced southward to the heights above 1500–2000 m on the northern slopes of the Taurus Range. The data summarized above collectively indicate that the present elevation of the Taurus Range began after the Late Miocene time. This view is supported further by the following isotope data: (1) the surface uplift rate in the Mut basin on the Taurides close to the southern edge of

Morphotectonic Development of Anatolia and the Surrounding Regions  47

Figure  2.30  Map showing zigzagging map pattern of the Taurus Mountains. The pattern was formed possibly under westerly directed compressional forces (black and red arrows). Red lines with black rim are the faults that cut the closures of the apexes and interrupt their continuity. These faults are oblique in nature, which has caused the southward penetration of the western blocks into the Mediterranean Sea (brown arrows) for more than 100 km (green arrows display approximate distance of lateral displacement). The brown curves indicate opposite sense of rotations of the regions around the Isparta angle. The white broken lines correspond to the transtensional provinces that were formed as a result of the tectonic deformations that affected the Taurus Mountains. The black ellipse shows location of Figure 2.31(c).

the central Anatolian plateau was identified to be varying between 0.25–0.37 mm/yr during the Late Miocene. But after the Late Miocene, the rate was more rapid. It was about 0.72–0.74 mm/yr until 1.62 m yr ago. From then on, it increased to 1.66 mm/yr [Schildgen et al., 2012a, 2014]. The figures display clearly that uplift rate has increased cyclically and significantly through time. The river morphology and profile projection also support cyclic uplift. According to Westaway et al. [2005], the relatively arid interior of southwestern Turkey has uplifted by ~400 m since the Middle Pliocene, whereas its southern part, closer to the Mediterranean Sea and with a much wetter climate, has uplifted ~900 m since the Middle Pliocene, and ~300 m since the Early Pleistocene. Dhont et  al. [1999] connected the rise of the Taurus with the northward subducting eastern Mediterranean oceanic lithosphere. Schildgen et al. [2012b, c] advanced the view further and suggested that the P wave seismic tomography is consistent with break off of the subducted slab under eastern Anatolia. They proposed also that not only the slab break off resulted in the elevation of the entire eastern Anatolia as suggested previously by Piromallo and Morelli [2003] and Keskin [2003], but also the rise propagated westward to central Anatolia. The westward limit of the broken slab under central Anatolia favors its significant role [Gans et al., 2009]. The arrival of

the Eratosthenes seamount at the collision zone may also be responsible from the increasing rate of the uplifting of the Tauride. However, there is no agreement on the mechanisms that are proposed to explain the higher rate of elevation of the western Taurides with respect to the central Anatolian plateau. Some of the previous studies favor the role of an approximately north‐south compression [Bertotti and Fernández‐Blanco, 2014], while some others favor a north‐south extension [Dhont et al., 1999]. The major morphotectonic pattern of the Taurus Range between the eastern and the western ends has not yet been fully treated and satisfactorily explained. The Taurus Mountains display a tight east‐west zigzagging buckled map pattern (Figs.  2.1 and 2.30). This pattern may mean that the Taurus have suffered an approximate east to west shortening deformation, and have accommodated this deformation by folding around vertical axes. Suess [1901] was the first scientist who pointed out that the zigzagging area around what we now call the Isparta Angle, is a region where Dinarides and Taurides form a center of syntaxis. The westerly directed compression caused trusting followed by an escape tectonics that generated a south‐southwest striking strike‐slip‐wrench fault system effecting more severely the apex regions of the buckles, that is, the Isparta Angle and the Andırın‐Misis fault zone in the west and east, respectively (the red lines

48  ACTIVE GLOBAL SEISMOLOGY

Figure 2.31  (a) Photo showing a view from the southern flank of the Taurus Mountains in the Antalya region displaying a set of step faults (a set of parallel, steeply dipping, closely spaced, normal faults. (b) An oblique aerial photograph from the central Taurus Mountain showing the flat mountaintop (the elevated erosional surface) and the river valley along which the deep carving dies out before reaching the plateau. (c) A Google photo from Taurus Mountain showing the region between the Gülnar Plateau (located between the towns of Silifke and Anamur) and the Mediterranean Sea, where the oblique slip faults may be clearly identified. The stream valleys (green lines) follow a zigzagging pattern under the strict tectonic control of the faults and, therefore, instead of following a shortest distant possible on a steep slope, they follow a long and diagonal path.

Morphotectonic Development of Anatolia and the Surrounding Regions  49

in Figs.  2.1 and 2.30). The southwest trending oblique slip faults (strike‐slip coupled with dip‐slip displacement) have played significant roles in the development of the present morphology of the mountains (see Fig. 2.31c). Both sides of the Isparta Angle were forced to rotate in the opposite direction when the shortening deformation reached an advance stage (Fig. 2.30), where a transtensional regime replaced the previous compressional regime. The rotation is partly coeval with the bending of the Isparta Angle. The previously formed thrusts were initially steepened and elevated the lands, and then they were collapsed when the southwest escape tectonic regime began (see the region bounded by the white broken lines in Fig. 2.30). The present basins (e.g., Aksu, Köprü, and Manavgat) began to form during this period, and they were developed partly above the older foreland flexural basins. Presently, in this region, two diverging trends of grabens met (Fig.  2.25b). But prior to this, the westerly vergent thrusts placed the eastern block above the western block (see the reverse faults in Fig. 2.25b). The thrusts were generated to the end of Miocene and possibly continued during the Early Pliocene [Şengör and Yılmaz, 1981; Akıncı et  al., 2003; Koçyiğit et  al., 2000b; Poisson et  al., 2003; Koçyiğit, 2005]. During the thrusting, the flexural foreland basins were developed in front of the thrusts. The oblique slip faults with pronounced strike‐slip components, which were generated under the southwest‐directed escape tectonics, have torn the entire width of the Taurus and extended to the Mediterranean Sea region in the eastern and the western parts of the Taurus (see Figs. 2.1 and 2.30; Robertson et  al. [1991]; Westaway et  al. [2005]; van Hinsbergen et  al. [2010, 2012]). Along the oblique faults, big lateral as well as dip‐slip displacements occurred (see Figs.  2.30 and 2.31c). Across the easterly located major fault (the Sarız‐Misis fault; KMF in Fig. 2.5), the Lower Paleozoic sedimentary rock representing the basement rocks of the Taurus Mountains is in contact with the Binbogă Metamorphic Massif in which the whole sequence including the Paleocene‐Lower Eocene (?) rocks are metamorphic [Yılmaz, 1993] (see northern part of Fig. 2.6). The present structural configuration of the Taurus system may be regarded as the result of an interaction between the Hellenic trench and the Anatolian Plate. The southwesterly motion of the Taurides, being the southern part of Anatolian Plate if restricted by an obstacle in the western termination, is expected to generate an approximately east‐west compressional deformation. A possible cause of a resistance may be the eastern end of the Hellenic trench. It is represented by the oblique fault system in the form of the Pliny and Strabo trenches and the Fethiye‐Burdur strike‐slip fault along the northeasterly extension on the land (Figs. 2.1 and 2.25b). If it could not keep pace with the rate of the southwestward advance of the Anatolian Plate, the motion along this fault system may have generated an approximately east‐west compressional and north‐

south extensional deformation in the Taurides. If this assumption is correct, then it may be concluded that the morphotectonics of the Taurus Range is determined by the relative rates of motions of the blocks [Kissel et  al., 1993] across the Strabo‐Pliny trenches [Peters and Huson, 1985], the Fethiye Burdur fault [Özbakır et  al., 2013]. Recently van Hinsbergen et al. [2012] kinematically restored the Aegean region through time and suggested that the opposite rotations of the western and southeastern Aegean regions caused up to 650 km of trench‐parallel extension between the southwest Peloponnesus and the Island of Rhodes. This rotation was accommodated along a transfer fault zone between the Menderes Massif and the western Taurus around the Isparta‐Burdur region. Another possible reason to accelerate the uplift rate of the Taurus may be the different circular rotations of the northern and the southern parts of the Anatolian Plate. During the anticlockwise rotation of the Anatolian Plate, the southern region follows a tighter and closer circular path compared to the northern regions (Fig. 2.2) and this possibly generates an east‐west shortening deformation in the western Taurus region. In summary, the Taurus Range has suffered an approximately east to west directed shortening deformation, which determined the present zigzagging pattern of the mountain ranges. This has no direct connection with break off and or retreat of the subducted slab and/or its backward retreat, which is assumed to have caused the elevation of the mountains to the present heights. As a result of the rapid elevation, the northern as well as the southern slopes of the mountain became unstable and began to descend by the step faults under the north‐south extension generated as response to the elevation. 2.6. THE NORTH AND EAST ANATOLIAN TRANSFORM FAULTS The North Anatolian transform fault (NATF) and the East Anatolian transform fault (EATF) are the most active tectonic elements of Anatolia that have formed during the NeoTectonic period [Şengör, 1979a, c] (Fig. 2.1). This is indicated by seismic activities [Yolsal‐Çevikbilen et  al., 2012] and associated morphotectonic features that formed along these fault zones [Gökçe et al., 2014, and references therein]. The transform faults collectively define the Anatolian Plate, which is moving away from the point of convergence of the two faults westward to accommodate northward advance of the Arabian Plate [Şengör, 1979c; and Şengör and Kidd, 1979] (Fig. 2.1). 2.6.1. The North Anatolian Transform Fault There are numerous publications on the nature and kinematic analyses of NATF starting with McKenzie [1972] and Şengör and Canitez [1982] (for detailed

50  ACTIVE GLOBAL SEISMOLOGY

Figure 2.32  The North Anatolian transform fault zone (NATF) and associated faults on the morphology map of Anatolia. Yellow lines are the faults that diverge from NATF: MF = Merzifon fault zone; LF = Laçin fault zone; YEF = Sungurlu fault zone or Yağmurlu‐Ezinepazarı fault zone; AF = Almus fault zone; TF = Tokat fault zone; CAFB = Central Anatolian fault zone; YGF = Yakapınar‐Göksun fault zone; MOF = Malatya‐Ovacık fault zone; NEAFZ = Northeast Anatolian fault zone; EADFB = Erzurum Aşkale Depression fault belt. White letters are the towns located along NATFZ: I = Ilgaz; T = Tosya; K = Kargı; HV = Havza; ER = Erbaa; NK = Niksar; RŞD = Reşadiye; SŞ = Suşehri; RF = Refahiye; ERZ = Erzincan. The green letters are depressions, small basins located along the NATF: I = Ilgaz; T = Tosya depression; K = Kargı basin; RR = Erbaa basin; SŞ = Suşehri basin; ERZ = Erzincan basin. Light green letters are sag ponds and lakes located along NATF. KPDL = Karapürçek Dam Lake; DDP = Dündarli Pond; AGL = Aşağı Gölyazı Lake; LL = Ladik Lake; UK‐TL = UluKöy‐ Taşova Lake; ER = Erbaa Lake; A‐ÇDL = Akçaağıl‐Çamlıca Dam Lake; GKL = Gölköy Lake; AYL =  Arpayazı Lake; GL = Gölova Lake. The red rectangle shows the location of Figure 2.34.

accounts on NATF, the reader is referred to the various publications of A. M. C. Şengör [Şengör, 1979c; Şengör et  al., 1985b, 2005, 2011, 2014; and references therein). Therefore, only its basic morphotectonic characteristics will be outlined here, and some of the problems that are still debated will be discussed briefly. NATF is a 1200 km long, dextral strike‐slip fault behaving as an intracontinental transform fault [Şengör, 1979c; Fichtner et  al., 2013]. It connects the East Anatolian Convergent‐Compressional Province with the West Anatolian‐Aegean Extensional Province (Fig. 2.1). It cuts Anatolia all along its length, stretches as a continuous fault zone from the Karlıova area in the east to the Dokurcun area of the Bolu Province in the west (Figs. 2.1, 2.32, and 2.33). Along the zone, it is represented by a distinct and nearly continuous single narrow valley, structurally resembling a rift (Fig.  2.32). Its continuation is interrupted in places by right‐stepping en echelon offsets of the master fault, which lead to the opening of local basins of dissimilar geometry as exemplified from east to west by the Erzincan, Gölova, Suşehri, Niksar, Taşova, Ladik, Havza‐Vezirköprü, Kargı, Tosya, Ilgaz, Yeniçağa, and Bolu depressions (Figs. 2.32 and 2.33). They are commonly located in the areas where the fault zone makes a

left bend and right step‐overs as exemplified by the Havza‐ Vezirkörü and Kargı basins (Fig. 2.32). As a consequence of the left bends, the regional strike of the fault zone turns gradually from northwest to east‐west (Figs. 2.1 and 2.32). There are some detailed studies on the local basins, for example, on the Niksar basin [Tatar et  al., 1995; Tatar, 1996; Tatar et al., 2012b; Barka et al., 2000], the Erbaa basin [Barka et al., 2000; Koçyiğit, 1996], the Merzifon‐ Suluova basin [Rojay, 1993], the Suşehri basin [Koçyiğit, 1989; Polat et al., 2014], and the Erzincan basin [Barka and Gülen, 1989; Koçyiğit and Rojay, 1992; Över et  al. 1993; see also Bozkurt, 2001b; Şengör et  al., 2005; and references therein]. But the lack of seismic data on the local basins makes it difficult to know whether they formed as a typical pull‐apart basin (e.g., Niksar and Taşova‐Erbaa basins [Barka et  al., 2000]; Fig.  2.32), a rhomb graben basin, or a negative flower [Polat et  al., 2014]. In some of depressions (e.g., the Reşadiye area and the Suşehri, and the Düzce basins), a young master fault cuts the basin in the middle and forms a thin pressure ridge in the center of the young alluvial plane (Figs. 2.34 and 2.35a). Seismically, the central fault is more active compared to the basin boundary faults. From the Dokurcun area of the Bolu region westward, NATF

Morphotectonic Development of Anatolia and the Surrounding Regions  51

Figure  2.33  Major branches of NATF on the morphology map of the Marmara region. NBNATF and SBNATF = northern and the southern branches of the North Anatolian transform faults, respectively; WBFP = western boundary fault zone of the Pontides; EBF = Eskişehir‐Bursa fault zone; PPB = Pamukova pull‐apart basin; SL= Sapanca Lake and depression; Adapazarı P = Adapazarı plain; Düzce Dep = Düzce depression; KR = Karasu Town; MRV = North Anatolian fault along the Mudurnu River valley; DK = Dokurcun; Ulu D and SM D = Uludağ and Samanlıdağ mountains, respectively.

Figure 2.34  Anostomasing map pattern of NATF along the Kelkit River valley around Reşadiye Town (for the location of the area see Fig.  2.32). Villages: SK = Saraykışla; YT = Yenituraç; ÇP = Çayırpınar; KŞK = Karşıkent; AP = Altıparmak; ÇB = Çambalı; ÇBL = Çavuşbeyli; KL = Köklü; GK = Gülkonak.

52  ACTIVE GLOBAL SEISMOLOGY

Figure 2.35  Photos from NATF. (a) A pressure ridge in the middle of the Suşehri basin. The branches of the NATF bound the lensoidal narrow hill. (b) An example to the zone of cataclastic deformation in radiolarite-pelagic limestone unit in NATF zone, west of the Erzincan Region. (c) An oblique areal photo showing anastomosing pattern of faults within the NATF zone between the Umurbey and Evrenli villages near the Gulf of Gemlik (see Fig. 2.33 for the location of the Gulf of Gemlik in the Marmara region). The green shaded surface is the steeply dipping fault plain. (d) A close‐up view of anastomosing strike‐slip faults in a rock quarry near the Gazanfer Bilge Meslek Yüksek Okulu (The Gazanfer Bilge Polytechnic School), located in Karamürsel Town of the Izmit Bay region.

bifurcates into two main strands as the northern branch and the southern branch (Fig.  2.33). The depressions formed along these branches are wider and thus support the view of Şengör et al. [2005] that NATF is a westerly widening fault zone. It is the widest around the Marmara Sea region (Fig.  2.33). Large depressions have formed along the northern (e.g., Izmit Bay and the Sea of

Marmara basin) as well as the southern branch (e.g., the Pamukova depression, the Iznik Lake; Fig. 2.33). Major structural and morphotectonic characteristics of NATF may be summarized as follows: 1. It is a right‐lateral strike‐slip fault. Within the zone there are a number of short or long faults collectively forming an anastomosing (less commonly parallel) pattern

Morphotectonic Development of Anatolia and the Surrounding Regions  53

(Figs. 2.34 and 2.35a–d; Şaroğlu [1988]). Along the anastomosing fault strands, the faults have collided with one another, and jostled away the intervening blocks. They push the neighboring blocks and cause their rotations [Dhont et al., 1998]. Therefore, within the fault zone, strikes of the faults are widely dispersed (Fig. 2.34). The dip‐slip and lateral‐slip components vary greatly from block to block [Hubert‐Ferrari et al., 2002]. Contrary to the essentially right‐lateral strike‐slip fault character, some of individual faults of the NATF display left‐lateral offset [Hancock and Barka, 1981]. This may in places be a false impression due to the lower rates of motion along these faults with respect to the surrounding fault blocks. However, this remains a problem and needs to be studied in detail in a number of localities. 2. NATF gives its own imprints on the morphology [Herece and Akay, 2003] (Figs. 2.34 and 2.36a, b) and the local and regional drainage network (Fig. 2.36a, b), the rivers flowing from the highlands toward the Black Sea trend roughly perpendicular to the main fault valley. But where they meet the NATF zone, their courses are diverted locally to become subparallel to the trend of the fault (Fig. 2.36a). 3. Across the fault, a considerably wide (1–5 km wide depending on the rock types) mylonitic zone is observed (Fig. 2.35b). Nearer to the fault plane, rocks display an increasingly more intense form of mechanical crushing, and texture of the rocks varies from protomylonite to ultramylonite. 4. The fault plane is commonly steep. The dip angle varies from subvertical to 50°. 5. NATF zone cuts and displaces continuity of the hills and mountains abruptly. Plunges of the parallel hills are sharply truncated along the fault‐bounded depressions. 6. Within the fault zone, tectonic effects commonly overwhelm the erosional effects. The morphology gives away presence of the fault zone with the following ­features: all along its length, NATF displays a typical strike‐slip fault morphology, which has formed as a consequence of lateral and vertical displacements. The common morphological features are sag ponds or sag depressions (Fig. 2.32), shutter, linear, and pressure ridges (Figs. 2.33, 2.35a, and 2.36a and b), offset or deflected streams (Fig.  2.36a), captured or dammed streams, linear benches along valley walls, offset river or marine terraces (particularly along the southern side of Izmit Bay), aligned notches (along the northeastern side of the Marmara Sea). NATF is characterized by frequent (one in 25 yr as an average) 6 ≤ M ≤ 7.5 earthquakes (in the 1999 Marmara and the 1992 Erzincan earthquakes, M was 7.5 and 6.5, respectively), which are separated by quiet periods of about 100–150 yr intervals [Barka, 1996; Ambraseys and Jackson, 1998; Parsons, 2004]. The seismic record reveals that the earthquakes along NATF may be successfully interpreted in terms of a

Coulomb Failure Model [Parsons et  al., 2000], every earthquake concentrates the shear stress at the western tips of the broken segments leading to westward migration of large earthquakes [Stein et al., 1997; Stein, 1999]. Age and total offset of the NATF are still widely debated. On the total displacement, estimates vary widely from 25–45 km [Barka and Hancock, 1984; Barka and Gülen, 1988; Şaroğlu, 1988; Koçyiğit, 1989, 1990] to 85 km. An 85 km displacement was estimated on various grounds: based on the field data [Seymen, 1975], kinematic analyses (assuming the slip rate has been constant during the last 5 m yr period) by Armijo et al. [1999], on the morphological ground [Hubert‐ Ferrari et al., 2002]; using large rivers valleys (80 ± 15 km) and structural markers (Pontide suture, 85 ± 25 km; Tosya‐ Vezirköprü basins, 80 ± 10 km). Hubert‐Ferrari et al. [2009] showed also that a 4 Ma old volcano along the eastern part of the NAFZ was displaced nearly 50 km. Yılmaz et al. [1993b] measured the offset from the two regions: (1) in the Suşehri region, across the fault, where the Eocene nap front is displaced, the offset is about 50 km; (2) across the Erzincan basin, the major water divide (the crest line) was displaced about 45–50 km (Fig. 2.14). Slip rate along the NATF has been calculated from different segments of the fault zones, which is stated to be varying from 5–10 mm/yr [Barka, 1992] to 30–40 mm/yr [Taymaz et  al., 1991]. Based on the GPS measurements, McClasky et al. [2000] and Meade et al. [2002] estimated the rate as 15–20 mm/yr. Similarly, Tatar et al. [2012a] state that the slip rate of the NAFZ increases westward within about 400 km from 16.3 ± 2.3 mm/yr to 24.0 ± 2.9 mm/yr. Kozacı et  al. [2007] and Hubert‐Ferrari et  al. [2002] deduced slip rates based on geochronology methods using cosmogenic Cl and 14C dating, to be about 18 ± 5 mm/yr. They state also that the results are consistent with the NATF seismic deformation rate, which ranges between 10 and 30 mm/yr. [Jackson and McKenzie, 1988; Westaway, 1994]. A 22 ± 3 mm/yr rate has been monitored by the GPS measurement during the last 10 yr period [Straub and Kahle, 1997; Straub et al., 1997; McClusky et al., 2000]. Different views have been proposed on the time of initiation of NATF, varying from the Late‐Middle Miocene [Şengör et al., 2005; Şengör et al., 2014] to the extremely young ages. According to Şengör et al. [2005], the present northern branch as a localized fault zone reached the Marmara region about 200,000 ka ago [Şengör et al., 2014, and references therein]. A critically important question is when the fault zone was localized to form the present narrow and discreet fault belt (Fig. 2.36b). The fault‐controlled local depressions along NATF are known to have generated in a wide time frame starting possibly from the Lower‐Middle Miocene. Some of the sediment depositions of basins between Vezirköprü and Suşehri (Fig. 2.32) began in the Early‐Middle Miocene Period [Seymen, 1975; Yılmaz and Tüysüz, 1984; Yılmaz et  al., 1993b; Koçyiğit, 1989,

54  ACTIVE GLOBAL SEISMOLOGY

Figure 2.36  (a) An oblique aerial photo showing the two main branches of NATF: NBNATR = the northern branch of the North Anatolian transform fault and SBNATF = the southern branch of the North Anatolian transform fault in the Marmara region. Note the lateral offset and sudden changes of the course of the Sakarya River within the fault zone. Linear ridges are clearly identified around Sapanca Lake. See (b) located to the immediate west of (a). (b) Google photo showing the main fault zones of NATF (NBNATF; the straight red line) and the faults that formed prior to the localization of NATF in the present narrow belt. The older faults (broken red lines) are cut and displaced by the centrally located main fault in the Izmit Lake area. Note that the older faults make an acute angle with the main fault zone.

Morphotectonic Development of Anatolia and the Surrounding Regions  55

1990]. However, in many places, there is no direct link to associate the generations of these local depressions with the present NATF zone. Geology of these basins reveals that they had formed long before the development of NATF. Therefore, it may be that the basins in a later period were incorporated into the NATF zone, and thus they owe their origins to some other tectonic events. The Karlıova basin represents the eastern corner of the Anatolian Plate, where NATF and EATF intersect [Şengör and Kidd, 1979; Tutkun and Hancock, 1990; Şaroğlu and Yılmaz, 1991]. The oldest sediments that are confined to the Karlıova depression, although not precisely dated, are not older than the Pleistocene. The faults bordering the depression, which are coeval with sediments that were deposited in the depression, cut and displaced the Lower-Middle Pliocene sediments, which are not restricted to the boundaries of the Karlıova basin. They crop out in all of eastern Anatolia with similar facial characteristics. These data may be interpreted to mean that the westward escape of the Anatolian Plate bounded by the two transform faults is younger than the Middle Pliocene [Şaroğlu and Yılmaz, 1991]. The Karlıova basin sediments are represented primarily by linear alluvial fans, poorly lithified and poorly sorted conglomerates, derived from the fault‐induced elevations, which define the boundaries of the depression. They get progressively older going away from the point of convergence of EATF and NATF. The younger fans rest partly on the older ones. The conglomerates have not yet been precisely dated. Comparing them with the neighboring depressions, they may be estimated to be Plio‐Quaternary in age. 2.6.2. The East Anatolian Transform Fault The East Anatolian transform fault is a 700 km long, 1–30 km wide, northeast trending, left‐lateral strike‐slip fault zone. It defines the southeastern boundary of the Anatolian Plate, and thus represents a megashear zone between the different plates (Figs. 2.1 and 2.37; McKenzie [1976]). It stretches from the Karlıova depression in the north to the northern end of the Karasu depression near the Maraş region in the south (Fig. 2.37; Perinçek and Çemen [1990]; the Karlıova depression and the southwesterly extension of EATF in the Amanos Mountains are briefly summarized in the related sections). The morphotectonic features that are observed all along the EATF zone display very similar features to that of NATF, and are detailed in a recent account by Duman and Emre [2013]. The southern [H. Yılmaz et al., 2006] and the northern parts [Aksoy et al., 2007; Çelik, 2008; and references therein) of EATF are described also in considerable detail. Therefore, they will not be repeated here. As is the case for NATF, the EATF also forms a riftlike valley along most of its extension, and also wider local depressions (Figs. 2.37 and 2.38). The left step‐overs are common. In the Hazar Lake areas of the Elazığ region (HL in Fig.  2.37), EATF displays a different pattern,

Figure 2.37  The East Anatolian transform fault zone on the morphology map. EATF = East Anatolian transform fault; NATF = North Anatolian transform fault; BZSM = Bitlis‐Zagros suture mountains; FT‐ZI = frontal thrust and the zone of imbrication of the Bitlis suture mountains; FFTB = foreland fold and thrust belt; ERZD = Erzincan depression along NATF; KJ = Karlıova junction, where NATF and EATF converge; MTJ = Maraş triple junction; ASF = Asi graben fault; DSF = Dead Sea fault; ADTJ = Amik depression; KF= The Karasu and Karasu Fault; BDC = Bingöl Caldera. Black letters are the cities and towns: B = Bingöl; P = Palu; E = Elazığ; M = Malatya; GA = Gaziantep; HL = Hazar Lake; MO = Muş depression.

where it consists of four to five subparallel main faults within a 3–10 km wide belt [Fig. 2.38a]. Each one of the fault branches consists of several smaller faults. Collectively, they form locally developed, long and short depressions, and intervening parallel hills, indicating that the fault segments, in addition to the major strike‐slip offset, have considerable vertical slip displacement. The wide depressions do not resemble classical pull‐apart basins or a rhomb graben. They commonly display more complex structural patterns representing possibly superimposed structural features that formed during the long history of developments. In places, a combined role of different structures such as a negative flower and a pull‐ apart basin were described [Lyberis et al., 1992]. The structural pattern of the Hazar Lake and Gölbaşı Lake areas, on the other hand, may be compared to two subparallel rhombohedric (lensoidal) depressions, intervened by a horst that formed possibly as a pressure ridge (Fig. 2.38). In places, where the EATF cuts across an ophiolitic assemblage (e.g., a large serpentinite body), its sharp linear appearance commonly disappears, and the fault zone is dispersed in the rocks to form an extremely wide shear zone consisting of numerous small faults, joints, and cracks as is the case in the north of Pazarcık, and also in the area, where the EATF cuts the southeast Anatolian suture zone (Fig. 2.37 and 2.38 inset). The rocks turn to an intensely mylonitized sheared serpentinite.

56  ACTIVE GLOBAL SEISMOLOGY

Figure 2.38  Morphotectonic map of southern part of EATF. The lakes that align along EATF are the sag ponds. The pull‐apart nature of the lake depressions is revealed by the rhombohedric geometry, defined by the boundary faults. The yellow lines are subsidiary faults with respect to the main branch of EATF. Note that some of the faults are cut and displaced by the main fault. Çc = Çağlayan Cerit Town; OF = Ophiolitic rocks on the fault zone, where the traces of the fault as a single, continuous line disappears. KSG = Karasu graben, DSF = Dead Sea fault,The settlements: TK = Türkoğlu, KM; KahramanMaraş, Pz = Pazarcık, GB = Gölbaşı, DŞ = Doğanşehir, Ad = Adıyaman, Mt = Malatya, D = Doğanyol, CL = Çelikhan, EL = Elazığ, SV = Sivrice, G = Gölcük-Sivrice Lake, P = Palu. White ­rectangle shows location of the inset map.

Aksoy et al. [2007] estimated 4 mm lateral offset along the Sivrice faults (one of the main components of EATF in the northern region). McClusky et  al. [2000] and Reilinger et  al. [2006], on the other hand, based on ­geodetic data estimated about 9 and 9.9 mm a‐1 slip rates in the northern part of the fault, which decreases to 7 mm in the southern sector, and about 4.5–5 mm near DSF. 2.7. THE MARMARA REGION The Marmara region refers to northwestern Turkey. It is situated among the Aegean, Thrace, Black Sea, and the western Anatolian regions (Figs. 2.1 and 2.39). Therefore, it is located in a tectonically complex area, where the north‐south extensional regime of the Aegean‐western Anatolian region and the northern plate boundary of Anatolia as represented by NATF intersect (Fig.  2.1). Such complexity forms a variety of structural elements, which is reflected in the morphology (Fig.  2.39). In the middle of the region is located the Marmara Sea basin, which separates two morphologically different lands, a rather smooth topography in the north and a mountainous terrain in the south. In the following paragraphs, an attempt is made to use the morphological features as clues to decipher successive stages of the morphotectonic development of the region.

Major morphotectonic components of the Marmara region, which are documented in detail by Yılmaz et al. [2010], may be classified as follows (Fig. 2.40): 1. The North Anatolian transform fault zone 2. The Thrace‐Kocaeli peneplain (and the Istanbul plateau) 3. The Marmara Sea basin 4. The Bursa-Balıkesir plateau With respect to the Marmara Sea basin, in the north is located the Thrace‐Kocaeli peneplain. It stretches from the Karasu depression of the Bolu Province in the east (WBFP in Fig. 2.33; Yiğitbaş et al. [2004]) to the Trace region in the west (Fig. 2.39). The most prominent active tectonic entity of the region is NATF. It cuts the whole length of the Marmara region and stretches to the Aegean Sea (Figs. 2.1, 2.39, 2.40). As an active tectonic element, it has significant morphological control in the region [Armijo et  al., 2005; Gasperini et  al., 2012; Şengör et  al., 2014]. NATF bifurcates into two strands in the Marmara region as the southern and the northern branches (Figs.  2.33 and 2.40). Between them rises the Armutlu High (the Armutlu Peninsula; Figs. 2.33, 2.39, 2.40). The northern branch follows a trend along the Sapanca Lake (SL in Figs. 2.33 and 2.39), and is buried under the ̇ waters of the Marmara Sea in the Izmit Bay (Figs. 2.39

Figure 2.39  Morphotectonic map of the Marmara region. Light green lines are the ridges that separate the subbasins (white lines with double arrows indicate long axes of the subbasins) within the Marmara Sea basin. Straight white lines with red glows are crest lines that correspond commonly to the horst axes, which separate the intervening grabens (the river and stream valleys). The thick black lines refer to the different trends of the structures in the western and eastern sides of the southern Marmara region. The thick white broken line separating the two regions of the opposite trend is the region where the wrench fault zone along which the eastern block moves southwesterly direction, partly above the western block. The concave yellow broken lines are normal faults. The red thick arrows indicate the right lateral shear that generated the subbasins. GP = Galipoli Peninsula; DS = Dardanelle Strait; SM = Samanlıdağ Mountain; UD = Uludağ Mountain; KD = Kazdağ Mountain; ML = Manyas Lake; UL = Ulubat Lake; IL = Iznik Lake; BP = Bursa plain; IP: Inegöl plain; YP = Yenişehir plain; NBNATF Northern branch of the North Anatolian transform fault; SBFZ = The southern boundary Fault of The Black Sea Basin; Ist H = Istanbul horst; Ist = City of Istanbul. The white broken square refers to the region shown in Figure 2.47a.

Figure 2.40  Morphotectonic subdivisions of the Marmara region (inspired by Yılmaz [2010] and Emre et al. [1998]). NATF = North Anatolian transform fault; NB and SB correspond to the northern and southern branches of NATF in the Marmara region, respectively.

58  ACTIVE GLOBAL SEISMOLOGY

Figure 2.41  (a) Morphotectonic and bathymetric map of the Marmara Sea and the surrounding region. The white line is the topographic profile across a north‐south section. White broken line represents the erosional marker as a reference surface before its fragmentation. The numbers indicate the fragments of the flat‐lying surface standing at different heights, below (−) and above the sea level. Pale green lines are major faults striking obliquely to the main axis of the Istanbul horse (IST H). Red arrows indicate significant vertical displacements by the major faults, which elevated the Uludağ horst and descended the Izmit Bay trough. The white circle shows approximate locations of the photos displayed in Figure 2.42. BP = Bursa plain; YP = Yenişehir plain; Armutlu P = Armutlu Peninsula; B = Bosporus. Straight red lines in the Sea of Marmara display approximate location of NATF. (b) Schematic north‐ south cross section from Uludağ Mountain in the south to Thrace‐Kocaeli peneplain in the north, along the white broken line, showing structures of the Marmara region that formed after the Pliocene. The region is extending north‐south (black arrows). As a consequence of this, the Uludağ has risen as a core complex (purple arrow). The low angle listric normal faults are flattened downward and possibly adjoin a master fault. A large lake of the Late Miocene–Early Pliocene age was fragmented, and some local depressions (BP = Bursa plain; YP = Yenişehir plain; ̇ IL = Iznik Lake) were formed, and they were separated by younger horsts; the Samanlı Dağ Range (SD) and Armutlu High (AP). The blue thick lines represent the lake deposits, which were separated from one another by the east‐west trending mountains, which are presently trapped within the local basins.

and 2.41). NATF extends along the northern edge of the Marmara Sea basin forming a steep slope in front of the shelf (Fig. 2.39). Within the Marmara Sea, continuation of NATF may be traced in the bathymetric and seismic profiles [Gökaşan et al., 2003; Yılmaz et al., 2010; Şengör et al., 2014; and references therein]. Farther west, NATF cuts through the Marmara Sea basin extending to the Aegean Sea region along the Saroz Bay structural depression (Figs. 2.39 and 2.40). Before entering into the Aegean Sea, it cuts across Ganos Mountain of the Gelibolu (Galipoli) Peninsula (Figs. 2.39 and 2.40). NATF reaches to the northern Aegean shear zone in the Aegean Sea

(Fig. 2.1; Şengör et al. [2005]). From thereon, it changes the trend to southwestward, and defines the northeasterly trending steep boundary of the northern Aegean basin. In this area it is known as the North Aegean fault zone [Papanikolaou et al., 2002, 2006]. The southern branch extends along the Mudurnu valley (Fig. 2.33; Yılmaz et al. [1981, 1982]) and the southern edge of the Iznik Lake (IL in Fig. 2.39). Farther west it stretches to the Gulf of Gemlik (Fig. 2.33 and Fig. 2.39). With respect to NATF, the areas located to the north and the south display different morphological characters (for the role of the southern branch of NATF on the

Morphotectonic Development of Anatolia and the Surrounding Regions  59

morphological development of the southern shore and the adjacent land regions, the reader is referred to Yılmaz et al. [2010]). The Marmara Sea is an oval‐shaped, small (11,350 km2) inland sea (Figs. 2.1 and 2.39). There is an unexpectedly deep trough reaching to 1370 m depth in the central part of the sea. The shelves of the Marmara Sea are about −100 m deep, and pass to gently and steeply dipping slopes, which display two different morphological styles as the straight and concave slopes (Figs.  2.39 and 2.41). The steep slopes commonly display straight map pattern. Most of the northeastern slopes belong to this category, where the shelf is truncated abruptly by NATF, and descends steeply down to −1000 (Fig.  2.39; Görür and Elbek [2013]). Most of the southern slopes and part of the northwestern slopes belong to the concave and low‐angle slope category. The concave slopes are also common on the land areas (Figs. 2.39 and 2.41). They correspond commonly to the traces of listric normal faults, which determined the initial form of the Marmara Sea basin, because the early sediments of the sea were controlled by these faults. They were developed during the Late Miocene–Early Pliocene period (Fig. 2.41b). The sea sediments of that period (the Para‐Tethyan Sea sediments: Sakınç et  al. [1999]; Şengör et al. [2014]) laterally transit to the coeval lake deposits on the adjacent land areas [Görür et  al., 1997; Elmas and Meriç, 1998; Sakınç et al., 1999]. Using multichannel reflection profiles and high‐resolution bathymetric and chirp profiles, Şengör et  al. [2014] recently identified the faults in the Marmara Sea basin, which cut the surface and older faults, and estimated a 55 km offset along NATF zone. A topographic section across the Marmara Sea shows clearly that a flat‐lying erosional surface is abruptly truncated and displaced by the normal faults displaying concave map pattern (Fig.  2.41). As a result of this, the remnants of the flat‐lying surface lie presently at different depths and heights around the shelf areas, varying from −200 to 200 m (Fig. 2.41a). The eastern part of the Marmara Sea is commonly delimited by linear coastlines, which correspond to the strike‐slip faults striking northeast and northwest (Fig. 2.39). They collectively define lozenge‐shaped depressions around Izmit Bay [Gökaşan et al., 2003; Cormier et al., 2006; Dolu et al., 2007], and some subbasins within the central Marmara Sea region (Fig. 2.39). As reflected clearly in the bathymetry (Figs. 2.39, 2.40, and 2.41), there are also long, deep (>1100 m), and narrow troughs in the central part of the Marmara Sea, formed within the broad zone of NATF (for detailed structural as well as bathymetrical data and the analyses associated with NATF in the Marmara Sea, the reader is referred to Şengör et al. [2014]. The linear strike‐slip fault morphology is presently more prominent than the curvilinear normal faults on the land as well as the subsea areas. The former faults cut and

offset the latter faults and the associated older and smoother topography (Figs.  2.42 and 2.43). Under the strong influence of the strike‐slip faults, the young morphology exhibits a variety of features formed in relation to the faults. Among these, offset streams and offset crest lines are readily identified (Fig. 2.43). The Marmara Sea is delimited along both ends of its long axes by two linear mountain ranges, the Armutlu and Ganos highs (Fig.  2.40). They are situated within NATF zone, and bounded by the branches of NATF. Therefore, the highs were regarded as the coeval morphotectonic entities with NATF. The Armutlu High was elevated as a pressure ridge under the compressive forces generated by the reverse‐slip components of the two branches of NATF (Figs.  2.40 and 2.44a). The reverse‐ slip component of the strike‐slip faults is seen clearly on the fault planes (Fig. 2.45a, b). Some of the oblique faults of the same character played an active role on the recent uplift of the Hersek restraining bend [Özaksay et  al., 2010] and its step‐over [Forte et  al., 2014] (HR in Fig. 2.43). The uplift rate is estimated to be about 3.5 m/ kyr as an average for the last 4250 yr period, along the bend of the northern branch of the NATF near Karamürsel [Özaksoy et al., 2010]. A significant vertical displacement of about 800 to ̇ 1000 m across NATF in Izmit Bay and Armutlu Mountain may be estimated from the displacements of the Upper Miocene–Lower Pliocene lacustrine sediments and the overlying flat‐lying erosional surface (Fig.  2.44a, b; the fault is therefore postdating the Pliocene). A similar amount of a vertical displacement is seen from the bathymetry across NATF, where it cuts the northern shelf of the Marmara Sea and forms the steep slope (the fault plane; Figs. 2.39, 2.41a, and 2.42). The Armutlu Peninsula as a horst block is gently back‐tilted to the south toward the southern branch of the North Anatolian transform fault (Fig. 2.44a, b). The northern slope of the Armutlu High is thus steeper. The Marmara Sea is a deep trough. The 3D basement tomography of the Marmara trough [Bayrakçı et  al., 2013] displays four subbasins (their depths vary from −800 to −1200 m; Fig.  2.39) separated by diagonally northeast‐trending ridges, which are about −600 deep. The seismic data reveal also that the basin fill of the Pleistocene and Quaternary ages was formed under the close tectonic control of the oblique faults [Beck et  al., 2007; Becel et al., 2010]. The Ganos High is viewed as a restraining bend of the Anatolian Plate, which formed in response to its southwesterly rotation [Şengör, 1979c; Şengör et al., 1985]. The GPS vectors (Fig. 2.2) are subparallel to the trend of the Galipoli Peninsula (compare Fig. 2.2 with Figs. 2.39 and 2.40) implying that the extent and elongation of this morphotectonic entity was developed in close harmony with the motion of the Anatolian Plate.

60  ACTIVE GLOBAL SEISMOLOGY

Figure 2.42  Photos showing faults on the northern flanks of Uludağ Mountain, Bursa Province of the Marmara region. (a) A panoramic view of the northern slope of Uludağ Mountain. A flat lying surface at the top is the elevated remnant erosional surface. The low‐angle slopes at the higher level, which collectively define a planar surface, correspond to the low‐angle normal fault (the detachment fault) that elevated the mountain as a core complex. The red broken lines on the high‐angle slopes corresponds to younger, steeply dipping normal faults that cut the low‐angle fault. (b) An easterly view of the fault plains of the steep faults from the city of Bursa displaying well developed flatirons. (c) A view of the northern flank of Uludağ Mountain from the Bursa Province shows flatirons of the steeply dipping normal faults. (d) A close up view of the fault plain of the low‐angle normal fault from the Hünkar Köşkü area, south of the city of Bursa. Most of the fault plain structures may be observed on the fault plain; the slicken lines as shown by the red arrows define top to the north sense of shear.

The Thrace‐Kocaeli peneplain is a low and flat land, above which dendritic drainage was developed (Figs. 2.33 and 2.41). The youngest rock unit lying below the peneplain is the Upper Miocene–Lower Pliocene lacustrine sediments [Sakınç et al., 1999; Emre et al., 1998; Yılmaz et al., 2010; and references therein]. The peneplain may be correlated in time and space with the flat‐lying erosional surface that is recognized throughout the western Anatolian region. In fact, the erosional marker surface of the western Anatolia may be traced northward to the surrounding land areas of the Marmara Sea (see Figs. 2 and 3 in Yılmaz et al., 2010]. Therefore, the two surfaces may

be regarded as the identical stratigraphic markers, remaining as one continuous erosional plane until the opening of the Marmara Sea basin. The seismic studies conducted in the Marmara region also reveal an erosional stratigraphic marker underneath the Marmara Sea and the Black Sea shelves (Figs. 2.41 and 2.46; Demirbağ et al., [1999]; Dolu et  al. [2007]). Above the peneplain have been deposited Plio‐Quaternary sediments [Yılmaz, 2007; Yılmaz et al., 2010]. The Istanbul plateau represents the central part of the Thrace–Kocaeli peneplain, and is bounded by the two right‐lateral strike‐slip fault zones in the north and the

Morphotectonic Development of Anatolia and the Surrounding Regions  61

Figure 2.43  Morphotectonic map of Izmit Bay and surrounding areas. The broken black lines at the horizon show traces of the fragmented and elevated erosional surface on the east‐west trending hills that were developed during the rejuvenated north‐south extension in the Late Miocene–Early Pliocene period. The yellow curvilinear lines represent major normal faults that formed during the same stage of the north‐south extension. These structures and the associated morphological features were later displaced by the younger oblique faults with dominant strike‐slip components that strike northeast and nothwest (red lines), and together they form a conjugated map pattern. A good example may be given from the Yalova area, where the strike‐slip faults (the fault zone that displays a typical anastomosing pattern) displaced the east‐west trending hills left laterally. The red broken line within Izmit Bay show approximate locations of the major faults of NATF [after Gökaşan et al., 2003]. HR = western side of the Hersek restraining bend (or step‐over, Özaksoy et al. [2010], Forte et al. [2014]).

south (Figs.  2.39, and 2.40; Yılmaz et  al. [2010]). The northern fault runs along the coastline of the Black Sea and therefore it is also known as the southern boundary fault of the Black Sea basin (SBF in Figs. 2.39 and 2.40; Yılmaz [2007]; Yılmaz et al. [2010]). It consists of a number of faults, which display an en echelon pattern (Figs.  2.33, 2.40, and 2.47a). Collectively, they define the present straight coastline. The block, bounded by the two fault zones, has been displaced obliquely, elevated as a horst (Fig. 2.46a), and also rotated in an anticlockwise sense under a dextral shear regime, which is generated between the two major right‐lateral strike‐slip faults (Figs. 2.46a and 2.47a; Yılmaz [2007]; Yılmaz et al. [2010]). The right‐lateral offset along SBF is supported by the multibeam bathymetry data, which displays that the Bosporus Channel in the Black Sea shelf has been right laterally displaced across the fault [Lericolais et al., 2002]. Therefore, it follows a sinusoidal path before reaching the canyon head. SBF is not seismically very active. It resists the motion of the region along the NATF. This generates a dextral shear, which forces the horst block to rotate in a counterclockwise sense. As a consequence, a northwest and northeast trending conjugated pair of structures were

formed, and they have played a major role in the development of the present morphology of the region (Fig. 2.47a). The stream valleys and the hills align along the major structures (Fig. 2.47a). As a result, the streams follow a long and diagonal path to reach the surrounding seas. The Bosporus Channel was opened using the conjugated pair of structures (Fig. 2.47a). This view is supported by seismic data, which display that the structures that are seen on the lands extend to the Bosporus and control also the bathymetry [Yılmaz and Sakınç, 1990; Gökaşan et al., 2006; see Fig. 13 in Yılmaz et al., 2010]. According to the following morphological data, uplifting of the Istanbul horst occurred rather recently. (1) Four sets of elevated terraces are readily identified at different heights as hanging terraces on both sides of the Bosporus [Erinç, 1977; Özşahin and Ekinci, 2014]. (2) The stream valleys on the both flanks of the Bosporus are in incipient stage (Fig.  2.47a and b), and the headword erosion has not yet reached to the surrounding flatlands lying at a height of only about 200–300 m to drain the rivers to the Bosporus. Instead the main rivers on the lands that surround the Bosporus are drained either

62  ACTIVE GLOBAL SEISMOLOGY

Figure 2.44  Morphotectonic map of the Armutlu high and surroundings. NBNATF and SBNATF are the northern and southern strands of NATF. Between the two faults, the Armutlu Peninsula thrust out and rose as a pressure ridge. The yellow arrows display southerly tilted erosional surfaces. (b) Photo showing major structural and morphological elements of the area shown above. The red square refers to the Thrace‐Kocaeli peneplain. The yellow square refers to the steeply dipping fault. The red broken arrows show the southerly tilted, planar erosional surfaces.

to the Black Sea or to the Marmara Sea (Fig. 2.47a). (3) Along the flanks of the Bosporus, three to four different slope angles are recognized from the top down to the sea level. They display angular connection with one another (Fig.  2.47b). This suggests that the time elapsed since their development is yet too short for erosional forces to obliterate their angular slopes. Collectively, the morphological data indicate that the plateau has been uplifted recently and probably episodically.

The fault‐bound shelf along the Black Sea was descended by SBF for more than 500 m (Fig. 2.46a). The shelf was above the sea level during the last Glacial‐Holocene period [Ryann, 2007], and the sediments are strongly affected by the sea level changes, which, in turn, were controlled by climatic changes [Lericolais et al., 2010]. The land to the south of the Marmara Sea is the Balıkesir plateau (or the Bursa‐ Balıkesir plateau), which rises steadily from about 300 to 800 m from the west to

Morphotectonic Development of Anatolia and the Surrounding Regions  63

Figure 2.45  Photographs showing oblique faults with significant strike‐slip and reverse‐slip components (yellow and red arrows) in rock quarries near the Karamürsel area, south of Izmit Bay. Note the anastomosing pattern of the small faults within the fault zone. As a result of severe semiductile deformation, the rocks along the fault planes display well‐developed lineations (which reveal also the oblique character of the fault) and poor fissility.

the east. The Balikesir plateau displays different morphological features, compared to the Istanbul plateau. It  is represented by subparallel hills and depressions (Fig.  2.39). They commonly correspond to horsts and grabens, which trend northwest and northeast in the eastern and western parts of the plateau, respectively (Fig. 2.39). The horsts have commonly flattop representing the remnants of the flat‐lying erosional surface elevated above the horsts (Fig.  2.43). The structures of different trends meet in the central area, where the dividing line between the two sectors corresponds to a major structural axis, a hinge line of a wrench fault, which appears to have accommodated anticlockwise rotation of northwestern Anatolia as a tectonic bend (Fig. 2.39). With respect to the wrench fault, the eastern block thrust westerly over the western block [Schindler, 1997; Mueller et al., 2007] to accommodate the counterclockwise rotation of the Anatolian Plate (Fig.  2.2), which induced an east‐west shortening deformation [Kaya et al., 2004]. In the Bursa‐Balıkesir plateau, there are two significant peaks. These are the Uludağ and Kazdağ mountains (horsts) (KD and UD in Figs.  2.33 and 2.39). Both of them rise steeply from the surrounding lowlands to over 2000 m heights (Figs. 2.39, 2.41a, and 2.42). Their western flanks are relatively smooth where the Upper Miocene–Lower Pliocene lacustrine limestone is observed

to have been elevated on the shoulder of the horsts to over 1000 m. This observation alone indicates that the final elevation of the mountains is a young event, occurring after the Early Pliocene. In accordance with the field observation, isotope data obtained from the morphotectonic elements support this and reveal that it occurred in a 0.5 and 1.3 my period and most probably around 0.8 my ago [Demoulin et al., 2013]. The structural data indicate further that both of the mountains have been uplifted as a core complex [Okay and Satır, 2000; Yaltırak and Ceyhan, 2011]. A set of subparallel normal faults defines the northern flanks of the mountains (Fig. 2.33). They are flattened downward possibly to connect with a major detachment fault (Fig. 2.41b). The curvilinear normal faults, formed in association with a regional north‐south extension, displaced the flat‐ lying erosional marker surface. As a result of this, the fragments butt against the faults that displace them (Figs.  2.41b and 2.43). The sediments deposited within the depressions reveal that following the fragmentation of the interconnected larger lakes (see Chapter  8), isolated small lakes remained during the Pleistocene time [Emre et al., 1997; Elmas and Meriç, 1998; Kazancı et al., 2004; Efe et al., 2011] (Fig. 2.41). Some of the lakes dried up later, and have lost their lake character [Mater et  al., 2003; Kazancı et al., 2004]. But some others continued to

64  ACTIVE GLOBAL SEISMOLOGY

Figure 2.46  (a) Schematic block diagram showing the Istanbul horst bounded by the two major strike‐slip fault zones with minor dip‐slip component [modified from Yılmaz et al., 2010, Fig. 9]. The fault A corresponds to the North Anatolian transform fault (NATF). SBF is the southern boundary fault zone of the Black Sea basin (fault B). The blue line refers to the present sea level. The brown vertical arrow displays about 500 m dip‐slip displacement of the Thrace‐Kocaeli peneplain across the fault. In the Black Sea shelf area, the flat‐lying surface is identified from seismic data [Demirbağ et al., 1999], beneath 200 m thick unconsolidated young sediment piles. The peneplain stands at the height of about 300 m on the land. (b) Seismic section showing location (red circle) of NATF in the Izmit Bay area [after Gökaşan et  al., 2003]. Along the fault a narrow trough was developed. Note the flower structure that formed within the fault zone. Most of the faults identified on the both sides of the major fault correspond commonly to the faults, formed before NATF was localized to the present narrow zone (depression).

exist to the present (see UL, ML, and IL in Figs. 2.39 and 2.41a). The terraces, which surround the lakes, show clearly that the lake level dropped since then [Kazancı et al., 2004]. As a consequence of this, four to six hanging lake terraces are recognized at different heights up to 100 ̇ m from the present lake level in the Iznik Lake area. A number of different views have been proposed for the development of the Marmara basin, mostly based on the seismic and bathymetric data [Ergün and Özel, 1995; Wong et al., 1995; Okay et al., 1999; Aksu et al., 2000; LePichon et al., 2001; Şengör et al., 2005]. The views are discussed and compared at length by Aksu et al. [2000] and Şengör et  al. [2005]. In addition to the seismic and bathymetry data, extensive field data that we obtained from the adjacent land areas have led us to envisage the following history of evolution. A regionwide flat‐lying erosional surface was established on the entire western part of

Turkey, including the Marmara region, to the end of Late Miocene [Yılmaz et  al., 2000; Yılmaz et  al., 2010]. Following the development of the erosional surface, the north‐south extension was rejuvenated in the Aegean‐ western Anatolian region [Yılmaz et  al., 2000], and east‐west–trending horst and graben structures began to form after the Early Pliocene. The approximately east‐ west striking curvilinear listric normal faults of the Marmara region were developed during this period. They are common on both sides of the Marmara Sea on the land as well as in the sea (Figs.  2.33, 2.39, and 2.41). These young faults fragmented the smooth lowland, and Plio‐Quaternary conglomerates were deposited in the newly developed east‐west trending long and narrow depressions (Fig.  2.41b). The normal faults that are mainly responsible from the curvilinear topography are not seismically very active today. Their morphological

Morphotectonic Development of Anatolia and the Surrounding Regions  65

Figure  2.47  (a) Morphotectonic map of the Istanbul region. Red lines are faults. The yellow lines display faults and/or joint dominating regions, which have formed zones of weaknesses and thus controlled development of the oblique‐trending streams. These northwest‐trending structures with respect to the Marmara and Black Sea coastlines were generated under the dextral shear regime, which forced the Istanbul horst to rotate between the two major boundary fault zones in a counterclockwise sense (red curvilinear arrows). Along the zone of weaknesses of the conjugated set of structures, the Bosporus (B) was opened as a zigzagging sea strait (white zigzagging line). Ist = City of Istanbul; PI = Princes Islands, ÇL = Çekmece Lagoon. (b) A northerly view of Bosporus. The peneplain is clearly identified as the broken red line at the horizon). The broken white lines display the low, moderate, and steep slopes, which show an angular connection with one another.

distinction was considerably obliterated (Fig.  2.43). The concave faults, which defined a centrally located region of subsidence, extended far beyond the present boundary of the Marmara Sea. The basin began to be localized to the present size in later periods. This is indicated by the sedimentological data, which show progressively shrinking boundaries of a shallow sea basin [Elmas and Meriç, 1998; Meriç et al., 2009]. The structural and stratigraphic data

reveal also that, while the center of the Marmara basin was subsiding, the surrounding lands were gradually elevated to form the bordering plateaus (Fig.  2.41b). This phase lasted until NATF entered the Marmara region. NATF was initiated in East Anatolia, and gradually penetrated into northwestern Anatolia [Şengör et  al., 2005]. Before the formation of the present narrow fault zone, a conjugated set of northwest and northeast striking oblique slip faults with dominant strike‐slip component began to develop under a regional dextral shear (Fig. 2.43; Şengör et  al. [2005]; Yılmaz et  al. [2010]). The oblique faults are observed in much wider areas than the present limits of NATF (Figs. 2.43 and 2.46b). The conjugated set of oblique faults collectively defines a number of small lozenge‐shaped depressions within the Marmara basin (Fig.  2.39) and the surrounding land areas (Figs.  2.33 and 2.43). Long axes of the four lozenge‐ shaped subbasins within the Marmara Sea basin and the intervening ridges vary from the east‐west trend of the Marmara Sea. They fan away from east to west (Fig. 2.39). Trends of these structures are also compatible with a dextral shear (Fig. 2.39). The bathymetric and seismic data together with the distribution of the earthquake epicenters reveal that the active tectonics of the Marmara Sea are mainly controlled by the narrow NATF zone [Ş engör et al., 2014]. Other strike‐slip faults, which define some of the boundaries of the Marmara Sea and the subbasins, display little seismic activities compared to the NATF. The seismic ̇ data [Imren et al., 2001; Şengör et al., 2014] and the field data [Yılmaz et  al., 2010] show further that NATF cuts and offsets all the other faults (Figs. 2.43 and 2.46b) and is therefore is a younger fault zone. The morphological and structural data are in close agreement with the model elaborated by Şengör et  al. [2005, 2014] for the tectonic evolution of the NATF zone in the Marmara region. According to this model, an intracontinental transform fault begins as a broad shear zone and becomes gradually converted into a single strand structure. This may also be summarized as the gradual localization of the shear zone through time. Accordingly, NATF entered the Marmara region as a wide shear zone and affected larger areas during the earlier stage. The width of the shear zone progressively narrowed in time, and was finally confined to a continuous lineament as is seen today along a linear trough (the narrow zone of NATF is represented as the negative flower structure in the seismic section; Fig. 2.46b). During the late stage, the faults of the narrow zone cut and displaced the older faults (Fig. 2.46b). Since it became a narrow fault zone, it defined the present plate boundary. This is a critical time also for the morphotectonic evolution of the Marmara region. From this time onward, the northern and southern regions with respect to

66  ACTIVE GLOBAL SEISMOLOGY

NATF began to behave semi‐independently. The southern region began to elevate with a higher rate and has formed the present Bursa‐Balıkesir plateau. The northern region has remained a lowland and retained most of its previous morphological features including the erosional surface, known presently as the Thrace‐Kocaeli peneplain. Penetration of NATF as a wide shear zone into the Marmara region has not yet been precisely dated. It occurred possibly during Pleistocene time [Forte et  al., 2014], because the oblique faults of this stage cut and offset the lake deposits and thus are younger than the Late Miocene–Early Pliocene. Partly consolidated coarse conglomerates and interfingering sandstones that were deposited in the young depressions are Pleistocene and Quaternary in age [Rückert‐Ülkümen et al., 2006]. Based ̇ on the kinematic data, Imren et  al. [2001] and Şengör et  al. [2011, 2014] suggest a 200 ka for the entrance of

NATF as a narrow fault zone into the Marmara Sea basin. But core data indicate that a steady‐state deformation is continuing in the NATF zone during at least the last 500 ka period [Sorlien et al., 2012]. 2.8. WESTERN ANATOLIA Western Anatolia is the eastern part of the Aegean Extensional Province. Major characteristics and problems of the region are the subjects of another chapter in this book (see Chapter 6). Therefore, mainly its morphotectonic features are discussed here. Western Anatolia is a composite tectonic entity consisting of the Sakarya continent, the Tauride‐Anatolide platform (the northern part of this platform is represented by the Menderes Massif) and the Izmir‐Ankara suture, which separates the two continental fragments (Fig. 2.48; Şengör and Yılmaz [1981]).

Figure 2.48  Geology map showing major tectonic components of western Anatolia [modified from Yılmaz, 1988]. A, C, D, Iz, and M: Cities of Aydın, Çanakkale, Denizli, Izmir, Manisa, and Muğla. GG = Gördes graben; DG = Demirci graben; SG = Selendi graben; UG = Uşak‐Ulubey graben; BEG = Bergama graben; GDG = Gediz graben; SoG = Soma graben; SG = Söke graben; ÖG: Ören graben; YG: Yatağan graben; KT = Kale‐Tavas basin; BH = Bozdağ horst; DP = Dilek Peninsula; SM = Sipil Mountain; Ln and LNF = Lycian nappes on the Taurus Mountains

Morphotectonic Development of Anatolia and the Surrounding Regions  67

The tectonic evolution of the western Anatolian region is discussed at length in a number of studies (see Chapter 6, this volume; see also Şengör and Yılmaz, 1981; Yılmaz, 1981a; Yılmaz, 1990; Yılmaz et  al., 1995; Okay, 1996; Ersoy et al., 2014]. These amalgamated units formed a base, above which sediments were deposited starting from the Early Miocene to the present and formed three tectostratigraphic rock units as the lower, middle, and upper units (Fig. 2.49). Approximately east‐west trending, about 10 grabens dominate the morphology in western Anatolia. They are intervened by thin and long horsts formed under the ongoing extension (Fig.  2.50). Normal fault zones running subparallel to the graben axes border the grabens. Along the faults, seismic activity is intense (Fig. 2.50). Motions on the faults confirm the extension, which affects the region from Bulgaria in the north to the Hellenic arc in the south [McKenzie, 1972; McKenzie and Yılmaz, 1991]. The seismic tomography images display a southward roll back in the northerly subducting eastern Mediterranean oceanic lithosphere [Meulenkamp et  al., 1988; Spakman et al., 1988; Facenna et al., 2006]. This is regarded to have generated a north‐south extensional regime in the upper plate. The present morphology of western Anatolia formed simultaneously with the development of the middle and the upper units. Therefore, emphasis will be given to this period. The middle unit was formed during the Late ̇ Miocene–Early Pliocene period [Iztan and Yazman, 1990;

Figure 2.49  Generalized stratigraphic section of western Anatolia [modified from Iztan and Yazman, 1990; and Yılmaz et al., 2000].

Figure 2.50  Seismotectonic map of western Anatolia [modified from Dewey and Şengör, 1979]. Red arc = nonvolcanic outer arc; green arc with yellow glow = extinct volcanic arc; yellow arc with green glow = active volcanic arc. Brown arrow displays lateral motion along the Pliny and Strabo trenches. NASZ = North Aegean shear zone. Brown dotted areas are young graben depressions.

68  ACTIVE GLOBAL SEISMOLOGY

Yılmaz et al., 2000; Gürer et al., 2001; Sarıca, 2000; Yılmaz, 2002]. It is a widespread unit of western Anatolia cropping out in the entire region from Thrace in the north stretching to the Mediterranean Sea areas in the south. The middle unit is represented commonly by white limestone and alternating fine‐grained sediments such as siltstone, claystone, and fine‐grained sandstone, which were deposited in a low‐energy lake environment (for the references on the age and environment of deposition of the middle unit see Chapter 6, this volume). This reveals that interconnected lakes invaded Anatolia during the Late Miocene–Early Pliocene period [Benda, 1971; Benda et al., 1974; Steinenger et  al., 1987 and 1996; Görür et al., 1997; Benda and Meulenkamp, 1990; Alçiçek, 2010]. It also indicates that the present irregular topography, represented by abruptly terminated highs and lows (Fig. 2.50), had not yet formed in this period. An exception to this smooth topography was the Bozdag horst (also a dome) of central western Anatolia (BH in Figs. 2.48 and 2.51a), which represented a highland. The horst began to form as a core complex [Bozkurt and Park, 1994; Hetzel et al., 1998] between the two low‐angle normal faults (Fig. 2.51a, b), while the surrounding lakes still survived. A red, coarse conglomerate and sandstone unit was formed during that period around the dome as lateral fan deposits derived from the elevating Bozdağ horst [Iztan and Yazman, 1990; Yılmaz et al., 2000]. The low‐angle normal faults are evaluated as major breakaway (detachment) faults [Verge, 1995; Hetzel et al., 1995; Gessner et al., 2001; Catlos and Çemen, 2005], which cut the entire thickness of the lithosphere. The fault plane is exposed extensively on both sides of the Bozdağ horst (Fig. 2.51a, b). According to the isotope ages, exhumation of the horst took place between 8 and 5 m yr interval [Hetzel et al., 1995, 1998; Lips et al., 2001; Bozkurt, 2001a; Sözbilir et al., 2011; and references therein]. The paleontological age determinations (e.g., based on the gastropods; Emre [1998]) and the sporomorph associations [Ediger et  al., 1996] from the surrounding red unit [Iztan and Yazman, 1990; Emre et  al., 1998; Yılmaz et  al., 2000] (Fig. 2.51a–c) also support the Late Miocene age span. As a result of the constant supply of material derived from the elevating Bozdağ horst, the coarse conglomerate reached up to 750 m of thickness, and in many places separated from the metamorphic rocks of the Menderes Massif by the low‐angle normal fault (Fig.  2.51b, c). Along the fault plane, the red units are steeply (>45 degrees) back‐tilted (Fig.  2.51c), which reveals that the low‐angle fault has a listric fault character. The red conglomerate wraps round the Bozdağ horst (Fig.  2.51a). Away from the horst the conglomerate passes laterally to fine‐grained red sandstone and mudstone, and farther away to the white lacustrine limestone and marl, suggesting that the lakes surrounded the horst. According to paleontological data, the lakes survive during

the Early Pliocene period [Görür, 1988a; Yılmaz et al., 2000], and then disappeared before the Middle to Late Pliocene time [Hakyemez et al., 2013]. Partly contemporaneous with the lake phase, a severe phase of denudation generated a flat‐lying erosional surface above the western Anatolian Neogene succession, including the lacustrine sediments at the top of the sequence [Yılmaz et al., 2000]. This is a superimposed erosional surface, which probably began to develop when the Aegean region was elevated as a giant land mass during the Oligocene, and formed episodically until the early Pliocene. Remnants of the erosional surface crop out extensively in western Anatolia and therefore may be used as stratigraphic marker to distinguish the later geological events and younger major morphological features. The marker surface may be observed at different heights from sea level to the top of the horsts indicating that the present horsts and grabens postdate the phase of erosion. Abrupt vertical displacements of the flat‐lying erosional surface may be used as a geomorphological clue to identify younger (post‐ Pliocene) faults. The upper unit is generally confined to the present graben depressions, where major rivers flow and drain the western Anatolian lands to the Aegean Sea (Figs. 2.48, 2.50, and 2.52). Therefore, it is composed essentially of fluvial sediments and fan deposits. They consist of semilithified, coarse, buff‐colored, poorly sorted, alluvial conglomerate with minor sandstone, sourced from the hills that border the grabens. They are Quaternary in age, and rest unconformably on the lower and middle units (Fig. 2.49; Ünay et al. [1995]; Ünay and Göktay [1999]; Yılmaz et al. [2000]; Sarıca [2000]; Beccaletto and Steiner [2005]; Hakyemez et al. [2013]). Average thickness of the upper unit is about 250 m, varying between 100 m and 500 m. According to the seismic data (unpublished and published data of the Turkish Petroleum Corporation; Iztan and Yazman, 1990; Yazman et  al., 1998; Yılmaz et  al., 2000], thickness of the east‐west trending graben fill rarely exceeds 500 m, only where the lower and/or the middle units are trapped within the graben depressions. Presently, the fills of the present grabens are exposed on the graben shoulders, where they were elevated by the graben‐bounding faults, which form steep linear hills (Fig. 2.52a, b). The bedding planes and the ancient river terraces elevated on the faulted blocks are also back‐tilted due to the rotational movement along the listric normal faults (Figs.  2.52a–c). The back‐tilting increases to the horst axis. This reveals that the faulting gets progressively younger going away from the grabens. An explicit example may be seen in the southern flank of the Kazdağ horst (Figs. 2.53a, b; Yılmaz and Karacık [2001]). The east‐west trending horst and grabens structures of  western Anatolia occurred relatively recently, and the morphological data, which support this view, may be

Morphotectonic Development of Anatolia and the Surrounding Regions  69

Figure  2.51  (a) A panoramic oblique aerial photo from the northern flank of the Bozdağ horst (and dome). FP = North facing (black arrow), low‐angle, normal fault (the detachment fault); red line defines a well‐exposed fault plane. The brown shaded areas are the red conglomerates of Upper Miocene age that wrap round the dome. The white broken line at the top of the horst is the crest line, which displays a smoothened topography above the horst. A mature fault scarp is seen also along the northern flank of the horst (yellow lines curving from the top of the hill to a younger and steeper fault (black broken line). From the fault downward, the slope is steep (the steeply dipping yellow lines), and the smoothened topography disappears. A morphological unconformity is seen between the slopes of different dips. The green square refers to the location of the photo in (b). The yellow lines also refer to the direction of the seismic section that is shown in (d). (b) Photo showing westerly view of the perfectly planar, low‐angle normal fault in the northern flank of the Bozdağ horst. The red broken line defines the fault separating metamorphic rocks of the Menderes Massif from the red conglomerate unit of Upper Miocene age (the little hill on the right). The green square refers to location of (c). Note that no stream has yet developed above the fault plane, indicating that the fault plain has only recently been exposed to erosion. (c) A close up view of the low‐angle normal fault. Around the fault, a north‐facing, low‐angle brittle to semibrittle extensional shear zone (top‐to‐north sin‐tectonic fabric) is identified. The fault separates metamorphic rocks on the footwall from the red coarse clastic rocks on the hanging wall. Due to the back‐tilting, bedding planes of the red conglomerate butt against the fault plane (red line). Within the shear zone, lozenge‐shaped textures were developed in the metamorphic rocks. They were formed between the conjugated pairs of shear planes, which define also asymmetric boudins, shear foliation, and shear band foliation. The striated slip surfaces indicate top‐to‐north sense of the extension. (d) Seismic section across the flank of the horst showing the steeply dipping normal fault (red line) that cut and displaced the low‐angle normal fault (yellow lines) [modified from Çiftçi and Bozkurt, 2009].

listed as follows: (1) streams along the slopes of the horst blocks are yet in an incipient stage (Fig. 2.54a–d; Erinç [1954]; Paton [1992]); (2) the hanging valleys and champagne glass valley form are common features (Fig. 2.54a, d); (3) incision generates canyonlike valley profile near the graben floor but upward along the valley the profile

gets wider and shallower (Fig.  2.54c), indicating that headword erosion has not yet reached the plateau level; (4) on the plateau, stream and river valleys display wide, shallow and meandering profiles and deep carving has not formed yet (Fig.  2.54b); and (5) a direct, shortcut connection between the drainage network on the plateaus

70  ACTIVE GLOBAL SEISMOLOGY

Figure 2.52  (a) A photo from the southern edge of the Gediz graben showing the graben bounding listric normal faults (yellow arrows) and back‐tilted fault blocks (white arrows). (b) Photo showing southern edge of the Söke Graben near the ancient city of Myletos. Fault planes are bounded by white broken lines and marked by north westerly dipping arrows. (c) A set of graben boundary normal faults along the southern edge of the Gediz graben. The brown‐shaded surfaces are the fault planes. The white arrows indicate back‐tilting. The green lines show the smooth initial slop that was later cut and displaced by the younger faults, which led to the development of the present graben basin. The graben fill consisting mainly of a buff colored, semilithified conglomerate of Quaternary age, is exposed on the faulted blocks all along the graben shoulder. (d) Faulted erosional surface (black broken arrows) lying along the northern slope of the Bozdağ horst. A rollover anticline was formed as a result of the development of the listric normal fault (red line). The yellow lines are bedding planes of the Late Miocene continental red beds.

and the surrounding deep graben depressions is not thoroughly established [Erinç, 1954]. The western Anatolian grabens are commonly asymmetrical (Fig. 2.52a) (the reader is referred to Yılmaz et al. [2000] and Bozkurt [2001]). One margin of the grabens is structurally more active as revealed by more prominent morphological features. The active sides along the bigger grabens of the region as in the case of the Büyük Menderes and Gediz River grabens are the flanks that border the Bozdağ horst (Figs. 2.51a). The major faults, although observed along the whole length of the grabens, are segmented on a short length scale (Fig. 2.52c). Each segment is commonly shorter than 15 km. The graben bounding steeply dipping normal faults are younger than the low‐angle detachment normal fault. The former cuts and displaces the latter and forms steep slopes (Fig. 2.51a, d) [Koçyiğit et al., 1999a, b; Koçyiğit, 2005; Rojay et al., 2005; Yılmaz et al., 2000; Bozkurt and Mittwede, 2005; Hakyemez et al., 2013, Gürer et al 2016]. As a result of this, a morphological unconformity was developed along the flanks of the Bozdağ horst

(Figs. 2.51a, 2.52c, d), and this may be used as evidence supporting a cyclic extension [Bozkurt and Mittwede, 2005; Koçyiğit, 1999a, 2000; Koçyiğit, 2005; Rojay et al., 2005; and also Chapter 2 this volume for a review of the topic]. The seismic data show that the low‐angle fault flattens downward under the graben depression (Fig. 2.51d; Bozkurt and Çiftçi [2009–2010]). The low‐angle detachment fault separates rocks of different metamorphic grades. Using the metamorphic grade difference (about 3 to 4 Kb), Bozkurt and Oberhansli [2001] estimate approximately 9–12 km dip‐slip displacement across the fault. There is an ongoing controversy on the develop­ ment of the present grabens, whether the extensional regime continued since it began or survived as pulses. According to one view, the extension began during the Late Oligocene–Early Miocene and has continued uninterruptedly to the present, approximately during the last 25–30 m yr period [Seyitoğlu and Scott, 1991, 1992, 1996; Catlos et al., 2008, 2011; Black et al., 2013; and references therein]. According to another view,

Morphotectonic Development of Anatolia and the Surrounding Regions  71

Figure 2.53  (a) A photo from the southern flank of Kazdağ Mountain. The red lines are the step (normal) faults along which stairlike morphology was developed. Dip angles of the fault planes get progressively less steep approaching the horst axis. The white broken square shows the location of the photo in (b). (b) Steeply dipping normal (step) faults (red lines). The white arrows indicate the back‐tilting on the faulted blocks.

the grabens evolved episodically. They formed mainly in two major phases, which generated two distinct structural styles. The first phase of extension resulted in the exhumation of the high‐grade metamorphic rocks in the footwall of the low‐angle detachment fault(s). The second phase began much later, when the extension was rejuvenated, possibly during the Late Miocene–Early Pliocene, and still continuing today. The high‐angle normal faults and the present graben depressions were developed during late stages of this phase in the Quaternary [Koçyiğit et  al., 1999a, 2000a; Bozkurt, 2001b; Boccaletti and Steiner, 2005; Bozkurt and Rojay, 2005b]. There are some north‐south trending horst‐graben pairs in western Anatolia that formed during the Early Miocene time [Seyitoğlu, 1997; Altunkaynak and Yılmaz, 1999; Yılmaz et  al., 2000; Yılmaz and Karacık, 2001; Güney et al. 2002; Ersoy et al., 2010; see also Chapter 2, this volume]. The horst block that is located between the Gediz graben and the Simav graben (Fig.  2.48) is also divided into a number of north‐northeast to south‐southwest trending cross grabens (sensu lato). From the west to the east, these are the Gördes, Demirci, Selendi, and

Uşak‐Ulubey grabens (GD, CG, SG, UG in Fig.  2.48). The graben fill was deposited under a strict tectonic control of the graben bounding, en echelon, north‐northeast trending, steeply dipping oblique faults [Koçyiğit et al., 1999a; Koçyiğit, 2005; Yılmaz et al., 2000; Hakyemez et al., 2013], which display left‐lateral strike‐slip and dip‐ slip components. The graben sediments and intercalated volcanic rocks, ejected along the graben bounding faults, are Early to Middle Miocene in age (18–14 m yr; Seyitoğlu [1997]; Yılmaz et al. [2000]; Purvis et al. [2005]). Therefore, the north‐south trending grabens are older than the east‐ west trending grabens of the post–Late Miocene–Early Pliocene age (see Chapter 2, this volume, for a discussion on the age controversy of the grabens of the two contrasting trends). The east‐west trending grabens are commonly bounded by pure normal faults in the interior areas of the western Anatolian region. Approaching the Aegean Sea, the graben trend deviates from east‐west to northeast‐northwest (Figs. 2.48, 2.50, and 2.55a), where the graben bounding pure normal faults (a perfect example to the normal faults is observed along the northern flank of Sipil Mountain near Manisa City, Fig. 2.55c) give way to the oblique slip faults, displaying strike‐slip and dip‐slip components (Fig.  2.55b, d) [Gürer et  al., 2001; Sümer et al., 2013; Özkaymak et al., 2013; Seghedi et al., 2015; Gürer et  al., 2016]. The oblique faults, where they converge, define westerly enlarging triangular tectonic wedges (e.g., BEG and DP in Figs. 2.48 and 2.55a), which escape away from the point of convergence toward the Aegean Sea. To this, the Samson Mountain (the Dilek Peninsula) in the Söke area may be given as an example. The Kuşadası fault in the north and the Priene Sazlı‐fault in the south bound the tectonic wedge (Fig. 2.55a; Gürer et al. [2001]). The Priene–Sazlı fault morphologically is the most prominent fault of the region. It forms the boundary between the Priene‐Milet‐Söke graben (the name given to the western part of the Büyük Menderes graben) and the Samson Mountain horst (Dilek Peninsula in Fig. 2.55a), which rises steeply to over 1200 m heights from the graben floor (about 15–20 m high). The fault zone consists of a number of en echelon faults, along which offset streams are frequently observed (Fig. 2.55b). Total lateral displacement across the fault is inferred to be more than 25 km [Gürer et al., 2001]. The Priene‐Sazlı fault has been active throughout Pleistocene and Quaternary [Gürer et al., 2001; Sümer et al., 2013]. The steep and polished fault plane is commonly exposed (Fig.  2.55d). The slicken lines on the fault plane trend N25–50, and plunge 40–45° northeast. More than 500 m dip slip is estimated due to the vertical displacement of the Upper Miocene–Lower Pliocene lacustrine sediments that are observed at the top of the horst block, which

72  ACTIVE GLOBAL SEISMOLOGY

Figure 2.54  (a) Photo showing a southerly view of Soma graben. The areas defined by purple broken lines are the flatirons of the steeply dipping graben‐bounding normal fault. The white lines across the fault show the abruptly steepening slopes. A hanging valley (white curving line) was developed along the stream valley adjacent to the fault plane. The black broken line at the background displays the trace of the flat lying marker erosional surface that has been elevated above the horst. The equivalent erosional surface on the northern horst is displayed by the broken red square at the forefront of the photo. Note that the erosional surface is on the Miocene lacustrine sequence. (b) Photo showing a representative morphological profile from the northern flank of east‐west trending horst of western Anatolia. A canyonlike valley forms at the low level. Headword erosion along the valley is commonly in an incipient stage, and therefore the valley dies out before reaching the top of the horst. This suggests that the time elapsed since the horst was elevated is not long for a morphological equilibrium to be established between top of the horst and surrounding graben depressions. (b) A broad and shallow stream valley is commonly observed (yellow line) at the top of the horsts of western Anatolia. Despite the high altitude of the horsts (elevated above 800–1000 m; see [a]), deep carving has not yet developed in the valleys due to the fairly recent rise of the horst. The valleys commonly preserve the inherited meandering profile (green line). A resistant uppermost layer that is seen at the background of the photo corresponds to the Upper Miocene–Lower Pliocene lacustrine white limestone succession, above which a well‐developed flat‐lying erosional surface is also recognized. The lacustrine unit is presently seen at different heights varying from the sea level to the top of the horsts. This evidence alone indicates that the present horst‐graben structures postdate the early Pliocene. (c) Photo showing a representative morphological profile from the flank of an E-W trending horst of the Western Anatolia. A canyon like valley forms at the low level. Headword erosion along the valley is commonly in an incipient stage, and therefore the valley dies out before reaching to the top of the horst. (d) As a result of the recent elevation of the horsts, hanging valleys are commonly observed along their flanks as exemplified from the northern flank of Kazdağ Mountain. The champagne glass form is commonly noted along the hanging valleys.

indicates also that the faulting is younger than the early Pliocene. The lateral alluvial fans, developed in front of the fault in the western areas near the Aegean Sea, are Late Pliocene in age [Gürer et al., 2001]. The younger faults in the fault zone cut and elevated the fans with respect to one another and the graben floor (Fig.  2.55b). Some raised beaches are also observed on the upthrown blocks. The

eastern part of the fault is apparently younger, because the lateral alluvial fans formed in front of the fault are either in an incipient stage or absent. This indicates further that the time elapsed since the rise of this part of the horst is too short to generate an alluvial fan. The ancient city of Priene that is located on a downthrown block next to the fault, suffered many severe

Morphotectonic Development of Anatolia and the Surrounding Regions  73

Figure 2.55  (a) A simplified geology map of the Priene‐Milet or Söke graben (SG in Fig. 2.48) and the surroundings, representing the northeast trending western end of the east‐west trending Büyük Menderes graben. C and D refer to the locations of the photos in (c) and (d). M and P = locations of the ancient cities of Miletus and Priene; KF = Kuşadası fault; PSF = Priene‐Sazlı fault; SM = Samson Mountain horst. The westward extrusion (white arrow) of the tectonic wedge, bounded by the two oblique faults (KF and PSF) with significant strike‐slip components in (d), resulted in the Dilek Peninsula. Different senses of movement of the southern and northern parts of the graben region are indicated by the two curves rotating in the opposite direction, which generated northwest‐southeast trending extension (small red arrows) possibly in association with the anticlockwise sense of rotation of the Anatolian Plate. (c) The Manisa fault, a beautifully exposed pure normal fault, which is observed all along the northern flank of Sipil Mountain (see M in Fig. 2.48 for location of the fault) forming the southern boundary of the Gediz graben. Most of the fault plane structures (e.g., comb fractures, gutters, corrugations, tool marks, slicken lines, and colluvium) are clearly observed on the fault plane. (b) The Priene‐Sazlı fault near Atburgazı (C  on the map above). The fault runs between the low‐angle slope (above the alluvial plain) and steep slope (above the metamorphic rocks). Branches of the fault display en echelon fault pattern. The stream network flowing from the hills changes the trend abruptly across the fault zone due to the left‐lateral displacement. The pair of white arrows indicates the shear sense. The small vertical white arrow shows a fault scarp indicating the vertical displacement observed in the alluvial deposits. (d) Fault plane of the Priene‐Sazlı fault near Sazlı Village. The oblique lying slicken lines are clearly observed on the fault plane.

74  ACTIVE GLOBAL SEISMOLOGY

historical earthquakes [Yönlü et  al., 2010]. The fault plane solution of one of the recent earthquakes supports the field data that the motion along the fault was oblique, displaying a significant strike‐slip component [Yönlü et  al., 2010; Eyidoğan and Jackson, 1985].The kinematic analyses of the faults of the Söke graben region [Gürer et al., 2001; Sümer et al., 2013] reveal also that westward escape of the fault‐bounded triangular tectonic wedges is compatible with the clockwise rotation of the Anatolian Plate. Approaching the Aegean Sea, the Anatolian Plate turns to a south‐southwest direction (Fig. 2.2), and this forces the blocks to protrude away from the point of convergence. As a result of this, triangular depressions were formed where the faults intersect. The high rate of extension in western Anatolia resulted in high heat flow [Öztürk et al., 2006], which is associated with the stretched and thinned continental crust [Tirel et al., 2004]. Generation of the Aegean extension has been attributed to different causes. Among these may be listed (1) collapse of the western Anatolian orogen [Dewey, 1988; van Hinsbergen et al., 2009]; (2) westward escape of the Anatolian Plate [Şengör, 1979a, 1987; Philippon et  al., 2014]; (3) the slab retreat or slab  tear of the northward subducting eastern Mediterranean Ocean floor [LePichon and Angelier, 1979 and 1981; Agostini et al., 2008, 2010]. For a discussion on the topic see Şengör [1987] and Şengör et al. [1984], Wortel and Spakman [2000], Spakman and Wortel [2004], Bozkurt and Mittwede [2005], Jolivet et al. [2013], and Y. Yılmaz chapter 6, this volume. 2.9. CONCLUDING SUMMARY Anatolia is one of the most actively deformed continental regions in the world as indicated by the intense seismic activity and GPS data. The cause of this is the post‐late orogenic deformation following the continent‐ continent collision along the Bitlis‐Zagros belt. The northward advance of Arabia is still continuing in a fast rate of about 15–20 mm/yr. While some tectonic domains are rising fairly rapidly, some others are descending or moving laterally causing development of new landforms. The fast plate motions are manifested in the present morphology of Anatolia. The present major morphological entities of Anatolia, for example, the peripheral mountains, the eastern Anatolian high plateau, and the western Anatolian east‐west trending horst‐graben structures, are young morphotectonic elements. They began to form during the transitional period from the PaleoTectonic regime to the NeoTectonic regime in the eastern part of Anatolia, and then the whole of Anatolia and the surroundings began to be morphologically restructured with the beginning of the Neotectonic era in an increasing rate. The NeoTectonic regime started when the two trans-

form faults, namely the North Anatolian transform fault and the East Anatolian transform fault, were intersected in the Karlıova Junction of eastern Anatolia. Collectively, they defined a discreet lithospheric entity known as the Anatolian Plate, which began to move westward. From this time onward, the compressional forces deforming eastern Turkey have not been solely accommodated within the volume of the east of Anatolia. They began to be effectively partitioned. Before the NeoTectonic regime began (e.g., during the Early and the Middle Miocene periods) morphology of Anatolia was distinctly different from what it is today. The western part of Anatolia was a highland similar to an east‐west elongated archlike elevation with respect to the surrounding sea realms. The Mediterranean and the Black Sea were present in the north and the south as the remnant seas. The present eastern Anatolian high plateau was under a shallow sea environment. The highland extended from the Greek mainland and the Balkan region in the west to the edge of the east of Turkey in the east. Late Miocene is a critical period in Anatolia, because major changes in the morphotectonic history of development began during that period in eastern Turkey. The brief outline of the changes may be summarized as follows: the Bitlis suture mountains of southeastern Anatolia began to form as a result of consecutive stages of collisional events. They gradually eliminated the NeoTethyan Ocean and its dependencies including some interarc (back arc?) basins, and finally the remnant basins [Yılmaz, 1993]. This chain of events took place over a long period from the late Cretaceous to Middle Miocene. As a result of this, a giant nappe stack that had formed in the north of the Arabian Plate collided as the final phase of collisional events with Arabia and welded to it during the Middle Miocene. Despite the collision, northward advance of the Arabian Plate continued. It began to squeeze the orogenic belt, and initially eliminated a remaining sea from the zone of collision. Then it began to be accommodated locally by tightening and internally rearranging the orogenic belt and the surrounding areas. In a later period, the compressional forces began to disturb the entire area of eastern Turkey severely. At an initial phase, all the eastern regions of Turkey began to rise as a coherent block. As a consequence of this, the sea that covered the eastern Anatolian region gradually retreated. This occurred from the Middle Miocene to the Late Miocene. Interconnected lakes replaced the sea environment and covered nearly the entire eastern region during the late Miocene as demonstrated by a gradual transition from sea sediments to lake sediments (regional regression). The interconnected lakes disappeared after the Late Miocene. Some local depressions surrounded by the fold‐fault‐bounded highs began to form starting from Early Pliocene time. The anticline and syncline structures correspond to the hills and depressions indicating that

Morphotectonic Development of Anatolia and the Surrounding Regions  75

the initial morphology still remains in the region revealing also that the time elapsed since their development is yet too short for morphological inversions to occur. Within these separate continental basins, different successions were deposited. As a result of the continuing shortening deformation, the crust is thick (commonly more than 38 km) under the eastern Anatolian plateau. In the central part of eastern Anatolia, it is thicker compared to the surrounding areas due possibly to orthogonal shortening. This area forms a thrust‐bound central, domical high. The morphological elements and their trends fan away from this high to the east and to the west. The North Anatolian transform fault and the East Anatolian transform fault were formed when the continuing compressional forces could no longer be accommodated within the volume of eastern Turkey. They intersect in the Karlıova Junction. It is a roughly triangular depression that formed as a consequence of the westward motion of the Anatolian Plate, which is protruding away from the converging jaws of the plates. It started transferring some north‐south compression mostly to the west and only a small fraction of that to the northeast. It began possibly during the Pleistocene, because the sediment fill that is confined to the Karlıova basin dates back to this time. Eastern Anatolia is underlain by an accretionary complex, which formed as a result of the gradual elimination and then the total consumption of the NeoTethyan Ocean, which was located between the Pontide arc in the north and the Arabian Plate and some continental slivers that were rifted off of Arabia in the south. This occurred from the Late Cretaceous to the Late Eocene period. Elimination of the oceanic environment along the subduction zones generated a wide belt of ophiolitic mélange association, which in the following period behaved as a cushion between the converging plates and began to absorb the compression. Consequently, the accretionary complex was shortened, thickened, and elevated to 2000 m heights, and formed a buffer, which did not allow direct contact of the approaching continental plates (the Turkic type orogen of Şengör et  al. [2008]). The northern and southern peripheries of eastern Turkey that are underlain by a common continental crust, responded to the increasing compression more severely, and thrust to the north and the south, and have risen steadily and possibly episodically with a higher rate with respect to the central region, and reached to elevations above 3500 m heights. Kaçkar Mountain is 3932 m high and Buzuldağ Mountain is 4135 m high on the northern and the southern mountain ranges, respectively. Therefore, the entire eastern Anatolian plateau may be regarded as a giant ramp basin with respect to the peripheral mountains. Probably coinciding with this period, a critical event occurred in the oceanic slab that subducted

under eastern Anatolia. Break off of the subducting slab occurred (this is seismically identified at depths of about −500 km), and the mantle lithosphere was possibly delaminated allowing the hot asthenosphere to be underplated the overthickened continental lithosphere. This is manifested by initiation of a volcanic activity on the northern part of eastern Anatolian region about 11 m yr ago. The volcanic activity migrated toward southern regions and intensified in the central part about 6 m yr ago. The southward migration of the volcanic activity continued to the present due possibly to the slab steepening. These events accelerated the expansion and the rise of the overlying plate. The timing of morphotectonic development and the young volcanism of the eastern Anatolian high plateau were simultaneous with the evolution of central Anatolia. In fact, there was no morphotectonic barrier between the two regions. This indicates that the two regions were parts of the same morphotectonic entity. This suggests further that the broken slab observed under eastern Anatolia extended westward to central Anatolia. (It does not extend farther west. It is not connected to the eastern Mediterranean subducting slab [van Hinsbergen et al., 2009].) Compressional components of the north‐south shortening deformation decrease steadily going away from the eastern Anatolia as exemplified from the gradual westward decrease of the height of the Pontides from 4000 m down to 1000 m. NATF and EATF play significant roles in the distribution of the stress. A major culmination, the Niğde‐Kırşehir Massif, representing a core complex, occupies the central part of central Anatolia. As a rigid body, it helps to accommodate a significant portion of the stress that forces central Anatolia to deform by rotating internally and externally as revealed by the paleomagnetic data. Therefore, the massif is divided into some rigid, fault‐bound blocks, and also surrounded by the faults that separate it from the young depressions. The massif has played a significant role in rearranging the old structures, generating a set of new structures, and therefore has determined morphotectonic development of the central part of central Anatolia. Some of the major faults of central Anatolia that splay off of NATF curve along a southwest direction. They are mostly transpressional in character, and accommodate part of the motions of the Anatolian Plate within the volume of central Anatolia. The major faults in the western part of central Anatolia, on the other hand, are commonly transtensional (e.g., the Tuz Gölü fault). Going westward away from this region, the transtensional province of western Central Anatolia is replaced by the extensional province of western Anatolia. There is no distinct morphotectonic boundary between the two regions in the northern areas. The northwest‐striking Eskişehir fault (known also as the Eskişehir‐Bursa fault) with a dominant strike‐slip component stretches to the Bursa plain, where it is replaced by

76  ACTIVE GLOBAL SEISMOLOGY

east‐west‐striking normal faults, which have elevated the Uludağ Mountain above 2500 m height. In the southern areas, on the other hand, the Fethiye‐Burdur strike‐slip fault zone located along the extension of Pliny‐Strabo trenches, which extends to the apex zone of the Isparta Angle, may be viewed as a morphotectonic boundary. Morphotectonic evolution of western Anatolia has followed a different path compared to that of eastern Anatolia. The sea environment left western Anatolia at the end of the Late Eocene–Oligocene time, and the region has remained land until now. Following a long absence, the sea incursion began into the region along a north‐south central axis of the present Aegean Sea region starting from the Late Miocene. The surrounding lands were covered by interconnected lakes, which were not far above sea level. The sea and lake sediments of that period show transitions in the north and the south of western Anatolian regions. The lake period survived during Early Pliocene time. Although the east and west of Anatolia were covered by shallow lake environment at about the same period during the Late Miocene, the major difference between them is that, while the eastern region was elevating, the eastern region was subsiding under the north‐south compressional and extensional regimes, respectively. The relative motions of the two regions may be compared to the behavior of a seesaw. Central Anatolia acted as an axis between them showing morphotectonic transitions to the neighboring regions. The western Anatolian highland subsided (collapsed!) at least twice as a result of approximately north‐south extension. The first phase occurred during a brief period sometime during the Late Oligocene—Early Miocene.The second phase began at the end of the Late Miocene—Early Pliocene time and has continued interruptedly until now. The present east‐west trending horst and graben structures that dominate the present landscape began to form during later stages of the second phase of extension in the Quaternary. Their young character is clearly identified in the morphological features; the major valleys and the bordering hills correspond to the grabens and horsts, respectively. They represent the initial morphology that still remains. Approaching the Aegean Sea, the graben trend changes from east‐west to northeast‐northwest, where the graben bounding pure normal faults gives way to oblique slip faults, displaying strike‐slip and dip‐slip components. The oblique faults, where they converge, define westerly enlarging triangular tectonic wedges, which escape away from the point of convergence toward the Aegean Sea. The westward escape of the tectonic wedges is compatible with the clockwise rotation of the Anatolian Plate. Interactions of different tectonic parameters appear to have played active roles in the generation of the north‐ south extension and the consequent morphology of west-

ern Anatolia. Among these may be mentioned (1) collapse of the orogen (the western Anatolian crust is claimed to have exceeded 60 km of thickness during the postcollisional convergence); (2) retreat of the northward subducting eastern Mediterranean oceanic slab under the Aegean‐Anatolian Plate; and (3) westward escape of the Anatolian Plate. Particularly, two spatially and temporally associated events, the escape and the retreat, may have played determining roles during this development of the present morphological profile of the region. The retreat of the down‐going slab is commonly regarded as a factor solely responsible for the extension of the overlying Aegean‐Anatolian Plate. The latest phase of the retreat approximately corresponds to the westerly movement of the Anatolian Plate. This leads to anticipation of a possible genetic connection between the two events. The escape of the Anatolian Plate may have influenced or caused retreat of the subducting slab. The Marmara region is a morphotectonically more complex province of Turkey, where the north‐south extensional regime of the Aegean‐western Anatolian region and the northern boundary of Anatolian Plate as represented by NATF intersect. NATF entered the Marmara region as a broad shear zone, possibly during the Pleistocene. It has put its own imprint in the morphology as the most influential structural agent. Prior to this, north‐south extension effectively deformed the entire Marmara region in a similar style as is seen in western Anatolia today. The width of the shear zone progressively narrowed in time, and finally was confined to a continuous lineament as a linear trough. Since it became a narrow fault zone, it defines a plate boundary, along which the Anatolian Plate began moving westward. This is a critical time also for the morphotectonic evolution of the Marmara region. From this time onward, the northern and southern regions with respect to the NATF began to  behave semi‐independently. The southern region responded to the ongoing north‐south extension more effectively and elevated to higher altitudes. ACKNOWLEDGMENTS This morphotectonic synthesis is based on decades‐ long field studies and mostly geological maps that I, together with my former students, prepared in Anatolia. MTA, TPAO, and TUBITAK of Turkey supported different parts of the research projects. I am grateful to these institutes for their long and enthusiastic support. As a result, hundreds of sheets of maps on the scale of 1/25,000 were produced representing a majority of the major segments of Anatolia from the north to the south, from the west to the east. During the preparation of the manuscript, we discussed many problems with a number of colleagues. They make a long list. Among them is Celal

Morphotectonic Development of Anatolia and the Surrounding Regions  77

Şengör at the forefront. Their contributions are greatly acknowledged. Many colleagues and friends at TPAO were most helpful in allowing us to use some of the seismic data and geological maps that we collectively produced. I sincerely thank Prof. Şafak Altunkaynak and some graduate students, Ö. Kamacı, A. Ünal, Tanyel Baykut, and Z. Çalışkanoğlu, who were most helpful in the preparation of the figures and manuscript. I also thank my wife, Prof. Tarcan Yılmaz, for her patience, understanding, and support throughout the preparation of the book. REFERENCES Agostini, S., J. G. Ryan, S. Tonarini, and F. Innocenti (2008), Drying and dying of a subducted slab: Coupled Li and B isotope variations in Western Anatolia Cenozoic volcanism, Earth Planet. Sci. Lett., 272, 139–147. Agostini, S. C., Doglioni, F. Innocenti, P. Manetti, and S. Tonarini (2010), On the geodynamics of the Aegean rift, Tectonophysics, 488, 7–21. Akinci, O., A. H. F. Robertson, A. Poisson, and E. Bozkurt (2003), Special issue on the Isparta Angle, SW Turkey, Geological J., 38, 195–234. Aksoy, E., Z. Murat, and A. Koçyiğit (2007), Lake Hazar basin: A negative flower structure on the East Anatolian fault system (EAFS), SE Turkey 2007, Turkish J. Earth Sci., 16, 319–338. Aksu, A., T. Calon, and R. N. Hiscott (2000), Anatomy of the North Anatolian fault zone in the Marmara Sea, western Turkey: Extensional basins above a continental transform, GSA Today, 10(6), 3–7. Aksu, A. E., T. J. Calon, J. Hall, and D. Yaşar (2005), Origin ̇ and evolution of the Neogene Iskenderun basin, northeastern Mediterranean Sea, Marine Geol., 221, 161–187. Akyüz, S. H., E. Altunel, V. Karabacak, and C. C. Yalçıner (2006), Historical earthquake activity of northern part of the Dead Sea fault zone, southern Turkey, Tectonophysics, 426, 281–293. Alchalbi, A., M. Daoud, F. Gomez, S. McClusky, R. Reilinger, M. Abu Romeyeh, A. Alsouod, R. Yassminh, B. Ballani, R. Darawcheh, R. Sbeinati, Y. Radwan, R. Al Masri, M. Bayerly, R. Al Ghazzi, and M. Barazangi (2010), Crustal deformation in northwestern Arabia from GPS measurements in Syria: slow slip rate along the northern Dead Sea Fault, Geophys. J. Int., 180, 125–135. Alçiçek, H. (2010), Stratigraphic correlation of the Neogene basins in Southwestern Anatolia: Regional palaeogeographical, palaeoclimatic and tectonic implications, Palaeogeogr. Palaeoclimatol. Palaeoecol., 291, 297–318. Altın, B. T. (2003), “Aladağlar uzerinde (Ecemiş Çayı Aklanı) buzul ve karst jeomorfolojisi” (On the Aladağlar, the Ecemiş ̇ River Valley, glacial and karstic morphology), Istanbul ̇ Üniversitesi Sosyal Bilimler Enstitüsü Istanbul, Doktora Tezi, 513 s. (Doctorate Thesis). Altın, T. B. (2005), Development of drainage and terrace systems in Ecemiş valley, NE of Mediterranean region, Proc. Int. Symp. Geog., 5–8 June 2007, Kemer, Antalya‐Turkey, 31–43.

Altunkaynak, S., and Y. Yılmaz (1999), The Kozak magmatic complex, western Anatolia, J. Volcanol. Geotherm. Res., 85, 211–231. Ambraseys, N. N., and J. A. Jackson (1998), Faulting associated with historical and recent earthquakes in the Eastern Mediterranean region, Geophys. J. Int., 133, 390–406. Angus, D. A., D. C. Wilson, and E. N. Sandvol (2006), Lithospheric structure of the Arabian and Eurasian collision zone in eastern Turkey from S‐wave receiver functions, Geophys. J. Int., 166, 1335–1346. Armijo, R., B. Meyer, A. Hubert, and A. Barka (1999), Westward propagation of the North Anatolian fault into the northern Aegean: Timing and kinematics, Geology, 27(3), 267–270; doi: 10.1130/0091‐7613(1999) 0272.3.CO;2. Armijo, R., N. Pondard, B. Meyer, G. Uçarkuş, B. Mercier de Lépinay, J. Malavieille, S. Dominguez, M. A. Gustcher, S. ̇ Schmidt, C. Beck, N. Çağatay, Z. Çakir, C. Imren, K. Eriş, B. Natalin, S. Özalaybey, L. Tolun, I. Lefèvre, L. Seeber, L. Gasperini, C. Rangin, O. Emre, and K. Sarikavak (2005), Submarine fault scarps in the Sea of Marmara pullapart (North Anatolian Fault): Implications for seismic hazard in Istanbul, Geochem. Geophys. Geosyst., 6, Q06009. Arpat, E., and F. Şaroğlu (1975), Türkiyede önemli bazı tektonik olaylar, Türkiye Jeoloji Kurumu Bült., 18, 91–101. Aydar, E., H. E. Çubukçu, E. Şen, and L. Akın (2013), Central Anatolian plateau Turkey: Incision and peleoaltimetry recorded from volcanic rocks, Turkish J. Earth Sci., 22, 739–746. Banks, C. J., A. G. Robinson, and M. P. Williams (1997), Achara Trialeti fold belt and the adjacent Riouni and Kartli foreland basins, Republic of Georgia, in AAPG Memoir 68: Regional and Petroleum Geology of the Black Sea and Surrounding Region, edited by A. G. Robinson, 331–346. Barazangi, M., E. Sandvol, and D. Seber (2006), Structure and tectonic evolution of the Anatolian plateau in eastern Turkey, in Post Collisional Tectonics and Magmatism in the Mediterranean Region and Asia, edited by Y. Dilek and S. Pavlides, 463–473, Geological Society of America Special Papers, 409. Barka, A., and R. Reilinger (1997), Active tectonics of the Eastern Mediterranean region: Deduced from GPS, neotectonic and seismicity data, Ann. Geophys., 40(3), 587–610. Barka, A. A. (1992), The North Anatolian fault zone, Ann. Tectonicae, 6, 164–195. Barka, A. A. (1996), Slip distribution along the North Anatolian fault associated with the 349 large earthquakes of the period 1939 to 1967, BSSA, 86(5), 1238–1254. Barka, A. A., and L. Gülen (1988), New constraints on age and total offset of the North Anatolian Fault Zone: Implications for tectonics of the eastern Mediterranean region, Middle East Tech. Univ. J. Pure Appl. Sci., 31, 39–63. Barka, A. A., and L. Gülen (1989), Complex evolution of the  Erzincan Basin (Eastern Turkey), Tectonophysics, 11, 275–283. Barka, A. A., and P. L. Hancock (1984), Neotectonic deformation patterns in the convex‐northwards arc of the North Anatolian fault, Geological Society, London, Special Publications, 17(1), 763–773; doi: 10.1144/GSL.SP.1984.017.01.61

78  ACTIVE GLOBAL SEISMOLOGY Barka, A. A., S. Akyüz, H. H. A. Cohen, and F. Watchorn (2000), Tectonic evolution of the Niksar and Taşova‐Erbaa pull‐apart basins, North Anatolian Fault Zone: Their significance for the motion of the Anatolian block, Tectonophysics, 322, 243–264. Bartol, J., R. Govers, and M. J. R. Wortel (2012), Mantle delamination as the cause for the Miocene–Recent evolution of the Central and Eastern Anatolian Plateau, Geophys. Res. Abstr., 14, EGU General Assembly 2012, Vienna, Austria, EGU2012–EGU11778. Bayrakçı, G., M. Laigle, A. Bécel, A. Hirn, T. Taymaz, S. Yolsal‐ Çevikbilen, and SEISMARMARA Team (2013), 3‐D basement tomography of the northern Marmara trough by a dense OBS network at the nodes of a grid of controlled source profiles along the North Anatolian Fault, Geophys. J. Int., 194, 1335–1357. Beccaletto, L., and C. Steiner (2005), Evidence of two‐stage extensional tectonics from the northern edge of the Edremit Graben, NW Turkey, Geodinamica Acta, 18, 283–297. Bécel, A., M. Laigle, B. De Voogd, A. Hirn, T. Taymaz, S. Yolsal‐Çevikbilen, and H. Shimamura (2010), North Marmara trough architecture of basin infill, basement and faults, from PSDM reflection and OBS refraction seismics, Tectonophysics, 490, 1–14 Beck, C., B. Mercier de Lépinay, J. L. Schneider, M. Cremer, N. Çağatay, E. Wendenbaum, S. Boutareaud, G. Ménot, S.  Schmidt, O. Weber, K. Eris, R. Armijo, B. Meyer, N.  Pondard, M. A. Gustcher, Marmaracore Cruise Party, J. L. Turion, L. Labeyrie, E. Cortijo, Y. Gallet, H. Bouqueral, N. Görür, A. Gervais, M. H. Castera, L. Londeix, A. de Resseguier, and A. Jaouen (2007), Late Quaternary co‐seismic sedimentation in the Sea of Marmara’s deep basins, Sed. Geol., 199, 65–89. Benda, F. (1971), Principles of the subdivision of the Turkish Neogene, Newsletter. Stratig., 1, 23–26. Benda, F., R. Innocenti, F. Mazzuoli, Radicati, and P. Steffens (1974), Stratigraphic and radiometric data of the Neogene in northwest Turkey, Zeit. Deutschen Geo. Gesellschaft, 125, 183–193. Benda, L., and J. E. Meulenkamp (1990), Bio stratigraphic correlations in the Eastern Mediterranean Neogene 9, Sporomorph associations and event stratigraphy of the Eastern Mediterranean Neogene, Newsletter. Stratig., 23, 1–10. Benda, L., J. E. Meulenkamp, R. R. Schmidt, P. Steffens, and J. W. Zachariasse (1997), Bio stratigraphic correlations in the Eastern Mediterranean Neogene. 2. Correlation between sporomorph associations and marine microfossils from the Upper Oligocene‐ Lower Miocene of Turkey, Newsletter. Stratig., 6, 1–22. Bertotti, G., and D. Fernández‐Blanco (2014), The regional geology of the south Turkey‐north Cyprus domain: New perspective and consequences for hydrocarbon plays, AAPG Int. Conf. Exhib.,14–17 Sept. 2014, Istanbul, Abstracts and Programme.30 and Abstract CD. Bird, P. (1979), Continental delamination and the Colorado Plateau, J. Geophys. Res., 84, 7561–7571. Bird, P. (2003), An updated digital model of plate boun­ daries, Geochem. Geophys. Geosys., 4(3), 1027; doi: 10.10292001GC000252.

Black, K. N., E. J. Catlos, T. Oyman, and T. Demirbilek (2013), Timing Aegean extension: Evidence from in situ U‐Pb geochronology and cathodoluminescence imaging of granitoids from NW Turkey, Lithos, http://dx.doi.org/ 10.1016/j.lithos.2013.09.001. Blackwell, B. A., J. A. Florentine, O. Tüysüz, U. Tarı, Ş. C. ̇ Genç, J. K. Imren, and C. Blickstein (2012), Slipping up ESR dating mollusks from marine terraces in Hatay Province, Turkey, Geol. Soc. Am., Abstracts with Program, 44, 296. Blackwell, B. A. B., J. Florentine, O. Tüysüz, U. Tarı, Ş. C. ̇ Genç, C. Imren, and M. Kim (2011), An uplifting idea: ESR dating marine terraces in Hatay province, Turkey, Geolog. Soc. Am., Abstracts with Programs, 43, 273. Boccaletti, L., and C. Steiner (2005), Evidence of two‐stage extension from the northern edge of the Edremit Graben, NW Turkey, Geodinamica Acta, 18(3–4), 283–297. Boulton, S. J. (2007), New insights into the northern Dead Sea Fault Zone (Karasu Rift and Hatay Graben), Southern Turkey, American Geophysical Union, Fall Meeting 2007, abstract #T42B‐02. Boulton, S. J. (2013), Tectonic development of the southern Karasu Valley, Turkey: Successive structural events during basin formation, in Geological Development of Anatolia and the Easternmost Mediterranean Region, edited by A. H. F. Robertson, O. Parlak, and U. C. Ünlügenç, 531–546, Geological Society, London, Special Publications, 372. Boulton, S. J., and A. H. F. Robertson (2007), The Miocene of the Hatay area, S. Turkey: Transition from the Arabian passive margin to an under filled foreland basin related to closure of the Tethys Ocean, Sed. Geol., 198, 93–124. Boulton, S. J., and A. H. F. Robertson (2008), The Neogene‐ Recent Hatay Graben, south central Turkey: Graben formation in a setting of oblique extension (trans tension) related to post collisional tectonic escape, Geol. Mag., 145, 800–821. Bozkurt, E. (2000), Timing of extension on the Büyük Menderes graben, western Turkey and its tectonic implications, Geological Society, London, Special Publications, 173, 385–403. Bozkurt, E. (2001a), A Late Alpine evolution of the central Menderes Massif, western Turkey, Int. J. Earth Sci., 89, 728–744; doi: 10.1007/s005310000141. Bozkurt, E. (2001b), NeoTectonic of Turkey: A synthesis, Geodinamica Acta, 31, 3–30. Bozkurt, E., and B. Rojay (2005b), Episodic, two‐stage Neogene extension and short‐term intervening compression in western Turkey: Field evidence from the Kiraz Basin and Bozdağ Horst, Geodinamica Acta, 18(3), 299–316. Bozkurt, E., and N. B. Çiftçi (2009), Structural evolution of the Gediz Graben, SW Turkey: Temporal and spatial variation of the graben basin, Basin Res., 22, 846–873; doi: 10.1111/ j.1365‐2117.2009.00438.x. Bozkurt, E., and R. G. Park (1994), Southern Menderes Massif: An incipient metamorphic core complex in western Anatolia, Turkey, J. Geol. Soc. London, 151, 213–216. Bozkurt, E., and R. Oberhänsli (2001), Menderes Massif (western Turkey): Structural, metamorphic and magmatic evolution: A synthesis, Int. J. Earth Sci., 89, 679–708; doi: 10.1007/s005310000173. Bozkurt, E., and S. K. Mittwede (2005), Introduction: Evolution of Neogene extensional tectonics of western Turkey, Geodinamica Acta, 18, 153–165.

Morphotectonic Development of Anatolia and the Surrounding Regions  79 Boztuğ, D., R. Jonckheere, G. A. Wagner, and Z. Yeğingil (2004), Slow Senonian and fast Paleocene–Early Eocene uplift of the granitoids in the Central Eastern Pontides, Turkey: Apatite fission‐track results, Tectonophysics, 382, 213–228. Carter, J. M. L., P. P. Hanna, A. M. Ries, and P. Turner (1991), Tertiary evolution of the Sıvas Basin, Central Turkey, Tectonophysics, 195, 29–46. Catlos, E. J., and I. Çemen (2005), Monazite ages and the evolution of the Menderes Massif, western Turkey, Int. J. Earth Sci., 94, 204–217 Catlos, E. J., C. B. Baker, S. S. Sorensen, I.̇ Çemen, and M. Hançer (2008), Monazite geochronology, magmatism, and extensional dynamics within the Menderes Massif, Western Turkey, Donald D Harrington Symposium on the Geology of the Aegean, Conf. Series: Earth and Environmental Science, 2, 01. Catlos, E. J., C. B. Baker, S. S. Sorensen, L. Jacob, and I. Çemen (2011), Linking micro cracks and mineral zoning of detachment‐exhumed granites to their tectonomagmatic history: Evidence from the Salihli and Turgutlu plutons in western Turkey (Menderes Massif), J. Struct. Geol., 33, 951–969. Çelik, H. (2008), Doğu Anadolu Fay sisteminde Sivrice Fay zonunun Palu‐Hazar Gölü (Elazığ) arasındaki bölümünde atımla ilgili yeni arazi bulgusu (New field data on the displacement of the Palu Lake Hazar part of the Sivrice Fault zone of the East Anatolian fault system) (Elazığ, E. Turkey), Science, Eng. J. Fırat Univ., 20(2) 305–314. Çemen, İ., M. C. Göncüoğlu, and K. Dirik (1999), Structural evolution of the Tuzgölü basin in central Anatolia, Turkey, J. Geol., 107, 693–706, N.B., Bozkurt, E., 2010. Structural evolution of the Gediz Graben, SW Turkey: temporal and spatial variation of the graben basin, Basin Res., 22, 846–873. Çemen, I.,̇ C. Helvacı, and E. Y. Ersoy (2014), Cenozoic extensional tectonics in western and central Anatolia, Turkey: Introduction, Tectonophysics, 635, 1–5, doi. org/ 10.1016/ j.tecto.2014.09.004. Çiner, A., U. Doğan, G. Go, C. Yıldırım, N. Akçar, and S. Ivy‐ Ochs (2015), Quaternary uplift rates of the Central Anatolian Plateau, Turkey: insights from cosmogenic isochron‐burial nuclide dating of the Kızılırmak River terraces, Quaternary Sci. Rev., November 2014; doi: 10.1016/j.quascirev.2014.10.007. Çinku, M., B. Ülker, N. Kaya, E. Öksüm, Y. Yılmaz, A. U. Özbey, and N. Orbay (in prep.), The tectonic history of the Niğde‐Kırşehir Massif and the Tauride since the Late Mesozoic: Paleomagnetic evidence for two‐phase orogenic curvature in the central Anatolia, Tectonics, (Jan.). Çinku, M., M. Hisarli, A. M. Hirt, F. Heller, T. Ustaömer, N. Kaya, E. Öksüm, and N. Orbay (2014), Evidence of late Jurassic oroclinal bending in North‐Central Anatolia: Paleomagnetic results from Mesozoic and Cenozoic rocks along the Izmir Ankara Erzincan Suture Zone, Geol. Soc. London Spec. Issue, Paleo magnetism in Fold and Thrust belts: New perspectives, edited by B. O. Urcia (in press). Collins, A. S., and A. H. F. Robertson (2003), Kinematic evidence for late Mesozoic‐Miocene emplacement of the Lycian Allochthone over the western Anatolide belt, SW Turkey, Geological J., 38, 295–310.

Copley, A., and J. Jackson (2006), Active tectonics of the Turkish‐Iranian Plateau, Tectonics, 25, 6006. Cormier, M. H., L. Seeber, C. M. G. McHugh, A. Polonia, N. Çağatay, Ö. Emre, L. Gasperini, N. Görür, G. Bortoluzzi, E. Bonatti, W. B. F. Ryan, and K. R. Newman (2006), North Anatolian Fault in the Gulf of Izmit (Turkey): rapid vertical motion in response to minor bends of a nonvertical­continental transform, J. Geophys. Res., 111, B04102 Cosentino, D., T. F. Schildgen, P. Cipollari, C. Faranda, E. Gliozzi, N. Hudáčková, S. Lucifora, and M. R. Strecker (2012), Late Miocene surface uplift of the southern margin of the central Anatolian plateau, Central Taurides, Turkey, Geol. Soc. Am. Bull., 124, 133–145. De Graciansky, P. C. (1972), Recherches géologiques dans le Taurus Lycien occidental, Thèse Doctorat d’état, Univ. Paris Sud, Orsay, France. De LaMotte F., A. Poisson, A. Aubourg, and H. Temiz (1995), Chevauchements post‐Tortoniens vers l’ouest puis vers le sud au coeur de l’angle d’Isparta (Taurus, Turquie): Conséquences géodynamiques, Bull. Soc. Géol. France, 166(1), 59–67. De Mets, C., R. G. Gordon, D. F. Argus, and S. Stein (1994), Effects of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions, Geophys. Res. Lett., 21, 2191–2194. De Mets, C. R. G., G. Gordon, and D. F. Argus (2010), Geologically current plate motions. Geophys. J. Int., 181, 1–80. Demir, T., A. Seyrek, R. Westaway, D. Bridgland, and A. Beck (2008), Late Cenozoic surface uplift revealed by incision by the River Euphrates at Birecik, southeast Turkey. Quaternary Int., 186, 132–163. Demir, T., A. Seyrek, R. Westaway, H. Guillou, S. Scaillet, A. Beck, and D. R. Bridgland (2012), Late Cenozoic regional uplift and localized crustal deformation within the northern Arabian platform in southeast Turkey: investigation of the Euphrates terrace staircase using multidisciplinary techniques, Geomorphology, 165–166, 7–24. Demirbağ, E., E. Gökaşan, F. Y. Oktay, M. Şimşek, and H. Yüce (1999), The last sea level changes in the Black sea: evidence from the seismic data, Marine Geol., 157, 249–265. Demoulin, A., B. Altın, and A. Beckers (2013), Morphotectonic age estimate of the last phase of accelerated uplift in the Kazdağ area (Biga Peninsula NW Turkey), Tectonophysics, 608, 1380–1393. Dewey, J. F. (1988), Extensional collapse of orogens, Tectonics, 7, 1123–1139. Dewey, J. F., M. R. Hempton, W. S. F. Kidd, F. Şaroglu, and A. M. C. Şengör (1986), Shortening of continental lithosphere: The Neotectonic of eastern Anatolia‐a young collision zone, in collision tectonics, Geological Society Special Publication, 19, 3–36. Dhont, D., J. Chorowicz, T. Yürür, and O. Köse (1998), Polyphased block tectonics along the North Anatolian fault in the Tosya basin area (Turkey), Tectonophysics, 299(1–3), 213–227. Dhont, D., L. Chorowicz, and T. Yürür (1999), The Bolkar mountains (central Taurides, Turkey): A Neogene extensional thermal uplift? Geol. Bull. Turkey, 42, 69–87.

80  ACTIVE GLOBAL SEISMOLOGY Dirik, K. (2001), Neotectonic evolution of the northwestward arched segment of Central Anatolian fault zone, Central Anatolia, Turkey, Geodinamica Acta, 14, 147–158. Dirik, K., and M. Göncüoğlu (1996), Neotectonic characteristics of central Anatolia. Int. Geol. Rev., 38, 807–817. Dirik, K., and O. Erol (2000), Tuzgölü ve civarının tektonomorfolojik evrimi (Tectonomorphologic evolution of Tuzgölü and surrounding area). Haymana‐Tuzgölü‐Ulukışla Basenleri Uygulamalı çalışma. Türkiye Petrol Jeologları Derneği, Özel Sayı 5, 27–46 (in Turkish with English abstract). Dirik, K., and O. Erol (2003), Tuzgölü ve civarının tektonomorfolojik evrimi, Orta Anadolu‐Türkiye. TPJD Özel sayı, 27–46. Dirik, K., M. C. Göncüoğlu, and H. Kozlu (2001), Ecemiş Fay Zonu orta kesiminin (Sultansazlığı‐Tuzlagölü arası) stratigrafisi ve tektoniği, Niğde Üniv. Müh. Fakültesi. Yayını, 73–90. Dolu, E., E. Gökaşan, E. Meriç, M. Ergin, T. Görüm, H. Tur, B. Ecevitoğlu, N. Avşar, M. Görmüş, F. Batuk, B. Tok, and O. Çetin (2007), Quaternary evolution of the Gulf of Izmit (NW Turkey): A sedimentary basin under control of the North Anatolian fault zone, Geo‐Mar. Lett., 27, 355–381. Duman, T. Y., and Ö. Emre (2013), The East Anatolia fault, geometry, segmentation and jog characteristics, Geological Society Special Publication, 2013; doi: 10.1144/SP372.14. Ediger, V. Ş., Z. T. Batı, and M. Yazman (1996), Paleo‐palynology of possible hydrocarbon source of the Alaşehir–Turgutlu area in the Gediz graben (western Anatolia), Turkish Assoc. Petrol. Geol. Bull., 9, 11–23. Efe, R., A. Soykan, I. Urebal, and S. Sönmez (2011), Reviewing the geomorphological and Neotectonic features of the Gönen basin, NW Turkey, Procedia Soc. Behav. Sci., 19, 716–725. Ekstrom, G. A. M., N. N. Dziewonski, N. Matternovskaya, and M. Nettles (2005), Global seismicity of 2003: Centroid moment‐tensor solutions for 1087 earthquakes, Phys. Earth Planet. Inter., 148, 327–351. Elliott, J. R., A. Copley, R. Holley, K. Scherer, and B. Parsons (2013), The 2011 Mw 7.1 Van (Eastern Turkey) Earthquake, J. Geophys. Res., 118, 1619–1637. Elmas, A., and E. Meriç (1998), The seaway connection between the Sea of Marmara and Mediterranean: Tectonic development of Dardanelles, Int. Geol. Rev., 40 (2), 144–163. Elmas, A., and Y. Yılmaz (2007), Development of on oblique subduction zone Tectonic evolution of the Tethys suture zone in southeast Turkey, Int. Geol. Rev., 45/9, 827–841. Emre, O., N. Kazancı, T. Erkal, S. Görmüş, N. Görür, I.̇ Kuşcu, M. Karabıyıkoğlu, and M. Keçer (1997), Uluabat‐Manyas Göllerinin Oluşumu ve Güney Marmara’nın Kuvaterner ̇ evrimi, TÜBITAK‐Üniversite‐MTA Ulusal Deniz Araştırmaları Programı WORKSHOP III (2‐3 Haziran 1997), Bildiri Özleri Kitabı, 23–27, Ankara. Emre, O., O. Tüysüz, and C. Yildirim (2009), Uplift of Pontide orogenic belt since the late Miocene, Abstract Book, Second International Symposium on the Geology of the Black Sea Region, MTA Congress Centre, Ankara, 5–9 Oct. MTA Pub. Emre, O., T. Erkal, A. Tchepalyga, N. Kazancı, M. Keçer, and E. Ünay (1998), Neogene‐Quaternary evolution of the eastern Marmara region, northwest Turkey, Doğu Marmara

bölgesinin Neojen‐Kuvaternerdeki evrimi, Bull. Mineral Res. Explor. Inst. Turkey, 120, 119–145. Emre, O., Y. Tamer, S. Duman, and S. Özalp (1912–1913), Türkiye Diri fay Haritası (Active Fault Map of Turkey), Maden Tetkik ve Arama Genel Müdürlüğü Yayını, Ankara. Emre, T. (1990), Sart Mustafa (Salihli)‐Adala–Dereköy (Alaşehir) arasının jeolojisi ve Gediz Grabeni’nin yapısına bir yaklaşım, TÜBİTAK, TBAG‐732 / YBAG–0001 project, 65 s. Emre, T. (1998), Gediz Grabeni (Salihli‐Alaşehir arası) karasal tortullarının yaşıyla ilgili yeni bulgular. Türkiye Jeoloji Kurumu Bildiri Özetleri, Ankara, 34–35. Erinç, S. (1954), Orta Ege Bölgesi’nin Jeomorfolojisi, MTA Raporu No: 2217, Ankara (yayımlanmamış). ̇ Erinç, S. (1977), Istanbul Boğazı ve Çevresi (Doğal Ortam: ̇ Etkiler ve Olanaklar): Istanbul Üniversitesi Coğrafya Enstitüsü Dergisi, Sayı, 20–21, 1–24. Ergün, M., and E. Özel (1995), Structural relationship between the Sea of Marmara basin and the North Anatolian Fault Zone, Terra Nova, 7(2), 278–288. Erol, O. (1963), Geomorphology of Asi River Delta and Sea and River Terraces in the Fourth Era, A.Ü.D.T.C.F. Yayınları Sayı, 1, 148, Ankara. Erol, O. (1991), Geomorphological evolution of the Taurus Mountains, Turkey, Zeitschr Für Geomorf. N. F. Supp. Bd. 82, Dho 99‐109. Ersoy, E. Y., C. Helvacı, and H. Sözbilir (2010), Tectono‐stratigraphic evolution of the NE–SW trending superimposed Selendi basin: Implications for Late Cenozoic crustal extension in western Anatolia, Turkey, Tectonophysics, 488, 210–232. Ersoy, E. Y., I.̇ Çemen, C. Helvacı, Z. Billor (2014), Tectono‐ stratigraphy of the Neogene basins in western Turkey: Implications for tectonic evolution of the Aegean extended region, Tectonophysics. Ertek, T. A., and A. Aytaç (2001), “Karadeniz Kıyılarımızda Denizel Taraçaların Korelasyonu,” (Correlations of the marine terraces along the Black Sea Coast of Turkey) ̇ (in  Turkish), Türkiye Kuvaterneri Çalıştayı, Istanbul, Türkiye, 21–22 Mayıs 2001, ss.70–74. Eyal, Y. (2011), The Syrian Arc Fold System: Age and rate of folding, Geophys. Res. Abstr. 13, EGU 2011‐7401, 2011 EGU General Assembly 201. Eyidoğan, H., and J. A. Jackson (1985), A seismological study of normal faulting in the Demirci, Alaşehir and Gediz earthquake of 1960–1970 in western Turkey: Implications for the nature and geometry of deformation in the continental crust, Geophys. J. R. Astron. Soc., 81, 569–607. Facenna, C., O. Bellier, J. Martinod, C. Piromallo, and V. Regard (2006), Slab detachment beneath eastern Anatolia: A possible cause for the formation of the North Anatolian fault, Earth Planet. Sci. Lett., 242, 85–97. Fernández‐Blanco, D., G. Bertotti, T. Cassola, and S. Willett (2012), Neogene vertical tectonics of the south margin of the Central Anatolia plateau in relation to Cyprus Arc subduction, Abstract T33B‐2663, presented at 2012 Fall Meeting, American Geophysical Union, San Francisco, CA, and 3–7 December. Fernández‐Blanco, D., G. Bertotti, and A. Çiner (2013), Cenozoic tectonics of the Tuz Gölü basin (central Anatolia

Morphotectonic Development of Anatolia and the Surrounding Regions  81 Plateau, Turkey, in Late Cenozoic Evolution of the Central Anatolian Plateau, edited by A. Çiner, M. Strecker, and G. Bertotti, 715–738, J. Earth Sci., 2. Fichtner, A., E. Saygin, T. Taymaz, P. Cupillard, Y. Capdeville, and J. Trampert (2013), The deep structure of the North Anatolian fault zone, Earth Planet. Sci. Lett., 373, 109–117. Forte, E., M. Sugan, A. Del Ben, L. Pıpan, L. Gasperini, and H. Kurt (2014), Multidisciplinary analyses to understand tectonic activity and evolution of the North Anatolian Fault in the ̇ Hersek peninsula (Izmit Gulf, Turkey), Bolletino di Geofisica Teorica Applicata, 55, 589–616. Gans, C. R., S. L. Beck, G. Zandt, C. B. Biryol, and A. A. Özaçar (2009), Detecting the limit of slab break‐off in central Turkey: New high‐resolution Pn tomography results, Geophys. J. Int., 179, 1566–1577. Garfunkel, Z. (1981), Internal structure of the Dead Sea leaky transform (rift) in relation to plate kinematics, Tectonophysics, 80, 81–108. Garfunkel, Z. (1997), History of the Dead Sea basin, in Structure and Tectonics of the Dead Sea Basin (The Dead Sea, The Lake and Setting), edited by T. M. Niemi, Z. V. Avraham, and J. R. Gat, 36–56, Oxford Monographs on Geology and Geophysics No. 96, Oxford Univ. Press. Garfunkel, Z., I. Zak, and Z. Ben‐Avraham (1996), The structure of the Dead Sea basin. Tectonophysics, 266, 155–176, 8042. Garfunkel Z., Z. Ben‐Avraham, and K. Alisa (2014), Dead Sea Transform Fault System: Reviews. Modern Approaches in Solid Earth Sciences. Springer. 365 p Gasperini, L., A. Polonia, F. Del Bianko, P. Favali, G. Marinaro, and G. Etiope (2012), Cold seeps, active faults and the earthquake cycle along the North Anatolian Fault system in the Sea of Marmara (NW Turkey) Bolletino di Geofisica Teorica ed Applicata 53, 371–384. Gessner, K., U. Ring, C. Johnson, R. Hetzel, C. W. Passchier, and T. Gungör (2001), An active bivergent rolling‐hinge detachment system: Central Menderes metamorphic core complex in western Turkey, Geology, 29, 611–614; doi: 10.1130/0091‐7613(2001)0292.0.CO;2. Girdler, R. W. (1990), The Dead Sea transform fault system, Tectonophysics, 180, 1–13. Göğüş, O. H., and R. N. Pysklywec (2008), Mantle lithosphere delamination driving plateau uplift and synconvergent extension in eastern Anatolia, Geology, 36, 723–726. Gök, R., E. Sandvol, N. Türkelli, D. Seber, and M. Barazangi (2003), An attenuation in the Anatolian and Iranian plateau and surrounding regions, Geophys. Res. Lett., 30(24). Gök, R., M. E. Pasyanos, and E. Zor (2007), Lithospheric structure of the continent‐continent collision zone: eastern Turkey, Geophys. J. Int., 169, 1079–1088. Gökaşan, E., H. Tur, B. Ecevitoğlu, T. Görüm, A. Türker, B. Tok, and H. Birkan (2006), Istanbul Boğazı deniz tabanı morfolojisini denetleyen etkenler: son buzul dönemi sonrası aşınma izlerinin kanıtları (Factors controlling the sea floor morphology of the Strait of Istanbul: Evidences of an erosional event after last glacial maximum), Yerbilimleri, 27, 143–161. Gökaşan, E., T. Ustaömer, C. Gazioğlu, Z. Y. Yücel, K. Öztürk, H. Tur, B. Ecevitoğlu, and B. Tok (2003), Morpho‐tectonic evolution of the Marmara Sea inferred from multi‐beam bathymetric and seismic data, Geo‐Mar. Lett., 23, 19–33.

Gökçe, O., K. Tüfekçi, and Ş. Gürboğa (2014), Yüzey faylanması tehlikesinin değerlendirilmesi ve fay sakınım bantlarının oluşturulması. Başbakanlık Afet ve Acil Durum Yönetimi Başkanlığı (AFAD), Ankara. Gökten, E., K. Gilbert, and M. Meydan (2013), The kinematic significance of rotation‐related deformation features in a fault‐defined wedge associated with the North Anatolian Fault, central Turkey, J. Geodinamica, 65, 228–243. Görür, N. (1988a), Triassic to Miocene Paleogeography Atlas of Turkey, MTA Publication, Ankara. Görür, N. (1988b), Timing of opening of the Black Sea Basin, Tectonophysics, 147, 247–262. Görür, N., and Ş. B. C. Elbek (2013), Tectonic events responsible for shaping the Sea of Marmara and its surrounding region, Geodinamica Acta, 26, 1–11. Görür, N., A. M. C. Şengör, R. Akkök, and Y. Yılmaz (1983), Pontidler’de NeoTetisin kuzey kolunun açılmasına ilişkin sedimentolojik veriler (Sedimentological data on the opening of the northern branch of the Neo‐Tethys in the Pontides), Türkiye Jeoloji Kurumu Bült., 26, 11–20. Görür, N., F. Oktay, İ. Seymen, and A. M. C. Şengör (1984), PaleoTectonic evolution of the Tuzgölu Basin complex, Central Turkey: Sedimentary record of a neo‐Tethyan ­closure, Geological Society, London, Special Publications, 17, 467–482. Görür, N., N. Çağatay, M. Sakınç, K. Sümengen, C. Şentürk, C. Yaltırak, and A. Tchepalyga (1997), Origin of the Sea of Marmara as deduced from the Neogene to Quaternary paleo‐ geography, Int. Geol. Rev., 39, 342–352. Görür, N., O. Tüysüz, and A. M. C Şengör (1998), Tectonic evolution of the central Anatolian basins, Int. Geol. Rev., 40, 831–850. Güney, A. Y. Yılmaz, E. Demirbağ, B. Ecevitoğlu, S. Arzuman, and I.̇ Kuşçu (2002), Reflection seismic study across the continental shelf of Baba Burnu promontory of Biga Peninsula, NW Turkey, Marine Geol., 176, 75–85 Gürer, O. F., M. Bozcu, K. Yılmaz, and Y. Yılmaz (2001), Neogene basin development around Söke‐Kuşadasi (western Anatolia) and its bearing on tectonic development of the Aegean region, Geodinamica Acta, 14, 57–69. Gürer, O. F., and Y. Yılmaz (2002), Geology of the Ören and surrounding areas SW Anatolia, Turkey J. Earth Sci., 11, 1–13. Gürer, F., E. Sangu, M. Özburan, and H. Sinir (2016), PlioQuaternary kinematic development and paleo stress pattern of the Edremit Basin, Western Turkey, Tectonophysics, doi:10.1016/tecto.2016.05.007. Gürsoy, H., O. Tatar, J. D. A. Piper, A. Heimann, and L. Mesci (2003), Neotectonic deformation linking the east Anatolian and Karataş‐Osmaniye intracontinental transform fault ̇ zones in the Gulf of Iskenderun, southern Turkey, deduced from paleomagnetic study of the Ceyhan‐Osmaniye volcanics, Tectonics, 22, 1067; doi: 10.1029/2003TC001524. Gürsoy, H., O. Tatar, J. D. A. Piper, F. Bulut, Z. Akpinar, B. Huang, A. P. Roberts, and B. L. Mesci (2011), Paleomagnetic study of the Kepezdağ and Yamadağ volcanic provinces, central Turkey: Neogene tectonic escape and block definition in the east‐central Anatolides, J. Geod., 51, 308–326. Hakyemez Y., F. Göktaş, and T. Erkal (2013), Gediz Grabeni’nin Kuvaterner Jeolojisi ve evrimi, Türkiye Jeoloji Bülteni, 56.

82  ACTIVE GLOBAL SEISMOLOGY Hall, J., A. E. Aksu, T. J. Calon, and D. Yaşar (2005), Varying tectonic control on basin development at an active micro plate margin: Latakia basin, eastern Mediterranean, Marine Geol., 221, 15–60. Hancock, P. L., and A. A. Barka (1981), Opposed shear senses inferred from Neotectonic mesofracture system in the North Anatolian Fault zone, J. Struct. Geol., 3, 383–392. Harris, B. (1995), The book entitled Russian Style Formation Evaluation, London, Petrophysical Society and Geological Society Publ., 367. Hayward, A. B. (1984a), Sedimentation and basin formation related to ophiolite nappe emplacement, Miocene, SW Turkey, Sed. Geol., 40, 105–129; doi: 10.1016/0037‐ 0738(84)90042‐3. Hayward, A. B. (1984b), Miocene clastic sedimentation related to the emplacement of the Lycian Nappes and the Antalya Complex, S.W. Turkey, in Geological Evolution of the Eastern Mediterranean, edited by J. E. Dixon and A. H. F. Robertson, 287–300, Geological Society Special Publication, 17; doi: 10.1144/GSL.SP.1984.017.01.21. Herece, E., and E. Akay (2003), Kuzey Anadolu Fayı (KAF) Atlası (the North Anatolian Fault Atlas), Maden Tetkik Arama Genel Müdürlüğü Özel Yayın, Ser. 2, Ankara [IV], 2009.05.016. Hetzel, R., U. Ring, C. Akal, and M. Troesch (1995), Miocene NNE directed extensional unroofing in the Menderes Massif, southwestern Turkey, J. Geological Society, 152, 639–664; doi: 10.1144/gsjgs.152.4.0639. n. domain. Hetzel, R. L., O. Romer, and C. W. Candan (1998) Passchier, Geology of the Bozdağ area, central Menderes Massif, SW Turkey: Pan‐African basement and Alpine deformation, Geo. Rundschau, 87, 394–406. Hubert‐Ferrari, A., G. King, J. van der Woerd, I. Villa, E. Altunel, and R. Armijo (2009), Long‐term evolution of the North Anatolian fault: New constraints from its eastern termination, in Collision and Collapse at the Africa‐Arabia‐Eurasia Subduction Zone, edited by D. J. J. van Hinsbergen et  al., 133–154, Geological Society Special Publication, 311. Hubert‐Ferrari, A., R. Armijo, G. King, B. Meyer, and A. Barka (2002), Morphology, displacement, and slip rates along the North Anatolian Fault, Turkey, J. Geophys. Res., 107(B10), 2235. Hüsing, S. K., W. J. Zachariasse, D. J. J. Van Hinsbergen, W. Krijgsman, M. Inceöz, M. Harzhauser, O. Mandic, and A. Kroh (2009), Oligocene‐Miocene basin evolution in SE Anatolia, Turkey: Constraints on the closure of the eastern Tethys gateway, Geological Society, London, Special Publications 2009, 311, 107–132; doi: 10.1144/SP311.4. ̇ Imren, C., X. Le Pichon, C. Rangin, E. Demirbag, B. Ecevitoğlu, and N. Görür (2001), The North Anatolian Fault within the Sea of Marmara: a new interpretation based on multi‐channel seismic and multi‐beam bathymetry data, EPSL, 186, 159–173. Işseven, T., and O. Tüysüz (2006), Palaeomagnetically defined rotations of fault bounded continental blocks in the North Anatolian Shear Zone, North Central Anatolia, J. Asian Earth Sci., 28, 469–479. ̇ Iztan, H., and M. Yazman (1990), Geology and hydrocarbon potential of the Alaşehir (Manisa) area, Western Turkey, IESCA, Int. Earth Science. Congress, On Aegean Regions, ̇ 1–6 October 1990, Izmir, Proc., 1, 327–338.

Jackson, J., and D. McKenzie (1988), The relationship between plate motions and seismic moment tensors, and the rates of active deformation in the Mediterranean and Middle East, Geophys. J., 93, 45–73. Jaffey, N., and A. H. F. Robertson (2001), New sedimentological and structural data from the Ecemiş Fault Zone, southern Turkey: implications for its timing and offset and the Cenozoic tectonic escape of Anatolia, J. Geological Society London, 158, 367–378. Jaffey, N., and A. Robertson (2005), Non‐marine sedimentation associated with Oligocene–Recent exhumation and uplift of the Central Taurus Mountains, S Turkey, Sed. Geol., 173, 53–89. Jolivet, L., C. Facenna, B. Huet, L. Labrousse, L. Le Pourhiet, O. Lacombe, E. Lecomte, E. Burov, Y. Denèle, J. P. Brunn, M. Philippon, A. Paul, G. Salaün, H. Karabulut, C. Piromallo, P. Monié, F. Gueydan, A. Okay, R. Oberhänsli, A. Pourteau, R. Augier, L. Gadenne, and O. Driussi (2013), Aegean tectonics: Strain localization, slab tearing and trench retreat, Tectonophysics, 597–598, 1–33. Kahle, H. G., C. Straub, R. Reilinger, S. McClusky, R. King, K. Hurst, G. Veis, K. Kastens, and P. Cross (1998), The strain rate field in the eastern Mediterranean region, estimated by repeated GPS measurements, Tectonophysics, 294, 237–252. Kahle, H. G., M. Cocard, Y. Peter, A. Geiger, R. Reilinger, A. A. Barka, and G. Veis (2000), GPS‐derived strain rate field within the boundary zones of the Eurasian, African, and Arabian Plates, J. Geophys. Res., 105, 23353–23370. Karabacak, V., and E. Altunel (2013), Evolution of the northern Dead Sea Fault Zone in southern Turkey, J. Geod., 65, 282–291. Karabacak V., E. Altunel, M. Meghraouı, and H. Akyüz (2010), Field evidences from northern Dead Sea Fault Zone (South Turkey): New findings for the initiation age and slip rate, Tectonophysics, 480, 172–182. Kaya, E. G., S. Ünay, S. Saraç, A Eichhorn, A. Hassenruck, S. Knappe, Pekdeğer, Mayda, and Halitpaşa (2004), Trans­ pressive zone: Implications for an Early Pliocene compressional phase in central western Anatolia, Turkey, Turkish J. Earth Sci., 13, 1–13. Kazancı, N., S. Leroy, Ö. Ileri, Ö. Emre, M. Kibar, and S. Öncel (2004), Late Holocene erosion in NW Anatolia from sediments of Lake Manyas, Lake Ulubat and the southern shelf of the Marmara Sea, Turkey, Catena, 2, 277–308. Keskin, M. (2003), Magma generation by slab steepening and breakoff beneath a subduction accretion complex: an alternative model for collision‐related volcanism in Eastern Anatolia, Turkey, Geophys. Res. Lett., 30(24), 8046, http:// dx.doi.org/ 10.1029/2003GL018019. Keskin, M. (2007), Eastern Anatolia: A hotspot in a collisional zone without a mantle plume, in Plates, Plumes, and Planetary Processes, edited by G. R. Foulger and D. M. Jurdy, 693–722, Geological Society of America Special Papers, 430,. Keskin, M., A. V. Chugaev, A. Lebedev, V. Sharkov, V. Oyan, and O. Kavak (2012), The geochronology and origin of mantle sources for late Cenozoic intraplate volcanism in the frontal part of the Arabian Plate in the Karacadağ Neovolcanic area of Turkey. Part I, The result of isotope‐geochronological studies, J. Volcanol. Seis., 6352–360.

Morphotectonic Development of Anatolia and the Surrounding Regions  83 Keskin, S., K. Pedoja, and O. Bektaş (2011), Coastal uplift along the eastern Black Sea coast: New marine terrace data from eastern Pontides, Trabzon (Turkey) and a review, J. Coast. Res., 27, 63–73 Ketin, I.̇ (1966), Tectonic units of Anatolia, Bull. Mineral Res. Explor. Inst. Turkey, 66, 23–34. Kissel, C., O. Averbuch, D. Frizon de LaMotte, O. Monod, and S. Allerton (1993), First paleomagnetic evidence for a post‐ Eocene clockwise rotation of Western Taurides thrust belt east of the Isparta reentrant (Southwestern Turkey), Earth Planet. Sci. Lett., 117, 1–14. Koçyiğit, A. (1989), Suşehri basin: An active fault‐wedge basin on the North Anatolian fault zone, Turkey, Tectonophysics, 167, 13–29. Koçyiğit, A. (1990), Tectonic setting of the Gölova basin, total offset of the North Anatolian fault zone, eastern Pontide, Turkey, Ann. Tectonic., 4, 155–170. Koçyiğit, A. (1996), Superimposed basins and their relations to the recent strike‐slip fault zone: a case study of the Refahiye superimposed basin adjacent to the North Anatolian transform fault, northeastern Turkey, Int. Geol. Rev., 38, 701–713. Koçyiğit, A. (2000), Güneybatı Türkiye’nin Depremselliği, in: BADSEM 2000–Batı Anadolu ‘nun Depremselliği ̇ Simpozyumu, Proceedings, 24–27 Mayıs 2000, Izmir, 2000, 30–39 (in Turkish with English abstract). Koçyiğit, A. (2005), The Denizli graben‐horst system and the eastern limit of western Anatolian continental extension: basin fill, structure, deformational mode, throw amount and episodic evolutionary history, SW Turkey, Geodinamica Acta, 18(3–4), 167–208. Koçyiğit, A., and A. Beyhan (1998), A new intracontinental transcurrent structure: The Central Anatolian fault zone, Turkey, Tectonophysics, 284 (3–4), 317–336. Koçyiğit, A., A. A. Özaçar, and M. Cihan (2000b), Batı Anadolu horst graben sisteminin doğu uzantısı ve Isparta açısı ile ilikişi nedir? “Fethiye‐Burdur Zonu” olarak bilinen yapının tektonik niteliği, Tektonik Araştırma Grubu 4, Toplantısı, ATAG‐5, 16–17 Nov., 4–5. Koçyiğit, A., A. Beyhan, K. Dirik, and M. C. Göncüoğlu (1996), Neotectonic characteristics of the Central Anatolia, Int. Geol. Rev., 38, 807–817. Koçyiğit, A., E. Ünay, and G. Saraç (2000a), Episodic graben formation and extensional NeoTectonic regime in west Central Anatolia and the Isparta Angle: A case study in the Akşehir‐Afyon graben, Turkey, in Tectonics and Magmatism in Turkey and the Surrounding Area, edited by E. Bozkurt, J.  A. Winchester, and J. D. A. Piper, 405–421, Geological Society, London, Special Publications, 173. Koçyiğit, A., and F. B. Rojay (1992), Erzincan basin and 1992/3/13–15 earthquakes: An active composite pull‐apart basin on the North Anatolian Fault Zone, Turkey, International Workshop: Work in Progress on the Geology of Türkiye, Keele, England, 9–10 April. Koçyiğit, A., H. Yusufoğlu, and E. Bozkurt (1999a), Evidence from the Gediz graben for episodic two‐stage extension in western Turkey, J. Geological Society, London, 156, 605–616. Koçyiğit, A., H. Yusufoğlu, and E. Bozkurt (1999b), Reply to “Discussion on evidence from the Gediz graben for episodic

two‐stage extension in western Turkey,” J. Geological Society, London, 156 (1999), 1240–1242. Koral, H., J. Kronfeld, N. Avşar, V. Yanko, and J. C. Vogel ̇ (2001), Major recent uplift in Iskenderun Bay, Turkey, Radiocarbon, 43, 2B, 957–963. Kozacı, Ö., J. F. Dolan, R. Finkel, and R. Hartleb (2007), Late Holocene slip rate for the North Anatolian Fault, Turkey, from cosmogenic 36Cl geochronology: implications for the constancy of fault loading and strain release rates, Geology, 35, 867–870. Lei, J., and D. Zhao (2007), Teleseismic evidence for a break‐off subducting slab under Eastern Turkey, Earth Planet. Sci. Lett., 257, 14–28 Le Pichon, X., and J. Angelier (1979), The Hellenic Arc and trench system: A key to the Neotectonic evolution of the eastern Mediterranean area, Tectonophysics, 60, 1–42. Le Pichon, X., and J. Angelier (1981), The Aegean Sea, Phil. Trans. R. Soc. London A, 300, 357–372. Le Pichon, X., A. M. C. Şengör, E. Demirbağ, C. Rangin, C. ̇ Imren, R. Armijo, N. Görür, N. Çağatay, B. Mercier de Lepinay, B. Meyer, R. Saatçılar, and B. Tok (2001), The active main Marmara fault, Earth Planet. Sci. Lett., 192 (4), 595–616. Lericolais, G., et al. (2002), Recent canyon heads evidenced at the Bosporus outlet, EOS Trans. AGU, Fall Meet. Suppl. 83, Abstract, 71B‐0409. Lericolais, G., F. Guichard, A. Morigi, A. Minereau, I. Popescu, and S. Radan (2010), A post younger Dryas Black Sea Regression identified from sequence stratigraphy correlated to core analyses and dating, Quaternary Int., 225, 199–209. Lips, A. L. W., D. Cassard, H. Sözbilir, and H. Yılmaz (2001), Multistage exhumation of the Menderes Massif, western Anatolia (Turkey), Int. J. Earth Sci., 89, 781–792; doi: 10.1007/s005310000101. Lucifora, S., F. Cifelli, F. B. Rojay, and M. Mattei (2013), Paleomagnetic rotations in the Late Miocene sequence from the Çankırı Basin (Central Anatolia, Turkey): The role of strike‐slip tectonics, Turkish J. Earth Sci., 22, 778–792. Lüdecke, T., T. Mikes, B. Rojay, and A. Mulch (2013) Stable isotope‐based reconstruction of Oligo‐Miocene Paleo environmental and Paleohydrology of Central Anatolian lake basins (Turkey), Turkish J. Earth Sci., 22, 793–819. Lüdecke, T., T. Mikes, F. B. Rojay, M. A. Cosca, and A. Mulch (2013), Stable isotope‐based reconstruction of Oligo‐Miocene Paleo environmental and paleohydrology of Central Anatolian lake basins (Turkey), Turkish J. Earth Sci., 22, 793–819. Lybéris, N. (1988), Tectonic evolution of the Gulf of Suez and the Gulf of Aqaba, Tectonophysics, 153, 209–220. Lybéris, N., T. Yürür, J. Chorowicz, E. Kasapoğlu, and N. Gündoğdu (1992), The East Anatolian fault: An oblique collision belt, Tectonophysics, 204, 1–15. Mahmoud, Y., F. Masson, M. Meghraoui, Z. Çakır, A. Alchalbi, ̇ H. Yavaşoğlu, O. Yönlü, M. Daoud, S. Ergintav, and S. Inan (2013), Kinematic study at the junction of the East Anatolian fault and the Dead Sea fault from GPS measurements, J. Geod., 67, 30–39. Masson, F., Y. Mahmoud, Z. Çakır, H. Yavaşoğlu, M. Meghraoui, A. Alchalbi, and S. Ergintav (2010), GPS characterization of the triple junction between Arabia, Africa and Anatolia: New measurements in SE Turkey and in NW Syria,

84  ACTIVE GLOBAL SEISMOLOGY Wegener 2010, 15th General Assembly of Geodinamica Acta, 53, Wegener, Programme and Book of Abstracts, 14–17 September 2010, Geodesy Department of Kandilli Observatory and Earthquake Research Institute of Boğaziçi University, Istanbul, Turkey, 47. Mater, B., M. Uludağ, I.̇ Cürebal, and C. Yıldırım (2003), Uluabat‐ Manyas gölleri ve yakin çevresinin jeomorfolojik gelişim modellemesi. Türkiye Kuvaterner Çalıştayı Bildiriler Kitabı, 29–30 ̇ Mayıs 2003, Avrasya Yerbilimleri Enstitüsü, ITÜ. McClusky, S., S. Balassanian, A. Barka, C. Demir, S. Ergintav, L. Georgiev, O. Gürkan, M. Hamburger, K. Hurst, K. Kahle, K. Kastens, G. Kekelidze, R. King, V. Kotzev, O. Lenk, S. Mahmoud, M. Mishin, M. Nadariya, A. Ouzounis, D Paradissis, Y. Peter, M. Prilepin, R. Reilinger, I. Sanli, H. Seeger, A. Tealeb, M. N. Toksöz, and G. Veis (2000), Global positioning system constrains on plate kinematics and dynamics in the eastern Mediterranean and Caucasus, J. Geophys. Res., 105 (B3), 5695–5719. McKay, G. A., and A. H. F. Robertson (2013), Upper Miocene‐ Pleistocene deformation of the Girne (Kyrenia) Range and Dar Dere (Ovgos) lineaments, northern Cyprus: Role in collision and tectonic escape in the easternmost Mediterranean region, in Geological Development of Anatolia and the Easternmost Mediterranean Region, edited by A. H. F. Robertson, O. Parlak, and U. C. Ünlügenç, 421–445, Geological Society, London, Special Publications, 372. McKenzie, D. (1972), Active tectonic of the Mediterranean region, Geophys. J. R. Astron. Soc., 30, 109–185; doi: 10.1111/j.1365‐246X.1972.tb02351. McKenzie, D., and Y. Yılmaz (1991), Deformation and volcanism in western Turkey and the Aegean, Bull. Tech. Univ. Istanbul, Spec. Issue on Tectonics, 44, 345–373. McKenzie, D. P. (1976), East Anatolian fault: Major structure in eastern Turkey, Earth Planet. Sci. Lett., 29 (1), 189–103. Meade, B. J., B. H. Hager, S. C. McClusky, R. E. Reilinger, S. Ergintav, O. Lenk, A. Barka, and H. Özener (2002), Estimates of seismic potential in the Marmara Sea region from Block models of secular deformation constrained by global positioning system measurements, Bull. Seism. Soc. Amer., 86, 1238–1254. Meriç, E., A. Nazik, N. Avşar, B. Alpar, S. Ünlü, and E. Gökaşan ̇ (2009), Evidences of a possible Marmara Sea‐ Iznik lake connection in Quaternary: determination of ostracods and forȧ minifers in the recent sediments of the Iznik lake (Bursa‐NW Turkey), Istanbul Üniversitesi Yerbilimleri Dergisi, 221–19. Meulenkamp, J. E., M. J. R. Wortel, W. A. Van Wamel, W. Spakman, and E. H. Strating (1988), On the Hellenic subduction zone and the geodynamical evolution of Crete since the late middle Miocene, Tectonophysics, 146, 203–215; doi: 10.1016/0040‐1951(88)90091‐1. Mueller, S., H. G. Kahle, and A. Barka (2007), Plate tectonic situation in the Anatolian‐Aegean region, in Active Tectonics of Northwestern Anatolia‐The Marmara Polyproject: A Multidisciplinary approach by Space‐Geodesy, Hydrology, Geothermics and Seismology, edited by C. Schindler and M. Pfister, 13–28, Vdf Hochschulverlag, Zurich. ̇ Nairn, S. P., A. H. F. Robertson, U. C. Ünlügenç, N. Inan, and K. Taşlı (2012), Tectonostratigraphic evolution of the Upper Cretaceous‐Cenozoic Central Anatolian basins: An integrated model of diachronous ocean basin closure and continental

collision, in Geological Development of Anatolia and the Easternmost Mediterranean Region, edited by A. H. F. Robertson and U. C. Ünlügenç, Geological Society, London, Special Papers, 372; first published online 8 October 2012, http://dx.doi.org/ 10.1144/SP372.9. Okay, A., and I. M. Satır (2000a), Coeval plutonism and metamorphism in a latest Oligocene metamorphic core complex in northwest Turkey, Geological Mag., 37, 495–516. Okay, A. I., E. Demirbağ, H. Kurt, N. Okay, and I.̇ Kuşçu (1999), An active, deep marine strike‐slip basin along the North Anatolian Fault in Turkey, Tectonics, 18(1), 129–147. Okay, A. I., M. Satır, H. Maluski, M. Siyako, P. Monie, R. Metzger, and S. Akyüz (1996), Paleo and Neo‐Tethyan events in northwest Turkey: Geological and geochronological constraints, in The Tectonic Evolution of Asia, edited by A. Yin and M. Harrison, 420–441, Cambridge University Press. Över S., K. Ş . Kavak, O. Bellier, and S. Özden (2004), Is the Amik basin a triple junction area? Analyses of SPOT XS imagery and seismicity, Int. J. Remote Sens., 25, 193857–193872. Över, S., O. Bellier, A., Poisson, J., Andrieux, and Z. Tutkun (1993), Late Cenozoic fault kinematics within basins along the central North Anatolian fault zone (Turkey), C. R. Acad. Sci., Ser. II, 317(6), 827–833. Özaçar, A. A., G. Zandt, H. Gilbert, and S. L. Beck (2010), Seismic images of crustal variations beneath the East Anatolian Plateau (Turkey) from teleseismic receiver functions, in Sedimentary Basin Tectonics from the Black Sea and Caucuses to the Arabian Platform, edited by M. Sosson, N. Kaymakçı, R. A. Stephenson, F. Bergerat, V. Atrostenko, 485–496, Geological Society of London, Special. Publication, 340. Özaçar, A. A., H. J. Gilbert, and G. Zandt (2008), Upper mantle discontinuity structure beneath East Anatolian Plateau (Turkey) from receiver functions, Earth Planet. Sci. Lett., 269, 427–435. Özaksay, V., O. Emre, C. Yıldırım, A. Doğan, S. Özalp, and F. Tokay (2010), Sedimentary record of late Holocene seismicity and uplift of Hersek restraining bend along the North Anatolian ̇ Fault in the Gulf of Izmit, Tectonophysics, 487, 33–45. Özbakır, A., A. M. C. Şengör, M. J. R. Wortel, and R. Govers (2013), The Pliny‐Strabo trench region: A large shear zone resulting from slab tearing, Earth Planet. Sci. Lett., 375, 188–195. Özdoğan, T., O. Kaya, I. Açıkbaş, I.̇ Bahtiyar, and M. Siyako (2011), The Miocene Lice basin of southwestern Turkey: An example of a shallow marine to non marine foreland basin, Achaean to Anthropocene, 2011, Geolog. Soc. Am. Annual Meeting and Exposition Abstracts, 146–148. Özeren, M. S., and W. E. Holt (2010), The dynamics of the eastern Mediterranean and eastern Turkey, Geophys. J. Int., 183(3), 1165–1184. Özkaymak, C., H. Sözbilir, and B. Uzel (2013), Neogene‐ Quaternary evolution of the Manisa basin: Evidence for variation in the stress pattern of the Izmir‐Balıkesir transfer zone, western Anatolia, J. Geod., 65 (2013), 117–135. Özşahin, E., and D. Ekinci (2014), How was the Anatolian side of Istanbul formed? A geomorphologic assessment (NW Turkey), Procedia‐Soc. Behav. Sci., 120(19), 404–413. Özsayın, E. A., and K. Dirik (2007), Quaternary activity of the ̇ Cihanbeyli and Yeniceoba Fault zones, Inönü‐Eskiş ehir fault system, central Anatolia, Turkish J. Earth Sci., 16, 471–492.

Morphotectonic Development of Anatolia and the Surrounding Regions  85 Özsayin, E., and K. Dirik (2011), The role of oroclinal bending in the structural evolution of the Central Anatolian Plateau: Evidence of a regional changeover from shortening to extension, Geol. Carpath., 62(4), 345–359. Özsayın, E., A. Çiner, B. Rojay, K. Dirik, D. Melnick, D. Fernández‐Blanco, G. Bertotti, T. F. Schildgen, Y. Garcin, M. R. Strecker, and M. Sudo (2013), Plio‐Quaternary extensional tectonics of the Central Anatolian Plateau: A case study from the Tuz Gölü basin, Turkey, Turkish J. Earth Sci., 22, 691–714. Öztürk, S., M. Destur, and I. M. Karl (2006), Heat Flow Map of Turkey, 1: 2 000 000, General Directorate of Mineral Research and Exploration, Department of Geophysical Exploration, Ankara, Turkey. Papanikolaou, D. M., M. Alexandri, and P. Nomikou (2006), Active faulting in the North Aegean Basin, in Post Collisional Tectonics and Magmatism in the Mediterranean Region and Asia, edited by Y. Dilek and S. Pavlides, 189–210, Geological Society of America Special Papers, 409. Papanikolaou, D. M., M. Alexandri, A. P. Nomikou, and D. Ballas (2002), Morphotectonic structure of the western part of the North Aegean basin based on swath bathymetry, Marine Geol., 190, 465–492. Parlak, O., V. Hoeck, H. Kozlu, and M. Delaloye (2004), Oceanic crust generation in an island arc tectonic setting, SE Anatolian orogenic belt (Turkey), Geological Mag., 141‐583‐603. Parsons, S., Toda, R. S. Stein, A. Barka, and J. H. Dietrich (2000), Heightened odds of large earthquakes near Istanbul: An interaction‐based probability calculation, Science, 288, 661–665. Parsons, T. (2004), Recalculated probability of M 7 earthquakes beneath the Sea of Marmara, Turkey, J. Geophys. Res., 109, B05304; doi: 10.1029/2003JB002667. Paton, S. M. (1992), Active normal faulting drainage patterns and sedimentation in southwestern Turkey, J. Geological Society London, 149, 1031–1044. Pearce, J. A., J. F. Bender, S. E. De Long, W. S. F. Kidd, P. J. Low, Y. Güner, F. Şaroğlu, Y. Yılmaz, S. Moorbath, and J. G. Mitchell (1990), Genesis of collision volcanism in eastern Anatolia, Turkey, J. Volcanol. Geotherm. Res., 44, 189–229. Perinçek, D., and I. Çemen (1990), The structural relationship between the East Anatolian Fault and Dead Sea fault zones in southern Turkey, Tectonophysics, 172, 331–340. Peters, J. M., and W. J. Huson (1985), The Pliny and Strabo trenches (eastern Mediterranean): Integration of seismic reflection data and Sea Beam bathymetry maps, Marine Geol., 64, 1–17; doi: 10.1016/0025‐3227(85)90157‐4. Peyret, M., F. Masson, H. Yavaşoğlu, S. Ergintav, and R. Reilinger (2013), Present‐day strain distribution across a segment of the central bend of the North Anatolian Fault Zone from a Persistent‐Scatters InSAR analysis of the ERS and Envisat archives, Geophys. J. Int., 192, 929–945. Philippon, M., J. P. Brunn, F. Gueydan, and D. Sokoutis (2014), The interaction between Aegean back‐arc extension and Anatolia escape since Middle Miocene, Tectonophysics, 631, 176–188. Piper, J. D. A., H. Gürsoy, O. Tatar, M. E. Beck, A. Rao, F. Koçbulut, and B. L. Mesci (2010), Distributed Neotectonic deformation in the Anatolides Turkey: A paleomagnetic analysis, Tectonophysics, 488, 31–50.

Piper, J. D. A., O. Tatar, H. Gürsoy, B. Koç, F. Bulut, and B. L. Mesci (2006), Paleomagnetic analysis of NeoTectonic deformation in the Anatolian accretionary collage, Turkey, in Post Collisional Tectonics and Magmatism in the East Mediterranean Region, edited by Y. Dilek and S. Pavlides, 417–439, Geological Society of America Special Papers, 409. Pirazzoli, P. A., J. Lael, J. Borel, J. F. Saliége, O. Erol, I. Kayan, and A. Persol (1991), Holocene raised shorelines on the Hatay coasts Turkey: Palaeoecological and tectonic implications, Marine Geol., 96, 295–311. Piromallo, C., and A. Morelli (2003), P wave tomography of the mantle under the Alpine‐Mediterranean area, J. Geophys. Res., 108, 1–23. Poisson, A., R. Wernli, E. K. Sagular, and H. Temiz (2003), New data concerning the age of the Aksu thrust in the south of the Aksu valley, Isparta Angle (SW Turkey): Consequences for the Antalya Basin and the eastern Mediterranean, Geological J., 38, 311–327; doi: 10.1002/gj.958. Polat, A., O. Tatar, H. Gürsoy, Ç. Yalçıner, and A. Büyük Saraç (2014), Two‐phased evolution of the Suşehri Basin on the North Anatolian Fault Zone, Turkey, Geodinamica Acta, 25, 132–145. Piromallo, C., and A. Morelli (2003), P wave tomography of the mantle under Alpine‐Mediterranean area, J. Geophys. Res., 108(B2), 2065. Piromallo, C., and V. Regard (2006), Slab detachment beneath eastern Anatolia: A possible cause for the formation of the North Anatolian Fault, Earth Planet. Sci. Lett., 242, 85–97. Price, S., and S. Scott (1994), Fault block rotations at the edge of a zone of continental extension: Southwest Turkey, J. Struct. Geol., 16, 381–392. Purvis, M., A. H. F. Robertson, and M. S. Pringle (2005), 40Ar‐39Ar dating of biotite and sanidine in tuffaceous sediments and related intrusive rocks: Implications for the early Miocene evolution of the Gördes and Selendi basins, W Turkey, Geodinamica Acta, 18, 239–253; doi: 10.3166/ ga.18.239–253. Reilinger, R., S. McClusky, D. Paradises, S. Ergintav, and P. Vernant (2010), Geodetic constraints on the tectonic evolution of the Aegean region and strain accumulation along the Hellenic subduction zone, Tectonophysics, 488, 22–30; doi: 10.1016/j.tecto.2009.05.027. Reilinger, R., S. McClusky, P. Vernant, S. Lawrence, S. Ergintav, R. Çakmak, H. Özener, F. Kadirov, I. Guliev, R. Stepanyan, M. Nadariya, G. Hahubia, S. Mahmoud, K. Sakr, A. ArRajehi, D. Paradissis, A. Al‐Aydrus, M. Prilepin, T. Guseva, E. Evren, A. Dmitrotsa, S. V. Filikov, F. Gomez, R. Al‐Ghazzi, and G. Karam (2006), GPS constraints on continental deformation in the Africa‐Arabia‐Eurasia continental collision zone and implications for the dynamics of plate interactions, J. Geophys. Res., 111, No. B5, B05411. Reilinger, R., S. C. McClusky, M. B. Oral, R. W. King, M. N. Toksöz, A. A. Barka, Kınık, O. Lenk, and I. Sanlı (1997), Global positioning system measurements of present‐day crustal movements in the Arabia‐ Africa‐Eurasia plate collision zone, J. Geophys. Res., 102, 9983–9999. Rızaoğlu, O., V. Parlak, F. Höck, W. E. Koller, Hames, and Z. Billor (2009), Andean‐type active margin formation in the eastern Taurides; Geochemical and geochronological

86  ACTIVE GLOBAL SEISMOLOGY e­vidence from the Baskil Granitoid (Elazığ, SE Turkey), Tectonophysics, 473, 188–207. Robertson, A. H. F., A. Poisson, and O. Akıncı (2003), Developments in research concerning Mesozoic–Tertiary Tethys and NeoTectonic in the Isparta Angle, SW Turkey, Geological J., 38, Spec. Iss. 3–4, 195–234. Robertson, A. H. F., P. D. Clift, Degnan, and G. Jones (1991), Palaeogeographical and PaleoTectonic evolution of the eastern Mediterranean NeoTethys Paleogeography, Paleoecology, 87, 289–343. Robertson, A. H. F., U. C. Ünlügenç, I. Nurdan, and K. Taşlı (2004), The Misis‐Andırın complex: Mid‐Tertiary subduction/accretion and mélange formation related to closure and collision of the South Tethys in southern Turkey, J. Asian Earth Sci., 22, 413–453. Rojay, B., A. Heimann, and V. Toprak (2001), Neotectonic and volcanic characteristics of the Karasu fault zone (Anatolia, Turkey): The transition zone between the Dead Sea transform and the East Anatolian fault zone, Geodinamica Acta, 14, 197–212. Rojay, B., V. Toprak, C. Demirci, and L. Süzen (2005), Plio‐ Quaternary evolution of the Küçük Menderes Graben southwestern Anatolia, Turkey, Geodinamica Acta, 18, 317–331. Rojay, F. B. (1993), Tectonostratigraphy and Neotectonic Characteristics of the Southern Margin of Merzifon‐Suluova Basin (Central Pontides, Amasya), Ph.D. Thesis, Middle East Technical University, Ankara. Rückert‐Ülkümen, N., T. Kowalke, R. Matzke‐Karasz, W. Witt, and E. Yiğitbaş (2006), Biostratigraphy of the ParaThetyan ̇ Neogene at Yalova (Izmit‐Province, NW‐Turkey), Newsletter. Stratig., 42 (1), 43–68. Ryann, W. B. F. (2007), Status of the Black Sea flood hypothesis, 2007, in The Black Sea Flood Question, edited by Yanko‐ Homback et al., 63–88, Springer. Sakınç, M., C. Yaltırak, and F. Oktay (1999), Palaeogeographical evolution of the Thrace Neogene basin and the Tethys– ParaTethys relations at northwestern Turkey (Thrace), Paleogeogr. Paleoclimatol. Paleoecol., 153(1–4), 17–40. Sandvol, E., N. Türkelli, E. Zor, R. Gök, T. Bekler, C. Gürbüz, D. Seber, and M. Barazangi (2003), Shear wave splitting in a young continent‐continent collision: An example from eastern Turkey, Geophys. Res. Lett., 30(24), 2003. Sarıca, N. (2000), The Plio‐Pleistocene ager of Büyük Menderes and Gediz grabens and their tectonic significance on N‐S extensional tectonics in western Anatolia: Mammalian evidence from continental deposits, Geological J., 35, 1–24. Sarıkaya M. A., C. Yıldırım, and A. Çiner (2015), No surface breaking on Ecemiş fault, central Turkey, since Late Pleistocene (~64.5 ka); new geomorphic and geochronological data from cosmogenic dating of offset  alluvial fans, Tectonophysics, 649 (March); doi: 10.1016/j.tecto.2015.02.022. Şaroğlu, F. (1985), Doğu Anadolu ‘nun Neotektonik Dönemde ̇ Jeolojik ve Yapısal Evrimi. Doctorate Thesis, Istanbul ̇ Üniversitesi, Fen Bilimleri Enstitüsü, Istanbul. Şaroglu, F. (1988), Age and offset of the North Anatolian Fault, Middle East Tech. Univ. J. Pure Appl. Sci., 31, 65–79. Şaroğlu, F., and U. Emre (1987), Karacadağ volkanitlerinin genel özellikleri ve Güneydoğu Anadolu otoktonundaki yeri: Türkiye 7. Petrol Kongresi Bildiriler Kitabı, 384–391.

Şaroğlu, F., and Y. Güner (1981), Doğu Anadolu ‘nun j­ eomorfolojik gelişimine etki eden öğeler; jeomorfoloji, ­tektonik, volkanizma ilişkileri: Türkiye Jeoloji, Kur. Bült., 24/2, 39–50.rfr. Şaroğlu, F., and Y. Yılmaz (1984), Doğu Anadolu ‘nun neotektoniği ve ilgili magmatizması: Türkiye Jeol. Kur. Ketin Simpozyumu Bildiriler Kitabı, 149–162. Şaroğlu, F., and Y. Yılmaz (1986), Doğu Anadolu’da neotektoniğin jeolojik gelişime başlıca etkileri, Türkiye Jeoloji Kurultayı 1986 Bildiri Özleri Kitabı, 5. Şaroglu, F., and Y. Yılmaz (1987), Geological evolution and basin models during neotectonic episode in eastern Anatolia, Bull. Mineral Res. Explor., 107, 61–83, Ankara. Şaroğlu, F., and Y. Yılmaz (1991), Geology of the Karlıova region: Intersection of the North Anatolian and the East Anatolian transform faults, Bull. Tech. Univ. Istanbul, Spec. Issue on Tectonics, 44/1, 475–493. Şaroğlu, F., O. Emre, and I. Kuşçu (1992), Türkiye Diri Fay Haritası (1/ 1 000000 scale) Active fault map of Turkey, General Directorate of Mineral Research and Exploration, Ankara. Şaroğlu, F., O. Emre, and I. Kuşçu (2001), Ecemiş Fayı ve deprem potansiyeli (The Ecemiş Fault and earthquake potential of the Ecemiş fault zone), Workshop I Proceedings, Niğde, 20–30, Şaroğlu, F., Y. Güner, W. S. F. Kidd, and A. M. C. Şengör (1980), NeoTectonic of Eastern Turkey: New evidence for crustal shortening and thickening in a collision zone, EOS, Trans. AGU, 61 (17), 360. Schemmel, F., T. Miles, B. Rojay, and A. Mulch (2013), Toward stable isotope paleo‐altimetry of Central Anatolia: A perspective from modern meteoric waters, Amer. J. Sci., 313, 61–80. Schildgen, T. F., C. Yıldırım, D. Cosentino, and M. R. Strecker (2014), Linking slab break‐off, Hellenic trench retreat, and uplift of the central and eastern Anatolian plateaus, Earth‐ Science Rev., 128, 147–168. Schildgen, T. F., D. Cosentino, A. Caruso, A. C. Buchwald, C. Yıldırım, S. A. B. Bowring, H. Rojay, H. Echtler, and M. R. Strecker (2012b), Surface expression of eastern Medi­ terranean slab dynamics: Neogene topographic and structural evolution of the southwest margin of the Central Anatolian Plateau, Turkey, Tectonics, 31, TC2005; doi: 10.1029/2011TC003021, 2012. Schildgen, T., D. Cosentino, and C. Yildirim (2012), Linked uplift of the central and eastern Anatolian plateaus through slab break‐off and upper mantle flow, American Geophysical Union, Fall Meeting, 14b‐04. Schildgen, T. F., D. Cosentino, B. Bookhagen, S. Niedermann, C. Yıldırım, H. P. Echtler, and M. R. Strecker (2012a), Multiphased uplift of the southern margin of the Central Anatolian plateau, Turkey: A record of tectonic and upper mantle processes, Earth Planet. Sci. Lett., 317–318, 85–95. Schindler, C. (1997), Geology of northwestern Turkey: Results of the Marmara Poly‐Project, in active tectonics of northwestern Anatolia, the Marmara Polyproject, in A Multidisciplinary Approach by Space‐Geodesy, Hydrology, Geothermics and Seismology, edited by C. Schindler and M. Pfister, 329–373,Vdf Hochschulverlag, Zurich. Şengör, A. M. C. (1979c), The North Anatolian transform fault: Its age, offset and tectonic significance, J. Geological Society London, 136, 269–282.

Morphotectonic Development of Anatolia and the Surrounding Regions  87 Şengör, A. M. C. (1979a), Türkiye’nin Neotektonik Esasları (Principles of the NeoTectonics of Turkey) (Vol. 2), Ankara: Türkiye Jeoloji Kurumu Yayınları Serisi. Şengör, A. M. C. (1979b), Mid Mesozoic closure of Permo‐ Triassic Tethys and its implications, Nature, 279, 590–593. Şengör, A. M. C. (1987), Cross‐faults and differential stretching of hanging walls in regions of low‐angle normal faulting: Examples from western Turkey, in Continental Extensional Tectonics, edited by M. P. Coward, J. F. Dewey, and P. L. Hancock, 575–589, Geological Society Special Publication, 28, Geological Society, London. Şengör, A. M. C. (1987), Cross‐faults and differential stretching of hanging walls in regions of low‐angle normal faulting: examples from western Turkey, in Continental Extensional Tectonics, edited by M. P. Coward, J. F. Dewey, and P. L. Hancock, 575–589, Geological Society Special Publication, 28, Geological Society, London. Şengör, A. M. C. (1990), A new model for the Late Mesozoic tectonic evolution of Iran and implication of the Oman Region, in  The Geology and Tectonics of the Oman Region, 797–831, Geological Society, London, Special Publications, 49. Şengör, A. M. C., and A. A. Barka (1992), Evolution of escape‐ related strike‐slip systems: Implications for distribution of collisional orogens, 29th International Geological Congress, Kyoto, Japan, Abstracts 1, 232. Şengör, A. M. C., and N. Canıtez (1982), The North Anatolian Fault, in Alpine Mediterranean Geodynamics, edited by H. Berkhemer and K. Hsu, 205–216, AGU. Şengör, A. M. C., and W. S. F. Kidd (1979), The post‐collisional ̇ tectonics of the Turkish–Iranian Plateau and a comparison with Tibet, Tectonophysics, 55, 361–376. Şengör, A. M. C., and Y. Yılmaz (1981), Tethyan evolution of Turkey: A plate tectonic approach, Tectonophysics, 75, 181–241. ̇ Şengör, A. M. C., C. Grall, C. Imren et al. (2014), The geometry of the North Anatolian transform fault in the Sea of Marmara and its temporal evolution: Implications for the development of intracontinental transform faults, Can. J. Earth Sci. 2014, 51(3); doi: 10.1139/cjes‐2013‐0160. ̇ Şengör, A. M. C., G. Uçarkuş, C. Imren, C. Rangin, X. Le Pichon, S. Özeren, and B. Natalin (2011), Broad shear zones and narrow strike‐slip faults in orogens and their role in forming the orogenic architecture: The North Anatolian fault as an active example, Joint Meeting Geo Munich, Fragile Earth, Geological Processes from Global to Local Scales Associated Hazards and Resources, 14‐2, A18, Munich, Germany, 4–7 September. Şengör, A. M. C., M. Satir, and R. Akkök (1984), Timing of tectonic events in the Menderes Massif, western Turkey: Implications for tectonic evaluation and evidence for Pan‐ African basement in Turkey, Tectonics, 3, 693–707. Şengör, A. M. C., N. Görür, and F. Şaroglu (1985b), Strike‐slip faulting and related basin formation in zones of tectonic escape: Turkey as a case study, in Strike‐Slip Deformation, Basin Formation, and Sedimentation, edited by K. T. Biddle and N. Christie‐Blick, 227–264, Dordrecht, Kluwer Academic. ̇ Şengör, A. M. C., O. Tüysüz, C. Imren, M. Sakınç, H. Eyidoğan, N. Görür, X. Le Pichon, and C. Rangin (2005), The North Anatolian fault, a new look, Ann. Rev. Earth Planet. Sci., 33, 37–112.

Şengör, A. M. C., S. Özeren, M. Keskin, M. Sakınç, A. D. Özbakır, and I. Kayan (2008), Eastern Turkish high plateau as a small Turkic‐type orogen: Implications for post‐collisional crust‐forming processes in Turkic‐type orogens, Earth Sci. Rev., 90(1–2), 1–48. Şengör, A. M. C., S. Özeren, T. Genç, and E. Zor (2003), East Anatolian high plateau as a mantle supported, north‐south shortened domal structure, Geophys. Res. Lett., 30 (24), 8045. Şengör, A. M. C., Y. Yılmaz, and I. Ketin (1980), Remnants of a pre‐Late Jurassic ocean in northern Turkey: Fragments of Permian‐Triassic Paleo‐Tethys? Geol. Soc. Am. Bull., 91, 599–609. Şengör, A. M. C., Y. Yılmaz, and I. Ketin (1982), Remnant of Pre‐Late Jurassic Ocean in northern Turkey: Discussion and reply, Geol. Soc. Am. Bull., 93, 929–936. Şengör, A. M. C., Y. Yılmaz, and O. Sungurlu (1985a), Tectonics of the Mediterranean Cimmerides; Nature and evolution of western termination of Paleo‐Tethys, in Geological Evolution of the Eastern Mediterranean, edited by J. E. Dixon and A. H. F. Robertson, 77–112, Geological Society, London, Special Publications, 17. Şengül, E., O. H. Göğüş, R. N. Pysklywec, T. Komut, and G. Houseman (2013), Lithospheric Removal and Dynamic Uplift/Subsidence of the Central Anatolian Plateau, AGU Fall Meeting, AGU Org. Poster T11B.2560. Seghedi, I., C. Helvacı, and Z. Peckskay (2015), Composite volcanoes in the south-eastern part of the I͘ zmir-Balıkesir Transfer Zone, western Anatolia/Turkey, Jour. Volc. Geoth. Res., 291, 72–85, doi.org/10.1016/j. jvolgeores. 2014.12.019. Seyitoğlu, G. (1997), Late Cenozoic Tectono‐sedimentary development of the Selendi and Uşak‐Güre basins: A contribution to the discussion on the development of E‐W and north trending basins in western Turkey, Geological Mag., 134, 163–175. Seyitoğlu, G., and B. Scott (1991), Late Cenozoic crustal extension and basin formation in west Turkey, Geological Mag., 128, 155–166. Seyitoğlu, G., and B. C. Scott (1996), The cause of N‐S extensional tectonics in western Turkey: Tectonic escape vs. back‐ arc spreading vs. orogenic collapse, J. Geod., 22, 145–153. Seyitoğlu, G., and B. Scott (1992), The age of the Büyük Menderes Graben (western Turkey) and its tectonic implications, Geological Mag., 129 (1992), 239–242. Seymen, I.̇ (1975), Kelkit Vadisi kesiminde Kuzey Anadolu Fay ̇ zonunun tektonik özelliği. Doctorate Thesis, Istanbul Teknik Üniversitesi, Maden Fakültesi. Seyrek, A., T. Demir, M. Pringle, S. Yurtmen, R. Westaway, D. Bridgland, A. Beck, and G. Rowbotham (2008), Late Cenozoic uplift of the Amanos mountains and incision of the middle Ceyhan River gorge, southern Turkey, Ar‐Ar dating of the Düziçi Basalt, Geomorphology, 97, 321–355. Seyrek, A., T. Demir, M. S. Pringle, S. Yurtmen, R. Westaway, A. Beck, and G. Rowbotham (2007), Kinematics of the Amanos fault, southern Turkey, from Ar/Ar dating of offset Pleistocene basalt flows: Transpression between the African and Arabian plates, Geological Society, London, Special Publications, 290, 255–284. Seyrek, A., T. Demir, R. Westaway, H. Guillou, S. Scaillet, T. S. White, and D. R Bridgland (2014), The kinematics of

88  ACTIVE GLOBAL SEISMOLOGY central‐southern Turkey and northwest Syria revisited, Tectonophysics, 618, 35–66. Silja, K., I. Hüsing, W. Zachariasse, J. J. Douwe, I. ̇ Vanhinsbergen, I. Woukrijgsman, M. I nceoğ lu, Harzhauser, and K. Andreas (2009), Oligocene‐Miocene basin evolution in SE Anatolia, Turkey: Constraints on the closure of the eastern Tethys gateway, in Collision and Collapse at the Africa‐Arabia‐Eurasia Subduction Zone, edited by D. J. J. Van Hinsbergen, M. A. Edwards, and R. Govers, 107–132, The Geological Society, London, Special Publications, 311,. Siyako, M., I.̇ Bahtiyar, I. Özdoğan, I.̇ Açıkbaş, and O. Kaya (2013), Batman çevresinde mostra veren birimlerin stratigrafisi, TPAO Arama Dairesi Arşivi Teknik Rapor No 5463 (unpublished). Sorlien, C. C., S. D. Akhun, L. Seber, M. Steckler, D. Shillington, H. Kurt, G. Çiftçi et al. (2012), Uniform basin growth over the last 500 ka, North Anatolian Fault, Marmara Sea, Turkey, Tectonophysics, 518–521, 1–16; doi: 10.1016/j, Tecto.2011.10.006. Sözbilir, H., B. Sarı, B. Uzel, O. Sumer, and S. Akkiraz (2011), Tectonic implications of transtensional supradetachment basin development in an extension‐parallel transfer zone: the Kocaçay Basin, western Anatolia, Turkey, Basin Res., 23, 423–448; doi: 10.1111/j.1365‐2117.2010.00496.x.38of40. Spakman, W., and M. J. R. Wortel (2004), Tomographic view on western Mediterranean geodynamics, in The TRANSMED Atlas, The Mediterranean Region from Crust to Mantle, edited by W. Cavazza, F. Roure, W. Spakman, G. M. Stampfli, and P. Ziegler, 31–52. Spakman, W., M. J. R. Wortel, and N. J. Vlaar (1988), The Hellenic subduction zone: A tomographic image and its geodynamic implications, Geophys. Res. Lett., 15(1), 60–63; doi: 10.1029/GL015i001p00060. Stein, R. S. (1999), The role of stress transfer in earthquake occurrence, Nature, 402, 605–609. Stein, R. S., A. A. Barka, and J. H. Dietrich (1997), Progressive failure on the North Anatolian fault since 1939 by earthquake stress triggering, Geophys. J. Int., 128, 594–604. Steinenger, F. F., F. Rogl, and Dermitzakis (1987), Report on the round‐table discussion: Mediterranean and ParaTethys correlations, Annales Institutes Geologicorum Publici Hungaria, 70, 397–421. Steinenger, F. F., W. A. Berggren, D. V. Kent, L. V. Bernor, Ş. Şen, and J. Agustl (1996), Circum‐Mediterranean Neogene (Miocene and Pliocene) marine continental chronology correlations of European mammal units, in The Evolution of Western Eurasian Neogene Mammal Faunas, edited by R. L. Bernor, V. Fahlbusch, and H. W. Mittmann, 7–45, Columbia University Press, New York. Straub, C., and H. G. Kahle (1997), Recent crustal deformation and strain accumulation in the Marmara Sea region, NW Anatolia, inferred from repeated GPS measurements, in Active Tectonics of Northwest Anatolia‐the Marmara Poly‐ Project, edited by C. Schindler and M. Pfister, 417–447, Hochschulverlag AG an der ETH Zürich. Straub, C. S., H‐G. Kahle, and C. Schindler (1997), GPS and geological estimates of the tectonic activity in the Marmara Sea region, NW Anatolia, J. Geophys. Res., 102, 27587–601.

Suess, E. (1901), Das Antlitz der Erde, v. III1 (Dritter Band. Erste Hälfte): F. Tempsky, Prag and Wien, and G. Freytag, Leipzig. ̇ Sümer Ö., U. Inci, and H. Sözbilir (2013), Tectonic evolution of the Söke basin: Extension‐dominated transtensional basin formation in western part of the Büyük Menderes Graben, western Anatolia, Turkey, J. Geod., 65, 148–175. Tan, O., and T. Taymaz (2006), Active tectonics of the Caucasus: Earthquake source mechanisms and rupture histories obtained from inversion of teleseismic body‐waveforms, in Post‐Collisional Tectonics and Magmatism in the Mediterranean Region and Asia, 531–578, Geological Society of America Special Papers, 409; doi: 10.1130/2006.2409(25). Tarı, U., O. Tüysüz, C. Genç et al. (2014), The geology and morphology of the Antakya graben between the Amik triple junction and the Cyprus arc, Geodinamica Acta, 26, Issue 1–2, 27–55. Tatar, O. (1996), Neotectonic structures indicating extensional and contractional strain within Pliocene deposits near the NW margin of the Niksar pull‐apart basin, Turkey, Turkish J. Earth Sci., 5 (1996), 81–90. Tatar, O., F. Poyraz, H. Gürsoy, Z. Çakir, S. Ergintav, Z. Akpınar, F. Koçbulut, F. Sezen, T. Türk, K. O. Hastaoğlu, A. B. Polat, L. Mesci, O. Gürsoy, E. Ayazlı, R. Çakmak, A. Belgen, and H. Yavaşoğlu (2012a), Crustal deformation and kinematics of the eastern part of the North Anatolian fault zone (Turkey) from GPS measurements, Tectonophysics, 518–521, 55–62. Tatar, O., F. Poyraz, H. Gürsoy, Z. Çakir, S. Ergintav, and Z. Akpınar (2012b), Paleomagnetic study of block rotations in the Niksar overlap region of the North Anatolian fault zone, central Turkey, Tectonophysics, 244, 251–266. Tatar, O., J. D. A. Piper, and H. Gürsoy (2000), Paleomagnetic study of the Erciyes sector of the Ecemiş fault zone: Neotectonic deformation in the eastern part of the Anatolian block, in Tectonics and Magmatism in Turkey and the Surrounding Area, edited by E. Bozkurt, J. A. Winchester, and J. D. A. Piper, 423–440, Geological Society, London, Special Publication, 173. Tatar, O., J. D. A. Piper, H. Gürsoy, A. Heimann, and F. Koçbulut (2004), Neotectonic deformation in the transition zone between the Dead Sea transform and the East Anatolian fault zone, southern Turkey: A paleomagnetic study of the Karasu rift volcanism, Tectonophysics, 385, 17–43. Tatar, O., J. D. A. Piper, H. Gürsoy, and H. Temiz (1996), Regional significance of Neotectonic counterclockwise rotation in Central Turkey, Int. Geol. Rev., 38, 692–700. Tatar, O., J. D. Piper, R. G. Park, and H. Gürsoy (1995), Paleomagnetic evidence for block large block rotations in the Niksar overlap area of the North Anatolian Fault Zone, Turkey, Tectonophysics, 244, 251–266. Taymaz, T., H. Eyidoğan, and J. A. Jackson (1991), Source parameters of large earthquakes in the East Anatolian fault zone (Turkey), Geophys. J. Int., Oxford, 106, 537–550. Taymaz, T., J. Jackson, and D. McKenzie (1991), Active ­tectonics of north and central Aegean Sea, Geophys. J. Int., 106, 433–490. Taymaz, T., Y. Yılmaz, and Y. Dilek (2007), The geodynamics of the Aegean and Anatolia: introduction, in The Geodynamics of the Aegean, Anatolia, edited by T. Taymaz, Y. Yılmaz,

Morphotectonic Development of Anatolia and the Surrounding Regions  89 and Y. Dilek, 1–16, Geological Society, London, Special Publications, 291. Tekeli, O., A. Aksay, B. M. Örgün, and A. Işık (1984), Geology of the Aladağ mountains, in Geology of the Taurus Belt, edited by O. Tekeli and M. C. Göncüoğlu, 143–158, Proceedings of the International Tauride Symposium, Mineral Research and Exploration Institute of Turkey (MTA) Publications. Tezel, T., T. Shibutani, and B. Kaypak (2013), Crustal thickness of Turkey determined by receiver function, J. Asian Earth Sci., 75, 36–45. Tirel, C., F. Gueydan, C. Tiberi, and J.-P. Brun (2004), Aegean crustal thickness inferred from gravity inversion, Geodynamical implications, Earth Planet. Sci. Lett., 228, 267–280; doi: 10.1016/j.epsl.2004.10.023. Tuna, D. (1973), VI Bölge Litostratigrafi birimleri adlamasının açıklayıcı raporu: TPAO, Arama Grubu, Rapor No.813. Tunoğlu, C. (1991), Orta Pontidler’de Devrekani Havzası’nın (Kastamonu kuzeyi) litostratigrafi birimleri, Suat Erk Jeoloji Simpozyumu, Bildiriler Kitabı, Ankara Üniversitesi, Fen Fakültesi, Ankara, 183–191. Türkelli, N., E. Sandvol, E. Zor, R. Gök, T. Bekler, A. Al‐Lazki, H. Karabulut, S. Kuleli, T. Eken, C. Gürbüz, S. Bayraktutan, D. Seber, and M. Barazangi (2003), Seismogenic zones in eastern Turkey, Geophys. Res. Lett., 30(24), 8039. Tutkun, S. Z., and P. L. Hancock (1990), Tectonic landforms expressing strain at the Karlıova continental triple junction (E Turkey), Ann. Tectonicae, 4, 182–195. ̇ Tüysüz, O., U. Tarı, S. C. Genç, C. Imren, B. A. B. Blackwell, J. Wehmiller, D. Kaufman, S. Altıok, M. Beyhan, D. Fleitmann, N. Lom, S. Üsküplü, O. Tekeşin, and J. A. Florentine (2013), Geology and morphology of the Antakya Graben between the Amik triple junction and the Cyprus Arc, in Abstract Volume, European Geosciences Union General Assembly 2013, Vienna, Austria, 7–12 April 2013; http://meetingorganizer. copernicus.org/EGU2013/EGU2013‐2064‐2.pdf. Umhoefer, P. J., D. L. Whitney, C. Teyssier, A. K. Fayon, G. Casale, and M. T. Heizler (2007), Yo‐yo tectonics in a wrench zone, central Anatolian fault zone, Turkey, in Exhumation Associated with Continental Strike‐Slip Fault, edited by S. M. Roeske, A. B. Till, J. C. Sample, and D. Foster, 35–57, Geological Society of America Special Papers, 434. Umhoefer, P. J., S. J. Maloney, B. Buchanan, J. R. Arrowsmith, G. Martinez-Gutiérrez, G. Kent, and T. Rittenour (2014), Late Quaternary faulting history of the Carrizal and related faults, La Paz region, Baja California Sur, Mexico, Geosphere, 10(3), 476–504. Ünay, E., and F. Göktaş (1999). Söke Çevresi (Aydın) Geç Erken Miyosen ve Kuvaterner yaşlı küçük memelileri: ön sonuçlar, Türkiye Jeoloji Bülteni, 42, 99–113. Ünay, E., F. Göktaş, H. Y. Hakyemez, M. Avşar, and O. San (1995), Büyük Menderes Grabeni’nin kuzey kenarındaki çökellerin Arvicolidae (Rodentia, Mammalia) faunasına dayalı olarak yaşlandırılması, Türkiye Jeoloji Bülteni, 38, 75– 80 (in Turkish with English abstract). Van Hinsbergen, D. J. J., and S. M. Schmid (2012), Map view restoration of Aegean‐West Anatolian accretion and extension since the Eocene, Tectonics, 31, TC5005; doi: 10.1029/2012TC003132, 2012.

Van Hinsbergen, D. J. J., M. A. Edwards, and R. Govers (2009), Geodynamics of collision and collapse at the Africa‐Arabia‐ Eurasia subduction zone, an introduction, in Collision and Collapse at the Africa‐Arabia‐Eurasia Subduction Zone D, edited by J. J. Van Hinsbergen and R. Govers, 1–7, Geological Society, London, Special Publications, 311. Van Hinsbergen, D. J. J., M. J. Dekkers, E. Bozkurt, and M. Koopman (2010), Exhumation with a twist: Paleomagnetic constraints on the evolution of the Menderes metamorphic core complex (western Turkey). Verge, N. J. (1995), Oligo‐Miocene extensional exhumation of the Menderes Massif, western Anatolia, Terra Abstracts, 7, 117. Walker, R. T. P., M. B. Gans, J. Allen, J. Jackson, N. Khalib, N. Marsh, and M. Zarrinkoub (2009), Late Cenozoic volcanism ̇ and rates of active faulting in eastern Iran, Geophys. J. Int., 177, 783–805. Warren, L., S. L. Beck, C. B. Biryol, A. A. Zandt, Özaçar, and Y. Yang (2013), Crustal velocity structure of central and eastern Turkey from ambient noise tomography, Geophys. J. Int., 194, 1941–1954. Westaway, R. (1994), Present‐day kinematics of the Middle East and eastern Mediterranean, J. Geophys. Res., 99 (1994), 12071–12090. Westaway, R. (2004), Kinematic consistency between the Dead Sea Fault Zone and the Neogene and Quaternary left‐lateral faulting in SE Turkey, Tectonophysics, 391, 203–237. Westaway, R., and J. Arger (1998), Kinematics of the Malatya‐ Ovacık Fault Zone, Third International Turkish Geology Symposium, METU‐Ankara, Abstracts, 197. Westaway, R., H. Guillou, S. Yurtmen, T. Demir, S. Scaillet, and G. Rowbotham (2005), Constraints on the timing and regional conditions at the start of the present phase of crustal extension in western Turkey, from observations in and around the Denizli region, Geodinamica Acta, 18, 209–238. Westaway, R., T. Demir, A. Seyrek, and A. Beck (2006), Kinematics of active left lateral faulting in southeast Turkey from offset Pleistocene river gorges: Improved constraint on the rate and history of relative motion between the Turkish and Arabian Plates, J. Geological Society London, 163, 149–164. Westaway, R. W. C., and J. Arger (2001), Kinematics of the Malatya‐Ovacik fault zone, Geodinamica Acta, 14, 103–131. Westrum, H. H. S. (1962), Das Geologisch Profile Des Aksu Dere Bei Giresun, Bayerische Akademie der Wissenschaften Matematisch‐ Natur. Klasse Abhandlungen Neuefolge, Heft 109, 23–58. Whitney, D. L., C. Teyssier, A. K. Fayon, M. A. Hamilton, and M. Heizler (2003), Tectonic controls on metamorphism, partial melting, and intrusion: Timing and duration of regional metamorphism and magmatism in the Niğde Massif, Turkey, Tectonophysics, 376(1), 37–60. Woodside, J. M., J. Mascle, T. A. C. Zitter, A. F. Limonov, M. Ergun, and A. Volkonskaia (2002), The Florence Rise, the western bend of the Cyprus Arc, Marine Geol., 185, 177–194. Wong, H. K., T. Luddman, A. Uluğ, and N. Görür (1995), The Sea of Marmara: A plate boundary sea in a tectonic escape regime, Tectonophysics, 244, 231–250. Wortel, M. J. R., and W. Spakman (2000), Subduction and slab detachment in the Mediterranean‐Carpathian Region, Science, 290, 1910–1917.

90  ACTIVE GLOBAL SEISMOLOGY Yaltırak, C., and A. Ceyhan (2011), Kuzey Bati Anadolu’da Oligo‐Miyosen çekirdek komplekslerinin (Kazdağ ve Uludağ) ve çevresinin eş zamanli jeolojik evrimi, 64, Türkiye Jeoloji Kurultayı, 25–29, Nisan 2011, Bildiri Özleri. Ankara (Geological Congress of Turkey 2011, Abstract CD). Yazman, M. K., A. Güven, Y. Ermiş, M. Yılmaz, I. Özdemir, Y. Akçay, U. Gönülalan, O. Tekeli, V. Aydemir, A. Sayılı, Z. ̇ Batı, H. Iztan, and O. Korucu (1998), Alaşehir Grabeni’nin ve Alaşehir‐1 prospektinin değerlendirme Raporu, TPAO Exploration Group, unpublished technical report, Ankara. Yegerova, T., E. Baranova, V. Gobarenko, and T. Yanovskaya (2013), Basin architecture and lithosphere structure of west‐ and East Black Sea basins from geophysical studies, Adopted from oral presentation at AAPG Europe. Ann. Conf. Kiev 2010, Search and Discovery Article 30153. Yetiş, C. (1984a), Ecemiş¸ Fay Kuşağının Jeotektonik Evrimi, Yerbilimleri 11, 1–12 (in Turkish with English abstract). Yetiş, C. (1984b), New observations on the age of the Ecemiş Fault, in Geology of the Taurus Belt, edited by O. Tekeli and M. C. Göncüoğlu, Mineral Research and Exploration Institute Special Publication, 1984, 386–394. Yetiş, C., and C. Demirkol (1984), Geotectonic evolution of the Ecemiş fault zone, Hacettepe Univ., Earth Sci., 11, 1–12. Yiğitbaş, E., and Y. Yılmaz (1997), Post‐late Cretaceous strike‐ slip tectonics and its implication on the southeast Anatolian orogen, Turkey, Int. Geol. Rev., 38, 818–831. Yiğitbaş, E., Y. Yılmaz, and S. C. Genç (1992), Güneydoğu Anadolu orojenik kuşağında Eosen nap yerleşmesi (The nappe emplacement on the Southeast Anatolian orogenic belt Arabian platform during the Eocene time) (in Turkish); Türkiye 9, Petrol Kongresi ve Sergisi, Bildiriler kitabı, Türkiye Petrol Jeologları (proceed), 307–319. Yiğitbaş, E., A. Elmas, A. Sefunç, and N. Özer (2004), Major neotectonic features of eastern Marmara region, Turkey: Development of the Adapazarı‐Karasu corridor and its tectonic significance, Geological J., 39, 179–198. Yıldırım, C., D. Melnick, P. Ballato, T. F. Schildgen, H. Echtler, A. E. Erginal, N. G. Kıyak, and M. R. Strecker (2013), Differential uplift along the northern margin of the Central Anatolian Plateau: Inferences from marine terraces, Quaternary. Sci. Rev., 81, 12–28. Yıldırım, C., T. F. Schildgen, H. Echtler, D. Melnick, and M. R. Strecker (2011), Late Neogene orogenic uplift in the central Pontides associated with the North Anatolian fault‐implications for the northern margin of the Central Anatolian Plateau, Turkey. Tectonics, 30, 5005, http://dx.doi.org/ 10.1029/2010TC002756. Yıldırım, M., and Y. Yılmaz (1991), Güneydoğu Anadolu orojenik kuşağının ekaylı zonu (The imbricated zone of the Southeast Anatolian orogenic zone. Turkish with English abstract), Türkiye Petrol Jeologları Derneği Bülteni, 307–319. Yıldız, A., V. Toker, Demircan, and S. Sevim (2003), Paleo environmental interpretation and findings of Pliocene‐Pleistocene nannoplankton, planktic foraminifera, trace fossil in the Mut basin, Yerbilimleri, 28, 123–144. Yılmaz, E., and O. Duran (1997), Güneydoğu Anadolu Bölgesi otokton ve allokton birimler Stratigrafi Adlama Sözlüğü ‘Lexicon‘ TPAO Araştırma Merkezi Grubu Başkanlığı Eğitim Yayınları No: 3.

Yılmaz, H., S. Over, and S. Özde (2006), Kinematics of the east Anatolian fault zone between Türkoğlu (KahramanMaraş) and Çelikhan (Adıyaman), eastern Turkey, Earth Planet Space, 58, 1463–1473. Yılmaz, Y. (1981a), Sakarya Kıtası Güney Kenarının tektonik evrimi (Tectonic Evolution of the Southern Edge of the Sakarya Continent) Istanbul, Yerbilimleri 1(1‐1), 35–52 (Turkish with English abstract). Yılmaz, Y. (1981b), Rift, Alakojen, Impaktojen ve Türkiye’den örnekler: Türkiye Jeoloji Kurumu Konferans Dizisi, 17, 52 s, Ankara. Yılmaz, Y. (1984), Geology of the Amanos mountains, Türkiye Petrolleri Anonim Ortaklıgı Rapor, (1920), 1–4. Yılmaz, Y. (1984a), Geology of Amanos Mountains (Vol. I‐4), T.P.A.O. Rap. No. 1920 (unpublished), Ankara. Yılmaz, Y. (1985), Geology of the Cilo Ophiolite: An ancient ensimatic island arc fragment on the Arabian platform, SE Turkey, Ofioliti, 10, 457–484. Yılmaz, Y. (1990), Allochthonous terrains in the Tethyan Middle East: Anatolia and surrounding regions, Phil. Trans. R. Soc. Lond. A, 331, 661–624 Yılmaz, Y. (1993), New evidence and model on the evolution of the southeast Anatolian Orogen, Geol. Soc. Am. Bull., 105, 251–271. Yılmaz, Y. (1997a), Geology of Western Anatolia, Theme 3, General Geology, in Active Tectonics of Northwestern Anatolia, The Marmara Polyproject, A Multidisciplinary Approach by Space‐Geodesy, Hydrology, Geothermics, and Seismology, Vdf Hochschulverlag, Zurich. Yılmaz, Y. (1997b), Tectonics of the East Anatolian‐Caspian regions, The Leading Edge, 16(6), 889–891. Yılmaz, Y. (2000), Ege Bölgesinin aktif tektoniği, Batı Anadolu ‘nun Depremselliği Simpozyumu (BADSEM 2000), 24–27 ̇ Mayıs 2000, Izmir, Bildiriler, 3–14. Yılmaz, Y. (2002), Tectonic evolution of western Anatolian extensional province during the Neogene and Quaternary, Geolog. Soc. Am., Abstracts with Programs, 34(6), 179. Yılmaz, Y. (2007), Morphotectonic Evolution of the Southern Black Sea Region and the Bosphorus Channel, in The Black Sea Flood Question: Changes in the Coastline, Climate and Human Settlement, edited by V. Yanko‐Homback, A. Gilbert, N. Panin, and P. Dolukhanov, 537–569, Springer, Dordrecht. Yılmaz, Y. (2011), Development of the Southeast Anatolian Orogen, Ipetgaz Symposium, 11–13 May 2011, Abs. Book, Turkish Petroleum Geologists Assoc. Ankara, 98. ̇ Yılmaz, Y., and M. Sakınç (1990), Istanbul Boğazı’nın jeolojik gelişimi üzerine düşünceler. Istanbul Boğazı Güneyi ve Haliç’in Geç Kuvaterner (Holosen) Dip Tortulları, E. Meriç ̇ (ed.), ITÜ Vakfı Yayınları, 99–105. Yılmaz, Y., and M. Yıldırım (1996), Geology of the Nappe region of the southeast Anatolian orogenic belt with particular emphasis on the metamorphic massifs (Güneydoğu Anadolu Orojenik Kuşağında Nap alanının ‘’metamorfik masiflerin” Jeolojisi ve Evrimi (Turkish with an English abstract), Turkish J. Earth Sci. (TUBITAK), 21–38. Yılmaz, Y., and O. F. Gürer (1996), Andırın (K. Maraş) dolayında Misis‐Andırın kuşağının jeolojisi ve evrimi (Geology and evolution of the Misis‐Andırın Belt around the

Morphotectonic Development of Anatolia and the Surrounding Regions  91 Andırın region (Kahraman Maraş), Turkish J. Earth Sci. (TUBITAK), 39–55. Yılmaz, Y., and O. Tüysüz (1984), Kastamonu‐Boyabat‐ Vezirköprü‐Tosya arasındaki bölgenin jeolojisi (Ilgaz‐Kargı Masiflerinin etüdü), MTA Report (unpublished MTA report). Yılmaz, Y., and Z. Karacık (2001), Geology of the northern side of the Gulf of Edremit and its tectonic significance for the development of the Aegean grabens, Geodinamica Acta, 14, 31–40. Yılmaz, Y., A. M. Gözübol, and O. Tüysüz (1982), Geology of an area in and around the northern Anatolian transform fault zone between Bolu and Akyazı; Multidisciplinary Approach to Earthquake Prediction Research, 2, 45–66,Viewig and Son, Braunsweight, Wiesbaden. Yılmaz, Y., O. Gürpınar, and E. Yiğitbaş (1988), Tectonic evolution of the Miocene basins at the Amanos Mountains and the Maraş region, Turkish Association of Petroleum Geologists Bulletin, 1(1), 52–72. Yılmaz, Y., E. Yiğitbaş, and S. C. Genç (1993a), Ophiolitic and metamorphic assemblages of southeast Anatolia, Tectonics, 12 (5), 1280–1297. Yılmaz, Y., F. Şaroğlu, and Y. Güner (1987), Initiation of the neomagmatism in East Anatolia, Tectonophysics, 34, 177–199. Yılmaz, Y., O. Gürpınar, and E. Yiğitbaş (1988), Tectonic evolution of the Miocene basins at the Amanos mountains and the Maraş region, Turkish Assoc. Petrol. Geol. Bull., 1, 52–72. Yılmaz, Y., Y. Güner, and F. Şaroğlu (1998), Geology of the Quaternary volcanic centers of eastern Anatolia, J. Volcanol. Geotherm. Res., 85, 173–210. Yılmaz, Y., C. Genç, F. Gürer, M. Bozcu, K. Yılmaz, Z. Karacık, S. Altunkaynak, A. Elmas (2000), When did the western Anatolian grabens begin to develop? in Tectonics and magmatism in Turkey and Surrounding Area, edited by E. Bozkurt, J. A. Winchester, and J. D. A. Piper, 353–384, Geological Society, London, Special Publications, 173. Yılmaz, Y., E. Gökaşan, and A. Y. Erbay (2010), Morphotectonic development of the Marmara Region, Tectonophysics, 488, 51–70. Yılmaz, Y., E. Yiğitbaş, M. Yildırım, and S. C. Genç (1992), Güneydoğu Anadolu Metamorphic Masiflerinin Kökeni (Origin of the metamorphic massifs of the Southeast Anatolian Orogenic Belt; Turkish with an English abstract), Türkiye 9, Petrol Kongresi Bildiriler kitabı, Proceed. of the 9 Petroleum Ann. Congress of Turkey, 296–307. Yılmaz, Y., E. Yiğitbaş, O. Tüysüz, and A. M. Gözübol (1981), Abant (Bolu)‐Dokurcun (Sakarya) arasında KAF zonunun kuzey ve güneyinde kalan tektonik birliklerin jeolojik evrimi, MTA Report 7085. Yılmaz, Y., H. S. Serdar, C. Genç, E. Yiğitbaş, O. F. Gürer, A. Elmas, M. Yıldırım, M. Bozcu, and O. Gürpınar (1997a), The geology and evolution of the Tokat massif, south‐central Pontides, Turkey, Int. Geol. Rev., 39, 365–382. Yılmaz, Y., O. Gürpınar, H. Kozlu, M. A. Gül, E. Yiğitbaş, M. Yıldırım, C. Genç, and M. Keskin (1985), Kahraman

Maraş kuzeyinin jeolojisi (Andırın‐Berit‐Engizek‐Nurhak‐ Binboğa‐Dağları), Geology of the region to the north of  KahramanMaraş (the Andırın‐Berit‐Engizek‐Nurhak‐ Binboğa Mountains), 3, Unpublished report of TPAO, report 2028, Ankara. Yılmaz, Y., O. Gürpınar, M. Yıldırım, S. C. Genç, A. Elmas, F. Gürer, N. Terzioğlu, and A. Çalişkan (1993b), Tokat Masifi doğusunun jeolojisi ve evrimi, TPAO Arama Grubu (Turkish), Unpublished report 3498. Yılmaz, Y., O. Tüysüz, E. Yiğitbaş, S. C. Genç, and A. M. C. Şengör (1997b), Geology and tectonic evolution of the Pontides, in Regional and Petroleum Geology of the Black Sea and Surrounding Region, edited by A. G. Robertson, 83–226, AAPG Memoir, 68. Yılmaz, Y., S. C. Genç, E. Yiğitbaş, M. Bozcu, and K. Yılmaz (1995), Geological evolution of the late Mesozoic continental margin of Northwestern Anatolia, Tectonophysics, 243, 155–171. Yılmaz, Y., S. C. Genç, O. F. Gürer, Z. Karacık, S.  Altunkaynak, M. Bozcu, K. Yılmaz, and A. Elmas (1999), Ege Denizi ve Ege bölgesinin jeolojisi ve evrimi, edited by N. Görür, 211–337, in Türkiye Denizleri, Devlet Planlama Teşkilatı,Türkiye Bilimsel ve Teknolojik Araştırma Kurumu yayını, Ankara. Yilmaz., Y., S. C. Genç, Z. Karacik, and S. Altunkaynak (2001), Two contrasting magmatic associations of NW Anatolia, and their tectonic significance, J. Geodynamics, 31, 243–271. Yolsal‐Çevikbilen, S., C. B. Biryol, S. Beck, G. Zandt, T. Taymaz, H. E. Adıyaman, and A. A. Özaçar (2012), 3‐D crustal ­structure along the North Anatolian fault zone in north‐ central Anatolia revealed by local earthquake tomography, Geophys. J. Int., 188(3), 819–849; doi: 10.1111/j.1365‐246X. 2011.05313.x. Yönlü, O., E. Altunel, V. Karabacak, S. Akyüz, and C. Yalçıner (2010), Offset archeological relics in the western part of the Büyük Menderes graben (western Turkey) and their tectonic implications, Geological Society of America Special Papers, 471, 30–41. Yurtmen, S., H. Guillou, R. Westaway, G. Rowbotham, and O.  Tatar (2002), Rate of strike slip motion on the Amanos fault (Karasu Valley, southern Turkey) constrained by K–Ar dating and geochemical analysis of Quaternary basalts, ­ Tectonophysics, 344, 207–246. Zankl, H. (1962), Geologisch Lagerstatenkundliche Unterscungen im Ostpontischen Gebirge. Bay. Akad. Wiss. Math‐Nat. Kl. Abh. No. F.109 Munchen. Zor, E., C. Gürbüz, N. Türkelli, E. Sandvol, D. Seber, and M.  Barazangi (2003), The crustal structure of the East Anatolian Plateau from receiver functions, Geophys. Res. Lett., 30(24), 8044; doi: 10.1029/2003GL018192, 2003. Zor, E., E. Sandvol, C. Gürbüz, N. Türkelli, D. Seber, and M.  Barazangi (2003), The crustal structure of the East Anatolian Plateau (Turkey) from receiver functions, Geophys. Res. Lett., 30(24), 8044. http://70.

3 Diversion of River Courses Across Major Strike‐Slip Faults and Keirogens A. M. Celâl Şengör

ABSTRACT Studies along the North Anatolian fault and other major active strike-slip faults and keirogens in the world have revealed complications in river offsets that cannot be explained simply by preexisting slope conditions and capture events. These seem to result from the presence of numerous lesser strike-slip faults parallel with the main displacement zone of a large strike-slip fault and from the structure and topographic evolution of synthetic and antithetic pull-apart basins. Some cuspate pull-apart-basin-bounding normal faults may give the mistaken impression of a river bending into a strike-slip fault because of numerous parallel faults. Other complications result from the presence of structures that predate the formation of a through-going main strike-slip fault. Although fluvial geomorphology is a very powerful tool in investigating strike-slip fault systems, it must be pursued parallel with careful structural mapping lest it prove misleading. 3.1. INTRODUCTION

or phacoids surrounded by the anastomosing branches visibly influence the topography creating whaleback ridges that in places may function as shutter ridges at the mouths of valleys consequent to the drainage of the main fault valley, sag ponds, pull‐apart basins that can be of various sizes and aspect ratios (contrary to the claim of the alleged scale independence of their aspect ratios by Aydın and Nur [1982]), and push‐up ridges that may be simple folds or thrust blocks. Whatever basins form along a strike‐slip fault zone, their floors may assume various slopes, both in direction and amount, depending on the geometry of the down‐dropping fault(s). It is a common observation that river courses do not always turn abruptly into a strike‐slip fault course, but begin bending toward it long before the mapped fault is encountered (Figs. 3.1 and 3.2). Some bend into the fault gently and then abruptly turn into the other direction, in places opposite to what one would expect from the offset of the fault (Fig. 3.3). Such aberrations have been commonly explained by stream capture, but they are also seen where there is no obvious capture to account for them. Some have offsets as expected from the sense of the fault deflecting them, but the observed deflection far exceeds

Diversion of river courses across strike‐slip faults is one of the indicators of sense of movement along the fault and has been used frequently in studies of active and young tectonics along these structures. Such offset indicators are commonly employed along fault strands assumed to be the expressions of single fault surfaces at depth [e.g., Barka and Gülen, 1989; Replumaz et al., 2001; Hubert‐Ferrari et  al., 2002; for textbook treatments, see Burbank and Anderson, 2001, 2012; Delcaillau, 2004; Bull, 2007]. However, no fault, of whatever scale, consists of a single surface. All faults consist of surfaces of slip anastomosing along the strike of a fault zone called “the fault.” When a fault zone is narrow, there is no harm in approximating it as a line in geomorphological studies. However, when its width exceeds a few kilometers (i.e., the North Anatolian shear zone: Şengör et  al. [2005]; Şengör and Zabcı [in press]; or the San Andreas fault system: Wallace [1990]), the motion of individual lozenges I͘TÜ Maden Fakültesi, Jeoloji Bölümü ve Avrasya Yerbilimleri ̇ Enstitüsü Ayazağa, Istanbul, Turkey

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 93

94  ACTIVE GLOBAL SEISMOLOGY

Figure 3.1  The North Anatolian keirogen and the courses of the major Anatolian rivers crossing it. All strike‐slip faults are green; green lines that are heavier represent parts of the active North Anatolian fault. Short, red lines are axial traces of folds that have formed as a consequence of the activity of the North Anatolian keirogen. Yellow patches represent basins that began forming during the medial Miocene, whereas those in gray represent basins that originated during the Quaternary. Red arrows are average trends of fold axial traces in any given area. A = Ankara; B = Bursa; E = Erzincan; I͘ = I͘ stanbul; K = Karlıova; OF = Ovacık fault; SF = Sungurlu fault. Names of rivers (in black). “Wrong” river diversion in the Gölova pull‐apart basin. The North Anatolian fault is right lateral, yet the major river following the highway shows left‐lateral diversion in much of the basin. (Image from Google Earth.) E = Elmalı/Peri Suyu (tributary of the Murat, ancient Arsanias, before the construction of the Keban Dam); F = Filyos (ancient Billaios); K = Kızılırmak (ancient Halys); Ka = Karasu (ancient Teleboas); S = Sakarya (ancient Sangarios); Su = Susurluk (ancient Macestus); Y = Yeşilırmak (ancient Halys).

Figure 3.2  The Karasu River gently bends into the North Anatolian fault east of the Erzincan pull‐apart basin (see Fig. 3.1 for location). (Image from Google Earth.)

the real offset of the fault measured from the displacement of stratigraphic units predating the fault (Fig. 3.4). Others define abrupt meanders while flowing near a major fault that are not seen elsewhere in the near landscape before they enter the fault zone or after they exit it (Fig. 3.5). Experience along the North Anatolian fault in Turkey has shown that such anomalies commonly have tectonic reasons and cannot be explained as products of a normal fluvial development on a stable landscape, as

has become well known all along the fault since the pioneering study by Erinç et al. [1961] around Gerede in the central‐western part of the fault. The purpose of this chapter is to explore some theoretical possibilities of river bends along strike‐slip fault zones. This is a preliminary report of an ongoing study along large active keirogens and for that reason I keep the references to a minimum. A fuller account will be published elsewhere.

Diversion of River Courses Across Major Strike‐Slip Faults and Keirogens  95

Figure 3.3  ‘Wrong’ river diversion in the Gölova pull-apart basin. The North Anatolian Fault is right lateral, yet the major river following the highway shows left-lateral diversion in much of the basin. Image from Google Earth.

Figure 3.4  The very large right‐lateral diversion of the Filyos far exceeding the real offset of the North Anatolian fault here. This immense diversion has been helped by a number of capture events caused by the activity of the North Anatolian fault [Erinç et al., 1961].

3.1.1. Gently Bending Rivers Figure  3.6a shows a river flowing across a series of fault blocks separated by parallel strike‐slip fault strands. There is zero offset along the faults (no fault situation)

and consequently the river exhibits no bends. Figure 3.6b shows a case where all faults have the same amount of right‐lateral offset. Ideally, the river should show the staircase pattern illustrated, but if the fault blocks are sufficiently narrow, then the offset will take the shape of

96  ACTIVE GLOBAL SEISMOLOGY A

D

B

E

C

F

Figure 3.5  The anomalous meanders of the Yeşilırmak north of the Taşova‐Erbaa pull‐apart basin. The Yeşilırmak also displays a remarkable wrong, that is, left‐lateral diversion as it leaves the pull‐apart basin. (Image from Google Earth.)

a single broken line as shown in Figure 3.6c. This pattern is rarely seen in nature, because rarely does such a set of parallel faults have the same amount of displacement along all its members. Figure  3.6d shows, by contrast, a series of faults with decreasing amount of offset toward the top of the page from a maximum. This is most commonly what is seen in nature. A series of faults take up displacement parallel with a main displacement zone. In this situation, the river begins bending as it approaches the main displacement zone (in Fig.  3.6d, simply the strike‐slip strand with the  largest displacement) and then abruptly crosses it (Fig. 3.6e). In the situation shown in Figure 3.6e, the curvature of bending is the largest where the displacement is the least.

Figure 3.6  Theoretical sketches of a river crossing a series of parallel strike‐slip faults: (a) Zero fault situation, in which there is no deflection of the course of the river; (b) all faults have the same sense and amount of offset; probable initial situation; (c) same as (b) except here the river has smoothed its course; (d) strike‐slip fault offsets, systematically, increase toward a main displacement zone (follow the figure); to note the increase in offset initial situation; (e) same as (d) but after the river has smoothed its course; (f) three stages in the production of a sinusoidal river offset across a sone of sliding blocks. This situation is seen in many of the major rivers illustrated in Figure 3.1.

Figure 3.6f shows the evolution of a system of faults whose displacements become larger as a principal ­displacement zone is approached. In such a case, the river will begin bending gently away from the principal displacement zone, will cross it with little apparent bending (because it will exploit the strike‐slip valley created by the principal displacement zone), and will bend again into its old course with a small curvature away from it. When encountered in nature, such a river geometry will betray the breadth of a shear zone the main displacement of which is taken up only by a few faults in its middle or close to its middle, although the bending of the river starts far away from the zone of main displacement because of the smaller parallel faults.

Diversion of River Courses Across Major Strike‐Slip Faults and Keirogens  97

In all the situations considered so far, the river was assumed to approach the shear zone at an angle of 90°. Where this is not the case, one of two situations may obtain: the offset along a fault zone may introduce a sharp bend into the course of the river making the approach angle to the main displacement zone greater, or the offset may make the approach much gentler. I call the first case, illustrated in Figure 3.7, the antithetic approach and the second, the synthetic approach (not illustrated in this chapter). In Figure 3.7a, a series of fault blocks with zero offset (no fault situation) is traversed by a river. In Figure 3.7b, the faults have all moved, with an increasing amount toward a main displacement zone toward the bottom of the page. The river should ideally acquire the serrated map view illustrated. This, however, is rarely seen and the

river generally creates a gently curving valley as shown in Fig. 3.7c and d. The question then becomes how the river does this. Figure  3.7e shows the same serrated map view as in Figure  3.7a and the small box is a portion of its valley illustrated in Figure  3.7f as a block diagram. The two principal valleys have flow directions opposite one another although located on the same consequent river. They would have joined outside the block diagram to allow the continuity of flow from right to left if abrupt turns in the river course would have been allowed. In the case illustrated, however, subsequent valleys have developed with greater gradient and therefore higher erosive power. If they capture each other, they may eventually connect the two segments of the main consequent river within the area of the block diagram and thus make it

A

B E

C F

D

Symmetric water gaps (because idealized)

G

Figure 3.7  Diagrams of an antithetic river crossing a zone of parallel strike‐slip faults: (a) A case of zero offset (no fault), where no offset of the river course is seen. (b) A case in which offset along faults increases toward a main displacement zone (toward the bottom of the page). (c) Same as (b); the serrated river course will likely smooth its way in the direction indicated by the dashed line given appropriate slope conditions. (d) Same as (b), but the river course has now smoothed itself. (e) Same as (b), but here the box shows the upper surface of the block diagram in (f). (f) A portion of the river course shown in the box in (e); here the developing subsequent valleys help to smooth the river course. (g) Same as (f), but here two additional strike‐slip faults cut across the block diagram rediverting the river course. Within a shear zone, many strike‐slip faults of diverse orientations with respect to the main shear zone (R, R’, P, and X shears) may cut across the river course and give it aberrant orientations. Such orientations may help to locate such secondary shears and explain the details of the change of course of the river.

98  ACTIVE GLOBAL SEISMOLOGY

avoid the abrupt turn caused by the strike‐slip faulting. However, there may be many more parallel faults and their effects may be to increase the curvature in the opposite sense as shown in Figure 3.7g. Parallel faults in a fault zone may create a very great complexity of river offsets. While such offsets may be useful in obtaining the sense of offset along major faults, they may very well be misleading in measuring the amount of offset along the entire shear zone, of which the principal displacement zone is only a member and perhaps only the latest member. River offsets in major fault zones must be measured where there is a suspicion that a zone broader than the fault may be affecting the course of a river. But such suspicions should be used only as a guide to where to map in great detail to see the previously unsuspected faults. Along the North Anatolian fault, Figure 3.2 shows some such major faults different from the main displacement zone of the North Anatolian fault. Some missing offsets considered in seismic studies may hide in such as yet unmapped parallel faults, which the rivers may give away. Now let us consider the case of homogeneous deformation along a strike‐slip shear zone, in fact an entire keirogen. Figure  3.8a shows an undeformed rectangle representing the map view of a region across which a river is flowing from upper left to the lower right corner. In Figure 3.8b, the rectangle is sheared right laterally with its upper edge being kept fixed. Notice that the course of the river must not only rotate, but also shorten during the shear. If the river bed had some competence contrast with its surroundings, it might have buckled as shown in Figure 3.8c, but this cannot happen as the river bed cannot have material coherence and such a competence contrast.

Figure 3.8  Diagrams of changes of course imposed on a river in a zone of homogeneous shear: (a) An undeformed zone in which a river flows from upper‐left corner to the lower‐right corner. (b) After a homogeneous distributed shear strain of 30° is imposed on the zone (upper edge is arbitrarily kept fixed). The green arrows show the rotation the river bed must undergo. Notice that the shear imposed must cause a shortening of the river bed along its thalweg. (c) Same as (b), but shows the river bed buckled under shortening. This is an impossible situation, because the river bed has no material coherence and no competence contrast with its surroundings and therefore cannot buckle. (d) The 30° shear shown in (b) has given rise to a series of anticlines that diverted the course of the river. (e) The shear in this case has given rise to a series of normal‐fault‐bounded blocks. Given the indicated slope conditions, the river will be diverted as shown. Note here that not only the presence of the  normal‐fault‐bounded blocks but also the slopes of the tops of the hanging walls determine the geometry of the river diversion. In such a situation, river diversion may be a valuable indicator of block tilting.

So, what would happen? Figure 3.8d shows the formation of a series of anticlines because of the shear and they might deflect the river course as shown depending on the original slope of the area. In Figure  3.8e, a series of parallel normal faults formed as a result of the shear and they deflect the river. In all such cases, whether anticlines or normal fault blocks deflect a river crossing a horizontal shear zone, the deflection will depend on the original slope of the region and the slopes created by the shear‐ related structures. In Figure 3.8e, for example, the slopes on top of the hanging walls of the normal faults are A

B

rot ate of d ori the e riv ntatio er b n ed

γ = 30°

orig ina of t l orien he r iver tation bed

sho rte riv ning er b of ed

C

D

E

Slopes on top of the hanging walls of the normal faults

Diversion of River Courses Across Major Strike‐Slip Faults and Keirogens  99

indicated by brown arrows and the river follows these slopes. In both cases, illustrated in Figure  3.8d and e, I  have ignored the possibility that the river may not change its course at all forming antecedent valleys across  the shear‐related structures.  Tietze [1878] has shown the  presence of such antecedent valleys in the Alborz Mountains in northern Iran, which owe their morphology to transpression [see Şengör, 1990]. 3.1.2. Pull‐Apart Basins and Their Role in Generating Complex River Displacements Rivers enter pull‐apart basins along major strike‐slip fault zones and they may display senses of offset contrary to the offset of the fault, or they may more or less offset in sympathy with the sense of offset of the main fault. If not considered carefully, such river offsets may mislead the student of the strike‐slip fault in question. Figure 3.9a shows a river offset in a pull‐apart basin along a right‐lateral strike‐slip fault. Notice that the floor of the pull‐apart basin dips to the right (i.e., synthetically to the displacement along the fault). I call such basins, whose floors dip in the direction of the sense

Apparent deflection larger than real offset

a

A

Apparent deflection opposite of real offset

B

Figure 3.9  Diagrams of river deflection in synthetic and antithetic pull‐apart basins. (a) River deflection in a synthetic pull‐ apart basin. (b) River deflection in an antithetic pull‐apart basin.

of displacement along the strike‐slip fault that created them, synthetic pull‐apart basins. Because the river shown will flow not only following the offset, but also following the basin floor dip, the apparent river displacement may be more than the displacement along the fault. In this case, the course of the river will be strongly dependent on the preexisting topography and/or the topography created by the strike‐slip/pull‐apart system outside the basin. In some cases, the river may be dammed against the basin margin (at location a, for example) and may create a lake before flowing onward. At that site, a compressional ridge may originate as a consequence of the pull‐apart tectonics depending on what Rodgers [1980] calls separation (distance between two master fault strands on both sides of a pull‐apart basin) and overlap (initial overlap of the two fault segments along their strike; see Şengör [1995], Figs. C, D, E) and contribute to the blocking of the river and may cause further apparent offset complications. Notice that the river in this figure is given an arrow indicating its direction of flow. Now imagine the same river flowing in the opposite direction. The apparent offset it indicates will be the opposite of the one shown in Figure 3.9a. We therefore need to specify the direction of flow of rivers entering synthetic or antithetic pull‐apart basins to evaluate the apparent offset of their courses. The following rules apply: 1. In both synthetic and antithetic pull‐apart basins, their rifts are said to face the direction of the dip of the main normal faults on which the basins opened. 2. In a synthetic pull‐apart basin, a river coming from the left of a person facing in the same direction as the rift (case illustrated in Fig. 3.9a), will show the same sense of offset for the generative strike‐slip fault. 3. In a synthetic pull‐apart basin, a river coming from the right of a person facing in the same direction as the rift (not illustrated) will show the opposite sense of offset for the generative strike‐slip fault. The same rules apply for the antithetic pull‐apart basins in right‐lateral strike‐slip faults. The reader may wish to generate the rules for left‐lateral faults for an exercise. Figure  3.9b shows the case of an antithetic pull‐ apart basin, in which the basin floor dips in the opposite direction of the displacement along the generative strike‐slip system. Here, depending on the dip of the basin floor and the fault separation and overlap, the river may be deflected in the opposite sense to the displacement of the fault. If the basin is filled to the brim, the geomorphologist may be misled by a cursory examination of the river offset. In Figure  3.10, a cuspate normal fault is shown to bound an antithetic pull‐apart basin. There, not only does the river show the “wrong” offset but it also gradually bends into the main strike‐slip fault, creating an

100  ACTIVE GLOBAL SEISMOLOGY

3.2. DISCUSSION AND CONCLUSIONS

impression that a number of parallel faults offset the river course as illustrated in Figure 3.6. Figure 3.11 shows how a “false displacement” of a river in an antithetic pull‐apart basin may grow. As the basin lengthens (from Fig.  3.11a to d), the river may become progressively more offset in the opposite sense to the fault offset. The lighter colored river course is what one would have expected from the fault offset had the pull‐apart basin not been there.

Although fluvial geomorphology is an extremely powerful tool to study strike‐slip fault offsets, the structural evolution of strike‐slip systems locally conspires to mislead the geomorphologist. Numerous parallel faults commonly make up a fault zone and the parallel faults with smaller displacements than that along a main displacement zone may be easy to miss. Every time a river bends into major strike‐slip fault system, one has to look carefully at (1) the slope conditions and (2) the structures around the major fault. These may be inherited or they may be a function of the origin of the major fault. No fault consists of a single surface of displacement. In case numerous faults are very closely bundled (within tens or hundreds of meters of each other), there is little harm in mapping them as a single fault in tectonic studies. But, if strike‐slip fault separations become large reaching kilometers in length scale, then mapping of each fault becomes a necessity to obtain a real displacement of the entire keirogen from rivers since the origin of the rivers, or, if the river predates the fault zone, since the origin of the fault zone. Mistaken assessments of the amount of strike‐ slip motion between the Anatolian Scholle [Şengör and Canıtez, 1982] and the Eurasian Plate have been made in the past, because one did not take into account the displacement along the entire North Anatolian shear zone [Şengör et al., 2005]. Şengör et al. [2005] have shown that the courses of the major Anatolian rivers crossing the North Anatolian shear zone are the most reliable indicators

Apparent deflection opposite of real offset

Figure 3.10  Illustration of river deflection in a pull‐apart basin with a curved surface trace of the bounding normal fault. Regional slope to the upper‐right corner is assumed greater than the slope cause by normal faulting.

A″

A′ A

B

C

D

B′

C′

D′

Figure 3.11  Diagrams showing growth of a false displacement of a river in an antithetic pull‐apart basin. As the basin lengthens (from A to D), the river may become progressively more offset in the opposite sense to the fault offset. The lighter colored river course (A’ to D’) is what one would have expected from the fault offset had the pull‐apart basin not been there.

Diversion of River Courses Across Major Strike‐Slip Faults and Keirogens  101

of its width, and in cases where the rivers are older than the shear zone, the total offset along it. Geometries of pull‐apart basins greatly complicate river offsets along strike‐slip faults and may even generate reverse river offsets without any capture event, or they may greatly exaggerate the estimate of the real fault displacement. A listric normal fault bounding a pull‐apart basin with a curved map trace may give the mistaken impression of a river bending into a major fault because of parallel fault segments (Fig. 3.10) as exemplified in Figure 3.6d and e. Geomorphology done from satellite images must be undertaken with a full knowledge of the 3D geology of a major strike‐slip fault. In the case of major keirogens, the San Andreas system in California or the Alpine fault in New Zealand are particularly prone to oversimplifications if studied only by geomorphology without regard to the underlying structural geology. Such cases have been encountered along the North Anatolian and the East Anatolian faults in Turkey, and this chapter is only a small part of an ongoing study with iterative theoretical and field investigations [e.g., Şengör and Zabcı, in press]. ACKNOWLEDGMENTS I thank Yücel Yılmaz for inviting this preliminary report for this book. The contents of this paper was first presented in a talk given on board the IFREMER research vessel Pourquoi Pas? during a cruise in the Sea of Marmara in November 2014 (Pierre Henry and A.  M. C. Şengör, cochiefs) and I am grateful to my ­audience for constructive comments. REFERENCES Aydin, A., and A. Nur (1982), Evolution of pull‐apart basins and their scale independence, Tectonics, 1, 91–105. Barka, A. and L. Gülen (1989), Complex evolution of the Erzincan basin (eastern Turkey), J. Struct. Geol., 11, 275–283. Bull, W. B. (2007), Tectonic Geomorphology of Mountains: A New Approach to Paleoseismology, Blackwell, Malden. Burbank, D. W., and R. S. Anderson (2001), Tectonic Geomorphology, Blackwell, Malden.

Burbank, D. W., and R. S. Anderson (2012), Tectonic Geomorphology, 2 ed., Blackwell, Malden. Delcaillau, B. (2004), Reliefs Tectonique Récente: Nouveau Précis de Géomorpohologie, Vuibert, Paris. Erinç, S., T. Bilgin, and M. Bener (1961), Gerede cıvarında akarsu şebekesi, I ṡ tanbul Üniversitesi Coğrafya Enstitüsü Dergisi, 6, 90–99. Hubert‐Ferrari, A., R. Armijo, G. C. P. King, B. Meyer, and A.  Barka (2002), Morphology, displacement, and slip rates along the North Anatolian fault, Turkey, J. Geophys. Res., 107, 2235. Replumaz, A., R. Lacassin, P. Tapponnier, and P. H. Leloup (2001), Large river offsets and Plio‐Quaternary dextral slip rate on the Red River fault (Yunnan, China), J. Geophys. Res., 106, 819–836. Rodgers, D. A. (1980), Analysis of pull‐apart basin development produced by en echelon strike‐slip faults, International Association of Sedimentology, Special Publication 4, 27–41. Şengör, A. M. C. (1990), A new model for the late Paleozoic‐ Mesozoic tectonic evolution of Iran and implications for Oman, in The Geology and Tectonics of the Oman Region, edited by A. H. F. Robertson, M. P. Searle, and A. C. Ries, Geological Society, London, Special Publication 49, 797–831. Şengör, A. M. C. (1995), Sedimentation and tectonics of fossil rifts, 53–117,.in Tectonics of Sedimentary Basins, edited by C. J. Busby and R. V. Ingersoll, Blackwell, Oxford. Şengör, A. M. C., and C. Zabcı (in press), The North Anatolian fault and the North Anatolian shear zone, in Landscapes and Landforms of Turkey, edited by C. Kuzucuoğlu and A. Ciner, Springer, Berlin. Şengör, A. M. C., and N. Canıtez (1982), The North Anatolian fault, 205–216, in Alpine‐Mediterranean Geodynamics, edited by K. J. Hsü, Geodynamics Series, 7, American Geophysical Union and the Geological Society of America, Washington, DC. ̇ Şengör, A. M. C., O. Tüysüz, C. Imren, M. Sakınç, H. Eyidoğan, N. Görür, X. Le Pichon, and C. Rangin (2005), The North Anatolian fault: A new look, Ann. Rev. Earth Planet. Sci., 33, 37–112. Tietze, E. (1878), Einige Bemerkungen über die Bildung von Querthälern, Jahrbuch der Kaiserlichen und Königlichen Geologischen Reichsanstalt (Wien), 28, 581–610. Wallace, R. E., ed. (1990), The San Andreas Fault System, California, USGS Professional Paper 1515.

Part II Neotectonics of the Aegean‐Western Anatolian Region

4 Effect of Slab‐Tear on Crustal Structure in Southwestern Anatolia: Insight From Gravity Data Modeling ̇ Rezene Mahatsente, Süleyman Alemdar, and Ibrahim Çemen

ABSTRACT The effect of an upwelling hot asthenospheric flow on the crust and upper mantle structure of southwestern Anatolia is assessed using gravity data modeling along north‑south transects perpendicular to the Hellenic and Cyprus trenches. The density models are based on terrestrial and satellite‑derived gravity data. The results of the gravity modeling, as constrained by results of receiver function and seismic tomography, show that the crust beneath southwestern Anatolia is relatively thin. The crustal thickness above the asthenospheric window, where the subducted African slab exhibits major lateral tears, ranges from 24 to 29 km. The location of the thinned crust coincides with high heat flow of magmatic centers in the Menderes Massif complex. The regions outside the asthenospheric window, however, show by far the largest crustal thickness (3–42 km). This leads to the conclusion that the observed crustal thinning in southwestern Anatolia may be partly attributed to thermal erosion induced by an upwelling hot asthenosphere and extensional tectonics related to the southwest retreating Hellenic trench and westward movement of the Anatolian Plate. 4.1. INTRODUCTION Western Anatolia, Turkey, and surrounding regions have experienced a series of continental collisions from the Late Cretaceous to the Eocene that led to the forma‑ ̇ tion of the Vardar‐Izmir‐Ankara‐Erzincan and Tauride suture zones as part of the Alpine‐Himalayan belt (Fig. 4.1). Following the collision, large‐scale Cenozoic postcollisional continental extension affected western and central Anatolia. This type of tectonic and petrologic succession may be “normal” for parts of collisional mountain belts due to gravitational collapse of overthick‑ ened crust, or, alternatively, extension may be driven by other tectonic forces such as lateral extrusion or subducting slab rollback. Western Turkey contains several prominent structural features indicative of large‐scale continental Cenozoic Department of Geological Sciences, The University of Alabama, Tuscaloosa, Alabama, USA

extensional tectonics. One of these is the Menderes metamorphic core complex (MMCC), which is bordered ̇ by Izmir‐Ankara suture zone to the north and the Lycian nappes to the south. It is separated into the northern, central, and southern segments by two east‐west trending grabens, the Alaşehir and Büyük Menderes grabens (Fig. 4.1). The region exhibits seismicity [Di Luccio and Pasyanos, 2007] and volcanism associated with large‐scale continental extension [Faccenna et  al., 2003; Pe‐Piper and Piper, 2006, 2007; Çemen et  al., 2006; Dilek and Altunkaynak, 2009; Gessner et  al., 2013; Jolivet et  al., 2013; Ersoy et  al., 2014]. The earthquake focal mecha‑ nisms in the region indicate normal faulting. Continuous GPS measurements in the region also indicate present‐day extension [McClusky et  al., 2000 and 2003; Reilinger et al., 2010]. Therefore, it has been accepted that western Turkey has been experiencing Cenozoic extensional ­tectonics and is considered one of the best examples of actively extending terrain in the world. However, the cause and exact timing of the initiation of the Cenozoic

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 105

106  ACTIVE GLOBAL SEISMOLOGY

Figure  4.1  Simplified geologic map of western and central Anatolia, Turkey, and the Aegean region showing major structural features and rock units. The black dashed lines show the locations of the 2.5‐D gravity models. EFZ = Ecemis fault zone; NAF = North Anatolian fault zone; VIAS = Vardar‐Izmir‐Ankara suture zone; FBFZ = Fethiye‐ Burdur strike‐slip fault zone; IA = Isparta angle; MMCC = Menderes metamorphic core complex [from Çemen et al., 2014].

extensional tectonics and associated extensional structures have remained controversial. Four major models have been proposed for the Cenozoic extensional tectonics in western Turkey: 1. The tectonic escape model proposes that the extension is a result of the westward escape of the Anatolian micro‑ plate through the dextral North Anatolian fault zone

(NAFZ) and sinistral East Anatolian fault zone (EAFZ), which were initiated about 5 Ma ago [Dewey and Şengör, 1979; Şengör et al., 1985; Çemen et al., 1999; Figs. 4.1, 4.2]. 2. Back‐arc spreading model suggests that the back‐arc extension related to the Hellenic‐Cyprian trench system produced western Anatolia in the Early Miocene [McKenzie, 1978; Le Pichon and Angelier, 1979; Fig. 4.1].

EFFECT OF SLAB‐TEAR ON CRUSTAL STRUCTURE IN SOUTHWESTERN ANATOLIA  107

Figure 4.2  Geotectonic setting of the Aegean‐Anatolian region on a digital elevation map. The white dashed lines show the locations of the 2.5‐D gravity models. Black arrows indicate the directions of plate motion. Red triangles are for volcanoes. The yellow circles indicate earthquake epicenters (events from 1985 to 2015; http://udim.koeri. boun.edu.tr). Bathymetric data are from ETOPO‐1 global relief model [Amante and Eakins, 2009]. KVF = Kula volcanic field; KAIVF = Kirka‐Afyon‐Isparta volcanic field; AM = Anaximander Mountains; ESM = Eratosthenes seamount; HB = Herodotus basin; CAV = central Anatolian volcanics; LB = Levantine basin.

3. Orogenic collapse model stipulates that the regional extension was initiated by the spreading and thinning of overthickened crust during the Late Oligocene–Early Miocene [Dewey, 1988; Seyitoğlu and Scott, 1996; Dilek and Whitney, 2000]. 4. Three‐stage continuous extension model combines three different extensional models into one model and suggests an uninterrupted three‐stage continuous exten‑ sion since the Late Oligocene [Çemen et al., 2006; Gessner et al., 2013; Ersoy et al., 2014]. The first three models propose different timing for the initiation of the Cenozoic extension in western Turkey. The three‐stage extension model basically combines these mechanisms into a continuous extension effecting western Anatolia since the Late Oligocene (Fig. 4.1). The extension begins in the Late Oligocene–Early Miocene possibly due to orogenic collapse. In Early Miocene to

present, back‐arc spreading and subduction rollback have been affecting the extension; continental escape was probably the cause of extension together with the slab rollback processes since Early Pliocene [Çemen et al., 2006; Gessner et  al., 2013; Ersoy et  al., 2014]. The extension, subsidence, and uplift in western Turkey may be promoted by postorogenic lithospheric removal in continental back‐ arc due to delamination of the mantle lithosphere [Komut et al., 2012; Göğüş, 2014]. Seismic tomography studies in the region indicate a low‐ velocity zone in the upper mantle [Wortel and Spakman, 2000; Piromallo and Morelli, 2003; van Hinsbergen et al., 2010; Biryol et al., 2011; Salaün et al., 2012]. This is inter‑ preted as a slow hot asthenospheric material rising through a vertical slab‐tear in the subducting African Plate [e.g., Spakman et  al., 1993; Piromallo and Morelli, 2003; Chang et al., 2010; Biryol et al., 2011; Salaün et al., 2012].

108  ACTIVE GLOBAL SEISMOLOGY

The slab‐tear (asthenospheric window) and possible detachment of the subducting African lithosphere (slab break off) are attributed to the differential retreat rates between the Hellenic and Cyprus trenches [Govers and Wortel, 2005; Wortel et  al., 2009; Özbakır et  al., 2013]. Other possible causes of slab‐tear are subduction of spreading ridge segment, transform faults, and weakness zones [Wortel and Spakman, 2000]. Southwestern Anatolia contains the Fethiye‐Burdur strike‐slip fault zone (FBFZ) and the compressional and extensional structural features in the Isparta angle (IA) area located to the south and southeast of the MMCC (Fig. 4.1). It has been proposed that the IA and the FBFZ are surficial expression attributed to the presence of the slab‐tear [Barka et al., 1995; Hall et al., 2014]. However, the effect of the slab‐tear, and hence the asthenospheric window on neotectonics, on the crust and upper mantle structure in western and southwestern Anatolia is not well understood. The presence of hot asthenospheric material in the upper mantle may have caused thermal erosion and structural segmentation on the overriding plate. Thus, part of the lithosphere, and hence the crust underlying the southwestern Anatolia region, may be modified by the presence of a partially molten‐buoyant asthenosphere. This is in addition to the observed crustal thinning caused by the extensional tectonics related to the southwest retreating Hellenic trench and westward movement of the Anatolian Plate [Zhu et  al., 2006; Gessner et al., 2013; Karabulut et al., 2013; Tezel et al., 2013]. But, the crustal response to the slab‐tear has not yet been proved. However, the high heat flow [>100 mW m−2; Dolmaz et al., 2005; Aydin et al., 2005; Erkan, 2014] in the Alaşehir and Büyük Menderes grabens may indicate a partially molten asthenospheric material beneath the grabens. The main purpose of this chapter is to examine the effects of the asthenospheric window on major crustal structures such as the MMCC including the Alaşehir and Büyük Menderes grabens in western Turkey and the upper mantle using gravity data modeling. The model is constrained by results from receiver function and seismic tomography [Biryol et al., 2011; Salaün et al., 2012; Tezel et al., 2013].

(Gravity Field and Steady‐State Ocean Circulation Explorer) satellite missions and is based on the WGS84 reference ellipsoid. The surface data over the oceans and lands are from the altimetry‐derived gravity anomalies and Earth Gravitational model (EGM2008), respectively [Andersen and Kundsen, 1998; Pavlis et al., 2012]. There are existing surface gravity data of varying quality in the Anatolian and Aegean regions. These data are included in the EGM2008 gravitational model [Pavlis et al., 2012]. Most of the data are contributed to the National Geospatial‐ Intelligence Agency (NGA) by external organizations or individuals. Satellite‐derived gravity data can provide a global gravity field with a spatial resolution of ~80 km [Förste et al., 2012; Mayer‐Gürr et al., 2012; Yi and Rummel, 2014]. The resolution becomes much higher (~10 km) when satellite‐ derived data are combined with surface gravity data [Pavlis et al., 2012; Förste et al., 2012, 2014]. Combined gravity field models are suitable for lithospheric‐scale gravity modeling and can be used to fill data gaps in regions where no terrestrial gravity data are available [Köther et al., 2012; Hosse et al., 2014; Gutknecht et al., 2014]. However, the quality and resolution of combined gravity‐field models depend on the availability of surface gravity data and topographic reliefs of the study area. Generally, the resolution is higher in regions where high quality surface gravity data are available and topographic relief is moderate [ 5.0). 4.4.2. Results and Discussions Three gravity models, portraying the crust and upper mantle structure of the Aegean‐Anatolian region, are shown in Figure 4.5 (see Figs. 4.1 and 4.2 for locations of the profiles). The models consist of the ocean, sediment, crust, lithospheric mantle, asthenosphere, and slab. The slab contains crustal bodies overlying the oceanic litho‑ sphere and subducts down to a depth of 165 km under the Anatolian and Aegean regions. Each model runs along a longitudinal line from 34°N to 40°N and spans approxi‑ mately 710 km, with a maximum depth of 165 km. The westernmost model (Fig.  4.5a) is located at 24°E and shows the structure of the Hellenic trench and the fore‐arc region. The central model (Fig. 4.5b) is at 29°E and crosses the Menderes Massif complex in western Anatolia, and its northern end coincides with the location of the North Anatolian fault zone (NAFZ). The easternmost model (Fig. 4.5c) is located at 33°E and depicts the deep struc‑ ture of the Cyprus trench and the fore‐arc region. The observed long‐wavelength gravity anomalies in the Aegean and Anatolian regions are well explained in terms of a subducting African slab along the Hellenic and Cyprus trenches (Fig. 4.5a and c) and an upwelling asthe‑ nospheric material beneath western Anatolia (Fig.  4.5b). The slab is continuous at depth for more than 160 km along the Hellenic and Cyprus trenches (Fig. 4.5a and c). However, the tip of the same slab, as determined from the earthquake hypocenters, is limited to a depth of 90 km in southwestern Anatolia (Fig. 4.5b; Boğaziçi University: http://udim.koeri.boun.edu.tr/zeqdb/indexeng.asp). The slab is terminated by an upwelling asthenospheric mate‑ rial between the eastern and western edges of the Hellenic and Cyprus trenches, respectively (Fig. 4.5b). The dip of the slab, as obtained from gravity modeling and

constrained by earthquake hypocenters, is ~15° (Fig. 4.5b), and this is in good agreement with results of surface wave tomography [Salaün et al., 2012]. The depth to the top of the asthenospheric material, as  determined from gravity modeling, ranges from 24 to  29 km below the Menderes Massif complex, and its east‐west dimension is ca. 280 km. The asthenospheric material, as deduced from its density value (3.27 g cm−3) and dimension, is most probably deep in origin (astheno‑ spheric and lithospheric mantle origin). This is consistent with the origin of magma in western Anatolia. The trace element and isotopic compositions of western Anatolian magmatism indicate magmas of asthenospheric and lithospheric mantle origin [Altunkaynak and Dilek, 2006]. The asthenospheric material may be related to the low‐velocity zone in the upper mantle imaged by seismic tomography. There is evidence of low‐velocity zone in the upper mantle (Pn  10 km (with respect to 42.6 km at t = 0) above this hinge since the “slab pull” produced by the hanging slab inherently drives the crustal shortening and thickening. Across the high‐elevation region, the crust thins to 40 km within the plateau gap likely due to the gravitational collapse of the previously thickened crust and/or weakening of the surface crust by heat transfer from the mantle. Note that although there is a plate convergence acting on the side boundary, the crustal response to this effect is not homogenously distributed along the model space. By t = 7.0 m.y., a broader area of mantle lithosphere delaminates/peels away from the crust and the crust

becomes entirely exposed to the hot sublithospheric mantle. Since slab detachment occurs under the delaminating hinge, the surface depression is recovered to less than ~ −500 m (Fig.  5.4b). The position of the surface depression has migrated near the model edge toward the opposite direction of the plate convergence velocity. Previous analogue modeling by Göğüş et al. [2011] finds that the behavior of the delaminating slab is similar to the retreating ocean slab in which the trench position migrates in the reverse direction of the plate convergence velocity. At this time, a larger area of surface is dominated by the plateau‐type elevation and this is interpreted to be the result of ongoing crustal shortening and the substitution of the hot (less dense) and dynamic mantle with the colder mantle lithosphere. Up to 3 km (with respect to t = 0, 42.6 km) of crustal thinning at x = 600 km (in the middle of the plateau) still persists although, again the plate convergence is actively pushing the lithosphere on the side boundary. The topography is

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  127

still considerably high (i.e., plateau elevation is more than 2 km), however, the crust has been thinned compared to its original thickness. Model predictions show that the imposed plate convergence is taken up on the edges of the plateau as the crust thickens to more than 55 km (Fig. 5.4b). The crustal extension that develops contemporaneous with the plate convergence has been suggested for various tectonic regions, especially in the Mediterranean (e.g., Apennines, Aegean, Alboran seas) and the geodynamic cause of such a process is still uncertain. We interpret that the delamination process operated in the terminal stage of the orogen cycle (postcollisonal) may drive anomalous topography and extension [Göğüş, 2015].

5.2.2.2. NUMMODEL‐2 In NUMMODEL‐2, the imposed convergence velocity is increased to Vp = 6 cm/yr and the rest of all the numerical parameters are kept the same as the previous experiment. Such higher plate convergence velocity in the delamination process may be considered as an approximation to the northward motion of the Indian Plate in which the uplift of the Tibetan plateau has been suggested to be driven by a combination of lithosphere delamination and plate‐shortening processes [Bird, 1978; Harrison et al., 1992; Ren and Shen, 2008]. Figure  5.5a shows the geodynamic model evolution after t = 2.9 m.y. and at t = 7.0 m.y. In the earlier time frame, mantle lithosphere delamination has already developed and this is followed by slab break‐off of the

A) Vp = 6 cm/year

B)

Crustal thickness at t = 2.9 m.y Surface topography at t = 2.9 m.y

5

40

moho at t = 0

Slab break-off after delamination

Topography (km)

4

45

3

50

2 55

1

60

0 –1

400

800 1200 x (km)

1600

Crustal thickness (km)

t = 2.9 m.y.

65 2000

t = 7.0 m.y.

40

moho at t = 0

Topography (km)

4

45

3

50

2 55

1

60

0 –1 400

800 1200 x (km)

1600

65 2000

Figure  5.5  (a) Geodynamic evolution of the numerical experiment (NUMMODEL‐2) in which convergence velocity of Vp = 6 cm/yr is imposed on the right boundary of the lithosphere. (b) Plots of surface topography and crustal thickness variation at 2.9 m.y. and 7.0 m.y. (Modified from Göğüş and Pysklywec [2008a]).

Crustal thickness (km)

5

Crustal thickness at t = 7.0 m.y Surface topography at t = 7.0 m.y

128  ACTIVE GLOBAL SEISMOLOGY

hanging mantle lithosphere. Compared to the previous model (with less convergence velocity) for the same time, the delamination and subsequent slab break‐off–detachment occurred more rapidly because higher plate convergence promoted more horizontal forcing onto the delaminating plate, therefore steepening and necking of the lithosphere occurs earlier. The delamination has developed relatively faster in this model and, while a broader area of lithospheric gap was expected to develop by t = 2.9 m.y., the shortening caused by the plate convergence effectively delimited the “delamination space”. For the same time, the maximum surface topography (~4 km) is 2 km more compared with the previous model, and most of the surface crust is now associated with positive surface elevation (Fig. 5.5b). The only localized surface subsidence (−1 km) is at the delaminating hinge where the partial break‐off of the postdelaminating slab occurs. This hinge location, represented by lower elevation, developed as a response to vertical pull through the sinking slab. Model predictions show widespread crustal shortening/thickening along the crustal domain and because of such deformation, the crust does not thin at any location with respect to its initial position at t = 0 (i.e., unlike 3 km of crustal thinning in the previous experiment). =  7.0  m.y., the accelerated plate convergence By t  controls the whole lithospheric realm and the crust ­ associated with the delamination space is relatively more shortened/thickened >55 km across the model. The higher amount of shortening and the asthenospheric mantle upwelling caused by the lithospheric delamination results in distinct plateau-like surface uplift (~5 km) at the lithospheric gap. Maximum crustal thickening is predicted to be as thick as 65 km on both margins of the plateau and the crust starts to thin down to 45 km on both edges of the model boundaries. We note that nowhere in the model space is crustal extension predicted as it is in NUMMODEL‐1. 5.2.2.3. Implications of the Model Results to the Last 13 m.y. Geodynamic Evolution of  the  East Anatolian Plateau The numerical experiments suggest that lithospheric delamination/removal contemporaneous with orogenic processes (i.e., vertical and horizontal tectonics) result in transitory patterns of surface (e.g., uplift/subsidence) and crustal (shortening/extension) anomalies. Furthermore, model predictions contain many of the significant components that can explain the last 13 m.y. geodynamic evolution of East Anatolian plateau by lithosphere delamination/ descent of the oceanic lithosphere in an orogenic setting. The East Anatolian plateau is represented by 2 km high average elevation and is bordered by the Bitlis suture front to the south and the Pontide arc to the north

(Fig.  5.6a). Geological studies suggest that the plateau has attained its high elevation in the last 13 m.y. [Şengör et al., 2003] about the same time as the presumed continental collision between the Arabian and the Eurasian plates [Okay et al., 2010]. Tectonic reconstructions by Şengör and Kidd [1979] suggest that the high plateau is compensated by thick lithosphere following plate shortening. However, seismological and petrological studies interpret that significant portions of the lithosphere have been removed beneath all of present‐day East Anatolia (Fig. 5.6b, c). Specifically, P‐wave tomography models [Al Lazki et al., 2004; Biryol et al., 2011], the Sn wave attenuation [Gök et al., 2007], and multifrequency waveform tomography work by Fichtner et  al. [2013] point out extensive low‐speed anomalies beneath East Anatolia. It has been interpreted that low‐speed anomalies may be induced by asthenospheric upwelling that occurs following the lithospheric removal‐ thinning process. Recent seismological (S‐receiver function) work by Kind et al. [2015] shows that the lithosphere‐asthenosphere boundary beneath all of Anatolia does not go deeper than 80–100 km and this prediction is in good agreement with the mantle upwelling hypothesis. Moreover, the crustal thickness under the plateau was formerly suggested to be as much as 55 km [Şengör and Kidd, 1979], however, receiver function work by Zor [2008], Ozacar et  al. [2008], Vanacore et  al. [2013], and integrated Pn analysis of Moho variation by Kömeç‐Mutlu and Karabulut [2011] suggest relatively thinner crust under the plateau (40–50 km; Fig.  5.6c). Corroborating the seismological interpretations, synthesized petrological work by Pearce [1990] and Keskin [2003] indicates that the widespread distribution of young (last 13 m.y.) volcanics in East Anatolia originated through decompression melting of upwelling asthenospheric mantle. Based on the recent and geological, geophyscial, and petrological observations, a slab steepening and break‐ off geodynamic hypothesis has been suggested by Şengör et al. [2003] and Keskin [2003, 2007] to account for all the anomalous tectonic features presented above (Fig.  5.7). According to the authors, the oceanic lithosphere subducted underneath the Pontides in the north of the East Anatolian plateau until the Oligocene. Following this subduction, the closure of the Neo-Tethyan Ocean occurred and the stacking of East Anatolian accretionary complex developed. Based on the proposed scenario, the northward subducting ocean slab started to peel away (delaminate) from the East Anatolian accretionary complex about 13 m.y. after reaching available buoyancy conditions. The peel‐away process of the ocean slab below the accretionary complex may have led to upwelling of the sublithospheric mantle into shallow crustal regions and widespread magmatism (migrating from north to south) and plateau uplift [Keskin, 2007; Şengör et  al.,

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  129

Figure 5.6  (a) Simplified geological map of the eastern Anatolia region showing Oligocene to recent volcanics. IAS represents Izmir‐Ankara suture zone. (b) Surface topography cross section (B‐B′) along 41°E. (c) Lithospheric scale cross section (B‐B′) along 41°E (see references in the text).

2003, 2008]. Geological evidence suggests the uplift of eastern Anatolia occurred nearly 13–11 m.y. ago [Şengör et al., 2003] in conjunction with the onset of magmatism in calc‐alkaline chemistry around the Erzurum‐Kars plateau (northern section of the plateau) [Keskin, 2003]. Both of these interpretations may be used as a critical time frame for the onset of the lithospheric instability that started about 13 m.y. (Fig. 5.7). In Figure 5.8, we compare our modeled surface topography and crustal thickness variation results for t = 7.0 m.y. against the present‐day surface topography and the Moho variation in the East Anatolian plateau. Results from the modeling indicate that the peel away of the dense

lithosphere (it is the oceanic lithosphere in this geological context) and plate convergence may cause surface uplift as a result of isostatic and dynamic effect of lithospheric removal comparable to the present‐day average elevation of East Anatolia along 42°E (Fig. 5.8a). It has been suggested that East Anatolia emerged from sea level ~11 m.y. ago, a timescale similar to that of modeled delamination events. Note that the short‐wavelength topographic features in the observed profile are related to geomorphologic processes not included in our models. The long‐wavelength plateau uplift of eastern Anatolia is consistent with the removal of mantle lithosphere across a ~500‐km‐wide zone. Comparison of the Moho variation

130  ACTIVE GLOBAL SEISMOLOGY

under the plateau by Zor [2008] against the modeled crust at t = 7.0 m.y. suggests that the crust is relatively thinner across the middle of the plateau; however, only in the latter case is the crust thickened at the plateau flanks. Several factors may account for this. The models do not include material transformations that could result in removal of

the lower parts of the thickened crust [Jull and Kelemen, 2001]. Also it is possible that anomalously thinner crust in the northern part of the Bitlis suture zone may be a result of post-removal of eclogitization of the lower crust (Fig. 5.8b). 5.3. LABORATORY EXPERIMENTS FOR SUBDUCTION TO DELAMINATION In this work, apart from investigating the surface response to delamination by numerical experiments, we have conducted a series of laboratory‐based lithospheric delamination experiments with preexisting ocean subduction. Our motivation is to explore the evolution of the orogenic cycle from subduction to continental delamination. We also test the influence of varying plate convergence forcing along the model boundary during the subduction‐ delamination process and determine how surface topography evolves in this process. Laboratory modeling has been a useful tool to understand the dynamics of subduction (references given in the following section) and it has advantages over numerical modeling since three dimensional aspects of coupled‐crust mantle evolution in the orogenic systems can be investigated. 5.3.1. Model Setup and Scaling Parameters of the Materials

Figure  5.7  Simplified geodynamic evolution of the eastern Anatolia region at t = 13, 10, and 2 m.y. to present [modified from Keskin, 2003, 2007; Şengör et al., 2003, 2008].

(km)

A)

B

B′

3

Observed topography

1

B) 35 (km)

x = 610 km

We used a rectangular Plexiglas tank with 25  cm (width) × 55 cm (length) × 25 cm (depth) and the expriment was scaled as upper mantle depth model of foreland plate (pro-plate) collision/subduction beneath a hinterland plate (retro-plate) (Fig. 5.9). The lowest layer/sublithospheric mantle is modeled by glucose syrup (ρslm = 1425 kg/m3, ηslm = 26.75 Pa s) [Espurt, et al., 2008; Guilliume et al., 2009] (Table 5.2). The 1.3 cm thick proplate mantle lithosphere is modeled by a

45 55

S

Modeled topography

N Observed crustal thickness (Zor, 2008)

Modeled crustal thickness

Figure 5.8  (a) Modeled and observed surface topography at t = 7.0 m.y. for the NUMMODEL‐1. See Figure 5.5 for the line of cross section, B‐B′. (b) Modeled and observed crustal thickness [Zor, 2008] and at t = 7.0 m.y. (Modified from Göğüş and Pysklywec [2008a]).

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  131

Figure  5.9  A 3D illustration of the laboratory model setup showing size of the Plexiglas tank and material domains/dimensions. Table 5.2  Scaling Relationship of the Experiment LABMODEL‐1 Parameter

Nature

LABMODEL‐1

g, gravitational acceleration

(m s ) 9.81 Thickness (m) 80,000 615,385 Lmodel/Lnature = 1.63 × 10−7 Density (kg m−3) 3300 3125 3052 Δρmodel/Δρnature = 0.46 Viscosity (Pa s) 1024 7.64 × 1019 7.42 × 1023 ηmodel/ηnature= 3.5 × 10−19 Characteristic Time (s) 3.16 × 1013 (1 Ma) Characteristic Velocity (m s−1) 3.81 cm yr−1

9.81

h, mantle lithosphere H, upper mantle Scale factor for length ρlit, mantle lithosphere ρslm, sublithospheric mantle ρrp, retro‐plate Scale factor for density ηlit mantle lithosphere ηslm, sublithospheric mantle ηrp, retro‐plate Scale factor for viscosity T, tmodel/tnaturea U, Umodel/Unatureb a

−2

0.013 0.10 1507 1425 1400 3.5 × 10s 26.75 2.6 × 105 148 4.1 × 10−5

 Here tmodel/tnature= (ΔρgL) nature/(ΔρgL)mode × ηmodel/ηnature = 4.72 × 10−12.  Here Umodel/Unature = (tnature × Lmodel)/(tmodel × Lnature) = 3.45 × 104.

b

Newtonian viscous silicone putty (Rhodrosil Gomme, PBDMS + galena fillers) [Funiciello et  al., 2004, 2006; Bellahsen et  al., 2005]. Silicone putty is a viscoelastic material that behaves viscously at experimental strain rates. That is, the experimental timescale, on the order of minutes (min), is always higher than the Maxwell relaxation time (order of seconds) [Weijermars and Schmeling,

1986]. For the experiment (LABMODEL‐1), the density and viscosity of the foreland plate (proplate) mantle lithosphere are ρml = 1507 kg/m3 and ηml = 3.5 × 105 Pa s, respectively (Table  5.2). The lower crust is modeled by glucose syrup (the same material as the sublithospheric mantle) to serve as a decoupling “weak layer” between the crust and the mantle lithosphere a (Fig.  5.9). The

132  ACTIVE GLOBAL SEISMOLOGY

upper crust and retroplate lithosphere are made up of viscous silicone putty (Rhodrosil Gomme, PBDMS + galena fillers) of similar viscosity to the mantle lithosphere of the pro-plate but of lower density (Table 5.2). These materials were used to simplify the stratified temperature‐dependent Earth rheological profile respecting standard scaling procedures of length, density, stress, and viscosity in a natural gravity field as described by Weijermars and Schmeling [1986]. Scaling relationships of the reference experiment (LABMODEL‐1) and general experimental parameters are listed in Table 5.2. The experiments were designed using the physical properties and the length scale of the mantle lithosphere material as the main constraints. For practical purposes, we define a length scale based on a model lithospheric thickness, lm = 0.013 m, which for a natural continental lithospheric thickness ln = 80,000 m gives a length scale ratio L = lm/ln = 1.63 × 10−7, where the n and m refer to nature and laboratory/model scales, respectively. Higher density silicone putty (Rhodrosil Gomme, PBDMS + galena filters) has a density of ρm = 1507 kg/m3, which models a natural density of proplate mantle lithosphere, ρn = 3300 kg/m3. This sets a density scale of P = ρm/ρn = 0.46. In these gravity‐driven experiments, at 1  g the gravity scale ratio is G = gm/gn = 1.0. The rheological properties of all ductile materials used in this study were measured with a concentric cylinder viscometer over a range of strain rates relevant to the experiments [Guillaume et  al., 2009]. The viscosity of dense mantle lithosphere/silicone putty was measured as 3.5 × 10 5 Pa s. Assuming the viscosity of η = 1024 Pa s for the natural oceanic mantle lithosphere defines a viscosity scale ratio M = (η)m/(η)n = 3.5 × 10−19. The timescale ratio for the experiments can now be defined as T = M/PLG = tm/ tn = 4.72 × 10−12. This scaling ratio gives that as one million years corresponds to 148 s in the laboratory model. The sidewalls and the bottom of the box have no slip boundary conditions and the top surface is a free boundary. Except the proplate, it is assumed that each plate is surrounded by weak fault zones (trench and transform faults). A 2 mm thick edge of Vaseline is applied to the plate contact margin of the retroplate to allow the proplate to slide underneath it. This prevents the plates from sticking and stopping the subduction process. Horizontal shortening is achieved by displacing a rigid indentor at a constant velocity perpendicular to the plate margin. The indentor extends only into the upper part of the glucose syrup and the syrup is free to move underneath. The experiments are isothermal, so the role of thermal convection is not taken into account. As such, the flow in the syrup is produced only by plate convergence and the internal dynamics of the subducting/colliding plates.

Further, the mantle flow is generated only by these internal dynamics and we did not consider any influence of larger scale “global” flow effects [Hager and O’Connell, 1978; Ricard et  al., 1991]. The closed bottom of the Plexiglas means that the models represent only upper mantle flow. This assumption has significant implications for the dynamics of the slabs in the subsequent modeling results where the bottom boundary of the box is modeled to be the impermeable barrier of the upper‐lower mantle discontinuity. Some justification for this assumption includes estimates of mean mantle viscosity that suggest there is an increase in viscosity of up to two orders of magnitude from the upper to lower mantle through this region [Mitrovica and Forte, 1997] that would inhibit flow through this level. Furthermore, the endothermic phase transformation of olivine at ~670 km depth may introduce buoyancy effects that may significantly hinder mass flux across the upper mantle–lower mantle boundary [Christensen and Yuen, 1984, 1985]. Oceanic lithosphere subduction is initiated in the experiments by forcing downward the leading edge of the mantle lithosphere into the glucose to a depth of 3 cm with an angle of approximately 45° (Fig. 5.9). After this, the buoyancy of the slab drives the subduction dynamics as the piston drives the proplate toward the retroplate. Previous studies show that faster plate convergence is associated with advancing plate subduction, whereas slower plate convergence is associated with retreating plate subduction [Bellahsen et  al., 2005; Schellart, 2005; Faccenna et al., 2007]. Our experiments demonstrate how the continental proplate behaves following this subduction phase. Both experiments were monitored using a sequence of side‐ and top‐view photographs taken during progression of the experiments. The surface topography evolution of the experiments is monitored with a topography scanner (Real Scan USB model 300 and the measurement uncertainity is ±0.5 mm). 5.3.2. Laboratory Experiment Results of Subduction to Delamination The role of preexisting ocean subduction in developing delamination tectonics has been tested in the experiments of LABMODEL‐1 and LABMODEL‐2. We test the effect of varying plate convergence velocities into the delamination process. 5.3.2.1. LABMODEL‐1 In this experiment, we modeled the evolution of an orogenic cycle from subuction to delamination while plate convergence velocity of Vp  =  0.25  cm/min is imposed. Based on the scaling ratio, the convergence velocity approximately corresponds to the 3.81 cm/yr in

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  133 A)

Top view

B)

x

x′

Top view

0

0

Retroplate

Crust

5 cm

5 cm

t = 33 min

t = 16 min

Mantle lithosphere

Side view

Side view Vp

Mantle flow

0

Vp

Crust Retroplate

0

t = 16 min

10 cm

Delaminating mantle lithosphere

t = 33 min

10 cm

C) 0.2 Migration of subsidence

Surface topgraphy (cm)

0

–0.2

–0.4

t = 16 min

Subsidence induced by slab pull

t = 33 min –0.6

0

5

10

15 x (cm)

20

25

30

35

Figure 5.10  (a) Top view and side view of the experiment LABMODEL‐1 at t =16 min and the imposed convergence velocity (Vp = 0.25 cm min −1). (b) Top‐view and side‐view of the experiment LABMODEL‐1 at t = 33 min and the imposed convergence velocity (Vp = 0.25 cm min −1). (c) Surface topography plots along X‐X′ cross section for 16 min, and 33 min. (Modified from Göğüş et al., [2011]).

nature (Table 5.2). The model begins with the subduction of the ocean plate into sublithospheric mantle at about 90° dip angle. By t = 8 min, the subducting slab reaches to the bottom of the Plexiglas tank and by t = 16 min, the subduction has evolved into the continental delamination as the dense mantle lithosphere peels away from the overlying crust (Fig. 5.10a). The top view and the side view of the model photography suggest that the ~5 cm of mantle lithosphere is replaced with the light‐colored sublithospheric mantle due to the delamination. The top view figure shows that the central part of the model is associated with more delamination than the edges of the mantle lithosphere because the upward return flow of the sublithospheric mantle under the proplate crust shows the entrainment of this flow.

The delamination of the model progresses as the model evolves and by t = 33 min, about 10 cm of the mantle lithosphere delaminated from the upper crust (Fig. 5.10b). Again, this process is visible from the top view and side view photographs as the light‐colored sublithospheric mantle intrudes beneath the convergent crust. At this time, 10 cm of the mantle lithosphere slab sinks into the mantle and flattens at the bottom of the tank. Although the plate convergence actively pushes the lithosphere to the pro (hinterland) direction, the retromotion (roll‐back) of the subducting slab opening a gap (basin) in the back‐arc may suggest that such roll‐back motion is kinematically similar to the subduction retreat process. However, this type of mantle lithosphere roll‐back may be considered as sub-crustal retreat since the process

134  ACTIVE GLOBAL SEISMOLOGY

occurs underneath the crust, rather than at the boundary between two plates. We tracked the surface topography evolution of LABMODEL‐1 by using a laser scanning technique where the laser machine is placed on top of the experiment (near the camera). Figure 5.10c shows the surface elevation for t = 16 min and t = 33 min along the X‐X′ cross section line. The zero topography line is somewhat arbitrary and we selected it to correspond to the topography at the left boundary of the retro(hinterland)plate. The precision of the surface topography measurements is  high enough that the profiles show even the shortwavelength variations that may be produced by the surface roughness. A semiperiodic variation of the surface

A)

wavelength topography is caused by the surface grid lines. We chose to plot the central (X‐X′) cross section for the average topography at different time frames, since the topography measurements on the boundaries may not represent simply the delamination related topography. By t = 16 min, the model shows the negative surface deflection at the retroplate (0–14.5 cm) ranging from zero at the left edge to ~0.13 cm. This subsidence (and plate tilting) develops as a negative dynamic topography and it is most likely driven by the response of the underlying mantle flow which the sinking slab in the model center produces. The proplate is represented by more subsidence (max −0.33 cm at x = 21 cm) and this is at the delaminating hinge location where the crust is pulled down by the

Top view

B)

Top view

0

x

0

x′ Retroplate

Crust

5 cm

5 cm

t = 27 min

t = 12 min Side view

Mantle lithosphere

side view Vp

Vp

Crust Mantle lithosphere Retroplate

t = 12 min

0

t = 27 min

10 cm

0

10 cm

C) 0.2 Migration of subsidence

Surface topgraphy (cm)

0.1

0

–0.1

t = 12 min

–0.2

t = 27 min –0.3

0

5

Subsidence induced by slab pull 10

15 x (cm)

20

25

30

Figure 5.11  (a) Top view and side view of the experiment LABMODEL‐2 at t = 12 min and the imposed convergence velocity (Vp = 0.5 cm min −1). (b) Top view and side view of the experiment LABMODEL‐2 at t = 33 min and the imposed convergence velocity (Vp = 0.5 cm min −1). (c) Surface topography plots along X‐X′ cross section for 12 min and 27 min. (Modified from Göğüş et al., [2011]).

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  135

delaminating slab. Such surface depression at the delaminating hinge is also predicted in the numerical experiments, however, plateau‐type uplift does not develop above the delamination because these laboratory models do not have thermal buoyancy effect of the sublithospheric mantle. By t = 33 min, the surface topography is associated with a wider zone of uplift due to the lithospheric shortening induced by the imposed plate convergence (Fig.  5.10c). The subsidence (−0.2 cm) at the delamination hinge still persists at x = 26 cm. We note that the migration of subsidence has moved ~5 cm on the reverse direction of the plate convergence and this suggests that the subcrustal retreat (delamination) develops in conjunction with orogenesis. 5.3.2.2. LABMODEL‐2 For this laboratory experiment, we increased the imposed convergence velocity to Vp = 0.5 cm/min and the rest of the model parameters are left the same as the previous experiment. At the beginning, as in the previous experiment, the mantle lithosphere subducts into the sublithospheric mantle at a high dip angle (70°–80°) and at t = 5 min, the subducted lithosphere hits the bottom of the box while it is also subducting beneath the retroplate (Fig.  5.11a). At t = 12 min, the ocean plate subduction still continues (an ocean basin still exists between two plates) as the slab has just reached the bottom of the tank. Unlike LABMODEL‐1, the subducting slab does not roll back, rather the slab drapes forward under the retroplate. By t = 27 min, the plate collision occurred between the two plates and the proplate crust is partially decoupled from the crust and overthrusts onto the retroplate (Fig.  5.11b). This suggests that the delamination of the mantle lithosphere from the overlying crust did not have sufficient time to occur, instead the partial decoupling of the crust from the mantle lithosphere occurs and a piece of the crust accretes to the retro(hinterland)plate. This type of crustal stacking onto the overriding plate is suggested for the nappe emplacement around the collision zones at the Alpine‐Himalayan orogenic systems (e.g., Swiss Alps; Schmid et  al. [1996]). The style of crustal accretion is termed flake tectonics by Oxburgh [1972]. Figure  5.11c shows the surface topography evolution (along X‐X′ cross section) for LABMODEL‐2 at times, t =12 min and t = 27 min. The zero topography line in the experiments is a reference level corresponding to the left edge of the proplate. In both time frames, there is a paired subsidence and uplift pattern developed owing to the underlying effect of the subducting slab and the crustal accretion, respectively. Note that such paired topography signatures moved towards the retroplate as both plates move together after the collision/coupling.

5.4. CONCLUSIONS AND DISCUSSION 5.4.1. Summary of the Numerical Model Predictions and Their Interpretation in the Context of the Eastern Anatolia Orogenic Region The results of the numerical experiments (NUMMODEL‐1 and NUMMODEL‐2) test the role of plate convergence on the geodynamic evolution of mantle lithosphere delamination/peel away of the lithosphere. Model predictions demonstrate varying crustal and surface response imposed by plate convergence and the results are in good agreement with the surface tectonics of the East Anatolian plateau. 1. Lithosphere delamination/descent of the ocean slab is associated with surface uplift as a result of the isostatic and dynamic effects of lithospheric removal. The uplift is enhanced and evened into a plateau by plate shortening. A pulse of (migrating) subsidence can develop as the delaminating lithosphere loads the lithosphere at the hinge zone. Detachment of this lithosphere changes the surface elevation variation near the delamination hinge or where the sinking slab is attached to the overlying crust [Göğüş and Pysklywec, 2008a]. Our modeled surface topography at t = 7.0 m.y. is consistent with the profile of the present‐day topography across eastern Anatolia at 42°E (Fig. 5.8a). Consequently, we attribute the anomalous plateau uplift in the region to the removal of the ocean lithosphere and slow plate shortening. Previous studies suggest that eastern Anatolia started to emerge from sea level ~11 m.y. ago and in the last 6–7 m.y. has experienced enhanced volcanic activity (with subduction and decompression melt related) and uplift [Keskin, 2003, 2007; Şengör et  al., 2003, 2008]. These events are also consistent with the proposed removal of the lithosphere beneath the plateau. 2. Delamination/peeling away lithosphere causes distinct zones of contraction/thickening (at the plateau/gap flanks) and extension/thinning (within the delamination gap, to the far side of the delaminating hinge) within the crust. In the presence of plate convergence, the contraction and thickening are amplified but still localized to the sites determined by the delamination. Notably, the results demonstrate that crustal extension and thinning can occur within an overall plate convergent regime, depending on the rate of plate shortening. The East Anatolian plateau is predominantly affected by plate shortening since the middle Miocene and the evidence for north‐ south crustal extension is not well recognized, though may be worthwhile to explore. GPS interpretations by Reilinger et al. [2006] suggest a small amount of localized extension in north‐northwest direction. Previous work by Göğüş and Pysklywec [2008a] interpreted their model results in the context of the synconvergent extension for

136  ACTIVE GLOBAL SEISMOLOGY

the geodynamic evolution of the East Anatolian orogenic region in which the plateau uplift, heating, and small amount of localized stretching may occur coevally. In the adjacent margin of East Anatolia, Copley and Jackson [2006] suggest oblique extension along north‐northwest and south‐southwest in northern Iran. Further, solid geological evidence is necessary to make direct comparison between the model outcomes and the regional tectonics, and in this work, we did not attempt to do such comparison because of many existing uncertainties in geological characteristics of the east Anatolia region. 5.4.2. Summary of the Laboratory Experiments In addition to the numerical experiments for delamination, we used laboratory‐based analogue modeling experiments to investigate the transition from ocean lithosphere subduction to continental collision and delamination. The following results suggest the role of plate convergence at experiments LABMODEL‐1 and LABMODEL‐2 may alter the surface response to the deep lithospheric delamination process. 1. Mantle lithosphere delamination can occur with slow plate convergence where the slab peels off/rolls back similar to a retreating ocean slab subduction. The results suggest that continental delamination may be a natural progression from ocean plate subduction and illustrate that the removal of mantle lithosphere does not necessarily require a significant density heterogeneity to initiate [Göğüş et al., 2011]. The regions of Alpine tectonics and Mediterranean basins [Royden, 1993; Seber et  al., 1996; Faccenna et al., 2001] may represent suitable examples of this transition. 2. The experiments show that when the plate convergence is higher, the mantle lithosphere is less prone to delaminate from the crust. With higher plate convergence, the consumed mantle lithosphere can drape forward instead. The proplate crust separates from the mantle lithosphere only at the collision zone and is overthrusted/accreted on top of the retroplate. Similar tectonic modeling results of flake tectonics are suggested for alpine collision tectonics associated with folded nappes such as the Swiss Alps [Oxburgh, 1972; Pfiffner et al., 2000]. The surface topography is high where the crustal accretion occurs, whereas right at the plate suture zone, there is subsidence as the downgoing mantle lithosphere pulls down the surface. 3. The evolution of surface topography with laboratory experiments suggests that there is an appreciable amount of surface depression caused by vertical forcing of the delaminating mantle lithosphere. The surface depression migrates as the mantle lithosphere delaminates underneath the crust. This finding is similar to numerical results where a free‐hanging delaminated slab, whether it is broken off or not, creates this type of migrating surface

trough. With the laboratory modeling results, there is broad isostatic uplift of the surface topography due to the imposed plate convergence. However, this is a relatively small magnitude uplift compared to the previous numerical models of delamination, in particular there is no well‐defined plateau uplift. This is a consequence of the absence of thermal effects in the analogue models so that the mantle material moving into the delamination gap is not sufficiently buoyant to raise a plateau. We note that the purpose of this chapter is to consider only the role of orogenesis (by imposing plate convergence velocities) in the course of lithosphere delamination (vertical forcing) by taking advantage of different modeling techniques, and eventually compare and contrast numerical model results against the observations in the eastern Anatolia orogenic region. We outline that two different modeling techniques may provide similar results in the large‐scale lithospheric peel‐away process (e.g., migrating surface subsidence) in the direction opposite of the imposed plate convergence. The investigation of various sizes, composition, and thermomechanical properties of the delaminating/peeling away of lithosphere and its surface response to such varying properties are not the focus of this contribution. Such model parameterizations and their associated crustal evolution patterns were discussed in the other geodynamic modeling studies [i.e., Morency and Doin, 2004; Göğüş and Pysklywec, 2008a; Pysklywec et al., 2010; Ueda et al., 2012]. ACKNOWLEDGMENTS Russell N. Pysklywec acknowledges funding from an NSERC Discovery Grant. Numerical calculations were done using a modified version of the SOPALE (2000) software. The SOPALE modeling code was originally developed by Phillup Fullsack at Dalhousie University with Chris Beaumont and his Geodynamics group. We also acknowledge support from the TUBITAK 2221 Visiting Scientist Program for facilitating collaboration for the work. The authors thank A. M. C. Şengor and an anonymous reviewer for constructive reviews of the manuscript. REFERENCES Al‐Lazki, A. I., E. Sandvol, D. Seber, M. Barazangi, N. Turkelli, and R. Mohamad (2004), Pn tomographic imaging of mantle lid velocity and anisotropy at the junction of the Arabian, Eurasian and African plates, Geophys. J. Int., 158(3), 1024–1040. Beck, S. L., and G. Zandt (2002), The nature of orogenic crust in the central Andes, J. Geophys. Res., 107, B10, 2230; doi: 10.1029/2000JB000124. Bellahsen, N., C. Faccenna, and F. Funiciello (2005), Dynamics of subduction and plate motion in laboratory experiments:

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  137 Insights into the plate tectonics behavior of the Earth, J. Geophys. Res., 110, B01401; doi: 10.01029/02044JB002999. Bird, P. (1978), Initiation of intracontinental subduction in the Himalaya, J. Geophys. Res., 83, 4975–4987. Bird, P. (1979), Continental delamination and the Colorado Plateau, J. Geophys. Res., 84, 7561–7571. Biryol, C., S. Beck, G. Zandt, and A. Ozacar (2011), Segmented African lithosphere beneath the Anatolian region inferred from teleseismic P‐Wave tomography, Geophys. J. Int., 184, p. 1037–1057. Burov, E. B., and A. B. Watts, (2006), The long‐term strength of continental lithosphere, “jelly sandwich” or “crème brûlée”?, GSA Today, 16, 4–10; doi: 10.1130/1052–5173(2006)016 2.0.CO;2 Channell, J. E. T., and J. C. Mareschal (1989), Delamination and asymmetric lithospheric thicken‐ ing in the development of the Tyrrhenian rift, in M. P. Coward et  al., eds., Alpine tectonics, Geological Society London Special Publication, 45, 285–302. Chen, W. P., and P. Molnar (1983), Focal depths of intracontinental and intraplate earthquakes and their implications for the thermal and mechanical properties of the lithosphere, J. Geophys. Res., 88, 4183–4214. Chiarabba, C., and G. Chiodini (2013), Continental delamination and mantle dynamics drive topography, extension and fluid discharge in the Apennines, Geology; doi:10.1130/ G33992.1. Christensen, U. R., and D. A. Yuen (1984), The interaction of a subducting lithospheric slab with a chemical or phase boundary, J. Geophys. Res., 89, 4389–4402. Copley, A., and J. Jackson (2006), Active tectonics of the Turkish‐ Iranian plateau, Tectonics, 25; doi:10.1029/2005TC001906, Elkins‐Tanton, L. T. (2005), Continental magmatism caused by lithospheric delamination, in Plates, plumes, and paradigms, edited by G. R. Foulger, J. H. Natland, D. C. Presnall, and D. L. Anderson, 449–461, Geological Society of America Special Paper, 388; doi: 10.1130/2005.2388(27). Espurt, N., F. Funiciello, J. Martinod, B. Guillaume, V. Regard, C. Faccenna, and S. Brusset (2008), Flat subduction dynamics and deformation of the South American plate: Insights from analog modeling, Tectonics, 27, TC3011; doi:10.1029/2007TC002175. Faccenna, C., A. Heuret, F. Funiciello, S. Lallemand and T. W. Becker (2007), Predicting trench and plate motion from the dynamics of a strong slab, Earth Planet. Sci. Lett., 257, 29– 36; doi: 10.1‐16/j.epsl.2007.02.016. Faccenna, C., T. W. Becker, F. P. Lucente, L. Jolivet, and F. Rossetti (2001), History of subduction and back‐arc extension in the central Mediterranean, Geophys. J. Int., 145, 809– 820; doi: 10.1046/j.0956‐540x.2001.01435. Fadil, A., P. Vernant, S. McClusky, R. Reilinger, F. Gomez, D. B. Sari, T. Mourabit, K. Feigl, and M. Barazangi (2006), Active tectonics of the western Mediterranean: Geodetic evidence for roll back of a delaminated subcontinental lithospheric slab beneath the Rif Mountains, Morocco, Geology, 34(7), 529–532. Fichtner, A., E. Saygin, T. Taymaz, P. Cupillard, Y. Capdeville, and J. Trampert (2013), The deep structure of the North Anatolian Fault Zone, Earth Planet. Sci. Lett., 373, 109–117.

Fullsack, P. (1995), An arbitrary Lagrangian‐Eulerian formulation for creeping flows and its applications in tectonic models, Geophys. J. Int., 120, 1–23. Funiciello, F., C. Faccenna, D. Giardini, and K. Regenauer‐ Lieb (2003.), Dynamics of retreat‐ ing slabs: 2. Insights from three‐dimensional laboratory experiments, J. Geophys. Res., 108(B4), 2207; doi: 10.1029/2001JB000896. Girbacea, R., and W. Frisch (1998), Slab in the wrong place: Lower lithospheric mantle delamination in the last stage of the Eastern Carpathian subduction retreat, Geology, 26, 611–614. Gleason, G.C., and J. Tullis (1995), A flow law for dislocation creep of quartz aggregates determined with the molten salt cell, Tectonophysics, 247, 1–23; doi: 10.1016/0040‐1951(95)00011‐B. Göğüş, O. H. (2015), Rifting and subsidence following lithospheric removal in continental back arcs, Geology, 43, 3–6; doi:10.1130/G36305.1. Göğüş, O. H., and R. N. Pysklywec (2008a), Mantle lithosphere delamination driving plateau uplift and synconvergent extension in eastern Anatolia, Geology, 36(9), 723–726; doi: 10.1130/G2498A.1. Göğüş, O. H., and R. N. Pysklywec (2008b), Near surface diagnostics of dripping or delaminating lithosphere, J. Geophys. Res., 113; doi:10.1029/2007JB005123, Göğüş, O. H., R. N. Pysklywec, F. Corbi, and C. Faccenna (2011), The surface tectonics of mantle lithosphere delamination following ocean lithosphere subduction: Insights from physical‐scaled analogue experiments, Geochem. Geophys. Geosyst., 12, Q05004; doi: 10.1029/2010GC003430. Gök, R., Pasyanos, M., and Zor, E. (2007), Lithospheric structure of the continent‐continent col‐ lision zone: Eastern Turkey, Geophys. J. Int., 169, 1079–1088; doi: 10.1111/j.1365‐246X.2006.03288.x. Gray, R., and R. N. Pysklywec (2012), Geodynamic models of mature continental collision: Evolution of an orogen from lithospheric subduction to continental retreat/delamination, J. Geophys. Res., 117. Guillaume, B., J. Martinod, and N. Espurt (2009), Variations of slab dip and overriding plate tectonics during subduction: Insights from analogue modeling, Tectonophysics, 463, 167– 174; doi:10.1016/j.tecto.2008.09.043. Hager, B. H., and. R. J. O’Connell (1978), Subduction zone dips and flow driven by the plates, Tectonophysics, 50, 111–134. Harrison, T. M., P. Copeland, W. S. F. Kidd, and A. Yin (1992), Raising Tibet, Science, 255, 1663–1670. Hirth, G., and D. L. Kohlstedt (1996), Water in the oceanic upper mantle: Implications for rheology, melt extraction and the evolution of the lithosphere, Earth Planet. Sci. Lett., 144, 93–108; doi: 10.1016/0012‐821X(96)00154‐9. Houseman, G. A., and L. Gemmer (2007), Intraorogenic extension driven by gravitational instability: Carpathian‐Pannonian orogeny, Geology, 35, 1135–1138; doi: 10.1130/G23993A.1. Houseman, G. A., D. P. McKenzie, and P. Molnar (1981), Convective instability of a thickened boundary layer and its relevance for the thermal evolution of continental convergent belts, J. Geophys. Res., 86, 6115–6132. Jull, M., and P. B. Kelemen (2001), On the conditions for lower crustal convective instability, J. Geophys. Res., 106, 6423–6446; doi: 10.1029/2000JB900357.

138  ACTIVE GLOBAL SEISMOLOGY Keskin, M. (2003), Magma generation by slab steepening and break off beneath a subduction accretion complex: An alternative model for collision‐related volcanism in Eastern Anatolia, Turkey, Geophys. Res. Lett., 30(24), 8046; doi: 10.1029/2003GL018019. Keskin, M. (2007), Eastern Anatolia: A hotspot in a collision zone without a mantle plume, in Plates, plumes and planetary processes, edited by G. R. Foulger and D. M. Jurdy, 693–722, Geological Society of America Special Paper, 430; doi: 10.1130/2007.2430(32). Kind, R., T. Eken, F. Tilmann, F. Sodoudi, T. Taymaz, T. Bulut, X. Xuan, B. Can, and F. Schneider (2015),Thickness of the lithosphere beneath Turkey and surroundings from S‐receiver functions, Solid Earth, 6, 971–984; doi:10.5194/se‐6‐971‐2015. Komec‐Mutlu, A., and H. Karabulut (2011), Anisotropic Pn tomography of Turkey and adjacent regions, Geophys. J. Int., 187, 1743–1758. Kosarev, G., R. Kind, S. V. Sobolev, X. Yuan, W. Hanka, and S. Oreshin (1999), Seismic evidence for detached Indian lithospheric mantle beneath Tibet, Science, 283, 1306–1309. Lachenbruch, A. H., and J. H. Sass (1977), Heat flow in the United States and the thermal regime of the crust, in The Earth’s Crust, Its Nature and Physical Properties, edited by J. G. Heacock, 626–675, American Geophysical Union, Geophysical Monograph Series, 20. Le Pourhiet, L. L., M. Gurnis, and J. Saleeby (2006), Mantle instability beneath the Sierra Nevada Mountains in California and Death Valley extension, Earth Planet. Sci. Lett., 251, 104–119; doi: 10.1016/j.epsl.2006.08.028. Levander, A., B. Schmandt, M. S. Miller, K. Liu, K. E. Karlstrom, R. S. Crow, C. T. A. Lee, and E. D. Humphreys (2011), Continuing Colorado plateau uplift by delamination style convective lithospheric downwelling, Nature, 472, 461– 465; doi:10.1038/nature10001. Mitrovica, J. X., and A.M. Forte (1997), The radial profile of mantle viscosity: Results from the joint inversion of convection and post‐glacial rebound observables, J. Geophys. Res., 102, 2751–2769. Morency, C., and M. P. Doin (2004), Numerical simulations of mantle lithospheric delamination, J. Geophys. Res., 109, B03410; doi: 10.1029/2003JB002414. Morency, C., M. P. Doin, and C. Desmoulin (2002), Convective destabilization of a thickened continental lithosphere, Earth Planet. Sci. Lett., 202, 303–320. Okay, A. I., M. Zattin, and W. Cavazza (2010), Apatite fission‐ track data for the Miocene Arabia‐Eurasia collision, Geology, 38, 35–38; doi: 10.1130/G30234.1. O’Reilly, S. Y., W. L. Griffin, Y. Poudjom Djomani, and P.  Morgan (2001), Are lithospheres forever? Tracking changes in subcontinental lithospheric mantle through time, GSA Today, 11, 4–9. Oxburgh, E. R. (1972), Flake tectonics and Continental collision, Nature, 239, 202–204. Özacar, A. A., H. Gilbert, and G. Zandt (2008), Upper mantle discontinuity structure beneath East Anatolian Plateau (Turkey) from receiver functions, Earth Planet. Sci. Lett., 269, 427–435; doi: 10.1016/j.epsl.2008.02.036. Pearce, J. A., J. F. Bender, S. E. De Long, W. S. F. Kidd, P. J. Low, Y. Guner, F. Saroglu, Y. Yilmaz S. Moorbath, and J. G.

Mitchell (1990), Genesis of collision volcanism in Eastern Anatolia, Turkey, J. Volcanol. Geotherm. Res., 44, 189–229. Pfiffner, O. A., S. Ellis, and C. Beaumont (2000), Collision tectonics in the Swiss Alps: Insight from geodynamic modeling, Tectonics, 19, 10651094. Platt, J. P., J.‐I. Soto, M. J. Whitehouse, A. J. Hurford, and S. P. Kelley (1998), Thermal evolution, rate of exhumation, and tectonic significance of metamorphic rocks from the floor of the Alboran extensional basin, western Mediterranean, Tectonics, 17, 671–689. Poudjom Djomani, Y. H., S. Y. O’Reilly, W. L. Griffin, and P. Morgan (2001), The density structure of subcontinental lithosphere through time, Earth Planet. Sci. Lett., 184, 605– 621. Pysklywec, R. N., C. Beaumont, and P. Fullsack (2002), Lithospheric deformation during the early stages of continental collision: numerical experiments and comparison with South Island, New Zealand, J. Geophys. Res., 107, 2133; doi: 10.1029/2001JB000252. Pysklywec, R. N., O. Göğüş, J. Percival, A. R. Cruden, and C. Beaumont (2010), Insights from geodynamical modeling on possible fates of continental mantle lithosphere: Collision, removal, and overturn, Canadian J. Earth Sci., 47(4), 541–563; doi:10.1139/E09‐043. Reilinger, R., et al. (1997), Global Positioning System measurements of present‐day crustal movements in the Arabia‐ Africa‐Eurasia plate collision zone, J. Geophys. Res., 102, 9983–9999. Reilinger, R., et  al. (2006), GPS constraints on continental deformation in the Africa‐Arabia‐Eurasia continental ­collision zone and implications for the dynamics of plate interactions, J. Geophys. Res., 111, B05411; doi: 10.1029/ 2005JB004051. Ren, Y., and Y. Shen (2008), Finite frequency tomography in southeastern Tibet: Evidence for the causal relationship between mantle lithosphere delamination and the north‐ south trending rats, J. Geophys. Res., 113; doi: 10.1029/ 2008JB005615. Reutter, K. G., P. Giese, and H. Closs (1980), Lithospheric split in the descending plate: observations from the northern Apennines, Tectonophysics, 64, 11–19. Ricard, Y., C. Doglioni, and R. Sabadini (1991), Differential rotation between lithosphere and mantle: A consequence of lateral mantle viscosity variations, J. Geophys. Res., 96, 8407–8415. Royden, L. H. (1993), Evolution of retreating subduction boundaries formed during continental collision, Tectonics, 12, 629–638. Saleeby, J., and Z. Foster (2004), Topographic response to ­mantle lithosphere removal, southern Sierra Nevada region, California, Geology, 37, 245–248. Schellart, W. P. (2005), Influence of the subducting plate velocity on the geometry of the slab and migration of the subduction hinge, Earth Planet. Sci. Lett., 231, 197–219; doi: 10.1016/j. epsl.2004.12.019. Schmid, S. M., O. A. Pfiffner, N. Froitzheim, G. Schnborn, and E. Kissling (1996), Geophysical geological transect and tectonic evolution of the Swiss‐Italian Alps, Tectonics, 15, 1036–1064. Seber, D., M. Barazangi, A. Ibenbrahim, and A. Demnati (1996), Geophysical evidence for lithospheric delamination

GEODYNAMICAL MODELS FOR CONTINENTAL DELAMINATION  139 beneath the Alboran Sea and Rif‐Betic mountains, Nature, 379, 785–790; doi: 10.1038/379785a0. Şengör, A. M. C., and B. A. Natal’in (1996), Turkic type Orogeny and its role in the making of the continental crust, Ann. Rev. Earth Planet. Sci., 24, 263–337. Şengör, A. M. C., and W. S. F. Kidd (1979), The post‐collisional tectonics of the Turkish‐Iranian Plateau and a comparison with Tibet, Tectonophysics, 55, 361–376. Şengör, A. M. C., M. S. Ozeren, M. Keskin, M. Sakinc, A. Ozbakir, and I. Kayan (2008), Eastern Turkish high plateau as a small Turkic‐type orogen: Implications for post‐collisional crust‐forming processes in Turkic‐type orogens, Earth‐Sci. Rev., 90, 1–48; doi:10.1016/j.earscirev.2008.05.002. Şengör, A. M. C., S. Ozeren, E. Zor, and T. Genc (2003), East Anatolian high plateau as a mantle‐supported, N‐S shortened domal structure, Geophys. Res. Lett., 30, 8045; doi: 10.1029/2003GL017858. Sobolev, S. V., and A. Y. Babeyko (2005), What drives orogeny in the Andes? Geology, 33(8); 617–620; doi: 10.1130/G21557.1. Sonder, L. J., and C. H. Jones (1999), Western United States extension: How the west was widened, Earth Planet. Sci., 27, 417–462.

Ueda, K., T. Gerya, and J‐P. Burg (2012), Delamination in collisional orogens: Thermomechanical modeling, J. Geophys. Res., 117; doi:10.1029/2012JB009144. Vanacore, E. A., T. Taymaz, and E. Saygin (2013), Moho structure of the Anatolian Plate from receiver function analysis, Geophys. J. Int., 193, 329–337. Weijermars, R., and H. Schmeling (1986), Scaling of Newtonian and non‐Newtonian fluid dynamics without inertia for quantitative modeling of rock flow due to gravity (including the concept of rheological similarity), Physics of the Earth and Planetary Interiors, 43, 316–330. West, J. D., J. M. Fouch, B. J. Roth, and L.T. Elkins‐Tanton (2009), Vertical mantle flow associated with a lithospheric drip beneath the Great Basin, Nat. Geosci., 2, 439–444; doi: 10.1038/ngeo526. Wortel, M. J. R., and W. Sparkman (2000), Subduction and slab detachment in the Mediterranean‐Carpathian region, Science, 290, 1910–1917. Zor, E. (2008), Tomographic evidence of slab detachment beneath eastern Turkey and the Caucasus, Geophys. J. Int., 175, 1273–1282.

6 Major Problems of Western Anatolian Geology Yücel Yılmaz

ABSTRACT Despite a mass of new data that have been collected during the last few decades, some major problems of western Anatolian geology still remain controversial. Among these, cause and timing of generation of the Menderes Massif and the magmatic associations, the north-south trending grabens and time of inception of the east-west grabens, are at the forefront. Almost all of these geological events appear to be genetically closely linked with the extensional tectonics, because the extension has played strong tectonic and geologic control on their development. Therefore, a first-order problem that needs to be addressed is whether the extensional regime is continually active, since it began possibly during the late Eocene time, or has it evolved in pulses. As a consequence of the nature of the problem, establishing a cross connection between the different events and evaluating them in time-space and regional geological perspective are critical. In this chapter, main geological entities of western Anatolia are reviewed under separate headings, the ongoing controversies around them are discussed first, and then some solutions are proposed. The data set reviewed on the major geological events as outlined above displays an apparent discontinuity in their development and thus support collectively the pulse-extension model. 6.1. INTRODUCTION The western Anatolian region is characterized by a number of approximately east‐west trending, subparallel, normal fault zones, which border a swarm of grabens and the intervening horst blocks (Fig. 6.1). As a consequence of this, there is an intense seismic activity, as evidenced by a number of instrumentally recorded and also historical earthquakes, roughly encircling the active faults. Motions on the faults confirm an extension approximately in a north‐south direction. Therefore, the western Anatolian and Aegean regions have long been known to represent a broad zone of extension stretching from Bulgaria in the north to the Hellenic arc in the south [McKenzie, 1972, 1978; Jackson and McKenzie, 1988; McKenzie and Yılmaz 1991; Taymaz, 1996]. ̇ Department of Geological Engineering, Istanbul Technical ̇ University, Istanbul, Turkey

The seismic data suggest that extension is accommodated by large faults during the big earthquakes [>6; Patton, 1992a, b; Eyidoğan and Jackson, 1985; Taymaz et  al., 2007; Yolsal‐Çevikbilen et al., 2012, 2014]. According to the fault plane solutions of the big earthquakes, the seismogenic layer in western Anatolia is thin [> 12 km; Saunders et  al., 1998; Karabulut et  al., 2003; Taymaz et  al., 2008]. This is supported further by the high heat flow [Öztürk et al., 2006]. In western Anatolia, two groups of rocks may readily be distinguished: (1) the Neogene and younger rocks and (2) the older rocks (Fig. 6.2). The former forms a common cover of no marine origin to the latter, which comprises a tectonic mosaic developed prior to the deposition of the Neogene units (Fig. 6.2). In the tectonic mosaic, the following tectonic zones are differentiated from the north to the south: the Sakarya continent, the Izmir‐Ankara suture, the Menderes Massif, and the Taurides (Fig. 6.2). They were amalgamated as a result of total demise of the Izmir‐Ankara branch of the

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 141

142  ACTIVE GLOBAL SEISMOLOGY

Figure 6.1  The Aegean Grabens, associated earthquakes and the fault plane solutions [modified after Dewey and Şengör,1979] Red arc; non volcanic arc, Yellow arc with green glow: active volcanic arc, Green arc with yellow glow; extinct volcanic arc. Brown arrow displays lateral motion along the Pliny and Strabo trenches. NASZ; The North Aegean Shear Zone. Brown dotted areas are young graben depressions and their sediment fill.

Tethyan Ocean and the consequent collision of the bordering continents during the Late Cretaceous–Eocene period [Şengör and Yılmaz, 1981; Yılmaz, 1981; Yılmaz et al., 1995]. The following evolutionary history may be summarized as illustrated in Fig. 6.3. The late Cretaceous marked the beginning of a convergent regime at all fronts in Turkey, and was particularly characterized by the emplacement of ophiolite nappes (Figs. 6.3b,c and Fig. 6.4). These nappes moved onto extensive carbonate platforms that began subsiding “en masse” with the onset of obduction (Fig. 6.4). The initial phase of metamorphism of the Menderes Massif may, in part, be ascribed to this subduction and the subsequent obduction events Fig. 6.3b, c, and Fig. 6.4). Following the obduction of the ophiolitic nappes onto the Tauride‐Anatolide platform, the platform began to be severally deformed, folded, and internally imbricated (Fig. 6.4). Starting from the Late Cretaceous, a thick (> 6 km) nappe stack formed. This was the result of accreted lithospheric units, piled by thrusting from the north to the south as a consequence of the total

elimination of the separating oceans [Şengör and Yılmaz, 1981; Jolivet and Brun, 2010]. The shortening deformation caused the continental crust and the lithosphere to have excessively thickened. As a result, the crust is assumed to have reached over 50 km in thickness [Le Pichon and Angelier, 1979; Jackson and McKenzie, 1988; Dewey and Şengör, 1979; Şengör et  al., 1984; Tirel et al., 2004; Stampfli and Hochard, 2009]. The crust is measured to be about 30–32 km in thickness now [Saunders et al., 1998; Taymaz et al., 2008; Zhu et al., 2006]. During the Late Eocene, the initial uplift of the Aegean metamorphic complexes occurred, and they were partly unroofed [Yılmaz 1997; Burchfiel et  al., 2003; Lacassin et  al., 2007]. During the Oligocene period, the whole Aegean region including western Anatolia, the Aegean Sea area, and the Balkan region collectively formed a highland and began to be effectively eroded [Siyako and Huvaz, 2007; Yılmaz et  al., 2010; Elmas et al., 2011] (Fig. 6.3d). Sometime during the Late Oligocene–Early Miocene, the entire area of western Anatolia together with the

Major Problems of Western Anatolian Geology  143

Figure  6.2  Geology map of western Anatolia showing its main tectonic components [modified from Yılmaz, 1988]. DG = Demirci graben; GG = Gördes graben; SG = Selendi graben; UG = Uşak‐Ulubey graben; BEG = Bergama graben; GDG = Gediz graben; BMG = Büyük Menderes graben SOG = Soma graben; DP = Dilek Peninsula; Sö G = Söke graben; BH = Bozdağ horst; ÖG; Ören Graben, YG; Yatağ an Graben, KT; The Kale-Tavas basin, LN; Lycian Nappes, LNF; Lycian nappe front , A, C, D, Iz and M; cities of Aydın, Çanakkale, Denizli, I͘ zmir, Manisa and Muğ la.

Menderes Massif suffered a semibrittle‐brittle, north‐ south compressional deformation. As a result of this, the whole region was shortened, internally imbricated, and thickened. In the Menderes Massif, thick‐skinned deformation occurred, and its tectonic components were structurally rearranged. The deeply buried metamorphic rocks (the core rocks) were thrust above the cover rocks (see Fig.  6.9 for the structural rearrangement of the Menderes Massif). This was followed by an extensional phase possibly as a result of the collapse of the orogen.

During the Late Miocene period, the western Anatolian domain subsided near sea level [Steinenger et  al., 1985; Görür, 1988; Görür et al., 1997; Popov et al., 2006; Yılmaz et  al., 2010; Elmas et  al., 2011] (Fig.  6.3h). A regional denudation that already began during the Late Eocene – Oligocene effectively eroded the region during the Late Miocene-early Pliocene period. It was accompanied by tectonic erosion. Under the collective influences of the two forces, the crust thinned and a regionally developed flat‐lying erosional surface formed above the successions,

144  ACTIVE GLOBAL SEISMOLOGY

Figure 6.3  Schematic cross sections illustrating consecutive stages of the tectonic evolution of western Anatolia [modified from Yılmaz, 1997]. (a) During the Cretaceous, the northern branch of the NeoTethyan Ocean was separating the Taurides and the Sakarya continent. The oceanic lithosphere that was subducting northward generated an Andean‐type volcanic arc on the Sakarya continent. (b) The ophiolitic slices began to be obducted on the Taurides. Following the total consumption of the ocean, further convergence of the bordering continents caused regional elevation, which eliminated the sea realms from the orogenic zone during the latest Cretaceous‐Paleocene period [Yılmaz, 1981]. (c) Following the head‐on collision, giant nappe packages consisting of the ophiolite nappe (the allochthonous units) at the top, underlain by metamorphic and nonmetamorphic slices of the Tauride platform that were sliced off of the leading edge of the Taurus (the para‐autochthonous units) were dragged under the nappes, which in turn, began to be transported on the Tauride starting from the later times of the Late Cretaceous (L. Late Cret.). The nappe transport continued possibly until the Late Eocene. The regional dynamothermal metamorphism of the Menderes Massif occurred during to this period. (d) One of the major regional extension events took place sometime during the Late Oligocene‐Early Miocene period. As a consequence, the metamorphic rocks were exhumed and partly unroofed. The Lower Miocene terrestrial sediments were deposited on the high‐grade metamorphic rocks and the adjacent ophiolitic rocks. (e) During the Late Miocene, a large part of western Anatolia was covered by interconnected lakes. This was partly contemporaneous with a severe phase of denudation that formed a regionwide flat‐lying erosional surface. A new extensional phase began at the end of Late Miocene. The Bozdağ dome was formed during this extensional stage (thin curving black line in the north symbolizes the detachment fault). (f) The present east‐west trending horst and graben structures formed during the Quaternary.

Major Problems of Western Anatolian Geology  145 Ophiolitic Melange Future Menderes Massif Lycian nappes

· Izmir-Ankara Suture

S

N Sakarya Continent

Taurides Eocene

Figure 6.4  A detail from Figure 6.3c showing geology and tectonics of western Anatolia during the Eocene time. The Tauride Mesozoic carbonate platform that was dragged under the southerly transported nappe packages, packages were folded and imbricated, and therefore were buried and severely deformed. As a result, the platform was buried under increasing depths to the south. The white rectangle refers to the region, which would later form the Menderes Massif.

including the Upper Miocene–Lower Pliocene strata (Fig. 6.3d). The following major tectonic units and geological events of western Anatolia are still debated. 1. The Menderes Massif: its origin, time and mechanism of the metamorphism, the main tectonic components, and their structural arrangements 2. The magmatic associations: age of development and the mechanisms of generation 3. The Neogene cover rocks: the tectonic regime, under which they were developed, and their tectostratigraphic divisions 4. The north‐south extensional regime; its time and mechanism of initiation and continuity A number of different views have been proposed on each one of these subjects Çemen et al (2014). Because models proposed by different authors were commonly incompatible with one another, we, as a team, undertook a major project and mapped the region extensively. The area mapped stretches from the Marmara region to the Mediterranean region. In the following years, we revisited the regions repeatedly to check and test the new views and ideas that appear in the literature. In the following paragraphs, the major problems are reviewed and discussed, and new solutions are proposed to some of the lingering problems. 6.2. MAJOR PROBLEMS OF WESTERN ANATOLIA 6.2.1. The Menderes Massif The Menderes Massif (first identified by Philipson [1918]) is the largest tectonic entity of western Anatolia. It occupies a 250 x 120 km region and is roughly elliptical in shape with a nearly north‐south (north‐northeast) trending long axis (Fig.  6.5). It is located between the Izmir‐Ankara suture in the north and the Taurides in

the south (Fig.  6.2). The Izmir‐Ankara suture extends to the Vardar suture in Greece [Şengör and Yılmaz, 1981]. The Major tectonic zones of the Greek mainland and Anatolia cannot be correlated one to one, for example, the Karakaya ophiolite of the Sakarya continent in Anatolia, and the Pindus and Pelagonian ophiolite in Greece do not extend to the other sides [for a discussion see Şengör and Yılmaz 1981; Dixon and Robertson, 1984; Gessner et al., 2011; van Hinsbergen et al., 2010]. Along the northern periphery of the Menderes Massif, there is a group of HP metamorphic rocks structurally above the high‐grade metamorphic rocks that are located in the center of the massif (Fig.  6.5), and these rocks are commonly identified as different tectonic entities or different zones, for example, the Tavşanlı zone–the Selçuk melange, the Dilek nappe, the Afyon zone–Ören unit [Okay, 1981, 2004; Sherlock et  al., 1999; Candan et  al., 2001, 2005; Önen and Hal, 2000; Güngör and Erdoğan, 2001; Pourteau et  al., 2010; Plunder et  al., 2013]. The metamorphic rocks extending along the Dilek Peninsula (Fig. 6.2) are regarded as the easterly continuation of the Cyclades Massif of the central Aegean regions, based on their geological and metamorphic characteristics, that is, both display similar HP metamorphism [Candan et  al., 2005; Rimmele et al., 2003a, 2005; Gessner et al., 2011; van Hinsbergen and Schmid, 2012; and the references therein]. In fact, the rocks cropping out along the northern periphery of the Menderes Massif envelope the core rocks of the massif, and extend to the south discontinuously at the same structural level, standing at the top of the massif (Fig.  6.5) [Bozkurt and Oberhansli, 2001; Rimmele et al., 2003b, 2005; Whitney and Bozkurt, 2002, 2008; Gessner et al., 2011]. There is no sharp boundary or clear distinction among the geographical domains or zones. Even defining the core and the cover may be difficult in many places as a result of the complex structural arrangements [e.g.,

146  ACTIVE GLOBAL SEISMOLOGY

Figure 6.5  Geology map of the Menderes Massif [modified from Candan and Dora, 1998]. The elliptical region that is defined by the pink broken line represents the main body of the Massif (Gneiss and schist dominant core of the Massif). Bluish green and dark green areas outside the ellipse represent allochthonous and para‐authochtonous units. The yellow elliptical region is the Bozdağ dome. It is separated from the North and the South by the low angle normal detachment faults (thin red curving lines). The black arrows indicate long axes of the Menderes Massif and the Bozdağ dome. BGM = Büyük Menderes graben; KMG = Küçük Menderes graben; GG = Gediz graben.

Öztürk and Koçyiğit, 1983; Candan et  al., 2007, 2011; Erdoğan and Güngör, 2004; Bozkurt and Oberhansli, 2001; Bozkurt and Mittwede, 2005; Whitney et  al., 2008; van Hinsbergen et al., 2010]. The Menderes Massif was initially evaluated by Şengör and Yılmaz [1981] and Şengör et  al. [1984] to be the metamorphic equivalent of the Taurus Mesozoic carbonate platform and its basement, because of their lithostatigraphic similarities and the ages. Therefore, the two

tectonic zones were regarded as a single tectonic entity, under the common name of the Tauride‐Anatolide platform implying that they were parts of the same platform prior to the metamorphism. This view refutes the earlier views lingering since Kober [1921] that they were correlated with the Pre‐Alpine Median Massifs of Europe and thus regarded as the old crystalline basement of Anatolia. It was therefore identified under the name of the Anatolides. In later years, the data obtained as a result

Major Problems of Western Anatolian Geology  147

Figure 6.6  Generalized stratigraphic section representing the sequence of the Menderes Massif prior to the metamorphism and the associated deformation [modified from Candan et al., 2011].

of the detailed mapping covering the entire Menderes Massif (1/500,000 scale MTA maps and their explanatory texts) [Boray, 1982; Konak, 2002; Konak and Şenel, 2002], and other detailed maps on the massif [e g., Candan and Dora, 1998; Rimmele et al., 2003a, b; Candan et al., 2005; Bozkurt and Park, 2004] supported the view that the Tauride and the Menderes Massif consist of similar successions. Irrespective of its present structural arrangement, the main body of the Menderes Massif is composed of essentially a Pan‐African basement association (550–520 my old) [Şengör et  al., 1984; Candan et  al., 2001, 2011; Erdoğan and Güngör, 2004; Bozkurt and Oberhansli, 2001; Oberhansli et al., 2010] and a cover sequence (Fig. 6.6). The former consists mainly of augen gneisses (orthogneisses), metagranites, and metagrabbros with eclogitic relicts [Bozkurt and Oberhansli, 2001; Oberhansli et  al., 2010; Candan et al., 2011]. The cover sequence is made up of schists, phyllites, and marbles. The schist envelope includes quartzofeldspatic schists, micaschists, amphibolite schists, and phyllites. The marble envelope is small in the northern and central areas, but large in the southern areas (Fig.  6.6). The gneisses followed upward by micaschist‐limestone intercalation passing up onto the lower grade rocks, which are made up of graphite bearing micaschist, phyllite, and marble. The marble envelope begins above a basal conglomerate, and is dominated by metalimestone, forming a thick succession.

The schist‐phyllite envelope is identified as the Paleozoic in age, while the marble and recrystallized limestone are mostly Mesozoic [Boray, 1982; Konak, 2002; Konak and Şenel, 2002; Şengör et al., 1984; Bozkurt and Oberhansli, 2001; Gessner et al., 2011; Candan et al., 2011]. The entire succession may simply be summarized as consisting of three major tectostratigraphic units: (1) a Pan‐African basement association of Late Pre‐Cambrian–Early Cambrian age. (2) Above this old basement there are two envelopes: a schist envelope of Paleozoic age [MTA Maps, Candan et al., 2011] (Fig. 6.6) followed by a marble envelope of Mesozoic age [MTA maps, Konak, 2002; Konak and Şenel, 2002; Candan et  al., 2011; and the references therein]. (3) These are tectonically overlain by nappes, which consist mainly of metaophiolite and the genetically associated rocks (Figs. 6.5 and 6.6) [Candan et al., 2011]. The Menderes Massif has suffered multiple phases of metamorphism and associated deformations [Bozkurt and Oberhansli, 2001], and thus the succession has lost the original ordering. The details of the successions that are observed in the complementary tectonic slices, when reconstructed may be displayed as shown in Fig. 6.6. The generalized stratigraphic sequence may be interpreted as follows: a transgression began above an old basement during the Early Paleozoic, and continued interruptedly to the Late Paleozoic. The land was close to sea level as represented by the quartzitic sandstone and accompanying felsic metavolcanic rocks (tuff ?) of possibly

148  ACTIVE GLOBAL SEISMOLOGY

Figure 6.7  Schematic cross section, detailed from Figure 6.3, showing tectonic window position of the Menderes Massif with respect to the nappes (the para‐autochthonous and allochthonous units). The broken red line refers to the erosional surface. The pink elliptical region refers to the northern part of the Taurides that formed the Menderes Massif. The related tectonic events are detailed in Figure 6.8.

Permo‐Carboniferous age [Boray et al., 1975; Çağlayan et al., 1980; Konak, 2002; Konak and Şenel, 2002]. A rifting began during the Triassic time as evidenced by the Triassic metabasal conglomerate, sandstone and accompanying metalavas [Akkök, 1983; Dora, 2011; Candan et al., 2011]. This apparently corresponds to the Tethyan rifting stage. The following thick marble sequence represents a passive continental margin, developed as the rifting stage advanced. At the top of the marble there is a red, thinly bedded, recrystallized pelagic limestone (Fig. 6.6) [Konak, 2002; Candan et al., 2011; Okay, 1989; Dora, 2011] indicating a rapid subsidence of the platform. This possibly corresponds to the onset of approaching ophiolitic nappes, when a foredeep (a flexural foreland basin) formed in the northern edge of the passive margin facing the Neo‐Tethyan Ocean. The evidence for this interpretation is a weakly metamorphosed flyschlike unit passing to olistostrome‐bearing, internally chaotic rocks that were deposited above the red pelagic sediments of Upper Cretaceous age (olistostrome in Fig. 6.6).The olistostrome‐bearing rocks are in turn tectonically overlain by the ophiolitic nappe pile [Candan 2001, 2011; Ersoy et al., 2005] (Fig.  6.6). Further support may be obtained from the northern peripheral zone of the Menderes Massif (e.g., from the Afyon zone). In this zone, the metamorphosed sequence is made up of three distinct layers: a Paleozoic succession, consisting of metaclastic rocks at the base, a thick carbonate sequence in the middle, which is Mesozoic in age. It is overlain by the flyschlike olistostrome‐bearing deposits, which in turn is tectonically overlain by an ophiolite nappe at the top. The zones have suffered HP metamorphism during the Late Mesozoic [Okay, 1981, 2004; Önen and Hal, 2000]. Ages of the HP metamorphism vary between 80 and 60 my [Okay, 2004], which reveals a close genetic connection between the HP metamorphism and the northerly subducting oceanic lithosphere of the Izmir Ankara branch of the NeoTethyan Ocean [Şengör and Yılmaz 1981; Yılmaz 1981; Yılmaz

1997; Yılmaz et  al., 1999a; 2000; Plunder et  al., 2013] (Figs.  6.3 and 6.8a). Undoubtedly the tectonic slices below the ophiolitic nappes represent the northern edge of the passive margin successions that were sliced and imbricated when dragged under the southwardly advancing ophiolite nappe (Figs. 6.3b to c, 6.4, 6.8a, b). The stacking and piling of the nappes are estimated to have continued to the end of Eocene (Fig. 6.8b). This was followed possibly by the first phase of extension and extensionally induced exhumation (the elevation and erosion) causing exposure of the metamorphic rocks for the first time during the Late Eocene–Early Oligocene (Fig. 6.3d). A temporally associated magmatism developed in the northern Marmara, Thrace, and Balkan regions (see Magmatic Associations, Section 6.2.2) may be regarded as supporting evidence for the extension. In the light of the successions that were previously summarized (Fig.  6.6), the Menderes Massive may be divided into the following major tectonic subdivisions from the bottom to the top: (1) a core, (2) a cover, (3) para‐autochthonous tectonic slices, and (4) nappes of the allochthonous units, the ophiolitic and associated rocks (Figs. 6.5, 6.6, and 6.7). In the center of the Menderes Massif rises an east‐west trending horst (the Bozdağ dome) (Fig. 6.5). With respect to the Bozdağ horst (the central domain), two geographic domains are usually distinguished as the northern and southern domains (Fig.  6.5). The three domains share many similar lithological and metamorphic features [Konak, 2002, the MTA maps of Menderes Massif [Candan, 2000; Bozkurt and Oberhansli, 2001; Gessner et al., 2011; Candan et al., 2011]. A north to south cross section across the Menderes Massif displays that the core of the massif crops out in a tectonic window with respect to the enveloping upper tectonic units, namely the ophiolitic nappes and the underlying para‐autochthonous thrust sheets (Fig. 6.7). The ophiolite at the top represents far‐travelled nappes,

Major Problems of Western Anatolian Geology  149

Figure 6.8  Schematic cross sections displaying consecutive stages of tectonic evolution of the Menderes Massif from Late Cretaceous to Late Oligocene‐Early Miocene period. Sc = Sakarya continent; BSC = basement of the Sakarya continent; OFM = Ophiolitic mélange; MOF = Mesozoic ophiolite of the Izmir‐Ankara Ocean; CP = Taurus carbonate platform; SE = Basement of the Taurus. (a) During the late Cretaceous NeoTethyan oceanic lithosphere subducting under the Sakarya Continent (dark green belt) underwent HP metamorphism. The ­continental slope sequence of the Tauride ( light green belt) also involved in the deep burial and underwent HP metamorphism together with the ophiolite. They were later rapidly elevated and formed the allochthonous (the ophiolite and associated rocks, ophiolitic mélange association, e.g., the Tavşanlı zone and the Selçuk mélange) and the para‐autochthonous (e.g., the Afyon zone and the Dilek nappe) nappe packages. (b)The major phase of the nappe transport on the Tauride began at the end of Late Cretaceous and lasted until the end of Middle Eocene. The parts of the Taurides under the southerly transported thick nappe pile (estimated to be greater than 6 km) were depressed, suffered severe ductile deformation, and underwent dynamothermal HT metamorphism and, when elevated, formed the Menderes Massif. (c) Sometime during the Oligocene‐Early Miocene, the entire western Anatolia together with the Menderes Massif suffered a semibrittle‐brittle, north‐ south compressional deformation. As a result of this, the whole region was shortened, internally imbricated, and thickened. In the Menderes Massif, thick‐skinned deformation occurred, and its tectonic components were tectonically rearranged. The deeply ­buried metamorphic rocks (core rocks) were thrust above the cover rocks (see Fig. 6.9 for the structural rearrangement of the Menderes Massif). The nonmetamorphic Tauride slices were imbricated and were thrust above the metamorphosed carbonate succession. See the light green and blue slices in the south of the section.

originated in the NeoTethyan Ocean [Şengör and Yılmaz, 1981] (Figs. 6.3a, 6.4, and 6.8b). Convincing evidence for the southerly transporting ophiolite above the Menderes Massif is the ages of the olistostrome deposits, eroded from the ophiolitic nappes, and shed into the flexural depressions formed in front of the nappes. They get progressively younger to the south from the Late Cretaceous to the Middle‐Late (?) Eocene as indicated by the paleon-

tological [Candan et al., 2011] and isotope age data [Lips et al., 2001]. The HP metamorphic rocks similar to the HP metamorphic rocks of the northern region are also identified from the southern periphery of the Menderes Massif (located to the south of the tectonic window) [Ring et  al., 1999, 2007a,b; Okay, 1989; Candan et  al., 2005; Gessner et  al., 2011]. This indicates that the para‐autochthonous thrust sheets, dragged under the

150  ACTIVE GLOBAL SEISMOLOGY

K. Menderes

S

Torbalı

Para-autochthon Gediz

B. Menderes

lim

iye

Se

C¸

ine

N

Allochthone C¸ ine

gˇ da

z

Bo

Bayındır

Boz

dagˇ

Sel

imiy

e

No scale

Figure 6.9  Schematic section across the Menderes Massif showing rearrangements of the different thrust sheets, based on the cross sections and descriptions from various publications including Ring et al. [1999], Candan et al. [2005], Gassner et al. [2011]. The geographical names (black and white) refer to approximate locations of the grabens.

ophiolitic nappes, also reached this region during the nappe transport (Figs. 6.3, 6.4, and 6.8). The scarcity of the marble envelope in the northern regions with respect to the southern regions (Fig. 6.5) may be interpreted to mean that part of the marble envelope detached from the more resistant schist and gneiss base as a decollement style, and was dragged and transported to the southern regions under the nappes. Along the southern boundary of the Menderes Massif there are also nonmetamorphic thrust sheets, made up mainly of the Mesozoic carbonate successions resting above the metacarbonate sequence of the Menderes Massif. These are identified commonly as the Lycian nappes (Figs. 6.7 and 6.8c) [de Graciansky, 1972; Şengör and Yılmaz, 1981; Güngör and Erdoğan, 2001; Ring et  al., 2001b; Regnier et  al., 2007]. The metasedimentary sequence below the Lycian nappes have some key metaboxite layers, which display HP met 500T, 14 KB of about 60–65 my old [Ring et  al., 2001b; Whitney et  al., 2008; Gessner et al., 2011]. The thrusting stage that imbricated the nonmetamorphic and metamorphic carbonate rock slices possibly occurred during the late phase of the orogenic development that affected the Menderes Massif probably during the Oligocene [Bozkurt and Satır, 2000] or Early Miocene [de Graciansky, 1972; Şengör and Yılmaz, 1981; Hayward, 1984b; Güngör and Erdoğan, 2001; Poisson et  al., 2003; Collins and Robertson, 2003; Rimmele et al., 2003a, b; Pourteau et al., 2010]. The southern and northern limits of the Lycian nappes and its boundaries with the Taurus units (sensu stricto) and the carbonate cover of the Menderes Massif have not yet been clearly identified and precisely established. In the central Menderes Massif, a number of major tectonic units (thrust sheets) were differentiated, based on the lithological ordering and the metamorphic grade, for example, the Bayındır, Çine, Bozdağ, and Selimiye units (Figs. 6.8c and 6.9) [Ring et al., 1999; Candan et al., 2005; Gessner et  al., 2011; and the references therein].

Their identities, structural positions, and therefore regional correlations from different parts of the massif are still widely debated. Based on some local observations (e.g., from the Bozdağ horst), the least metamorphosed tectonic unit such as the Rudist‐bearing metasediment (the Bayındır unit) of Upper Cretaceous age [Özer, 1998; Özer and Sözbilir, 2003] is observed at the bottom and the high‐grade metamorphic rocks, for example, the Selimiye unit stands at the top across a thrust plane (Fig. 6.9). This indicates simply that the Menderes Massif has undergone multiple structural rearrangements during a long history of evolution [Şengör et al., 1984; Bozkurt and Oberhansli, 2001; Ring et  al., 2001a; Gessner et  al., 2004]. In fact, Bozkurt and Oberhansli [2001] differentiated two pre‐ Alpine and three Alpine deformation stages. The present tectonic rearrangement has occurred in a late stage, after the massif experienced the major progressive and retrogressive penetrative metamorphic phases as indicated by abrupt changes of the metamorphic grades across the thrusts (Fig. 6.9). Apparently this was due to a basement‐ involved, thick‐skinned shortening deformation, during which the high‐grade metamorphic rocks (the core rocks) were thrust above the envelope rocks (Fig. 6.8c; such tectonic rearrangement corresponds to the out of sequence thrusts of Gessner et  al. [2001 a, b]). This is observed particularly in the northern region close to the Izmir‐ Ankara suture (Fig. 6.9). 6.2.1.1. Discussion Summary of the accounts given in the preceding paragraphs is that the Menderes Massif is a nappe stack (Figs. 6.7, 6.8, and 6.9). At the top there is a meta‐ophiolite (having a 90 my old granulite‐eclogite facies sole) and metamorphosed melange association (the Tavsanlı zone or Selçuk melange), represented by Mg garnet and omphacite‐bearing (450‐500 T‐20‐26 kb) HP metamorphic rocks (Figs. 6.7 and 6.10) [Okay, 1981; Whitney et al., 2011]. Structurally, below the metaophiolite are the

Major Problems of Western Anatolian Geology  151

Figure 6.10  Schematic cross section displaying development of different tectonic zones of the Menderes Massif. (1) Izmir‐Ankara ophiolite and ophiolitic mélange, deeply buried along the subduction zone and rapidly elevated, and thus formed HP metamorphic units (e.g., the Selçuk mélange and the Tavşanlı zone); (2) Continental slope of the Taurus, involved in the subduction zone and rapidly elevated, and thus formed HP metamorphic units, such as the Dilek nappe and the Afyon zone; (3) Taurus continental margin, moderately buried and imbricated, suffered Barrowian metamorphism and formed the main body of the Menderes Massif (e.g., the core and the cover). There are many tectonic slices in the Taurus range, which represent a transitional zone between the nonmetamorphic Taurus and the metamorphic rocks of the Menderes Massif.

­ etasediments; metamorphosed thinly bedded pelagic m limestone, metaflysch, and marble association, possibly representing the leading edge (the continental slope sequence) of the Tauride continental margin, facing the NeoTethyan Ocean in the north (Figs. 6.3, 6.7, and 6.8) [Şengör and Yılmaz, 1981; Yılmaz, 1981, 1987,1997; Yılmaz et  al., 1997]. This broadly corresponds to the Dilek nappe and the Afyon zone of Okay [1981, 2004]. They underwent HP metamorphism as represented by phengite, glocophane, epidote, and albite mineral assemblage (350T‐9Kb) [Okay, 1984; Candan et  al., 2011; Gessner et al., 2011] indicating that the northern edge of the Taurus was also involved in the subduction (Figs. 6.7 and 6.10). The allochthonous and the para‐autochthonous units represent the rocks that were rapidly buried along the subduction zone and then rapidly elevated. The main body of the Menderes Massif cropping out in the tectonic window (Figs.  6.5 and 6.7) represents also a nappe stack, which suffered imbrication and accompanying regional, dynamothermal HT Barrowian metamorphism up to the amphibolite grade (peak metamorphism: 550 T‐6‐8 Kb) (Figs. 6.4, 6.7, 6. 8 a, and Fig. 6.10). The retrograde metamorphic evolution of the Menderes Massif, at the greenschist facies condition was generally

associated with the ascent (exhumation) phases, while mylonitic fabrics penetrated into the rocks along the major fault zones [Bozkurt, 2001b; Çemen et  al., 2006]. Exhumation of the Menderes Massif has occurred in at least two major phases [Gessner et  al., 2001a and b; Bozkurt and Park, 2004; Catlos and Çemen, 2005; Ring et  al., 2007; Çemen et  al., 2006; van Hinsbergen and Schmid, 2012]. Evidence for one more stage of extension [Purvis and Robertson, 2005a, b] dating back to the late Eocene period [Burchfiel et  al., 2003; Lacassin et  al., 2007a,b] is also documented in a number of studies from the Aegean and the surrounding regions [Lips et al., 2000, 2001; Ring and Layer, 2003; Purvis and Robertson, 2004; Rimmele et al., 2005]. In the late stage of the extensions, the Bozdağ horst (dome) of the central Menderes Massif was elevated (Fig.  6.5). The dome is elliptical in shape with an east‐ west long axis. This is compatible with the ongoing north‐ south Aegean extension. The Bozdağ horst is separated in the north and the south by two extensional detachment faults (Fig.  6.5), which are symmetrical with respect to the dome axis. The elevation of the Bozdağ horst began during the Late Miocene time as evidenced by three sets of data: (1)

152  ACTIVE GLOBAL SEISMOLOGY

Figure 6.11  Schematic sections displaying the core complexes of western Anatolia that formed during the two different extensional phases. (a) Late Oligocene–Early Miocene detachment normal faults display top to the north stretching lineations. According to the age data the four major low‐angle normal faults, which led to the development of the major part of the core complex, formed during this period. The arrow refers to the region located between the Çine and Simav detachment faults. (b) The schematic section showing the younger core complex, which was exhumed between the two symmetrical faults, and added up to the D refers to detachments main body of the Menderes Massif as the Bozdağ dome. This extensional event started during the Late Miocene.

Synextensionally developed red coarse conglomerates were eroded from the elevated Bozdağ dome and shed into the surrounding depressions (Fig. 6.23). They yield Upper Miocene age (for a discussion on this subject see the Cover Rocks, Section  6.2.3, and the accompanying paper by Y. Yılmaz; Chapter 2, this volume). (2) The isotope data from the detachment zones that bound the Bozdağ horst yield narrow age span varying between 8 and 5 my [Gessner et al., 2001a and b, 2011; Çemen et al., 2006; Bozkurt et  al., 2011]. (3) Indirect supporting data may also be derived from the various kinematic analyses on the ongoing Aegean extension. They reveal that initiation of the extension (possibly the latest stage) is not older than the Late Miocene [McKenzie, 1972, 1978; Jackson and McKenzie, 1988; Kastens, 1991; Tirel et al., 2004; Westeway et  al., 1994, 2005]. During the Late Miocene period, the Bozdağ horst (Fig.  6.5) was the main morphological high in western Anatolia (see Cover Rocks, Section  6.2.3). The previously elevated western Anatolian and the Aegean land mass subsided close to sea level during this period [Yılmaz et al., 2000]. Exhumation of a much larger region along the low angle detachment faults, namely the Simav and Çine detachment faults located farther north and south with respect to the Bozdağ horst, occurred earlier (Figs.  6.5 and 6.11). This approximately elliptical province aligns in a north‐northeast–south‐southwest direction. The region was later divided into separate horsts during the development of the present horst‐graben structures

(Fig. 6.11). Exhumation of this region occurred between 30 and 20 my [Verge, 1995; Işık and Tekeli, 2001; Ring and Collins, 2005; Thomson and Ring, 2006; Gessner et  al., 2001a and b, 2011; Erkül, 2010; Bozkurt et  al., 2011]. A supporting evidence for this exhumation period is obtained from the field data that the Lower Miocene terrestrial sediments were deposited directly on the high‐ grade metamorphic rocks in the entire northern part of the Menderes Massif. The Lower Miocene sediments also were deposited and sealed the contacts between the Menderes Massif and the surrounding tectonic units (e.g., the ophiolitic rocks of the Izmir‐Ankara suture). The north‐northeast trending elliptical form of the main body of Menderes Massif was delineated during this period (Fig. 6.5). The sense of shear directions and the associated major lineation directions measured from different sectors of the Menderes Massif are indecisive yet in quantity and geographical coverage to yield clear connections between the different extensional phases and their dates (see van Hinsbergen et al. [2010], and van Hinsbergen and Schmid [2012], for a discussion on the available data), and to separate different extensional and compressional phases [see the ongoing debate on the southern Menderes Massif: e.g., Gessner et al., 2001a and b, 2004, 2011; and Bozkurt, 2007; Güngör, 1998; Erdoğan and Güngör, 2004; Bozkurt and Oberhansli, 2001]. An oblique fault with a more prominent dip‐slip component defines the western border of the Menderes Massif

Major Problems of Western Anatolian Geology  153

Figure  6.12  Schematic section displaying tectonics of western Anatolia during the Late Oligocene (?)–Early Miocene (?) period.

against the ophiolitic rocks. It strikes north‐northeast (Fig. 6.5), subparallel to the long axis of the Massif and the associated structural data derived from the contact zone. The long axial trend of the Menderes Massif and the associated structures are not compatible with a broadly north‐south continuous extension that was previously proposed [Gautier et al., 1999; Tirel et al., 2009; Jolivet and Brun, 2010]. Instead, they appear to comply more favorably with a west‐northwest–east‐southease (present position) extensional direction, unless western Anatolia was rotated counter‐clockwise for more than 45° between the Early and Late Miocene (the angle between the east‐ west trending long axis of the Bozdağ horst and the main north‐northeast–south‐southwest trending culmination of the massif (see Fig. 6.5). Van Hinsbergen et al. [2010] calculated about 25–30° counter‐clockwise rotation between the northern and southern edges of the western Anatolia (the eastern and western parts of the Menderes Massif). Even this amount of counter‐clockwise rotation does not satisfy the required amount of rotation. The three different major metamorphic events that affected the Menderes Massif occurred partly simultaneously. 1. HP metamorphic ages range from 80 my to 20 my [Okay, 2001, 2004; Candan et al., 2005; Ring et al., 2007; Whitney et  al., 2008; Gessner et  al., 2011; Jolivet et  al., 2013], from the north (e.g., the Tavsanlı zone; 80–70 my) to the southwest (e.g., from the Selçuk Blue schist and its southwesterly equivalent in the Cyclades; 42–32 my) [Jolivet et al., 1996, 2004; Ring et al., 2007b, Plunder et al., 2013] and farther south from the Cyclades Islands (19 my or younger). 2. HT metamorphic ages also display a wide span from 62 to 21 my [Şengör et al., 1984; Bozkurt and Satır, 2000; Bozkurt and Oberhansli, 2001; Whitney and Bozkurt,

2002, Ring et al., 2001a, Catlos and Çemen, 2005; Thomson and Ring, 2006; Lagos et al., 2007; Gessner et al., 2011]. It is commonly accepted that the first major phase corresponds to the Late Eocene, and is coeval with the inception of the magmatism in the region (see Section 6.2.2, Magmatic Associations). Therefore, the two simultaneous events may be linked with the delamination of the lithospheric mantle, which accelerated the HT metamorphism and generated the magmatism. The delamination of lithosphere from the crust increased the density of the subducting plate and thus triggered development of a rollback event. This, in turn, caused upwelling of the asthenosphere between delaminated lithosphere and crust, and thus heated the overriding plate. 3. The cooling ages of the metamorphic rocks range from 50 to 5 my [Gessner et al., 2001a and b; Lips et  al., 2001; Gessner et al., 2001a and b, 2011; Catlos and Çemen, 2005; Catlos et al., 2008; Purvis et al., 2005; Bozkurt et al., 2011]. Such a wide range of ages obtained from the three different metamorphic events indicates that the synorogenic deformation affecting the Menderes Massif and the Cyclades complex, farther west, survived to the Miocene period [Yılmaz et al., 2000; Rimmele et al., 2005; Yılmaz, 2010; Jolivet et  al., 2010; Chatzeras et  al., 2011; Papanikolau and Vassilakis, 2010; Yücel-Öztürk et al., 2015]. In fact, the regional geologic data derived from the entire Aegean–western Anatolian region support the view that the northern and the southern parts of western Anatolian and Aegean regions suffered a north‐south shortening deformation sometime during the Late Oligocene and Early Miocene period. The data for this is shown in Fig. 6.12, and may be listed as follows: 1. In the north, the Istranca Massif was thrust to the north and the south above the Oligocene sediments

154  ACTIVE GLOBAL SEISMOLOGY

[Şentürk and Karaköse, 1998; Şentürk et al., 1998; Siyako et al., 1989; Siyako, 2006; unpublished TPAO seismic data]. 2. The Lycian nappes transported southward (Figs. 6.4 and 6.8), and tectonically emplaced on the Upper Oligocene–Lower Miocene flexural flysch basin sequence, developed above the Beydağları Autochthon [de Graciancky, 1972; Gutnic et al., 1979; Hayward, 1984a,b; Güngör and Erdoğan, 2001; Collins and Robertson, 2003 (two different tectonic phases may be distinguished associated with the emplacement of the Lycian nappes onto the Taurus Beydağları relative autochthon; a south‐vergent thrusting occurred first [Bozkurt and Satır, 2000], and then a gravity sliding‐translation stage occurred during the extension phase that followed the thrusting). 3. The entire western Anatolia and the present Aegean Sea regions were elevated above the sea; a gradual transition from marine environment to continental environment is recorded in the sedimentary sequences in the northern Aegean, Marmara, and Thrace regions during the Oligocene period (TPAO unpublished reports, Şentürk et al., 1998; Siyako and Huvaz, 2007 and the references therein]. A short spell of north‐south compression phase is also claimed to have occurred to the end of Early Miocene–Middle Miocene transition [Koçyiğit, 1999a, b; Bozkurt and Rojay, 2005; Chatzeras et al., 2011]. The wide range of overlapping metamorphic ages may also be ascribed to repeated pulses of shortening (nappe transport, thickening, and the consequent burial [HT metamorphism] followed by collapse and exhumation [retrograde metamorphism]). The highly attenuated

Aegean lithosphere [Meulenkamp et al., 1988; Tirel et al., 2004; Spakman and Wortel, 1988, 2004], and the closely associated long oceanic slab, subducting under the Aegean‐Anatolian Plate, are expected to have caused plate boundary processes such as consecutive phases of extension and compression. The three groups of metamorphic ages appear to be broadly younging toward the south. If this correlation is valid, it may than be concluded that the exhumation front migrated from the north to the south between the Late Cretaceous and the Late Miocene periods. The Aegean extension is interpreted to have resulted from retreat of the subducting eastern Mediterranean oceanic lithosphere [Le Pichon and Angelier, 1979; Meulenkamp et  al., 1988; Tirel et al., 2004]. The retreats took place in pulses, which have possibly caused the migration of the exhumation front from the north to the south. The young front may be correlated with the Cretan and the Cyclades detachments in the Aegean. The exhumation of the Cretan nappe stack occurred between 24 and 12 Ma. On the Peloponnesus, the oldest sediment, deposited on the exhumed metamorphic rocks are Late Miocene (10.8 and 10.4 my; Van Hinsbergen and Schmid[2012])–Early Pliocene in age [Frydas, 1993]. The major phases of the exhumation and the associated extensions appear to be broadly associated with the periods of westward escape phases of Anatolia (Fig. 6.13) [Şaroğlu and Yılmaz 1991; Yılmaz et al., 1997]. The two events may well be genetically connected to one another as the cause and the consequence. The extensional phases of western Anatolia may in part be ascribed to the

Figure 6.13  Map showing Anatolian major suture zones [after Şengör and Yılmaz, 1981] and the ages of suturing [after Yılmaz et al., 1997]. The arrows indicate times of the westward escapes of Anatolia following the suturing [see Yılmaz et al., 1993, 1997; Yılmaz and Tüysüz, 1993; for the timing of the westward escapes].

Major Problems of Western Anatolian Geology  155

westward escape of Anatolia, which followed the major collisional events (Fig.  6.13). This may have caused the slab steepening and the consequent delamination of the mantle lithosphere from the crust allowing the asthenosphere inflow into the overriding plate (for a discussion see Section 6.2.2, Magmatic Associations). Therefore, the three events (escape, rollback, and delamination) may be genetically connected to one another, and the escape may be the major factor responsible from the other two events. The escape triggered the slab retreat, and, in turn, generated the delamination of the lithospheric mantle. 6.2.2. Magmatic Associations The Magmatic rocks cover large areas in western Turkey (Fig. 6.14) from the Thrace [Yılmaz and Polat, 1998] and the Marmara region [Genç and Yılmaz 1997] in the north

to the Bodrum Peninsula in the south [Genç et al., 2000]. They were formed in two phases, the early phase and the late phase [Yılmaz, 1989; Seyitoğlu et  al., 1997; Yılmaz et al., 2000, 2001; Erkül et al., 2005a,b, 2008]. During the early phase, granitic stocks and small plutons [Yılmaz 1989, 1990; Yılmaz et  al., 2000; Karacık et  al., 2008; Altunkaynak et al., 2012a,b], and intermediate and felsic volcanic rocks were extensively developed. Major characteristics of the early phase of the magmatic activity may be listed as follows: 1. In the entire western Anatolia, emplacement of plutonic and hypabyssal rocks and the edifices of the volcanic rocks was controlled commonly by approximately north‐ northeast–south‐southwest trending structures (Fig. 6.14) [Yılmaz et al., 2000; Purvis and Robertson, 2005b, Karacık et  al., 2007; Ersoy et  al., 2010; Erkül and Erkül, 2010; Karaoğlu et  al., 2010]. The lower Miocene terrestrial

Figure 6.14  Geology map of western Anatolia showing location of major granitic plutons (red), distribution of the volcanic associations of the early volcanic phase (the green unit), the Neogene terrestrial sediments and some major north‐south grabens (the yellow areas). Broken red lines display southerly migrated volcanic fronts. The numbers refer to the ages of the volcanic fronts. The green arrow indicate general trend of the migration between the Oligocene and late Miocene. The age data are derived from the following publications: For the Ayvacık‐ Ayvalık region, Aldanmaz et al. [2000], Ercan et al. [1996]; the Yuntdağ area, Ersoy et al. [2014]; the Bigadiç basin, Helvacı [1995], Erkül et al. [2005]; the Dikili Bergama area, Borsi et al. [1972], Aldanmaz et al. [2000]; the Soma‐Akhisar region, Ercan et al. [1996], Innocenti et al. [2005], Ersoy et al. [2014]; the Gördes basin, Seyitoğlu et  al. [1992], Purvis et  al. [2005]; the Foça‐AliAğa region, Altunkaynak [2010]; the Izmir region, Borsi et  al. [1972]; the Karaburun Peninsula, Borsi et al. [1972], Helvacı et al. [2009], Karacık et al. [2013]; the Torbalı area, Sözbilir et al. [2011]; the Samos‐Söke region, Ercan et al. [1985], PePiper and Piper [2007].

156  ACTIVE GLOBAL SEISMOLOGY

Figure 6.15  Geology map of the Kozak magmatic center [modified from Altunkaynak and Yılmaz, 1998]. The granite was emplaced in the country metamorphic rocks of the Paleozoic age (brown peripheral zone). The ring adjacent to the granite is the metamorphic aureole, formed around the granite body during the early Miocene time. The hypabyssal and volcanic rocks surround the granite (yellow and white rings) that are closely associated with granite in time and space.

sedimentary rocks of western Anatolia were commonly deposited within north‐northeast trending structural depressions lying subparallel to the volcanic highs. Due to the proximity to the volcanic centers, the sediments alternate with, and include, increasing amounts of volcanoclastics, pyroclastics (including ignimbrite flows; Karacık and Yılmaz, 1998], and lava layers (Figs. 6.21, 6.22). 2. Commonly, they display a cluster of plutonic and associated volcanic rocks. Shallow level granites and temporally and spatially associated hypabyssal and volcanic rocks collectively form a collapsed caldera environment (Figs.  6.15 and 6.16) [Yılmaz, 1989; Altunkaynak and Yılmaz, 1998, 2000; Karacık and Yılmaz, 1998; Genç et  al., 2000; Yılmaz et  al., 2001; F.Erkül et  al., 2009; Altunkaynak et al., 2012 a; Altunkaynak and Dilek, 2013]. A number of caldera‐type granites and associated volcanic centers have been mapped in the region as exemplified from the Ezine‐Bayramiç [Karacık and Yılmaz, 1998], Kozak [Altunkaynak and Yılmaz, 1998] (Fig. 6.15), and the Bodrum area [Yılmaz et al., 1999b; Genç et al., 2001; Altunkaynak and Yılmaz, 2000] (Fig. 6.16). 3. Intermediate and felsic rocks are dominant. The region is nearly devoid (Ersoy and Helvacı 2016) of the

true basaltic rocks in this period. Most of the volcanic centers are stratovolcanoes, which produced pulses of lavas and associated pyroclastic rocks evolving in compositions from basaltic andesite to latite and dacite (rarely to rhyolite) or vice versa. Basic magmas that were probably generated in the heterogeneous upper mantle, evolved extensively rising through the thickened crust. 4. The magmatic activities of this phase began in the northern regions during the Late Eocene time. The granitic rocks yield slightly older but overlapping ages with the associated volcanic rocks (see Ersoy et al. [2014] for the ages of the western Anatolian granites and the associated volcanic rocks). The granitic rocks were formed about 40–35 my ago [Bingöl et al., 1982; Delaloye and Bingöl, 2000; Karacık et al., 2008; Altunkaynak et al., 2012 a, b, Ersoy et al., 2014], and the volcanic rocks about 35–30 my ago [Yılmaz and Polat, 1998; Elmas et al., 2016]. The magmatic activity of this phase continued almost uninterruptedly to the Late Miocene [11–9 m; Ercan et al., 1985,1996; Aldanmaz et al., 2000; Agostini et al., 2008; Helvacı et al., 2009; F.Erkül et al., 2009; Dilek and Altunkaynak, 2009; Dilek and Sandvol, 2009; Altunkaynak et al., 2012a; Erkül and Erkül, 2010; Hasözbek et al., 2010; Prelevic et al., 2012; Ersoy and Helvacı 2016].

Major Problems of Western Anatolian Geology  157

Figure 6.16  Geology map of the Bodrum volcanic province displaying the Bodrum caldera, and distribution of the associated plutonic, hypabyssal, and volcanic rocks. The broken red and blue lines represent parts (I, II) of the caldera, displaced by oblique fault (thick red broken line), which cut and right‐laterally offset the northeastern part of the caldera. The southwestern half of the caldera is buried under the waters of the Aegean Sea. The black broken line is a small crater that was built on the southern shoulder of the Bodrum caldera (Turgut Reis caldera, see Altunkaynak and Yılmaz [2000]) [modified from Yılmaz et al., 1999b; Genç et al., 2000].

The ages obtained from the magmatic centers get younger across western Anatolia from the north to the south, and display an overlapping pattern (Fig. 6.14). The granite ages span from 40–50 my in the northern Marmara region to 12–10 in the south (the Bodrum Peninsula and the neighboring islands; Roberts [1992]; Altherr et  al. [1988]). This reveals that the magmatic‐volcanic front migrated steadily from the north to the south (noticed previously by Besang et al. [1977]; Ercan et al. [1979]; Bingöl et al. [1982]). It began in the Balkan‐Thrace and the northern Marmara region and extended to the south of the Biga Peninsula in the Early‐Middle Miocene, that is, Kozak, Evciler, Çataldag, Kestanbol, Ilica‐Samli, Eybek, Egrigoz, Alaçam, Koyunoba granites and the intimately associated volcanic rocks [Yılmaz, 1989; Ercan et al., 1996; Altunkaynak and Yılmaz, 1998, 1999; Genç, 1998; Yılmaz et  al., 2001; ̇ Özgenç and Ilbeyli, 2008; Erkül, 2010]. It reached the Izmir region about 13–10 my ago (e.g., Kiraz, 13 my; Kuşadası, 12 my; and Urla, 11 my), and finally extended to the Bodrum magmatic province about 10–9 my ago (Fig. 6.14). 5. The magmas have evolved from shoshonitic and high‐K calk‐alkaline to calk‐alkaline, and their compositions

form a cluster displaying a common character and origin (Fig. 6.17). The geochemical and isotope values display affinity to an Andean type island arc, indicating a source, metasomatically enriched from the subduction components; low Nb and TiO2 contents and hence low Nb/La, Ti/Nb ratios, distinct trace element pattern with anomalies at Nb and Ti, enrichment of LILE over HFSE and LREE resulting in high Ba/Nb, Th/Nb, Ba/La, K/Ti, and Th/La ratios. They are strongly depleted in highly compatible elements with respect to mantle, but enriched in highly compatible elements. Ti/Hf, La/Nb, Sr/Y, La/Yb ratios imply that the magma was not derived directly from the slab melting but modified from slab released fluids. Both lithospheric and asthenospheric mantle melts appear to have contributed to the source region [Fytikas et al., 1984; Innocenti et al., 2005; Yılmaz et al., 2000, 2001; Aldanmaz et  al., 2000; Altunkaynak and Dilek, 2006; Conticelli, 1998; Conticelli et  al., 2009; Prelevic et al., 2010b, 2012; Altunkaynak et al., 2012a, b; Seghedi et al., 2015]. The isotope values further indicate that the crustal components are also involved in the enrichment [Fytikas et  al., 1984; Pearce and Peate, 1995, Yılmaz

158  ACTIVE GLOBAL SEISMOLOGY 16 14

Phonolite

Na2O + K2O

12 10

Late Miocene basaltic lavas (EGA) Middle Miocene basaltic rocks Middle Miocene intermediate rocks Early Miocene porphyritic lavas Early Miocene (Kovacli) dykes Early Miocene ignimbrites Early Miocene porphyritic lavas

Tephriphonolite Foidite

A l k a l i n e Tephrite & Basanite

6

Trachyte & Trachydacite

4

0 40

EGA

Trachyandesite

Phonotephrite

8

2

DAB

Basaltic trachyandesite Trachybasalt Picro basalt

Basalt

45

50

Rhyolite

Dacite

S u b a l k a l i n e Basaltic andesite

55

Andesite

60 SiO2 wt.%

65

70

75

80

Figure 6.17  Compositional variations and clusters of the early (red elliptical area) and late phases (green dots) of the volcanic associations of western Anatolia [modified from Aldanmaz et al., 2000, Fig. 3].

1989; Aldanmaz et al., 2000; Yılmaz et al., 2001; Innocenti et al., 2005; Helvacı et al., 2009; Conticelli et al., 2009; Altunkaynak et al., 2010, 2012]. Based on the detailed geochemical studies on the granitic rocks, Altunkaynak et al. [2012a,b] and Altunkaynak and Dilek [2013] stated that the compositional variations observed are the result of open system processes (AFC; Assimilation-Cystallization, and/or MASH; melting, assimilation, storage, homogenization) as hitherto suggested [Yılmaz et  al., 2001] rather than a reflection of ­different compositions of crustal lithologies. More asthenosphere‐derived melts found their way to the surface with lesser degrees of crustal contamination. The late phase of the magmatic activity produced mainly basic volcanic rocks, which were virtually missing during the early phase. In contrast, the extensively developed granitic and felsic volcanic rocks of the previous phase were missing in this period. The basic volcanic rocks are sporadically developed and much less extensive. They form a distinctly different compositional cluster compared to the early phase (Fig. 6.17). The late phase of magmatic activity began during the late Middle Miocene, but developed extensively during the Late Miocene period. Activities of this volcanic phase survived until the Quaternary. The late volcanic phase does not display a southerly younging trend, occurred simultaneously in the entire region. Commonly, there is a time gap between the developments of the Early Phase and the Late Phase. The gap is wide in the northern regions. In the Thrace region, for example, it is about 20 my, but narrows toward the south. The two phases may be briefly overlapping around the southern volcanic centers [Prelevic et al., 2012].

Major characteristics of the late volcanic phase may be listed as follows: 1. It is observed throughout western Anatolia as small and scattered outcrops. 2. It is closely linked with the Late Miocene or younger structures that commonly trend east‐west in the west of western Anatolia (connected with the present graben structures, such as Ezine and Bergama, Fig.  6.21) and north‐south in the eastern sectors (the Isparta angle and in the regions farther north). 3. It is represented commonly by alkali‐olivine basalt, basanite, trachybasalt, and the associated basic rocks (SiO2, 42–50%; Yılmaz et al. [2001]; Fig. 6.17). 4. Geochemically, this group resembles the intraplate (continental) basic lavas with high MgO, Ni, Cr, and low Pb contents and contains mantle xenoliths [Bunbury, 1992, 2006] revealing its mantle source, but some of them display enrichment of LILE, HFSE, and LMREE, and low 87Sr/86Sr (< 0.74) and high 143Nd/144Nd ratios (0.512927–0.51298) [Patton, 1992a,b, Yılmaz 1989,1990, 2002, Yılmaz et  al., 2001; Aldanmaz et al., 2000; Innocenti et al., 2005; Altunkaynak and Dilek, 2006]. The trace element patterns imply that the source was variously enriched (metasomatized) by incompatible elements compared with the depleted mantle. 5. It is stated also that there may be a correlation between the composition of the volcanic rocks of this period and the amount of extension [Patton, 1992a]. There is another group of volcanic rocks in the Aegean‐ western Anatolian region that is geochemically different from the other two groups. This group is distinguished as high Mg ultrapotassic volcanic rocks (K2O > 5% and Mg > 2.5%), and were identified in various parts of western Anatolia [Savaşçın and Oyman, 1998; Yılmaz et al., 2001;

Major Problems of Western Anatolian Geology  159

Doglioni et al., 2002; Innocenti et al., 2005; Altunkaynak and Dilek, 2006; Ersoy and Helvacı, 2007; Ersoy et  al., 2008; Akal, 2008, 2012; Karacoğlu et  al., 2010; Prelevic et al., 2008, 2010a, b, 2012, 2015]. They crop out preferentially above the Menderes Massif, in and around its eastern and northeastern peripheries, and also along the north‐south Isparta‐Afyon trend as small volcanic cones. The High Mg ultrapotassic rocks show petrochemical similarities to the lamproites [Francalanci et  al., 2000; Innocenti et  al., 2005; Akal, 2008; Yılmaz K., 2010; Prelevic et  al., 2008, 2010a, b; Tommasini et  al., 2011; PePiper et al., 2013]; they are extremely enriched in large ion lithophile elements, radiogenic 87Sr/86Sr and 207Pb/204Pb [Prelevic et al., 2008, 2010; Conticelli et al., 2009], and are considered to have formed from initially depleted and then strongly enriched (by the subduction components) mantle [Zanetti et al., 1999]. 6.2.2.1. Discussion There are a number of petrochemically oriented studies on the magmatic associations of western Anatolia, starting from Washington [1894]. Different views have been proposed on the origin of the magmas particularly after the plate tectonic concept [see Aldanmaz et al., 2006; Prelevic et al., 2012, for a review]. Despite the rich data available on the source of the magmas, the mechanisms proposed on the magma generation vary greatly. They commonly consider active role of the following factors: (1) Major parameter is the subduction of the eastern Mediterranean oceanic lithosphere under the Aegean‐Anatolian Plate; (2) the irregular geometry of the slab; (3) the rate of the subduction; (4) the dip angle and the rollback (retreat) of the subducting slab; (5) the slab breakoff or tears (window) or segmentation of the slab allowing upwelling of hot asthenosphere [Prelevic et al., 2015]; (6) detachment or decoupling of the thickened lithosphere (lithospheric mantle separated from the thickened crust); and (7) presence, absence, or even the ­tectonic positions of the mantle wedge. One or more of the factors stated above were considered to have played a significant role for the magma generation [Wortel and Spakman, 1992; Roberts et al., 1992; de Boordes et al., 1998; Conticelli et al., 1998, 2009; Faccenna et al., 2003, 2004, 2006, 2010; Aldanmaz et al., 2006; Çoban and Flower, 2007; Rosenbaum et al., 2008; Agostini et al., 2008; Bailey et  al., 2009; Dilek and Altunkaynak 2009; Dilek and Sandvol, 2009; Ersoy et  al., 2010; Thomassini et al., 2011; Royden and Papanikolaou, 2011; Altunkaynak et al., 2012a,b; Altunkaynak and Dilek, 2013; Jolivet et al., 2013; Elmas et al., 2016; see also Dilek and Altunkaynak, 2009; Altunkaynak et  al., 2012a,b; Lustrino et  al., 2011; and Prelevic et  al., 2012, for the review of the subject]. However a critically important factor in evaluating the genesis of the magmas in the Aegean–western Anatolia region is

to know the geological‐tectonic constraints. Therefore, in the following lines some major factors, which played as controlling elements, are listed in evaluating the cause of the magma generation within time‐space reference, and the tectonic events and environment. 1. The magmatic activity began during the late Eocene time and continued nearly uninterruptedly to the present. 2. The orogeny and the associated major orogen across north‐south shortening lasted interruptedly to the Miocene (see Section 6.2.1, The Menderes Massif). 3. The extensions occurred at least in two major phases. The present graben‐horst system, which is assumed to be linked with the latest phase of the extension, did not begin before the Late Miocene time. 4. There is a close link between the early phase of magmas and the approximately north‐south trending extensional structures. 5. The magmatic rocks of the early phase display an Andean‐type arc geochemical signature. The magmas were formed from metasomatically enriched, heterogeneous lithospheric mantle source, and contaminated from the continental crust on the way to the surface. During this period, the crust was anticipated to be excessively thickened as a result of postcollisional convergence. 6. During the early phase, the magmatic front steadily migrated from the north to the south. The early phase is partly contemporaneous with the periods of the peak  metamorphism of the Menderes Massif (see Section 6.2.1). The two major and partly simultaneous events, the magmatism and metamorphism, lead one to assume that an extra heat was added to the source region by way of some of the mechanisms stated above. 7. The tomographic image of the eastern Mediterranean slab may be traced under the Aegean to depths of more than 800 km [Hinsbergen et al., 2010]. Based on the length of the subducting slab, the Aegean subduction is stated to be continuing since at least the late Cretaceous time [Spakman et al., 1988; Wortel and Spakman, 1992, 2000; Spakman and Wortel, 2004; Faccenna et al., 2006; Dilek and Sandvol, 2009; van Hinsbergen et  al., 2004, 2010; Hafkenscheid et al., 2006]. According to the tomographic images, more than two stages of rollback are identified along the whole length of the slab. The southward retreat of the subducting slab and the southward migration of the magmatic front appear to be temporally and spatially connected leading us to infer that the two events may be genetically connected, and the early phase of magmatic activity was generated in association with the subduction of the slab. It evolved in the suprasubduction environment and was influenced by one or more of the parameters that are listed above, (e.g., slab delamination and the consequent upwelling of the mantle to the subcrustal levels; see the discussion in Section 6.2.1). The late phase began during the period when the lithosphere and the crust

The Neogene cover rocks of western Anatolia are terrestrial deposits. They form three tectostratigraphic units as the lower, middle and upper units, which were developed in different tectonic environments, and thus are separated by regional unconformities (Figs.  6.18 and 6.19). The terrestrial units cover an age span from the Early Miocene to the present. The lower unit is Early to Middle Miocene in age. The volcanic rocks commonly accompany the sediments. Ages of the volcanic rocks, intercalated with the sediments, range from 23 to 16 my [Borsi et al., 1972; Ercan et al., 1979,1985, 1996; Yılmaz, 1989; Seyitoğlu, 1997; Yılmaz et  al., 2000; Aldanmaz, 2002; Purvis and Robertson, 2004, 2005; Erkül and Erkül, 2010; Ersoy et al., 2011; Göktaş, 2012]. The Early‐Middle Miocene rocks were commonly deposited within approximately north‐south trending structural depressions (Fig.  6.20). The trend deviates slightly from the north‐south either to north‐northeast or north‐northwest in the northern and the southern regions of western Anatolia, respectively (Fig. 6.20). The north‐south grabens are bounded commonly by oblique faults with major dip‐slip and subordinate strike‐slip components. The faults controlled the sediment deposition, and thus display fining sideward‐upward profile away from the fault zone [Yılmaz et  al., 2000]. Commonly, the north‐south grabens are delimited by the volcanic highs (horsts). Therefore, they lie subparallel to the volcanic axes (Fig. 6.14). Examples of the subparallel horst‐graben pairs may be given from the northern and southern sides of the Bergama graben, where the approximately north‐south trending grabens are bounded by the Ayvalık ignimbrite high and the Kozak magmatic center (Figs.  6.21, 6.22) [Karacık and Yılmaz, 1998, 2001; Temel et  al., 2004]. Studies on the

100–250

Upper unit

250–750

Low-energy sediments Middle unit

Late Miocene- Early Plioc.

UNCONFORMITY lacustrine deposits; white limestone, siltstone

250–>100

Lower unit Menderes Massif basement

6.2.3. The Cover Rocks

Fluvial deposits

Continental red beds

UNCONFORMITY Early-Middle Miocene

began to turn rapidly to the normal thickness (see the discussion on the Neogene cover rocks, Section  6.2.3). The geochemically distinct basic magmas of this volcanic phase were possibly generated due to the extensionally induced rapid decompressional melting from the previously permanently metasomatized heterogeneous mantle. The magmas were extruded commonly along the east‐west trending extensional structures. This volcanic phase began in the Late Miocene and prolonged to the Quaternary. The high Mg ultrapotassic volcanic rocks that were derived from the initially highly depleted and subsequently enriched mantle were formed favorably in a suprasubduction environment either in a fore arc [Flower and Dilek, 2003] or in a back arc setting [PePiper et al., 2013]. Their location at the time of development is critically important in assessing their tectonic position and the origin. However, this problem has not yet been tackled effectively.

Pliocene-Quat.

160  ACTIVE GLOBAL SEISMOLOGY

Lignite

Sedimentary rocks; Fine-grained continental deposits

Volcanic association UNCONFORMITY Metamorphic rocks

Figure 6.18  Generalized tectonostratigraphic section of western Anatolia [modified from Yazman et al., 1998, and Yılmaz et al., 2000]. The three tectostratigraphic units that are separated by major unconformities are observed throughout western Anatolia (see Fig. 6.19 for the data representing other areas of western Anatolia. See also T.Erkül et al. [2009] for the Bigadiç and Demirci basins, Ersoy et al. [2010], for the Selendi, Gördes, and Emet basins. See also a compilation of the data from various basins in Ersoy et al. [2014]).

north‐south grabens from the entire western Anatolia, for example, the Etili‐Küçükkuyu and the Ahmetli grabens in the northern region [Karacık and Yılmaz 1998; Yılmaz and Karacık, 2000]; the Demirci, Gördes, and Selendi grabens in the central western Anatolian region [Ercan et al., 1996; Seyitoğlu, 1997; Yılmaz et al., 2000; Purvis and Robertson, 2005 a, b; Ersoy et al., 2011]; and the Ören graben in the southern region [Gürer and Yılmaz, 2002] display collectively that they share many similar features. For example, contents of the graben fills (Fig. 6.19) may be summarized as alternating conglomerate, sandstone, siltstone, claystone, marl, and shale with alternating coal seams. They were deposited in lake and connecting fluvial environments. The middle unit is Late Miocene–Early Pliocene in age (Figs.  6.18, 6.19). The successions are represented

Major Problems of Western Anatolian Geology  161

Figure 6.19  Measured stratigraphic sections from different parts of western Anatolia [modified from Yilmaz, 2000]. LU, MU, and UU are the lower, middle, upper units, respectively.

primarily by white limestone, marl, and siltstone deposited in a low‐energy lake environment. They are the most extensive Neogene rocks in western Anatolia, and apparently were formed in interconnected lake basins [Benda, 1971; Benda et al., 1974; Akyürek and Soysal, 1983; Emre et al., 1998; Steinenger et al., 1987; Görür, 1998; Gentry and Heizmann, 1996; Alçiçek, 2010; see also Yılmaz in this volume for discussion on the age and environment of deposition of the middle unit]. The shallow lakes appear to have covered all of western Anatolia before the development of present east‐west horst‐graben structures. The lakes were penecontemporaneous with a phase of denudation, which formed a regionwide flat‐lying erosional surface.

The Upper Miocene–Lower Pliocene limestone sequence and the overlying erosional surface were fragmented after the Early Pliocene period, when the present east‐west horst‐graben system began to form. The young, steep faults of this period cut and displaced the older units and structures [Angelier et  al., 1981; Patton, 1992; Yılmaz et  al., 2000; Gürer at al., 2009; Purvis and Robertson, 2005a]. Commonly a clear steep scarp and a truncated spur easily identify the faults that bound the east‐west grabens. But the fault planes are generally difficult to see due to degrading and weathering. There are about 10 approximately east‐west oriented grabens in western Anatolia (Fig. 6.1). The grabens close eastward and enlarge westward. The best‐developed

162  ACTIVE GLOBAL SEISMOLOGY

Figure 6.20  Schematic map showing locations of approximately north‐south trending grabens of Early‐Middle Miocene age [modified from Yilmaz et al., 2000]. EG = Etili graben; AG = Ahmetli graben; YGG = Yenice‐Gönen graben; ZG = Zetindağ graben; GÇG = Gömeç graben; GG = Gördes graben; DG = Demirci graben; SG = Selendi graben; UG = Uşak‐Ulubey graben; ÇG = Çubukludağ graben; MKG = Mustafa Kemal Paşa graben; EVG = Evciler graben; TG = Cumaovası graben; ORG = Torbalı graben; ADG = Ahmetli-Aydın-Çine graben; YG = Yatağan graben; KG = Karacasu graben; SOG = Sarıca Ova graben.

grabens are Büyük Menderes, Gediz, Edremit, and Kerme (Fig. 6.1). Some of them are on shore, and some others extend into the offshore Aegean Sea. They are about 100–150 km long, and 5–15 km wide. The grabens are asymmetrical. In each graben, one margin is characterized by steeper topography, associated with surface breaks of the active faults. On the footwall blocks of the grabens, numerous closely spaced, steeply dipping normal faults are readily observed. The geomorphology associated with the grabens and the horst dominates the landscape of western Anatolia, and controls the major west‐flowing drainage system. The intervening horsts (> 1000 m) control lateral subsequent streams. On the grabens, there are some geomorphological studies [Erinç 1954, 1955; Erol, 1982; Ozaner and Bozbay, 1982; Bircan et al., 1983; Patton, 1992; Erkül and Hakyemez, 1993] and a number of geological studies [Yağm urlu 1987; Yağm urlu and Karaman, 1987; Roberts, 1988; Aksu et  al., 1990; Patton, 1992, 1995; İztan and Yazman, 1990; Cohen et al., 1995; Dart et al., 1995; Ediger et  al., 1996; Emre, 1996, 1998; Emre

et  al., 2005; Seyitoğlu, 1997; Seyitoğlu and Scott, 1994,1996; Gürsoy et al., 1998; Koçyiğit et al., 1999a, b; Yılmaz, 2000; Yılmaz et  al., 2000; Sözbilir, 2001; Kazancı et  al., 2009, 2011]. The graben fill is essentially  fluvial coarse sandstone and conglomerate of Quaternary age [Ünay et al., 1995; Ünay and Göktaş, 1999; Hakyemez et al., 1999, 2013; and the references therein). They are observed all along the graben shoulders, where they were elevated by the listric ­ ­normal faults. In the central part of western Anatolia, the tectonically active sides of the grabens are the sides adjacent to the major horsts (e.g., the Bozdağ horst that separates the Gediz graben and the Büyük Menderes graben; Figs. 6.1 and 6.5). Along the other sides of the grabens, the topography is subdued and the structures are less distinct. The amount of rotation of sediments on the footwall block about a horizontal axis is apparently related to the extension in the graben area. The east‐west grabens cut and displace the north‐south trending grabens [Koçyiğit et al., 1999a, b; Yılmaz et al., 2000; Sözbilir, 2001,2002; Bozkurt, 2003; Westeway et al.,

Major Problems of Western Anatolian Geology  163

Figure 6.21  Map showing major structures of the Kozak horst, the Bergama graben and the surrounding regions. A and B represent the maps displayed in Figures 6.22 and 6.15, respectively. C refers to the map in the inset. The thick red arrow indicates direction of the westward escape of the Çandarlı tectonic wedge, bounded by the two oblique fault zones (black arrows). A Dp = Altınova depression; Ö‐E Dp = Ören‐Eğiller depression; ZG = Zeytindağ graben; MH = Maruflar horst; DB = Değirmendere graben.

2004; Purvis and Robertson, 2004, 2005; Ersoy et al., 2011] (Figs. 6.14 and 6.20). Consequently, parts of the north‐ south grabens are observed on the adjacent young horst blocks commonly as hanging grabens (e.g., the Gördes, Demirci, and Selendi grabens are located on the horst block, which was elevated between the Gediz and Simav grabens. The Mustafa Kemalpaşa, Evrenli, and Sarıcaova grabens are elevated on the Bozdağ dome between the Büyük Menders and Gediz grabens; Fig.  6.20). Other parts along the trend were buried under the east‐west graben basins as trapped grabens.

6.2.3.1. Discussion There is continuing controversy on a number of aspects of western Anatolian sedimentary associations and the grabens. Major issues of the discussion are on the continuity and ages of the three tectostratigraphic units, and age and mechanisms of the associated graben development. Seyitoğlu and Scott [1992, 1996] claimed that the three tectonostratigraphic units that are outlined above form one continuous succession in the Gediz graben area, where the lower unit and middle unit (red conglomerates) belong to the same tectonic entity, and

164  ACTIVE GLOBAL SEISMOLOGY

Figure 6.22  Geology map of the area located to the west of the Kozak magmatic center (A in Fig. 6.21). The block diagram displays the map area during the Early‐Middle Miocene time, based on the major structural elements and main trends (arrows) of the horsts and grabens of the region [modified from Temel et al., 2004, Figs. 5 and 6].

are Lower Miocene in age. However, Figure  6.19 displays the tripartite rock‐stratigraphic divisions of all the western Anatolian regions. The Early Miocene age conflict has continued despite the new age data obtained from the same rock units in the following years [see Yılmaz et al., 2000, for the discussion]. For example, the Upper Miocene age for the red conglomerate of the middle unit that surrounds the Bozdağ horst is repeatedly emphasized based on micromammals [Steinenger et al.,1985; Sarıca, 2000], gastropods [Emre, 1996,1998], and sporomorph associations [Akgün and Akyol, 1987, ̇ 1999; Ediger et al., 1996; Batı et al., 1998; I nci, 1998]. A

clear angular unconformity is observed between the lower unit and middle unit in many places and particularly around the eastern plunge of the Bozdağ dome (Fig. 6.24) [Yılmaz et al., 2000; Bozcu, 2010]. There is apparent discontinuity between the deposition of the upper unit (the fill of the present grabens) and the older tectonic units that crop out around the graben depressions, in time, space, and tectonic setting. Compared to the underlying Upper Miocene lacustrine sediments of low‐energy environment, the present graben fill represents a very high‐energy depositional environment. Around the Bozdağ dome, where the two conglom-

Major Problems of Western Anatolian Geology  165

Figure  6.23  Photo showing the red clastic rocks of the Upper Miocene age transiting laterally to the coeval, white‐colored, lacustrine fine‐grained sediments. Photo from the northwestern part of the Gediz graben. The accompanying schematic block diagram displays the formation of the red clastic rocks, formed around the elevated Bozdağ dome as syntectonically developed lateral fan deposits.

erate units are in direct stratigraphic contact with one another (Fig. 6.19), conglomerates of the upper unit may also be differentiated from the underlying red conglomerate of middle unit due to 1. The Upper Miocene conglomerates have coarser pebbles (5 × 15 cm). 2. They radiate away from the Bozdağ dome as lateral fans (Fig. 6.23), indicating that elevation of the Bozdağ horst results from their development. The conglomerates wrap round the horst and surround its eastern plunge (Fig. 6.24) [Yılmaz et al., 2001; Bozcu, 2010]. 3. Thickness of the conglomerate unit is more than 750 m, revealing that constant supply of materials escorts continuous elevation of the horst. 4. There is a pronounced regional angular unconformity between the syndetachment, back‐tilted, and locally

folded (rollover anticlines are common) Upper Miocene red‐colored fanglomerates unit with the overlying, commonly horizontal, Pliocene‐Quaternary conglomerate that was elevated on the graben shoulders. 5. The Upper Miocene clastic rocks around the Bozdağ dome show fining upward profile [Yılmaz et al., 2000]. The clast orientations (paleocurrent directions) display dispersions from the Bozdağ horst to the surrounding lowlands. In contrast, the fluvial sediments of the east‐west grabens show paleocurrent directions in close harmony with the east‐west river profile [Hakyemez et al., 2013]. 6. The middle unit and the upper unit were deposited under different climatic environments, and this is reflected by colors. The present graben fill is commonly gray. The Upper Miocene conglomerate on the other hand is red, indicating that it was deposited in strongly oxidizing environment.

166  ACTIVE GLOBAL SEISMOLOGY

Fig 6.24  Geology map of the area around the eastern plunge of the Menderes Massif [modified from Bozcu, 2010] showing approximately northwest trending older grabens (purple lines) and the crosscutting east‐west, west‐­ northwest–east‐southeast striking faults (black and red lines) and the associated younger major horst (Buldan horst) and the grabens (thick black lines indicate the present graben axes). Note the Upper Miocene sediments wrap around the eastern plunge of the Menderes Massif (broken green curve), and also rest on the old graben fill with a distinct angular unconformity. Note also that the older grabens are hanging grabens on the Buldan horst with respect to the present grabens.

For the discontinuity in the Miocene and the complete Neogene successions in the entire western Anatolian region, the following data may also be listed: (1) the sequence of the lower unit has some bituminous black shale layers with source rock potential. Therefore, the grabens have long been the regions of interest for the petroleum companies [Iztan and Yazman, 1990; Yazman et al., 1998; Temel et al., 2004]. For this reason, detailed seismic data have been collected from the large grabens (hundreds of kilometers long 2D and 3D seismic data  from the Büyük Menderes, Gediz, and Edremit grabens) to find out thickness and lithological contents of the sedimentary sequences, and particularly to detect the lower Miocene successions. The unpublished and published data reveal that generally the upper unit (the present graben fill) is confined to the east‐west grabens, and it is thinner than  60 mm a –11 [Jackson and McKenzie, 1988] to ~2‐3‐4 mm a –1 [Main and Burton, 1989; Sellers and Cross, 1989; Patton, 1992; Westaway, 1994a and b; LePichon et al., 1995; Barka and Reilenger, 1997; Straube et  al., 1997; Reilenger et  al., 1997, 2010, 2006; Kahle et al., 2000; Mc Clusky et al., 2000; Westeway et al., 2005]. Using these rates, the time required to calculate the present volume of the grabens is not more than 5 my. This period corresponds to the development of one of the possible causes of the extension. The synextensional granites, formed in association with the exhumation, yield widely distributed and scattered whole rock and mineral ages ranging from about 60 my to 5 my [Gessner et al., 2001a and b; Işık and Tekeli, 2001; Işık et al., 2004; Ring and Collins, 2005; Catlos and Çemen, 2005; Çemen et  al., 2006; Bozkurt et  al., 2011]. The young ages broadly corresponds to the initiation of the westward escape of the Anatolian Plate along NATF (see discussion under the heading of the North Anatolian Transform Fault in Section 2.6.1.); also see Westeway et  al. [2005] and the references therein). The young granite ages support also that the extensionally induced exhumation of the crust continued during the Pliocene [Gessner et  al., 2001a and b] and even the Pleistocene time [Lips et al., 2001].

Major Problems of Western Anatolian Geology  171

The structural indicators display that the exhumation of the Menderes Massif with both, top‐to‐the north– north‐northeast or south–south‐southwest‐directed kinematics began during the Late Eocene–Oligocene time [Hetzel et  al., 1995a and b; Gessner et  al., 2001a and b; Bozkurt and Park, 2004; Bozkurt and Oberhansli, 2001; Whitney and Bozkurt, 2002; Rimmele et al., 2003a, b; Jolivet et al., 2004; Bozkurt and Rojay, 2005]. The old granite ages, on the other hand, correspond to the initial phase of the extension connecting possibly with the collapse of the orogene during the Late Eocene– Oligocene period. The styles of crustal extension in western Anatolia are observed in two distinct modes: (1) core complex development and (2) continental rift valley development. The core complexes in the Aegean are bounded by the low‐angle normal faults. The views on the ­possible role of the two different structural styles on the extension are also divided into two major groups. According to one group, both of the extensional ­structures (e.g., low‐angle normal faults and the steeply dipping faults that bound the present young, narrow graben) are the products of the same extensional regime, which is continuing uninterruptedly since it began in the Late Oligocene. The transitions from the high to low angles of the faults were formed in the different stages of the progressive deformation. It is stated also that the abrupt changes from high‐angle to low‐angle and low‐ angle to high‐angle faults may be explained by the rolling hinge model of Buck [1988,1991; see Çiftçi, 2013]. However, the data documented for the dates of the fault development are not tightly constrained to rule out the possibility that the low‐angle and high‐angle sets of faults formed during the two subsequent but separate faulting phases. Furthermore, a time gap is clearly recorded also in the history of the various other geological events. The data gathered from a wide range of researches that are discussed in the foregoing paragraphs, including the times of the metamorphism of the Menderes Massif, the two stages of the magmatic associations, and the three tectonostratigraphic units differentiated in the Neogene sedimentary sequences, collectively support the view that there is a discontinuity in their developments and therefore favor the pulse extension model. 6.3. CONCLUDING SUMMARY The western Anatolian region is characterized by a number of approximately east‐west trending graben basins that are bounded by subparallel, normal fault zones. The faults and the associated seismic activity reveal that the region is undergoing north‐south extension. The geological data indicate that prior to the development of

the ongoing extension, the region witnessed a continental collision between the Sakarya continent and the Taurides, which took place between the Late Cretaceous and Eocene period. Along the collisional zone, the Izmir‐ Ankara suture was formed. The continuing convergence following the collision shortened, thickened, and elevated the crust resulting in the Aegean–western Anatolian highland during the Late Eocene–Oligocene period. Terrestrial sediments began to be deposited as the common cover above the tectonically amalgamated basement mosaic starting from the Early Miocene time, and continuing interruptedly to the present. It is assumed, therefore, that the inception of the sediment deposition corresponds to the first major extensional stage. In western Anatolia, there are three major rock groups. A majority of the metamorphic rocks collectively forms the Menderes Massif. The plutonic and volcanic rocks form the magmatic associations, and three different tectonostratigraphic successions collectively form the cover sediments. On the origin of these rocks and their connections with the extensional tectonic regime, various views have been proposed. Therefore, these subjects have long been controversial. The Menderes Massif is a north‐northeast–south‐ southwest trending culmination of western Anatolia occupying the largest volume of the region between the Izmir‐Ankara suture and the Taurides in the north and the south, respectively. The sequences compiled from the complementary tectonic slices of the Menderes Massif display that they are similar to the order and lithologies of the Mesozoic and the underlying units of the Taurides. This suggests that the Menderes metamorphic rocks represent the northern part of the Taurus passive continental margin prior to the metamorphism. The Menderes Massif is a core complex formed in association with the development of the Aegean–western Anatolian extensional terrain. It is composed of thick thrust piles consisting of different tectonic entities. Among thrust slices are HP and HT metamorphic units, which underwent metamorphisms under different tectonic setting and regimes, and together they formed the present Menderes Massif. The metamorphic evolution of the Menderes Massif occurred over a long period, which covers an age span from the Late Cretaceous to Miocene. During this period, initially some HP metamorphic rocks were formed in the north, along the northerly subducting slab of the Izmir‐Ankara Ocean. Later, they were tectonically emplaced on the Tauride passive margin succession and transported southward from the Late Cretaceous to the Middle‐Late (?) Eocene period. Presently, they represent the highermost tectonic components of the Menderes Massif, and represent ­

172  ACTIVE GLOBAL SEISMOLOGY

allochthonous (the ophiolitic and associated rocks) and the para‐autochthonous thrust slices (representing the leading age of the Taurus continental margin, involved in the subduction). They are observed along the periphery of the Menderes Massif. The main body of the Menderes Massif is exposed in the central region as a tectonic window with respect to the peripherally located allochthonous and paraautochtonous units (the nappes). The main body of the Menderes Massif also consists of the thrust slices. Across the thrusts, the metamorphic grade is reversed. The least metamorphosed rocks are observed at the bottom while the high‐grade metamorphic rock stands at the top. The structural arrangements postdate the main metamorphism, and may be the result of basement‐involved shortening deformation, associated with the convergence, which continued to the Miocene. The isotope ages associated with the main metamorphism of the Menderes Massif (MMM; the Amphibolite Facies HT Barrowian metamorphism) show a wide variation between Eocene and Miocene. This suggests that the massif was possibly heated up repeatedly. Exhumation of the Menderes Massif has occurred at least in two main stages. They correspond to the Late Oligocene–Early Miocene and the Late Miocene– Early Pliocene periods. Evidence for an earlier stage of extension dating back to the Late Eocene period is also documented. During the Late Oligocene– Early Miocene exhumation stage, the north‐northeast trending major elliptical part of the Menderes Massif was elevated. The lower Miocene terrestrial sediments were deposited on the metamorphic rocks and the surrounding ophiolitic rocks, and sealed their contacts. The age data indicate also that each one of the metamorphic events that caused the HP, HT, and retrograde metamorphisms occurred partly simultaneously in a wide time span, and also display a broadly southward younging trend. A wide range of overlapping ages may be the result of repeated pulses of shortening (nappe transport, thickening, and the consequent burial followed by collapse and exhumation). This may also mean that during this period, the Menderes Massif and the western Anatolian region as a whole behaved like an accordion. The repeated collisional events, followed by consequent westward escape pulses of the Anatolia plates, appear to be timely associated with the deformation‐extension phases that are recorded in the Menderes Massif. The escape may have triggered the retreat of the downgoing slab and, in turn, it caused the delamination of the thickened lithosphere and reheated the crust.

In western Anatolia, magmatic activities were developed in two distinctly different stages and styles as the early phase and the late phase. During the early phase, an intermediate and felsic magmatic association began to develop, initially in the Thrace and the northern Marmara regions, along the north‐south trending magmatic centers. This occurred during the Late Eocene–Middle Miocene period. During this phase, the granitic magmas reached shallow levels in the crust. The temporally, spatially, and genetically associated hypabyssal and volcanic rocks escorted the plutonic rocks, and collectively formed a collapsed caldera environment. The geochemistry reveals that the early phase magmas show the Andean‐type arc affinity; they were derived from an enriched mantle. The isotopic signatures and trace element characteristics indicate that both lithospheric and asthenospheric mantle melts have contributed to the source region, and the continental crustal materials contaminated the magmas later. The magmas also underwent AFC processes and/or MASH, and thus differentiated on the way to the surface. The volcanic front steadily migrated to the south and reached to the Bodrum Peninsula of southwestern Anatolia around 10 my ago. The island arc affinity of the early phase possibly reflects its genetic connection with the northerly subducting eastern Mediterranean oceanic slab under the Aegean– western Anatolian Plate. The retreats of the subducting slab appear to be the cause of the southward migration of the magmatic front. The late phase began in the Late Miocene and produced basic volcanic rocks, and has continued to the present. This volcanic phase is temporally and spatially connected with the development of the present horsts and grabens in a later stage. The volcanic rocks of basic composition, which were virtually missing during the first magmatic phase, form a distinctly different compositional cluster compared to the early phase. Geochemically, this group is commonly alkaline and displays similar affinities to the magmas that formed in continental regions experiencing extension. Commonly there is a time gap between the developments of the two magmatic phases. The gap is wide in the northern regions. In the Thrace region, for example, the time gap between the early and the late phase is about 20 my. The time gap narrows toward the south. In the region, a third volcanic rock group known as high‐Mg ultrapotassic volcanic rocks is also differentiated. However, this group is not distinctly different from the other two groups in time and space. In western Anatolia, the cover rocks are terrestrial deposits of the Neogene age. The successions are made up

Major Problems of Western Anatolian Geology  173

of three tectonostratigraphic units separated by regional unconformities. The lower unit is Early Miocene–Middle Miocene in age, and consists essentially of lacustrine and fluvial detrital sediments. Alternating sandstone, claystone, marl, and shale represent them. Coeval volcanic and volcanoclastic rocks generally intercalate with the sediments. The middle unit is Late Miocene–Early Pliocene in age, and is represented by the white lacustrine limestone‐marl dominated successions that are the most extensive rocks in the western Anatolian region. The only exception to this general picture is the area around the Bozdağ dome, where a coarse, red conglomerate formed in connection with the elevation of the dome as a synextensional deposit. Away from the dome, the lateral fanglomerates transit laterally to the lake deposits. The upper unit is made essentially of fluvial conglomerates of Quaternary age, deposited within the present graben depressions. They crop out in the surrounding highs, which have been elevated on the graben shoulders during the course of the ongoing extension. The locally developed basic lavas are also built around the grabens and accompany the sediments. The times of the regional unconformities that separate the three tectonostratigraphic units correspond commonly to the beginning of a new phase of extension. Initiation of the north‐south extension in the Aegean– western Anatolia region has been variously attributed to different mechanisms including (1) collapse of the orogene, (2) westward escape of the Anatolian Plate, (3) rollback of the subducted slab beneath the Hellenic arc, and (4) combination of these events. Each one of the events listed seems temporally and spatially connected to one another and triggered one of the extensional phases. The data collected from the different groups of rocks and the major geological events of western Anatolia including the graben formations suggest that they evolved in pulses. ACKNOWLEDGMENTS ̇ I am grateful to many colleagues at the Istanbul Technical University and abroad who have contributed significantly to the development of the ideas documented here. Among these Dr. A. M. C. Şengör is at the forefront. I exchanged ideas and discussed a number of problems with Dr. I.̇ Çemen. I thank him sincerely. During the preparation of the manuscript, my young colleagues contributed significantly. For this I thank deeply Mrs. Omer Kamacı Tanyel Baykut, Alp Ünal, and Miss Zeynep Çalışkanoğlu. Two western Anatolian experts, Prof. Cahit Helvacı and Prof. Şafak Altunkaynak,

and one anonymous reviewer reviewed the manuscript and made many valuable comments, which helped to improve the work. To them I am grateful. REFERENCES Agostini, S., C. Doglioni, F. Innocenti, P. Manetti, and S. Tonarini (2010), On the geodynamics of the Aegean rift, Tectonophysics, 488, 7–21. Agostini, S., J. G. Ryan, S. Tonarini, and F. Innocenti (2008), Drying and dying of a subducted slab: Coupled Li and B isotope variations in western Anatolia Cenozoic volcanism, Earth Planet. Sci. Lett., 272, 139–147. Akal, C. (2008), K‐richterite‐olivine‐phlogopite‐diopside‐ sanidine lamproites from the Afyon Volcanic Province, Turkey, Geological Mag., 145, 1–16. Akal C., O. Candan, O. E. Koralay, R. Oberhansli, F. Chen, and D. Prellevic (2012), Early Triassic potassic volcanism in the Afyon zone of the Anatolides/Turkey: Implications for the rifting of the Neo‐Tethys, Int. J. Earth Sci., 101, 177–194. Akgün, F., and E. Akyol (1987), Akhisar (Çıtak) çevresi kömürlerinin palinolojik incelenmes, Türkiye Jeoloji Kurumu Bülteni, 30, 35–50. Akgün, F., and E. Akyol (1999), Palinostratigraphy of the coal bearing Neogene deposits in Büyük Menderes graben, Western Anatolia, Geobios, 32, 367–383; doi: 10.1016/ S0016‐6995(99)80013‐8. Akkök, R. (1983), Structural and metamorphic evolution of the northern part of the Menderes Massif: New date from Derbent area and their implication for the tectonics of the Massif, J. Geol., 91, 342–350. Aksu, A. E., T. Konuk, A. Uluğ, M. Duman, and D. J. W. Piper (1990), Quaternary tectonic and sedimentary history of Eastern Aegean Sea shelf area (Doğu Ege Denizi şelf alanının Kuvaterner’deki tektoniği ve tortul tarihçesi), Jeofizik, 4, 3–35. Akyürek, B., and Y. Soysal (1983), Biga Yarımadası Güneyinin (Savaştepe‐KırkağaÇ‐ Ayvalık) Temel Jeolojik Özellikleri, Mineral Res. Exploration Bull. Turkey, 95/96, 1–13. Alçiçek, H. (2010), Stratigraphic correlation of the Neogene basins in Southwestern Anatolia: Regional palaeogeographical, palaeoclimatic and tectonic implications, Palaeogeogr. Palaeoclimatol. Palaeoecol., 291, 297–318. Aldanmaz, E., J. A. Pearce, M. F. Thirlwall, and J. G. Mitchell (2000), Petrogenetic evolution of the late Cenozoic, post‐ collision volcanism in western Anatolia, Turkey, J. Volcanol. Geotherm. Res., 102, 67–95. Aldanmaz, E. (2002), Mantle source characteristics of alkali basalts and basanites in an extensional intracontinental plate setting, western Anatolia, Turkey: Implications for multi‐ stage melting, Int. Geol. Rev., 44(5), 440–457. Aldanmaz, E., N. Köprübaşı, O. F. Gurer, N. Kaymakçı, and A. Gourgaud (2006), Geochemical constraints on the Cenozoic, OIB‐type alkaline volcanic rocks of NW Turkey: Implications for mantle sources and melting processes, Lithos, 86, 50–76.

174  ACTIVE GLOBAL SEISMOLOGY Altherr, R., F. Henjes‐Kunst, A. Mathews, H. Friedrichsen, B. T. Hansen (1988), O‐Sr isotopic variations in Miocene granitoids from the Aegean: evidence for an origin by combined assimilation and fractional crystallization, Contrib. Mineral. Petrol., 100, 528–541. Altunkaynak, S., and Y. Dilek (2006), Timing and nature of  postcollisional volcanism in western Anatolia and ­geodynamic implications, in Postcollisional Tectonics and Magmatism in the Mediterranean Region and Asia, edited by Y. Dilek and S. Pavlides, 321–353, Geological Society of America Special Papers, 409. Altunkaynak, S., and Y. Dilek (2013), Eocene mafic volcanism in northern Anatolia: Its causes and mantle sources in the absence of active subduction, Int. Geol. Rev., 13(55), 1641–1659. Altunkaynak, Ş., and Yılmaz, Y. (1998), The mount Kozak magmatic complex, western Anatolia, J. Volcanol. Geotherm. Res., 85(1), 211–231. Altunkaynak, Ş., and Y. Yilmaz (1999), The Kozak Pluton and its emplacement, Geol. J., 34(3), 257–274. Altunkaynak, Ş., Y. Dilek, C. Ş. Genç, G. Sunal, R. Gertisser, H. Furnes, … and J. Yang (2012a), Spatial, temporal and geochemical evolution of Oligo–Miocene granitoid magmatism in western Anatolia, Turkey, Gondwana Res., 21(4), 961–986. Altunkaynak, S., and Y. Yılmaz (2000), The Tugutreis Strato volcano of the Bodrum, SW Anatolia, IESCA Proced., edited by Ö. Dora, 33–38. Altunkaynak, S., G. Sunal, E. Aldanmaz, S. C. Genç, Y. Dilek, H. Furnes, K. A. Foland, J. Yang, and M. Yıldız (2012b), Eocene granitic magmatism in NW Anatolia (Turkey) revisited: New implications from comparative zircon SHRIMP U‐Pb and 40Ar‐39Ar geochronology and isotope geochemistry on magma genesis and emplacement, Lithos, 155, 289–309. Altunkaynak, S., N. W. Rogers, and S. P. Kelley (2010), Causes and effects of geochemical variations in late Cenozoic volcanism of the Foça volcanic centre, NW Anatolia, Turkey, Int. Geol. Rev., 52, 579–607. Altunkaynak, S., Y. Dilek, S. C. Genç, G. Sunal, R. Gertisser, H. Furnes, K. A. Foland, and J. Yang (2012), Spatial, temporal and geochemical evolution of OligoMiocene granitoid magmatisin western Anatolia, Turkey, Gondwana Research, 21(2012A), 961–986. Angelier, J., J. F. Dumont, H. Karamanderesi, A. Poisson, Y. Şimşek, and Y. Yusal (1981), Analyses of fault mechanisms and expansion of southwestern Anatolia since the late Miocene, Tectonophysics, 75, T1–T9. Angelier, J., N. Lybéris, X. Le Pichon, E. Barrier, and P. Huchon (1982), The tectonic development of the Helennic Arc and the Sea of Crete: A synthesis, Tectonophysics, 86, 159–196. Armijo, R., B. Meyer, A. Hubert, and A. Barka (1999), Westward propagation of the North Anatolian fault into the northern Aegean: Timing and kinematics, Geology, 27(3), 267–270; doi: 10.1130/0091‐7613(1999)0272.3.CO;2.

Avigad, D., and Z. Garfunkel (1991), Uplift and exhumation of highpressure metamorphic terrains: The example of the Cycladic Blueschist belt (Aegean Sea), Tectonophysics, 188, 357–372; doi: 10.1016/0040‐1951(91)90464‐4. Bailey, J. C., E. S. Jensen, A. Hansen, A. D. J. Kann, and K. Kann (2009), Formation of heterogeneous magmatic series beneath North Santorini, South Aegean island arc, Lithos, 110, 20–36. Barka, A., and R. Reilinger (1997), Active tectonics of the Eastern Mediterranean region: deduced from GPS, neotectonic and seismicity data, Ann. Geofis. X2(3), 587–610, 67–270. Barka, A. A. (1992), The North Anatolian fault zone, Ann. Tectonicae, 6, 164–195. Barka, A. A., and C. Kadinsky‐Cade (1988), Strike‐slip fault geometry in Turkey and its influence on earthquake activity, Tectonics, 7, 663–684, 1988. Batı, Z., M. Yazman, and I.̇ Özdemir (1998), Alkaşehir grabeni ve Alaşehir‐1 prospektinin değerlendirme raporu, TPAO rapor no 3864, Anatolian Fault, Geophys. J. Int., 194, 1335–1357. Beccaletto, L., and C. Steiner (2005), Evidence of two‐stage extensional tectonics from the northern edge of the Edremit Graben, NW Turkey, Geodinamica Acta, 18, 283–297. Benda, F. (1971), Principles of the subdivision of the Turkish Neogene, Newslett. Stratig. 1, 23–26. Benda, F., R. Innocenti, F. Mazzuoli, F. Radicati, and P. Steffens (1974), Stratigraphic and radiometric data of the Neogene in  northwest Turkey, Zeit. Deutschen Geo. Gesellschaft 125(1974), 183–193 Besang, C., F. J. Eckhardt, W. Harre, H. Kreuzer, and P. Muller (1977), Radiometrische altersbestimmungen an neogenen eruptivgesteinen der Turkei, Geologisches Jahrbuch Reihe B 25, 3–36. Bingöl, E., M. Delaloye, and G. Ataman (1982), Granitic intrusions in Western Anatolia: a contribution to the geodynamic study of this area, Eclogae Geoogica. Helv., 75, 437–446. Bircan, E. Bozbay, S. Gökdeniz, A. T. Kozan, and F. Öğ d em (1983), Gediz Graben Sisteminin Jeomorfolojisi ve Genç Tektoniğ i , MTA Report, 1983 (in Turkish, unpublished). Black K. N., E. J. Catlos, T. Oyman, and M. Demirbilek (2013), Aegean extension: Evidence from in situ U‐Pb geochronology and cathodoluminescence imaging of granitoids from NW Turkey, Lithos. Boray, A. (1982), Selimiye‐Beşparmak yöresindeki (Muğla) Menderes Masifi kayalarının stratigrafisi: tartışma ve Yanıt. Türkiye Jeoloji Kurumu Bülteni, 25, 161–162. Boray, A., U. Akat, N. Akdeniz, Z. Akçören, A. Çağlayan, E. Günay, B. Korkmazer, E.M. Öztürk, H. ve Sav (1975), Menderes Masifinin güney kenarı boyunca bazı önemli sorunlar ve bunların muhtemel çözümleri, Cumhuryeti’n 50, Yılı Yerbilimleri Kongresi, 11–20. Borsi, J., G. Ferrara, F. Innocenti, and R. Mazzuoli (1972), Geochronology and petrology of recent volcanics in the eastern Aegean Sea (West Anatolia and Lesvos Island), Bull. Volcanol., 36, 473–496.

Major Problems of Western Anatolian Geology  175 Bozcu, M. (2010), Geology of Neogene basin of Buldan‐ Sarıcaova region and their importance in western Anatolia neotectonics, Int. G. J. Earth Sci., 99, 851–865. Bozkurt Çiftçi, N., and E. Bozkurt (2008), Folding of the Gediz Graben fill, SW Turkey: Extensional and/or contractional origin? Geodinamica Acta, 21(3), 145–167. Bozkurt, E. (2000), Timing of extension on the Büyük Menderes graben, WesternTurkey, and its tectonic implications, Geological Society London, Special Publications, 173, 385–403. Bozkurt, E. (2001), A Late Alpine evolution of the central Menderes Massif, western Turkey, Int. J. Earth Sci., 89, 728–744; doi:10.1007/s005310000141. Bozkurt, E. (2001a), Origin of NE‐trending basins in western Turkey, Geodinam. Acta 16, 61–81. Bozkurt, E. (2001b), Neotectonics of Turkey‐A synthesis, Geodinamica Acta, 31, 3–30. Bozkurt, E. (2007), Extensional v. contractional origin for the southern Menderes shear zone, SW Turkey: Tectonic and metamorphic implications, Geol. Magazine, 144(01), 191–210. Bozkurt E., and B. Rojay (2005), Episodic, two‐stage Neogene extension and short‐term intervening compression in western Turkey: Field evidence from the Kiraz basin abd Bozdağ Horst, Geodinamica Acta, 18/3, 299–316. Bozkurt, E., and H. Sözbilir (2004), Tectonic evolution of the Gediz Graben: Field evidence for 890 an episodic, two‐stage extension in western Turkey, Geological Mag., 141, 63–79. Bozkurt, E., and M. Satır (2000), The southern Menderes Massif (western Turkey): Geochronology and exhumation history, Geol. J., 35, 285–296. Bozkurt, E., and R. G. Park (1994), Southern Menderes massif: An incipient metamorphic core complex in western Anatolia, Turkey, J. Geological Society London, 151, 213–216. Bozkurt, E., and R. Oberhänsli (2001), Menderes Massif (western Turkey): Structural, metamorphic and magmatic evolution: A synthesis, Int. J. Earth Sci., 89, 679–708; doi: 10.1007/ s005310000173. Bozkurt, E., and S. K. Mittwede (2005), Introduction: evolution of Neogene extensional tectonics of western Turkey, Geodinamica Acta, 18, 153–165. Bozkurt, E., M. Satır, and C. Bugdaycıoglu (2011), Surprisingly young Rb/Sr ages from the Simav extensional detachment fault zone, northern Menderes Massif, Turkey, J. Geod., 52, 406–431. Bozkurt, E. B. (2003), Origin of NE trending basins in western Turkey, Geodinamica Acta, 16, 61–81. Buck, W. R. (1988), Flexural rotation of normal faults, Tectonics, 7, 959–973. Buck, W. R. (1991), Modes of continental lithospheric extension, J. Geophys. Res., 96, 20161–20178. Buğdaycıoğlu, Ç. (2004), Tectono‐Metamorphic evolution of the Northern Menderes Massif: Evidence from the Horst Between Gördes and Demirci Basins (West Anatolia, Turkey), Ph.D. Thesis, Middle East Technical University. Bunbury, J. M. (1992), The Basalts of Kula, Western Turkey, Unpublished Ph.D. Thesis, University of Cambridge.

Burchfiel, B. C., R. Nakov, and T. Tzankov (2003), Evidence from the Mesta half‐graben, S.W. Bulgaria, for the Late  Eocene beginning of Aegean extension in the Central  Balkan Peninsula, Tectonophysics, 375, 61–76; doi: 10.1016. Burchfiel, B. C., R. Nakov, N. Dumurdzanov, D. J. Papanikolaou, T. Z. Tzankov, T. Serafimovski, R. W. King, V. Kotsev, A. Todosov, and Nurce, B. (2008), Evolution and dynamics of the Cenozoic tectonics of  the  South Balkan extensional system, Geosphere, 4, 919–938;doi: 10.1130. Çağlayan, A., E. M. Öztürk, S. Öztürk, H. Sav, and U. Akat (1980), Menderes masifi güneyine ait bulgular ve yapısal yorum (New data on the southern part of the Menderes massif and a structural interpretation), Jeoloji Mühendisligi, 10, 9–17. Candan, O. and O. Ö. Dora (1998), Menderes Masifinin genelleþtirilmiþ jeoloji haritası, DEU Jeoloji Mühendisliði Bölümü ̇ Bornova‐Izmir (yayımlanmamış). Candan, O., O. Dora, R. Oberhänsli, M. Çetinkaplan, J. Partzsch, F. Warkus, and S. Dürr (2001), Pan‐African high‐ pressure metamorphism in the Precambrian basement of the Menderes Massif, western Anatolia, Turkey, Int. J. Earth Sci., 89(4), 793–811. Candan, O., E. Koralay, et  al. (2011), Supra Pan African unconformity between core and cover series of the Menderes Massif/Turkey and its geological significance, Precambrian Res., 184,1–23. Candan, O., E. Koralay, Ö. Dora, F. Chen, R. Oberhansli, M. Çetinkaplan, C. Akal, M. Satır, and O. Kaya (2007), Menderes Masifinin Pan‐Afrikan temelin stratigrafisi ve örtü–çekirdek serilerinin ilksel dokanak iliþkisi, Menderes Masifi ̇ Kolokyumu, Izmir, 8–14. Candan, O., M. Çetinkaplan, R. Oberhänsli, G. Rimmelé, and C. Akal (2005), Alpine high‐P/low‐T metamorphism of the Afyon Zone and implications for the metamorphic evolution of Western Anatolia, Turkey, Lithos, 84(1), 102–124. Candan, O., O. Ö. Dora, R. Oberhansli, M. Çetinkaplan, J. H. Partzsch, F. C. Warkus, and S. Dürr (2001), Pan‐African high‐pressure metamorphism in the Precambrian basement of the Menderes Massif, western Anatolia, Turkey, Int. J. Earth Sci. 89, 793–811. Candan, O., O. E. Koralay, C. Akal O. Kaya, R. Oberhänsli, O. Dora, and F. Chen (2011), Supra‐Pan‐African unconformity between core and cover series of the Menderes Massif/Turkey and its geological implications, Precambrian Research, 184(1), 1–23. Catlos, E. J., and I. Çemen (2005), Monazite ages and the evolution of the Menderes Massif, western Turkey, Int. J. Earth Sci., 94, 204–217. Catlos, E. J., C. B. Baker, S. S. Sorensen, İ. Cemen, and M. Hancer (2008), Monazite geochronology, magmatism, and extensional dynamics within the Menderes Massif, western Turkey, Donald D Harrington Symposium on the Geology of the Aegean, Conf. Series: Earth and Environmental Science 2, 01.

176  ACTIVE GLOBAL SEISMOLOGY Catlos, E. J., C. B. Baker, S. S. Sorensen, L. Jacob, and ̇ I. Çemen (2011), Linking microcracks and mineral zoning of detachment‐exhumed granites to their tectonomagmatic history: Evidence from the Salihli and Turgutlu plutons in western Turkey (Menderes Massif), J. Structural Geology, 33(5),951–969. Çemen, L., E. J. Catlos, O. Gogus, and C. Ozerdem (2006), Post‐collisional extensional tectonics and exhumation of the Menderes Massif in the Western Anatolia Extended Terrane, Turkey, in Postcollisional Tectonics and Magmatism in the Mediterranean Region, edited by Y. Dilek and S. Pavlides, 353–379, Geological Society of America, Special Papers, 409. Çemen, I͘ ., C. Helvacı, and E. Y. Ersoy (2014), Cenozoic extensional tectonics in western and central Anatolia, Turkey: Introduction, Tectonophysics, 635, 1–5; doi: org/10.1016/j. tecto.2014.09.004. Chatzaras, V., P. Xypolias, S. Kokkalas, and I. Koukouvelas (2011), Oligocene‐Miocene thrusting in central Aegean: Insights from the Cycladic island of Amorgos, Geol. J., 46, 619–636. Çiftçi, N. B., and E. Bozkurt (2007), Anomalous stress field and active breaching in relay ramps: A field example from Gediz Graben, S.W. Turkey, Geological Mag. 144, 687–699. Çiftçi, N. B., and Bozkurt, E. (2008), Folding of the Gediz Graben fill, SW Turkey: extensional and/or contractional origin? Geodinamica Acta, 21(3), 145–167. Çiftçi, N. B., and E. Bozkurt, (2009a), Evolution of the Miocene sedimentary fill of the Gediz Graben, SW Turkey, Sedimentary Geology, 216, 49–79. Çiftçi, N. B., and E. Bozkurt (2009b), Pattern of normal faulting in the Gediz Graben, S.W. Turkey, Tectonophysics, 473, 234–260. Çiftçi, N. B., and E. Bozkurt (2010), Structural evolution of the Gediz Graben, SW Turkey: Temporal and spatial variation of the graben basin, Basin Res., 22, 846–873. Çoban, H., and M. F. J. Flower (2007), Late Pliocene lamproites from Bucak, Isparta (southwestern Turkey): Implications for mantle wedge evolution during Africa‐Anatolian plate convergence, J. Asian Earth Sci., 29, 160–176. Cohen, H. A., C. J. Dart, H. S. Akyüz, and A. Barka (1995), Synrift sedimantation and structural development of the Gediz and Büyük Menderes graben, Western Turkey, J. Earth Society, London, 152, 629–638. Collins, A. S., and A. H. F. Robertson (2003), Kinematic evidence for late Mesozoic–Miocene emplacement of the Lycian Allochthon over the Western Anatolide Belt, SW Turkey, Geol. J., 38, 295–310. Conticelli, S. (1998), The effect of crustal contamination on ultrapotassic magmas with lamproitic affinity: Mineralogical, geochemical and isotope data from the Torre Alfina lavas and xenoliths, Central Italy, Chem. Geol., 149, 51–81. Conticelli, S., L. Guarnieri, A. Farinelli, M. Mattei, R. Avanzinelli, G. Bianchini, E. Boari, S. Tommasini, M. Tiepolo, D. Preleviç, and G. Venturelli (2009), Trace elements and Sr‐Nd‐Pb isotopes of K‐rich, shoshonitic, and calc‐alkaline magmatism of the Western Mediterranean

Region: Genesis of ultrapotassic to calc‐alkaline magmatic associations in a post‐collisional geodynamic setting, Lithos, 107, 68–92. Dart, C., H. A. Cohen, H. S. Akyüz, and A. Barka (1995), Basinward migration of rift‐border faults: Implications for facies distributions and preservation potential, Geology, 23 (1), 69–72. De Graciansky, P. C. (1972), Recherches géologiques dans le Taurus lycien occidental, Thèse Doctorat d’Etat, Univ. Paris Sud, Orsay, France. Delaloye, M., and E. Bingöl (2000), Granitoids from Western and Northwestern Anatolia: Geochemistry and modeling of geodynamic evolution, Int. Geol. Rev., 42(3), 241–268. Demoulin, A., B. Altın, and A. Beckers (2013), Morphotectonic age estimate of the last phase of accelerated uplift in the Kazdağ area (Biga Peninsula NW Turkey), Tectonophysics, 608, 1380–1393. Denele, Y., E. Lecomte, L. Jolivet, O. Lacombe, L. Labrousse, B. Huet, L. LePourhiet (2011), Granite intrusion in a metamorphic core complex: The example of the Mykonos Iaccolith (Cyclades, Greece), Technophysics, 501 (1‐4), 52–70; doi: 10.1016/j.tecto.2011.01.013. Dewey, J. F. (1988), Extensional collapse of orogens, Tectonics, 7, 1123–1139. Dewey, J. F., and A. M. C. Şengör (1979), Aegean and surrounding region: complex multiplate and continuum tectonics in a convergent zone, Geol. Soc. Am. Bull., 90, 84–92. Dewey, J. F., M. L. Helman, E. Turco, D. H. W. Hutton, and S. D. Knott (1989), Kinematics of the western Mediterranean, 265–283, Geological Society, London, Special Publications, 45; doi:10.1144/GSL.SP.1989.045.01.15. Dilek, T., and E. Sandvol (2009), Seismic structure, crustal architecture and tectonic evolution of the Anatolian– African Plate boundary and the Cenozoic Orogeni belts in the Eastern Mediterranean region,in Ancient Orogens and Modern Analogues, edited by J. B. Murphy, J. D. Keppie, and A. J. Hynes, Geological Society, London, Special  Publications, 327, 127–160; doi: 10.1144/SP327.8 0305‐8719/09. Dilek, Y., and S. Altunkaynak (2009), Geochemical and temporal evolution of Cenozoic magmatism in western Turkey: Mantle response to collision, slab breakoff, and lithospheric tearing in an orogenic belt, Geological Society, London, Special Publications, 311, 213–233. Dixon J. E., and A. H. F. Robertson (1984), The geological evolution of the eastern Mediterranean, Geological Society, London, Special Publications, 17. Doglioni, C., S. Agostini, M. Crespi, F. Innocenti, P. Manetti, F. Riguzza, and Y. Savaşçın (2002), On the extension in western Anatolia and the Aegean Sea, J. Virtual Explorer, 8, 169–183. Dora, Ö. (2011), Menderes Masifindeki Jeolojik Araştırmaların Tarihsel Gelişimi, MTA 142, 1–23. Ediger, V. Ş., Z. T. Batı, and M. Yazman (1996), Paleo‐palinology of possible hydrocarbon source of the Alaşehir‐Turgutlu area in the Gediz graben (western Anatolia), Turkish Association of Petroleum Geologist Bulletin, 9, 11–23.

Major Problems of Western Anatolian Geology  177 Efe, R., A. Soykan, I.̇ Cürebal, and S. Sönmez (2011), Reviewing the geomorphological and neotectonic features of the Gönene Basin (NW of Turkey), Procedia‐Soc. and Behav. Sci., 19, 716–725. Elmas, A., E. Koralay, O. Duru, and A. Schmidt (2016), Geochronology, geochemistry, and tectonic setting of the Oligocene magmatic rocks (Marmaros Magmatic Assemblage) in Gökçeada Island, northwest Turkey. Int. Geol. Review, doi: 10.1080/00206814.2016.1227941. Emre, T. (1990), Şart Mustafa (Salihli)‐Adala‐Dereköy (Alaşehir) arasının jeolojisi ve Gediz Grabeni’nin yapısına ̇ bir yaklaşım: TÜBITAK, TBAG ‐ 732/YBAG  – 0001 project, 65 s. Emre, T. (1996), Gediz Graben’nin tektonik evrimi (Tectonical evolution of the Gediz graben) Türkiye Jeoloji Bülteni, Geol. Bull. Turkey, 39(2), 1–18. Emre T. (1998), Gediz Grabeni (Salihli‐Alaşehir arası) karasal tortullarının yaşıyla ilgili yeni bulgular, Türkiye Jeoloji Kurumu Bildiri Özetleri, Ankara, 34–35. Emre, T., H. Sözbilir, and N. Gökçen (2005), Neogene‐ Quaternary stratigraphiy of the Kiraz‐Beydağ vicinity, Küçük Menderes graben, western Anatolia, Mineral Res. Exp. Bull., 132, 1–32. Ercan, E., A. Dinçel, S. Metin, A. Türkecan, and A. Günay (1978), Uşak yöresindeki Neojen havzalarının jeolojisi (Geology of the Neogene basins in Uşak region), Bull. Geol. Soc. Turkey, 21, 97–106. Ercan, E., A. Türkecan, A. Dinçel, and E. Günay (1983). Kula‐ Selendi (Manisa) dolaylarının jeolojisi (Geology of Kula‐ Selendi [Manisa] area), Jeoloji Mühendisliği, 17, 3–28. Ercan, T. (1979), Batı Anadolu, Trakya ve Ege adalar|ndaki Senozoyik volkanizması, Jeoloji Mühhendisliği Dergisi, 9, 23–46. Ercan, T., M. Satır, H. Kreuzer, A. Türkecan, E. Günay, A. Çevikbaş, M. Ateş, and B. Can (1985), Batı Anadolu Senozoyik volkanitlerine ait yeni kimyasal, izotopik ve radyometrik verilerin yorumu, Türkiye Jeoloji Kurumu Bülteni, 28, 121–136. Ercan, T., M. Satir, D. Sevin, and A. Türkecan (1996), Bati Anadolu’daki Tersiyer ve kuvaterner yaşli volkanik kayaçarda yeni Yapilan radyometrik yaş ölçümlerinin yorumu, MTA Dergisi, 119, 103–112. Erdoğan, B., and T. Güngör (2004), The problem of the core‐cover boundary of the Menderes Massif and an emplacement mechanism for regionally extensive gneissic granites, Western Anatolia (Turkey), Turk. J. Earth Sci., 13, 15–36. Erinç, S. (1954), Orta Ege Bölgesi’nin Jeomorfolojisi, MTA Raporu No: 2217, Ankara (yayımlanmamış). Erinç, S. (1955), Gediz ve Küçük Menderes deltalarının morfolojisi, 9, Coğrafya Meslek Haftası (22–29 Aralık l954) Tebliğler ve Konferanslar, Türk Coğ. Kur. Yay. No. 2, 33–66, İstanbul. Erkal, T., and Hakyemez, H.Y. (1993), Gediz Nehri Deltasının Kuvaterner jeoloji ve jeomorfolojisi. Türkiye Kuvaterneri Workshop Bildiri Özleri, 17–19, Kasım, ̇ I stanbul, 32–33.

Erkül, F. (2010), Tectonic significance of synextensional ductile shear zones within the early Miocene Alacamdag granites, northwestern Turkey, Geol. Mag., 147, 611–637; doi: 10.1017/ S0016756809990719. Erkül, F. C. Helvacı, and H. Sözbilir, (2005a) Evidence for two episodes of volcanism in the Bigadiç borate basin and tectonic implications for western Turkey, Geol. J., 40. Erkül, F., C. Helvacı, and H. Sözbilir, (2005b) Stratigraphy and geochronology of the Early Miocene volcanics in the Bigadiç borate basin, western Turkey, Turkish J. Earth Sci., 14. Erkül, F., C. Helvaci, Y. Ersoy, H. Sözbilir, F. Erkül, Ö. Sümer, and B. Uzel (2009), Geochemistry and 40Ar/39Ar geochronology of Miocene varing lithospheric mantle source, western Anatolia, J. Volcanol. Geotherm. Res., 181–202. Erkül, F. C., and S. T. Erkül (2010), Geology of the early Miocene Alaçamdað (Dursunbey‐Balikesýr) magmatic complex and implications for the western Anatolian extensional tectonics, Mineral Res. Exploration. Bull., 141, 1–25. Erkül, S. T., H. Sözbilir, F. Erkül, C. Helvaci, Y. Ersoy, and Ö. Sümer (2008), Geochemistry of I‐type granitoids in the Karaburun Peninsula, West Turkey: Evidence for Triassic continental arc magmatism following closure of the Palaeotethys, Island Arc, 17(3), 394–418. Erkül, T. S., F. Erkül, E. Bozkurt, H. Sözbilir, and C. Helvacı (2009), Geodynamic setting of the early Miocene Alaçamdag volcano‐plutonic complex based on petrologic, isotopic and geochronological data: Northwestern Turkey, 62, Geological Kurultay of Turkey, Abstract, 180–181. Erol, O. (1982), Batı Anadolu genç tektoniğinin jeomorfolojik sonuçları. Türkiye Jeol. Kur. Batı Anadolu’nun Genç Tektoniği ve Volkanizması Paneli, Ankara, 15–21. Ersoy, E. Y., C. Helvacı and H., Sözbilir (2010), Tectono‐stratigraphic evolution of the NE–SW trending superimposed Selendi basin: implications for Late Cenozoic crustal extension in western Anatolia, Turkey, Tectonophysics, 488, 210–232. Ersoy, E. Y., C. Helvacı, E. Bozkurt, F. Erkül, and, H. Sözbilir (2005), Oligocene (?)–Lower Miocene olistostromal sedimentary sequence structurally above the Menderes metamorphic core complex, Selendi basin, western Anatolia and its tectonic implications, International Symposium on the ̇ Geodynamics of Eastern Mediterranean, 15–8, Istanbul, Abstracts, 75. Ersoy, E. Y., I.̇ Çemen, C. Helvacı, and Z. Billor (2014), Tectono‐stratigraphy of the Neogene basins in Western Turkey: Implications for tectonic evolution of the Aegean Extended Region, Tectonophysics; doi:10.1016/J.tecto. 2014.09.002 Ersoy, E.Y., and C. Helvacı (2007), Stratigraphy and geochemical features of the Early Miocene bimodal (ultrapotassic and calc‐alkaline) volcanic activity within the NE‐trending Selendi basin, western Anatolia, Turkey, Turkish J. of Earth Science, 16, 117–139. Ersoy, E.Y., C. Helvacı, and M. R. Palmer (2010), Mantle source characteristics and melting models for the early‐ middle Miocene mafic volcanism in Western Anatolia:

178  ACTIVE GLOBAL SEISMOLOGY Implications for enrichment processes of mantle lithosphere and origin of K‐rich volcanism in post‐collisional settings, J. Volcanol. Geothermal. Res., 198, 112–128. Ersoy, E.Y., C. Helvaci, H. Sözbilir, F. Erkül, and E. Bozkurt (2008), A geochemical approach to Neogene‐Quaternary volcanic activity of western Anatolia: An example of episodic bimodal volcanism within the Selendi basin, Turkey, Chem. Geol., 265–282. Ersoy, E.Y., C. Helvacı, H. Sözbilir, F. Erkül, and E. Bozkurt (2010), “Superimposed” Selendi Havzası’nın stratigrafik revizyonu, Batı Anadolu Revised Stratigraphy of the Superimposed Selendi Basin, Western Anatolia, Tectonophysics, 488, 210–232. Ersoy, E.Y., C. Helvaci, and M. R. Palmer (2011), Stratigraphic, structural Tectono‐stratigraphy of the Neogene basins in western Turkey: Implications for tectonic evolution of the Aegean extended region and geochemical features of the NE–SW trending Neogene volcanosedimentary basins in western Anatolia– Implications for associations of supra‐ detachment and transtensional strike‐slip basin formation in extensional tectonic setting, J. Asian Earth Sci., 41, 159–183. Eyidoğan, H., and J. A. Jackson (1985), A seismological study of normal faulting in the Demirci, Alaşehir and Gediz earthquake of 1960–1970 in western Turkey: implications for the nature and geometry of deformation in the continental crust, Geophys. J. R. Astron. Soc., 81, 569–607. Faccenna, C., C. Piromallo, A. Crespo‐Blanc, L. Jolivet, and F. Rossetti (2004), Lateral slab deformation and the origin of the western Mediterranean arcs, Tectonics, 23, TC1012. Faccenna, C., L. Jolivet, C. Piromallo, and A. Morelli (2003), Subduction and the depth of convection of the Mediterranean mantle, J. Geophys. Res., 108(B2), 2099; doi: 10.1029/ 2001JB001690. Faccenna, C., O. Bellier, J. Martinod, C. Piromallo, and V. Regard (2006), Slab detachment beneath eastern Anatolia: A possible cause for the formation of the North Anatolian fault, Earth Planet. Sci. Lett., 242, 85–97. Faccenna, C., O. Riguzzi, F. Bellier, J. Martinod, and Y. Savasçin (2002), On the extension in western Anatolia and the Aegean Sea, J. Virtual Explorer, 8, 169–183. Flower, M., and Y. Dilek (2003), Arc‐trench rollback and forearc accretion: 1. A collision‐induced mantle flow model for Tethyan ophiolites, in Ophiolites in Earth History, edited by Y. Dilek and P. T. Robinson, 218, Geological Society Special Publication. Francalanci, L., F. Innocenti, P. Manetti, and M. Y. Savaşçın (2000), Neogene alkaline volcanism of the Afyon‐Isparta area, Turkey: Petrogenesis and geodynamic implications, Mineralogy and Petrology, 70, 285–312. Frydas, D. (1993), Über die Nannoplankton‐Stratigraphie des Pliozäns des SE‐Peloponnes, Griechenland, Neues Jahrb. Geol. Palaeontol. Monatsh, 4, 227–238. Fytikas, M., F. Innocenti, P. Manetti, R. Mazzuoli, A. Peccerillo, and L. Villari (1984), Tertiary to Quaternary evolution of volcanism in the Aegean region, in The Geological Evolution of the Eastern Mediterranean, edited by J. E. Dixon and A. H.

F. Robertson, 687–699, Geological Society, London, Special Publications, 17. Gautier, P., and J.‐P. Brun (1994), Crustal‐scale geometry and kinematics of late‐orogenic extension in the central Aegean (Cyclades and Evvia Island), Tectonophysics, 238, 399–424; doi: 10.1016/0040‐1951(94) 90066‐3. Gautier, P., J.‐P. Brun, R. Moriceau, D. Sokoutis, J. Martinod, and L. Jolivet (1999), Timing, kinematics and cause of Aegean extension: A scenario based on a comparison with simple analogue experiments, Tectonophysics, 315, 31–37. Genç, C. Ş., Ş. Altunkaynak, Z. Karacık, M. Yazman, and Y. ̇ Yılmaz (2001), The Çubukludağ graben, South of Izmir: Its tectonic significance in the Neogene geological evolution of the western Anatolia, Geodinamica Acta, 14, 45–55. Genç, C. Ş., Z. Karacık, S. Altunkaynak, and Y. Yılmaz (2000), Geology of magmatic complex in the Bodrum peninsula SW Turkey, IESCA Proced., edited by Ö. Dora et al., 63–68. Genç, C. Ş., Z. Karacık, Ş. Altunkaynak, and Y. Yılmaz (2001), Geology of a magmatic complex in the Bodrum Peninsula, SW Anatolia, IESCA Proceed, edited by O. Dora et  al., 63–68. Genç, Ş. C. (1998), Evolution of the Bayramiç magmatic complex, J. Volcanol. Geotherm. Res., 85 (1–4), 233–249. Genç, Ş. C., and Y. Yılmaz (1997), An example of post‐collisional magmatism: The Kızderbent volcanics (Armutlu peninsula, Turkey), Turkish J. Earth Sci. (TUBITAK), 6(1), 33–42. Genç Ş. C., and Y. Yılmaz (2000), Aliağa dolayının jeolojisi ve genç tektoniği (Geology and young tectonics of the Aliağa region), Batı Anadolu Depremselliği Sempozyumu (BAD‐ ̇ SEM) Bildiriler Kitabı, Izmir, 152–159. Gentry, A. W., and E. P. J. Heizmann (1996), Miocene ruminants of the central and eastern Tethys and Paratethys, 1996, in The Evolution of Western Eurasian Neogene Mammal Faunas, edited by R. L. Bernor, V. Fahlbusch, and H.‐W Mittmann, 378–391, Columbia University Press, New York. Gessner, K., A. S. Collins, U. Ring, and T. Güngör (2004), Structural nd thermal history of poly‐orogenic basement: U‐ Pb geochronology of granitoids rocks of the Menderes Nappes Anatolide belts western Turkey, J. Geol. Soc. London, 161, 93–10 Gessner, K., U. Ring, and T. Güngör (2011), Field guide to Samos and the Menderes Massif: Along‐Strike variations in the Mediterranean Tethyan orogen, Field Guides, 23, 1–52. Gessner, K., U. Ring, C. Johnson, R. Hetzel, C. W. Passchier, and T. Güngör (2001a), An active bivergent rolling‐hinge detachment system: Central Menderes ­m etamorphic core complex in western Turkey, Geology, 29(7), 611–614. Gessner, K., S. Piazolo, T. Güngör, U. Ring, A. Kröner, and C. W. Passchier (2001b), Tectonic significance of deformation patterns in granitoid rocks of the Menderes nappes, Anatolide belt, southwest Turkey, Int. J. Earth Sci., 89(4), 766–780.

Major Problems of Western Anatolian Geology  179 Glodny, J., and R. Hetzel (2007), Precise U‐Pb ages of syn‐ extensional Miocene intrusions in the central Menderes Massif, western Turkey, Geol. Mag., 144, 235–246. ̇ Göktaş, F. (2012), Kemalpaşa‐Torbalı (Izmir) havzası ile yakın çevresindeki Neojen‐Kuvaterner tortullaşması ve magmatizmasının jeolojik etüdü. Maden Tetkik ve Arama Genel Müdürlüğü Rapor No. 11575 (unpublished). Görür, N., ed. (1988), Triassic to Miocene Plaeogeography Atlas of Turkey, MTA publication Ankara. Görür, N., N. Çağatay, M. Sakınç, K. Sümengen, C., Şentürk, C., Yaltırak, and A. Tchapalyga (1997), Origin of the Sea of Marmara as deduced from the Neogene to Quaternary paleo‐geographic evolution of its frame, Int. Geol. Rev., 39, 342–352. Güngör, T. (1998), Stratigraphy and tectonic evolution of the Menderes Massif in the Söke‐Selçuk region, na. Güngör, T., and B. Erdogan (2001), Emplacement age and direction of the Lycian nappes in the Söke‐Selcuk region, western Turkey, Int. J. Earth Sci., 89, 874–882; doi:10.1007/ s005310000100.2010.06.002. Gürer, Ö. F., and Y. Yılmaz (2002), Geology of the Ören and surrounding areas SW Anatolia, Turkish J. Earth Sci., 11, 1–13. Gürer, O. F., N. Sarıca‐Filoreau, M. Ozburan, E. Sangu, and B. Dogan (2009), Progressive development of the Buyuk Menderes Graben based on new data, western Turkey, Geological Mag. 146, 652–673. Gürer, F., E. Sangu, M. Özburan, and H. Sinir (2016), PlioQuaternary kinematic development and paleo stress pattern of the Edremit Basin, Western Turkey, Tectonophysics, doi: 10.1016/tecto.2016.05.007. Gürsoy, H., H. Temiz, and O. Tatar (1998), Gediz grabeni güney kenarında güncel deformasyon verileri. Aktif Tektonik ̇ Araştırma Grubu 1, Toplantısı, Makaleler, 103–112, ITÜ Maden Fakültesi, 8–9 Aralık 1997. Gürsoy, H., O. Tatar, J. D. Piper, A. Koc¸ F. Bulut, Z. Akpinar, B. Huang, A. P. Roberts, and B. L. Mesci (2011), Palaeo­ magnetic study of the Kepezda˘g and Yamadag volcanic provinces, central Turkey: Neogene tectonic escape and block definition in the east‐central Anatolides, J. Geod., 51, 308–326. Gutnic, M., O. Monod, A. Poisson, and J. F Dumont (1979), Geologie des Taurides Occidentales (Turquie), Memoires de la Societe Geologique de France, 137. Hafkenscheid, E., M. J. R Wortel, and W. Spakman (2006), Subduction history of the Tethyan region derived from seismic tomography and tectonic reconstructions, J. Geophys. Res., 111, B08401; doi: 10.1029/2005JB003791. Hakyemez, H. Y., T. Erkal, and F. Göktaş (1999), Late Quaternary evolution of the Gediz and Büyük Menderes grabens, Western Anatolia, Turkey, Quaternary Science Reviews, 18, 549–554. Hakyemez, Y., F. Göktaş, and T. Erkal (2013), Gediz Grabeninin Kuvaterner Jeolojisi ve evrimi, Türkiye Jeoloji Bülteni, 56. Hasözbek, A., E. Akay, B. Erdoğan, M. Satır, and W. Siebel (2010), Early Miocene granite formation by detachment tectonics or not? A case study from the northern Menderes Massif (Western Turkey), J. Geod., 50, 67–80.

Hayward, A. B. (1984a), Sedimentation and basin formation related to ophiolite nappe emplacement, Miocene, SW Turkey, Sed. Geol., 40, 105–129; doi: 10.1016/0037‐ 0738(84)90042‐3. Hayward, A. B. (1984b), Miocene clastic sedimentation related to the emplacement of the Lycian Nappes and the Antalya Complex, S.W. Turkey, in Geological Evolution of the Eastern Mediterranean, edited by J. E. Dixon and A. H. F. Robertson, Geol. Soc. Spec. Publ., 17, 287–300; doi: 10.1144/ GSL.SP.1984.017.01.21. Hayward, A. B., and A. H. F. Robertson (1982), Direction of ophiolite emplacement inferred from Cretaceous and Tertiary sediments of an adjacent autochthon, the Bey Daglari, southwest Turkey, Geol. Soc. Am. Bull., 93, 68–75. Helvacı, C., E. Y. Ersoy, H. Sözbilir, F. Erkül, Ö. Sümer, and B. Uzel (2009), Geochemistry and 40Ar/39Ar geochronology of Miocene volcanic rocks from the Karaburun peninsula: Implications for amphibole‐bearing lithospheric mantle source, western Anatolia, J. Volcanol. Geotherm.l Res., 185, 181–202; doi: 10.1016/j.jvolg. Hetzel, R., C. W. Passchier, U. Ring, and Ö. O. Dora (1995a), Bivergent extension in orogenic belts: the Menderes massif (southwestern Turkey), Geology, 23 (5), 455–458. Hetzel, R., U. Ring, C. Akal, and M. Troesch (1995b), Miocene NNE‐directed extensional unroofing in the Menderes Massif, southwestern Turkey, J. Geol. Soc., 152, 639–664; doi:10.1144/ gsjgs.152.4.0639. Huet, B., L. Le Pourhiet, L. Labrousse, E. Burov, and L. Jolivet (2010), Post‐orogenic extension and metamorphic core complexes in a heterogeneous crust: The role of crustal layering inherited from collision, Application to the Cyclades (Aegean domain), Geophys. J. Int., 184, 611–625; doi: 10.1111/j.1365‐246X.2010.04849.x Ilkışık, O. M. (1995), Regional heat‐flow in western Anatolia using silica temperature estimates from thermal springs, Tectonophysic, 244, 175–184. Inci, U. (1998), Lignite and carbonate deposition in middle lignite succession of the Soma formation, Soma coalfield, western Turkey, Inter. J. Coal Geol., 37, 287–313. Innocenti, F., S. Agostini, G. DiVincenzo, C. Doglioni, P. Manetti, M. Y. Savascin, and S. Tonarini (2005), Neogene and Quaternary volcanism in Western Anatolia: Magma sources and geodynamic evolution, Marine Geol., 221, 397–421. Işik, V., and O. Tekeli (2001), Late orogenic crustal extension in the northern Menderes massif (western Turkey): Evidence for metamorphiccore complex formation, Int. J. Earth Sci., 89, 757–765; doi: 10.1007/s005310000105. Işık, V, O. Tekeli, and G. Seyitoğ lu (2004), The 40Ar/39Ar age of extensionalductile deformation and granitoid intrusion in the northern Menderes core complex: implications for the initiation of extensional tectonics in western Turkey, J. Asian Earth Sci., 23, 555–566. İztan, H. and M. Yazman (1990), Geology and hydrocarbon potential of the Alaşehir (Manisa) area, western Turkey, In Proceedings of International Earth Sciences Congress, Aegean Region, pp. 327–333.

180  ACTIVE GLOBAL SEISMOLOGY Jackson, J., and D. McKenzie (1988), The relationship between plate motions and seismic moment tensors, and the rates of active deformation in the Mediterranean and Middle East, Geophys. J. Int., 93(1), 45–73. Jolivet, L., and J. P. Brun, (2010), Cenozoic geodynamic evolution of the Aegean, Int. J. Earth Sci., 99, 109–138. Jolivet, L., B. Goffé, P. Monié, C. Truffert‐Luxey, M. Patriat, and M. Bonneau (1996), Miocene detachment on Crete and exhumation P‐T‐t paths of high‐pressure metamorphic rocks, Tectonics, 15(6), 1129–1153; doi: 10.1029/96TC01417. Jolivet, L., C. Faccenna, B. Goffe, E. Burov, and P. Agard (2003), Subduction tectonics and exhumation of high‐pressure metamorphic rocks in the Mediterranean orogens, Am. J. Sci., 303, 353–409. Jolivet, L., C. Faccenna, B. Huet, L. Labrousse, L. Le Pourhiet, O. Lacombe, E. Lecomte, E. Burov, Y. Denèle, J.‐P. Brun, M. Philippon, A. Paul, G. Salaün, H. Karabulut, C. Piromallo, P. Monié, F. Gueydan, A. Okay, R. Oberhänsli, A. Pourteau, R. Augier, L. Gadenne, and O. Driussi (2013), Aegean tectonics: strain localization, slab tearing and trench retreat, Tectonophysics, 597–598, 1–33. Jolivet, L., E. Lecomte, B. Huet, Y. Denèle, O. Lacombe, L. Labrousse, L. Le Pourhiet, and C. Mehl (2010), The North cycladic detachment system, Earth Planet. Sci. Lett., 289, 87–104; doi:10.1016/j.epsl.2009. 10.032. Jolivet, L., G. Rimmelé, R. Oberhänsli, B. Goffé, and O. Candan (2004), Correlation of syn‐orogenic tectonic and metamorphic events in the Cyclades, the Lycian Nappes and the Menderes Massif. Geodynamic implications, Bull. Soc. Geol. Fr., 175, 217–238; doi: 10.2113/175.3.217. Kahle, H. G., M. Cocard, Y. Peter, A. Geiger, R. Reilinger, A. A. Barka, and G. Veis, (2000), GPS‐derived strain rate field within the boundary zones of the Eurasian, African, and Arabian Plates, J. Geophys. Res., 105, 23353–23370. Karabulut, H., S. Özalaybey, T. Taymaz, M. Aktar, O. Selvi, and A. Kocaoğlu (2003), A tomographic image of the shallow crustal structure in the Eastern Marmara, Geophys. Res. Lett., 30(24). Karabulut, H., A. Paul, T. E. Ergün, D. Hatzfeld, M. Childs, and M. Aktar (2013), Long‐wavelegth undulations of the seismic Moho beneath the strongly stretched Western Anatolia, Jeophys. J. Int.; doi: 10.1093/gji/ggt100. Karacık, Z., and Y. Yılmaz (1998), The ignimbrite eruptions and the associated magmatic rocks of the Ezine area, northern Anatolia, J. Volcanol. Geotherm. Res., 85, 252–264. Karacık Z., and Y. Yılmaz (2001), Volcanism of the Dikili‐ Çandarlı high and the surroundings, western Anatolia, IESCA proceed, edited by O. O. Dora et al., 33–38. Karacık, Z., Y. Yılmaz, and J. A. Pearce (2007), The Dikili‐ Çandarlı Volcanics, Western Turkey: Magmatic Interactions as Recorded by Petrographic and Geochemical Features, Turkish J. Earth Sci., 16, 493–522. Karacık, Z., Y. Yılmaz, J. Pearce, and I. Ece (2008), Petrochemistry of the South Marmara granitoids, northwestern Anatolia, Int. J. Earth Sci. (Geol Rundscahu), 97, 1181–1200.

Karaoğlu, Ö., C. Helvacı, and Y. Ersoy (2010), Petrogenesis and 40Ar/39Ar geochronology of the volcanic rocks of the Uşak‐ Güre basin, western Türkiye, Lithos, 119, 193–210; doi: 10.1016/j.lithos.2010.07.001. Karaoğlu, Ö., J. Browning, M. Bazargan, and A. Gundmundsen (2016), History of western Anatolia during the exhumation of the Menderes core complex, Earth and Planet. Sci. Lett., 452, 157–170, doi: org/10.1016/j.epsl.2016.07.037. Kastens, K. A. (1991), Rate of outward growth of Mediterranean ridge acreationary complex, Tectonophysics, 199, 25–560; doi: 10.1016/0040‐1951(91) 90117. Kaya, O., E. Ünay, G. Saraç, S. Eichhorn, S. Hassenrück, A. Knappe, A. Pekdeğer, and S. Mayda (2004), Halitpaşa transpressive zone: Implications for an Early Pliocene compressional phase in central western Anatolia, Turkey, Turkish J. Earth Sci., 13, 1–13. Kayseri‐Özer, M. S., H. Sözbilir, and F. Akgün (2014), Miocene palynoflora of the Kocaçay aand cumaovası basins: a contribution to the synthesis of Miocene palynology, paleoclimate, and paleovegatation in western Turkey, Turkish J. Earth Sci., 23, 1301–1309; doi: 10.3906/yer. Kazancı, N., A. Gürbüz, and S. Boyraz (2011), Büyük Menderes Nehri’nin Jeolojisi ve Evrimi, Türkiye Jeol. Bült., 54 (1–2), 25–55. Kazancı, N., S. Dundar, M. C. Alçiçek, and A. Gürbüz (2009), Quaternary deposits of the Büyük Menderes Graben in Western Anatolia, Turkey: Implications for river capture and the longest Holocene estuary in the Aegean Sea, Marine Geol., 264, 165–176. Kazancı, N., S. Leroy, Ö. Ileri, Ö. Emre, M. Kibar, and S. Öncel (2004), Late Holocene erosion in NW Anatolia from sediments of Lake Manyas, Lake Ulubat and the southern shelf of the Marmara Sea, Turkey, Catena, 2, 277–308. Kissel, C., and C. Laj (1988), Tertiary geodynamical evolution of the Aegean arc: A paleomagnetic reconstruction, Tectonophysics, 146, 183–201. Kober, L. (1921), Der Bau der Erde, Gebrüder Borntraeger. Koçyiğit, A. (2005), The Denizli graben‐horst system and the eastern limit of western Anatolian continental extension: basin fill, structure, deformational mode, throw amount and episodic evolutionary history, SW Turkey, Geodinamica Acta, 18/3–4,167–208. Koçyiğit A., H. Yusufoğlu, and E. Bozkurt (1999a), Evidence from the Gediz graben for episodic two‐stage extension in western Turkey, J. Geological Society, London, 156 (3), 605–616. Koçyiğit, A., H. Yusufoğlu, and E. Bozkurt (1999b), Reply to “Discussion on evidence from the Gediz graben for episodic two‐stage extension in western Turkey," J. Geological Society, London, 156 (1999), 1240–1242. Komut, T., R. Gray, R. Pysklywec, and O. H. Göğüş (2012), Mantle flow uplift of western Anatolia and the Aegean: Interpretations from geophysical analyses and geodynamic modeling, J. Geophys. Res.: Solid Earth, 117(B11). Konak, N. (2002), Geological Map of Turkey in 1/500,000 Scale: Izmir Sheet. Publication of Mineral Research and Exploration Directorate of Turkey, Ankara.

Major Problems of Western Anatolian Geology  181 Konak, N., and M. Şenel (2002), Geological Map of Turkey in 1/500,000 scale: Denizli Sheet, Publication of Mineral Research and Exploration Directorate of Turkey, Ankara. Lacassin, R., N. Arnaud, P. H. Leloup, R. Armijo, and B. Meijer (2007), Exhumation of metamorphic rocks in N Aegean: The path from shortening to extension and extrusion, eEarth, 2, 51–63. Lacassin, R., N. Arnaud, P. H. Leloup, R. Armijo, and B. Meyer (2007a), Exhumation of metamorphic rocks in N Aegean: The path from shortening to extension and extrusion, eEarth Discussions, 2, 1–35. Lacassin, R., N. Arnaud, P. H. Leloup, R. Armijo, and B. Meyer (2007b), Synorogenic and postorogenic exhumation of ­metamorphic rocks in N Aegean, eEarth, 2, 51–63. Lagos, M., E. E. Scherer, F. Tomaschek, C. Munker, M. Keiter, J. Berndt, and C. Ballhaus (2007), High precision Lu‐Hf geochronology of Eocene eclogite‐facies rocks from Syros, Cyclades, Greece, Chem. Geol., 243, 16–35; doi:10.1016/j. chemgeo.2007.04.008. Le Pichon, X., and J. Angelier (1979), The Hellenic arc and trench system: A key to the Neotectonic evolution of the eastern Mediterranean area, Tectonophysics, 60, 1–42. Le Pichon, X., J. Angelier, M. F. Osmaston, and L. Stegena (1981), The Aegean Sea, Phil. Trans. R. Soc. London Ser. A., 300, 357–372; doi: 10.1098/rsta.1981.0069. Le Pichon, X., N. Chamot‐Rooke, S. Lallemant, R. Noomen, and G. Veis (1995), Geodetic determination of the kinematics of central Greece with respect to Europe: Implications for eastern Mediterranean tectonics. J. Geophys. Res.: Solid Earth, 100(B7), 12675–12690. Lips, A. L. W., D. Cassard, H. Sözbilir, and H. Yilmaz, (2001), Multistage exhumation of the Menderes Massif, western Anatolia (Turkey), Int. J. Earth Sci., 89, 781–792; doi: 10.1007/s005310000101. Lips, A. L. W., S. H. White, and J. R. Wijbrans (2000), Middle‐ late Alpine thermotectonic evolution of the southern Rhodope Massif, Greece, Geodin. Acta, 13, 281–292; doi: 10.1016/S0985‐3111(00)00042‐5. Lustrino, M., S. Duggen, and C. L. Rosenberg (2011), The central‐western Mediterranean: Anomalous igneous activity in an anomalous collisional tectonic setting, Earth‐Sci. Rev., 104, 1–40. Main, I. G., and P. W. Burton (1989), Seismotectonics and the earthquake frequency‐magnitude distribution in the Aegean area, Geophys. J. R. Astron. Soc., 98, 575–586. McClusky, S., S. Balassanian, A. Barka, C. Demir, S. Ergintav, I. Georgiev, O. Gürkan, M. Hamburger, K. Hurst, K. Kahle, K. Kastens, G. Kekelidze, R. King, V. Kotzev, O. Lenk, S. Mahmoud, M. Mishin, M. Nadariya, A. Ouzounis, D. Paradissis, Y. Peter, M. Prilepin, R. Reilinger, I. Sanli, H. Seeger, A. Tealeb, M. N. Toksöz, and G. Veis (2000), Global positioning system constrains on plate kinematics and dynamics in the eastern Mediterranean and Caucasus, J. Geophys. Res., 105(B3), 5695–5719. McKenzie, D., and Y. Yılmaz (1991), Deformation and volcanism in western Turkey and the Aegean, Bull. Tech. Univ. Istanbul., Spec. Issue on Tectonics, 44, 345–373.

McKenzie, D. P. (1970), The plate tectonics of the Mediterranean region, Nature, 226, 239–243. McKenzie, D. P. (1972), Active tectonic of the Mediterranean region, Geophys. J. R. Astrol. Soc., 30, 109–185; doi: 10.1111/j.1365‐246X.1972.tb02351.x. McKenzie, D. P. (1978), Some remarks on the development of the sedimentary basins, Earth Planet. Sci. Lett. 40, 25–32. McKenzie, D. (1979), Finite deformation during fluid flow, Geophys. J. Int., 58(3), 689–715. Meijer, P. T., and M. J. R. Wortel (1997), Present‐day dynamics of the Aegean region: A model analysis of the horizontal pattern of stress and deformation, Tectonics 16, 879–895. Oberhänsli, R., O. Candan, and F. D. H. Wilke (2010), Geochronological evidence of Pan‐African eclogites from the Central Menderes Massif, Turk. J. Earth Sci., 19, 431–447 Okay, A. I. (1981), Lawsonite zone blueschists and a sodic amphibole producing reaction in the Tavsanli region, ­northwest  Turkey, Contrib. Mineral. Petrol., 75, 179–186; doi: 10.1007/BF01166758. Okay, A. I. (1984), Distribution and characteristics of the northwest Turkish blueschists, in The Geological Evolution of the Eastern Mediterranean, edited by J. E. Dixon and A. H. F. Robertson, 455–466, Geological Society Special Publication 17; doi: 10.1144/GSL.SP.1984.017.01.33. Okay, A. I. (2007), Menderes Masifi: Nap paketimi yoksa stratigrafik bir istifmi? (Menderes Massif: A nappe pile or a stratigraphic sequence), Menderes Masifi Kollokyumu, Genişletilmiş Bildiri özleri kitabı, MTA yayını. Okay, A. I. (1989), Geology of the Menderes Massif and the Lycian nappes south of Denizli, western Taurides, Mineral Res. Expl. Bull., 109, 37–51. Okay, A. I. (2001), Stratigraphic and metamorphic inversions in the central Menderes massif: A new structural model, Int. J. of Earth Sci. (Geol Rundschau), 89, 709–727. Okay, A. I. (2004), Tectonics and high‐pressure metamorphism in northwest Turkey, Post‐Congress Excursion P01, 32nd International Geological Congress, Ital. Agency for Environ. Protect and for Tech. Serv., Florence, Italy. Okay, A. I. (2008), Geology of Turkey: A synopsis, Anschnitt, 21, 19–42. Önen, A. P., and R. Hall (2000), Subophiolite metamorphic rocks from NW Anatolia, Turkey. J. Metamorph. Geol., 18, 483–495. Öner, Z, and Y. Dilek (2013), Supradetachment basin evolution during continental extension: The Aegean province of  western Anatolia, Turkey, Geol. Soc. Am. Bull., 123, 2115–2141. Özacar, A. A., C. B. Biryol, S. Beck, G. Zandt, and N. Kaymakci (2010), Crust and upper mantle dynamics of Turkey inferred from passive seismology: implications of segmented slab geometry, in Proceedings of the Conference Tectonic Crossroads. Evolving Orogens of Eurasia‐Africa‐Arabia, edited by Y. Dilek and E. Bozkurt, Ankara, Turkey. Ozaner, S., and E. Bozbay (1982), Kula dolaylarının morfojenezi, genç tektoniği ve Gediz‐Alaşehir grabeni ile ilişkisi,

182  ACTIVE GLOBAL SEISMOLOGY 36, Türkiye Özer, S. (1998), Rudist bearing Upper Cretaceous metamorphic sequences of the Menderes Massif (western Turkey), Geobios, 31, 235–249, Özer, S., and H. Sözbilir (2003), Presence and tectonic significance of Cretaceous rudist species in the so‐called Permo‐Carboniferous Gktepe formation, central Menderes metamorphic massif, western Turkey, Int. J. Earth Sci., 92, 397–404; doi: 10.1007/s00531‐003‐0333‐z. ̇ Özgenç, I.,̇ and N. Ilbeyli (2008), Petrogenesis of the Late Cenozoic Eğrigöz Pluton in Western Anatolia, Turkey: implications for magma genesis and crustal processes, Int. Geol. Rev., 50, 375–391. Özkaymak, Ç., H. Sözbilir, and B. Uzel (2013), Neogene‐ Quaternary evolution of the Manisa Basin: Evidence for variation in the stress pattern of the Izmir‐Balıkesir Transfer Zone, western Anatolia, J. Geod., 65 (2013), 117–135. Öztürk, A., and A. Koçyiğit (1983), Menderes grubu kayalarının temel‐örtilişkisine yapısal bir yaklaşım (Selimiye‐Muğla), Bull. Geol. Soc. Turkey, 26, 99–106. Öztürk, S., M. Destur, and M. Karlı (2006), Heat flow map of Turkey, 1:2,000,000, General Directorate of Mineral Research and Exploration, Department of Geophysical Exploration, Ankara, Turkey. Papanikolaou, D. J., and E. Vassilakis (2010), Thrust faults and extensional detachment faults in Cretan tectono‐stratigraphy: Implications for middle Miocene extension, Tectonophysics, 488, .233–247; doi;10.1016/j.tecto.2009.06.024. Paton, S. (1992), Active normal faulting drainage patterns and sedimentation in southwestern Turkey, J. Geological Society London, 149, 1031–1044. Paton, S. M. (1992a), The Relationship Between Extension and Volcanism in Western Turkey, the Aegean Sea, and Central Greece, Ph.D. Thesis, Queens’ College Cambridge University. Pearce, J. A., and D. W. Peate (1995), Tectonic implications of the composition of volcanic arc lavas, Ann. Rev. Earth Planetary Sci., 23, 251–285. Pe‐Piper, G., Y. Zhang, D. J. W. Piper, and D. Prelević (2013), Relationship of Mediterranean type lamproites to large shoshonite volcanoes, Miocene of Lesbos, NE Aegean Sea, Lithos, 184, 281–299. Peter, Y., M. Prilepin, R. E. Reilinger, I. Sanlı, H. Seeger, A. Tealeb, M. N. Toksöz, G. Veis (2000), Global Positioning System constraints on plate kinematics and dynamics in the Eastern Mediterranean and Caucasus, J. Geophys. Res., 105, 5695–5720. Philippon, M., J. P. Brun, F. Gueydan, and D. Sokoutis (2014), The interaction between Aegean back‐arc extension and Anatolia escape since Middle Miocene, Tectonophisics, 631, 176–188. Phillipson, A. (1918), Kleinasien, Handbuch der Regionalen Geologie, edited by G. Steinmann and O. Wilckens, 5, Heidelberg. Plunder, A., P. Agard, C. Chopin, and A. I. Okay (2013), Geodynamics of the Tavşanlı zone, western Turkey: Insights into subduction/obduction processes, Tectonophysics, 608, 884–903.

Poisson, A., F. Yağmurlu, M. Bozcu, and M. Şentürk (2003), New insights on the tectonic setting and evolution around the apex of the Isparta Angle (SW Turkey), Geol. J., 38(3‐4), 257–282. Popov, S. V., I. G. Shcherba, L. B. Ilyina, L.A. Nevesskaya, N. P. Paramonova, S. O. Khondkarian, and I. Magyar (2006), Late Miocene to Pliocene palaeogeography of the Paratethys and its relation to the Mediterranean, Palaeogeogr. Palaeoclimatol. Palaeoecol., 238, 91–106; doi: 10.1016/j. palaeo.2006.03.020. Pourteau, A., O. Candan, and R. Oberhänsli (2010), High‐ pressure metasediments in central Turkey: Constraints on the Neotethyan closure history, Tectonics, 29; doi: 10.1029/ 2009TC002650. Prelević, D., S. F. Foley, R. Romer, and S. Conticelli (2008), Mediterranean Tertiary lamproites derived from multiple source components in postcollisional geodynamics, Geochim. et Cosmochim. Acta, 72(8), 2125–2156. Prelevic, D., A. Stracke, S. F. Foley, R. L. Romer, and S. Conticelli (2010b), Hf isotope compositions of Mediterranean lamproites: Mixing of melts from asthenosphere and crustally contaminated mantle lithosphere, Lithos, 119, 297–312. Prelević, D., C. Akal, R. L. Romer, and S. F. Foley (2010a), Lamproites as indicators of accretion and/or shallow subduction in the assembly of south‐western Anatolia, Turkey, Terra Nova, 22(6), 443–452. Prelevic, D., C. Akal, and S. F. Foley (2008), Orogenic vs anorogenic lamproites in a single volcanic province: Mediterranean‐ type lamproites from Turkey, IOP Conference Series: Earth and Environmental Science, 2, 012024. Prelevic, D., C. Akal, S. F. Foley, R. L. Romer, A. Stracke, and A. P. Bogaard (2012), Ultrapotassic‐mafic rocks as geochemical proxies for post‐collisional dynamics of orogenic lithospheric mantle: The case of southwestern Anatolia, Turkey, Petrology, 53, 1019–1055. Prelevic, D., C. Akal, R. L. Romer, R. Mertz-Kraus, and C. Helvacı (2015), Magmatic respons to slab tearing: constrains from the Afyon alkaline volcanic complex, western Turkey. Jour. Petrol., 1–36. doi: 10.1093/petrology/egv008. Purvis, M., and A. H. F. Robertson (2004), A pulsed extension model for the Neogene–Recent E‐W‐trending Alasehir graben and the NE‐SW trending Selendi and Gordes basins, western Turkey, Tectonophysics, 391, 171–201; doi: 10.1016/j. tecto.2004.07.011. Purvis, M., and A. Robertson (2005a), Miocene sedimentary evolution of the NE‐SW‐trending Selendi and Gördes basins, W Turkey: Implications for extensional processes, Sed. Geol., 174 (1–2), 31–62. Purvis, M., A. H. F. Robertson, and M. S. Pringle (2005b), 40Ar‐39Ar dating of biotite and sanidine in tuffaceous sediments and related intrusive rocks: Implications for the early Miocene evolution of the Gördes and Selendi basins, W Turkey, Geodinamica Acta, 18, 239–253; doi: 10.3166/ ga.18.239‐253. Régnier, J. L., J. E. Mezger, and C. W. Passchier (2007), Metamorphism of Precambrian Palaeozoic schists of the Menderes core series and contact relationships with  Proterozoic orthogneisses of the western Çine

Major Problems of Western Anatolian Geology  183 Massif, Anatolide belt, western Turkey, Geol. Mag. 144, 67–104. Reilinger, R., S. C. McClusky, M. B. Oral, R. W. King, M. N. Toksöz, A. A. Barka, I. Kınık, O. Lenk, and I. Sanlı (1997), Global positioning system measurements of present‐day crustal movements in the Arabia‐Africa‐Eurasia plate collision zone, J. Geophys. Res., 102, 9983–9999. Reilinger, R., S. C. McClusky, P. Vernant, S. Lawrence, S. Ergintav, S. Cakmak, R. Ozener, H. Kadirov, F. Guliev, I. Stepanyan, R. Nadariya, M. Hahubia, G. Mahmoud, S. Sakr, K. ArRajehi, A. Paradissis, D. Al‐Aydrus, A. Prilepin, M. Guseva, T. Evren, E. Dmitrotsa, A. Filikov, S. V. Gomez, F. Al‐Ghazzi, and G. Karam (2006), GPS constraints on c­ ontinental deformation in the Africa‐Arabia‐ Eurasia continental collision zone and implications for the dynamics of plate interactions, J. Geophys. Res., 111 (B5), B05411. Reilinger, R., S. McClusky, D. Paradissis, S. Ergintav, and P. Vernant (2010), Geodetic constraints on the tectonic evolution of the Aegean region and strain accumulation along the Hellenic subduction zone, Tectonophysics, 488, 22–30; doi: 10.1016/j.tecto.2009.05.027. Richardson‐Bunbury, J. M. (1996), The Kula volcanic field, western Turkey: The development of a Holocene alkali basalt province and the adjacent normal faulting graben, Geo. Mag., 133(3), 275–283. Rimmelé, G., L. Jolivet, R. Oberhänsli, and B. Goffé (2003a), Deformation history of the high‐pressure Lycian Nappes and implications for the tectonic evolution of SW Turkey, Tectonics, 22(2), 1007; doi: 10.1029/2001TC901041. Rimmelé, G., R. Oberhänsli, B. Goffé, L. Jolivet, O. Candan, and M. Cetinkaplan (2003b), First evidence of high‐pressure metamorphism in the “Cover Series” of the southern Menderes massif–Tectonic and metamorphic implications for the evolution of SW Turkey, Lithos, 71, 19–46; doi: 10.1016/S0024‐4937(03)00089‐6. Rimmelé, G., T. Parra, B. Goffé, R. Oberhansli, L. Jolivet, and O. Candan (2005), Exhumation paths of high‐pressure–low‐ temperature metamorphic rocks from the Lycian Nappes and the Menderes Massif (SW Turkey): A multi‐equilibrium approach, J. Petrol., 46(3), 641–669; doi: 10.1093/petrology/ egh092. Ring, U., and A. S. Collins (2005), U‐Pb SIMS dating of synkinematic granites: Timing of core complex formation in the northern Anatolide belt of western Turkey, J. Geological Society London, 162, 289–298. Ring, U., and P. W. Layer (2003), High‐pressure metamorphism in the Aegean, eastern Mediterranean: Underplating and exhumation from the Late Cretaceous until the Miocene to recent above the retreating Hellenic subduction zone, Tectonics, 22(3), 1022; doi: 10.1029/2001TC001350. Ring, U., A. P. Willner, and W. Lackmann (2001a), Stacking of nappes with different pressure‐temperature paths: An example from the Menderes nappes of western Turkey, Amer. J. Sci., 301, 912–944. Ring, U., J. Glodny, T. Will, and S. N. Thomson (2007b), An Oligocene extrusion wedge of blueschist‐facies nappes on Evia, Aegean Sea, Greece: Implications for the early exhuma-

tion of high‐pressure rocks, J. Geol. Soc., 164, 637–652; doi: 10.1144/0016‐76492006‐041. Ring, U., K. Gessner, T. Güngör, and C. W. Passchier (1999), The Menderes Massif of western Turkey and the Cycladic Massif in the Aegean: Do they really correlate? J. Geological Society London, 156, 3–6. Ring, U., P. W. Layer, and T. Reischmann (2001b), Miocene high‐pressure metamorphism in the Cyclades and Crete, Aegean Sea, Greece: Evidence for large‐magnitude displacement on the Cretan Detachment, Geology, 29, 395–398; doi: 10.1130/0091‐7613(2001)0292.0. Ring, U., T. Will, J. Glouny, C. Kumerics, K. Gessner, S. N. Thomson, T. Güngör, P. Monie, M. Okrusch, and K. Drüppel (2007a), Early exhumation of high‐pressure rocks in extrusion wedges: the Cycladic blueshist unit in the eastern Aegean, Greece and Turkey. Tectonics, 26, 175. Roberts, S. C. (1988), Active Normal Faulting in Central Greece and Western Turkey, Unpublished Ph.D. Thesis, University of Cambridge. Robert, U., J. Foden, and R. Varne (1992), The Dodecanese Province, SE Aegean: A model for tectonic control on potassic magmatism, Lithos, 28(3‐6), 241–260. Rojay, B., V. Toprak, C. Demirci, and L. Suzen (2005), Plio‐quaternary evolution of the Küçük Menderes graben southwestern Anatolia, Turkey, Geodinamica Acta, 18, 317–331. Rosenbaum, G., M. Gasparon, F. P. Lucente, A. Peccerillo, and M. S. Miller (2008), Kinematics of slab tear faults during subduction segmentation and implications for Italian magmatism, Tectonics, 27, TC2008; doi: 10.1029/2007TC002143. Royden, L. H., and D. J. Papanikolaou (2011), Slab seg­ mentation and late Cenozoic disruption of the Hellenic arc,  Geochem. Geophys. Geosyst., 12, Q03010; doi: 10.1029/2010GC003280. Sakınç, M., C. Yaltırak, and F. Oktay (1999), Palaeogeographical evolution of the Thrace Neogene Basin and the Tethys‐ Paratethys relations at northwestern Turkey (Thrace), Palaeogeogr. Palaeoclimatol. Palaeoecol., 153(1–4), 17–40. Sarıca, N. (2000), The Plio‐Pleistocene ager of Büyükmenderes and Gediz grabens and their tectonic significance on N‐S extensional tectonics in western Anatolia: Mammalian evidence from continental deposits, Geological J., 35, 1–24. Şaroğlu, F., and Y. Yılmaz (1991), Geology of the Karlıova region: Intersection of the North Anatolian and the East Anatolian transform faults, Bull. Tech. Univ. Istanbul, Spec. Issue on Tectonics, 44(1), 475–493. Saunders, P., K. Priestley, and T. Taymaz (1998), Variations in the crustal structure beneath western Turkey, Geophys. J. Int., 134, 373–389. Savaşçın, M. Y., and T. Oyman (1998), Tectono‐magmatic evolution of alkaline volcanics at the Kırka‐Afyon‐Isparta structural trend, SW Turkey, Turkish J. Earth Sci., 7, 201–214. Seghedi, I., C. Helvacı, and Z. Peckskay (2015), Composite volcanoes in the south-eastern part of the I͘ zmir-Balıkesir Transfer Zone, western Anatolia/ Turkey. Jour. Volc. Geoth. Res., 291, 72–85; doi.org/10.1016/j. Sellers, P. C. and P. A. Cross (1986), Wegener‐Medlas baselines determined using the Pseudo‐Short Arc Technique, In Proceedings Int. Conf. WEGENERMEDLAS Project.

184  ACTIVE GLOBAL SEISMOLOGY Şenel, M. (1997), Geological Map of Turkey, 1: 100,000 scale, Denizli‐J9 Quadrangle, and accompanying 18 page explanatory booklet, General Directorate of Mineral Research and Exploration, Ankara. Şengör, A. M. C. (1979), Türkiyenin neotektonik esasları (Neotectonics of Turkey, 2), Ankara: Türkiye Jeoloji Kurumu Yayınları Serisi. Şengör, A. M. C. (1987), Cross‐faults and differential stretching of hanging walls in regions of low‐angle normal faulting: examples from western Turkey, in Continental Extensional Tectonics, edited by M. P. Coward, J. F. Dewey, and P. L. Hancock, Geological Society, London, Special Publication, 28, 575–589. Şengör, A. M. C. and W. S. F. Kidd (1979), Post‐collisional tectonics of the Turkish‐Iranian plateau and a comparison with Tibet, Tectonophysics, 55(3‐4), 361–376. Şengör, A. M. C., and Y. Yılmaz (1981), Tethyan evolution of Turkey: A plate tectonic approach, Tectonophysics, 75, 181–241. Şengör, A. M. C., M. Satir, and R. Akkök (1984), Timing of tectonic events in the Menderes massif, western Turkey: Implications for tectonic evaluation and evidence for Pan‐ African basement in Turkey, Tectonics, 3, 693–707. Şengör, A. M. C., N. Gorür, and F. Şaroglu (1985b), Strike‐slip faulting and related basin formation in zones of tectonic escape: Turkey as a case study, in Strike‐Slip Deformation, Basin Formation, and Sedimentation, edited by K. T. Biddle and N. Christie‐Blick, 227–264, Dordrecht: Kluwer Academic. Şentürk, K. and C. Karaköse (1998), Çanakkale‐D2 Paftası, 1: 100 000 ölçekli açınsama nitelikli Türkiye jeoloji haritaları, No: 62, Mineral Research and Exploration, Ankara. Şentürk, K., M. Sümengen, I.̇ Terlemez, and C. Karaköse (1998), 1:100 000 ölçekli Açınsama Nitelikli Türkiye Jeoloji Haritaları Bandırma‐D4 Paftası, Ankara (Explanatory map of Turkey on the scale of 1/100,000; Bandirma sheet) Maden Tetkik ve Arama Genel Müdürlüğü, no. 64. Seyitoğlu, G. (1997), Late Cenozoic tectono‐sedimentary development of the Selendi and Usak‐Güre basins: a contribution to the discussion on the development of E‐W and north trending basins in western Turkey, Geological Mag. 134, 163–175. Seyitoğlu, G., and B. C. Scott (1994), Late Cenozoic basin development in west Turkey: Gördes basin tectonics and sedimentation, Newsletter. Stratig., 131, 133–142. Seyitoğlu, G., and B. C. Scott (1996), The cause of N‐S extensional tectonics in western Turkey: Tectonic escape vs back‐arc spreading vs orogenic collapse, J. Geod., 22, 145–153. Seyitoğlu, G., and B. Scott (1991), Late Cenozoic crustal extension and basin formation in west Turkey, Geol. Mag., 128, 155–166. Seyitoğlu, G., and B. Scott (1992), The age of the Büyük Menderes Graben (western Turkey) and its tectonic implications, Geol. Mag. 129 (1992) 239–242. Seyitoğlu, G., D. Anderson, G. Nowell, and B. Scott (1997), The evolution from Miocene potassic to Quaternary sodic magmatism in western Turkey: Implications for enrichment processes in the lithospheric mantle, J. Volcanol. Geotherm. Res., 76, 127–147.

Seyitoğlu, G., I.̇ Çemen, and O. Tekeli (2000), Extensional folding in the Alaşehir (Gediz graben, western Turkey), J. Geological Society London, 157, 1097–1100. Seyitoğlu, G., O. Tekeli, I.̇ Çemen, Ş. Şen, and V. Işık (2002), The role of the flexural rotation/rolling hinge model in the tectonic evolution of the Alaşehir graben, western Turkey, Geological Mag., 139, 15–26. Sherlock, S., S. Kelley, S. Inger, N. Harris, and A. I. Okay (1999), 40Ar/39Ar and Rb‐Sr geochronology of high‐pressure metamorphism and exhumation history of the Tavsanli Zone, NW Turkey, Contrib. Mineral. Petrol., 137, 46–58; doi: 10.1007/PL00013777. Siyako, M. (2006), “Lignitic sandstones” of the Trakya basin, Mineral Res. Exploration Bull., 132, 63–72. Siyako, M., K. Bürkan, and A. I. Okay (1989), Biga ve Gelibolu yaryımadalarının Tersiyer jeolojisi ve hidrokarbon olanaklary (Geology of the Biga abd Gelibolu Peninsulas and their hydrocarbon potentials), Türkiye Petrol Jeologları Derneği Bülteni, 1/3, 183–199. Siyako, M., and O. Huvaz (2007), Eocene stratigraphic evolution of the Thrace basin, Turkey, Sed. Geol., 198, 75–91. Sözbilir, H. (2001), Geometry of macroscopic structures with their relations to the extensional tectonics: Field evidence from the Gediz detachment, western Turkey, Turkish J. Earth Sci., 10, 51–67. Sözbilir, H. (2002), Geometry and origin of folding in the Neogene sediments of the Gediz graben, western Anatolia, Turkey, Geodinamica Acta, 15, 277–288. Sözbilir, H., B. Sarι, B. Uzel, Ö. Sümer, and S. Akkiraz (2011), Tectonic implications of transtensional supradetachment basin development in an extension‐parallel transfer zone: The Kocaçay Basin, western Anatolia, Turkey, Basin Research, 23(4), 423–448. Spakman, W., and M. J. R. Wortel (2004), Tomographic view on Western Mediterranean geodynamics, in The TRANSMED Atlas: The Mediterranean Region from Crust to Mantle, edited by W. Cavazza, F. Roure, W. Spakman, G. M. Stampfli, P. Ziegler, 31–52, Springer Berlin Heidelberg. Spakman, W., M. J. R. Wortel, and N. J. Vlaar (1988), The Hellenic subduction zone: A tomographic image and its geodynamic implications, Geophys. Res. Lett., 15(1), 60–63; doi: 10.1029/GL015i001p00060. Stampfli, G. M., and C. Hochard (2009), Plate tectonics of the Alpine realm, in Ancient Orogens and Modern Analogues, J. B. Murphy, J. D. Keppie, and A. J. Hynes, 89–111, Geological Society, London, Special Publications. Steininger, F. F., F. Roegl, and M. Dermitzakis (1987), Report on the round‐table discussion: “Mediterranean and Paratethys Correlations,” Annals of the Hungarian Geological Institute, 70, 397–421. Steininger, R. F., J. Senes, K. Kleemann, and F. Rögl, eds. (1985), Neogene of the Mediterranean Tethys and Paratethys, Stratigraphic Correlation Tables and Sediment Distribution Maps, 1 (Inst. Paleontol.) Vienna. Straub, C. S., H.‐G. Kahle, and C. Schindler (1997), GPS and geological estimates of the tectonic activity in the Marmara Sea region, NW Anatolia, J. Geophys. Res., 102(B12), 27587–27601.

Major Problems of Western Anatolian Geology  185 ̇ Sümer, Ö., U. I nci, and H. Sözbilir (2013), Tectonic evolution of the Söke basin: extension‐dominated transtensional basin formation in western part of the Büyük Menderes Graben, western Anatolia, Turkey, J. Geod., 65, 148–175. Taymaz, T. (1996), S‐P wave travel‐time residuals from earthquakes and lateral inhomogeneity in the upper mantle beneath the Aegean and the Hellenic trench near Crete, Geophys. J. Int.‐Oxford, 127, 545–558. Taymaz, T., O. Tan, and S. Yolsal (2008), Recent devastating earthquakes in Turkey and active tectonics of the Aegean and Marmara Seas, in Earthquake Monitoring and Seismic Hazard Mitigation in Balkan Countries, edited by E. Husebye, 47–55, NATO‐Science Series: IV, Earth and Environmental Sciences, 81, Springer; doi: 10.1007/978‐1‐ 4020‐6815‐7. Taymaz, T., Y. Yılmaz, and Y. Dilek (2007), The geodynamics of the Aegean and Anatolia: Introduction, in The Geodynamics of the Aegean, Anatolia, edited by T. Taymaz, Y. Yılmaz, and Y. Dilek, 1–16, Geological Society, London, Special Publications, 291. Temel, Ö., B. N. Çiftçi, N. Terzioğlu, and H. Sancay (2004), Edremit Körfezi dolayının jeolojisi ve Hidrokarbon olanakları (Geology of the Edremit Bay and the surroundings, and the hydrocarbon potential), TPAO Arama Dairesi Arşivi Rapor, No. 4534. Thomson, S. N., and U. Ring (2006), Thermochronologic evaluation of postcollision extension in the Anatolide orogen, western Turkey, Tectonics, 25. Tirel, C., F. Gueydan, C. Tiberi, and J.‐P. Brun (2004), Aegean crustal thickness inferred from gravity inversion, geodynamical implications, Earth Planet. Sci. Lett., 228, 267–280; doi: 10.1016/j.epsl.2004.10.023. Tirel, C., P. Gautier, D. J. J. Van Hinsbergen, and M.J.R. Wortel (2009), Sequential development of interfering metamorphic core complexes: Numerical experiments and comparison with the Cyclades, Greece, Geological Society, London, Special Publications, 311(1), 257–292. Tommasini, S., S. Conticelli, and R. Avanzinelli (2011), TheTh/ La and Sm/La conundrum of the Tethyan realm lamproites, Earth Planet. Sci. Lett., 301, 469–478. Ünay, E., and F. Göktas (1999), Söke Cevresi (Aydın) Gec Erken Miyosen ve Kuvaterner yaşlı küçük memelileri: Ön sonuçlar, Türkiye Jeoloji Bülteni, 42, 99–113. Ünay, E., F. Göktas, H. Y. Hakyemez, M. Avsar, and Ö. San (1995), Büyük Menderes Grabeni’nin kuzey kenarındaki cökellerin Arvicolidae (Rodentia, Mammalia) faunasına dayalı olarak yaslandırılması, Türkiye Jeoloji Bülteni, 38(7), 5–80 (in Turkish with English abstract). Uzel, B., H. Sözbilir, and Ç. Özkaymak (2012), Neotectonic evolution of an actively growing superimposed basin in western Anatolia: The inner bay of Iż mir, Turkey, Turkish J. Earth Sci., 21 (4), 439–471. Van der Meer, D. G., W. Spakman, D. J. J. van Hinsbergen, M. L. Amaru, and T. H. Torsvik (2010), Toward absolute plate motions constrained by lower mantle slab remnants, Nat. Geosci., 3, 36–40.

Van Hinsbergen, D. J. J. (2010), A key extensional metamorphic complex reviewed and restored: The Menderes Massif of western Turkey, Earth‐Sci. Rev., 102, 60–76; doi:10.1016/j. earscirev.2010.05.005. Van Hinsbergen, D. J. J., and S. M. Schmid (2012), Map view restoration of Aegean–West Anatolian accretion and extension since the Eocene, Tectonics, 31, TC5005; doi: 10.1029/2012TC003132, 2012. Van Hinsbergen, D. J. J., C. G. Langereis, and J. E. Meulenkamp (2005b), Revision of the timing, magnitude and distribution of Neogene rotations in the western Aegean region, Tectonophysics, 396, 1–34; doi: 10.1016/j. tecto.2004.10.001. Van Hinsbergen, D. J. J., E. Hafkenscheid, W. Spakman, J. E. Meulenkamp, and M. J. R. Wortel (2005a), Nappe stacking resulting from subduction of oceanic and continental lithosphere below Greece, Geology, 33, 325–328; doi: 10.1130/ G20878.1. Van Hinsbergen, D. J. J., E. Snel, S. A. Garstman, M., Marunteanu, C. G. Langereis, M. J. R. Wortel, and J. E. Meulenkamp (2004), Vertical motions in the Aegean volcanic arc: Evidence for rapid subsidence preceding in situ volcanism on Milos and Aegina, Marine Geol., 209, 329–345; doi:10.1016/j.margeo.2004.06.006. Van Hinsbergen, D. J. J., M. J. Dekkers, E. Bozkurt, and M. Koopman (2010), Exhumation with a twist: Paleomagnetic constraints on the evolution of the Menderes metamorphic core complex (western Turkey), Tectonics, 29; 10.1029/ 2009TC002596. Verge, N. J. (1995), Oligo‐Miocene extensional exhumation of the Menderes Massif, western Anatolia, Terra Abstracts, 7, 117. Verge, N. J. (2000), The tectonics and consequences of mid‐ Oligocene to earliest Miocene detachment‐facilitated, large magnitude NNE‐SSW horizontal extension/vertical thinning of the earlier tectonically thickened continental lithosphere of western Anatolia, International Earth Sciences Colloquium on the Aegean Region, Izmir, September 2000, Abstract, 256. Washington, H. S. (1894), Art. XV: On the Basalt of Kula, Amer. J. Sci.47,114–123. Westaway, R. (1994b), Present‐day kinematics of the Middle East and Eastern Mediterranean, J. Geophys. Res., 99 (1994), 12071–12090. Westaway, R. (2003), Kinematics of the Middle East and eastern Mediterranean updated, Turkish J. Earth Sci., 12(1), 5–46. Westaway, R. (2004), Kinematic consistency between the Dead Sea Fault Zone and the Neogene and Quaternary left‐lateral faulting in SE Turkey, Tectonophysics, 391, 203–237. Westaway, R. (1994a), Evidence for dynamic coupling of surface processes with isostatic compensation in the lower crust during active extension of western Turkey, J. Geophys. Res., 99, 20203–20223. Westaway, R., H. Guillou, S. Yurtmen, T. Demir, S. Scaillet, and G. Rowbotham (2005), Constraints on the timing and regional conditions at the start of the present phase of crustal

186  ACTIVE GLOBAL SEISMOLOGY extension in western Turkey, from observations in and around the Denizli region, Geodinamica Acta, 18(3‐4), 209–238. Whitney, D. L., and E. Bozkurt (2002), Metamorphic history of the southern Menderes massif, western Turkey, Geol. Soc. Am. Bull., 114(7), 829–838. Whitney, D. L., C. Teyssier, E. Toraman, N. C. A. Seaton, and A. K. Fayon (2011), Metamorphic and tectonic evolution of a structurally continuous blueschist‐to‐Barrovian terrane, Sivrihisar Massif, Turkey, J. Metamorph. Geol., 29, 193–212; doi: 10.1111/j.1525‐1314.2010.00915.x. Whitney, D. L., C. Teyssier, S. C. Kruckenberg, V. L. Morgan, and L. J. Iredale (2008), High‐pressure‐low‐temperature metamorphism of metasedimentary rocks, southern Menderes Massif, western Turkey, Lithos, 101, 218–232. Wortel, M. J. R., and W. Spakman (1992), Structure and dynamics of subducted lithosphere in the Mediterranean region, Proceedings of the Koninklijke Nederlandse Akademie van Wetenschappen, 95, 325–347. Wortel, M. J. R., and W. Spakman (2000), Subduction and slab detachment in the Mediterranean‐Carpathian region, Science, 290, 1910–1917. Yağmurlu, F. (1987), Salihli güneyinde üste doğru kabalaşan Neojen yaşlı alüvyonel yelpaze çökelleri ve Gediz Grabeni’nin tektonosedimanter gelişimi, Türkiye Jeoloji Bülteni, 33–40. Yağmurlu, F., and M. E. Karaman (1987), Kovada güneyinde yeralan linyit içerikli dağarasi Neojen havzalarinin jeolojik özellikleri [Geological features of the lignite‐bearing inttramontaneous basins (Neogene), south of Kovada Lake, Isparta, Turkey], Akdeniz Üniversitesi Isparta Mühendislik Fakültesi Jeoloji Mühendisliği Dergisi (Isparta), 3. Yalçıner, C. Ç., E. Altuner Maksim Bano, et  al. (2013), Application of GPR to normal faults in the Büyük Menderes Graben western, Turkish J. Geod., 65, 218–227. Yazman, M. K., A. Güven, Y. Ermis, M. Yılmaz, I.̇ Özdemir, Y. Akçay, U. Gönülalan, Ö. Tekeli, V. Aydemir, A. Sayılı, Z. Batı, H. Iztan, and Ö. Korucu (1998), Alaşehir Grabeni’nin ve  Alaşehir‐1 Prospektinin Değerlendirme Raporu, TPAO  Exploration Group, unpublished technical report, Ankara. Yılmaz, K. (2010), Origin of anorogenic “lamproite‐like” potassic lavas from the Denizli region in Western Anatolia Extensional Province, Turkey, Mineralogy and Petrology, 99, 219–239. Yılmaz, Y. (1981), Sakarya kıtası güney kenarının tektonik evrimi (Tectonic evolution of the southern edge of the Sakarya continent), Istanbul Yerbilimleri, 1(1–1), 35–52 (Turkish with English abstract). Yılmaz, Y. (1987), Old basement and ophiolite in mid‐Sakarya area, Sixth Colloquium on the Aegean Region, Sept.26–29, ̇ Izmir, edited by E. Dizdar, E. Nakoman, Piri Reis International Contribution Series Publ., No. 2, 699–704. Yılmaz, Y. (1989), An approach to the origin of young volcanic rocks of western Turkey, Tectonic Evolution of the Tethyan Region, edited by A. M. C. Şengör (Kluwer) 159–189. Yılmaz, Y. (1990a), Allochtonous terrains in the Tetyan Middle East: Anatolia and surroubnding regions, Phil. Trans. R. Soc. Lond., A, 331, 661–624.

Yılmaz, Y. (1990), Comparison of young volcanic associations of western and eastern Anatolia under compressional regime: A review, J. Volcanol. Geotherm. Res., 44, 69–87. Yılmaz, Y. (1997), Geology of Western Anatolia, Theme 3: General geology, in Active Tectonics of Northwestern Anatolia: The Marmara Polyproject, A Multidisciplinary Approach by Space–Geodesy, Hydrology, Geothermics, and Seismology, Vdf Hochschulverlag, Zurich. Yılmaz, Y. (2000), Ege Bölgesinin aktif tektoniği. Batı Anadolu’nun depremselliği Sempozyumu (BADSEM 2000), ̇ 24–27, Mayıs 2000, Izmir, Bildiriler, 3–14. Yılmaz, Y. (2002), Tectonic evolution of western Anatolian extensional province during the Neogene and Quaternary, Geological Society of America Abstracts with Programs, 34(6), 179. Yılmaz, Y. (2008), Main geological problems of western Anatolia and the significance of the Bodrum magmatic province, IOP Conference Series, Earth Environ Sci., 2. Yılmaz, Y., and A. Polat (1998), Geology and evolution of the Thrace Volcanism, Turkey. Acta Volcanologica, 10, 293–303. Yılmaz, Y., and Z. Karacık (2000), Geology of northern side of the Gulf of Edremit and its tectonic significance for the development of the Aegean grabens, Geodinamica Acta, 14, 312–343. Yılmaz, Y., C. Genç, F. Gürer, M. Bozcu, K. Yılmaz, Z. Karacık, S. Altınkaynak, and A. Elmas (2000), When did the western Anatolian grabens begin to develop? in Tectonics and Magmatism in Turkey and Surrounding Area, edited by E. Bozkurt, J. A. Winchester, J. D. A. Piper, 353–384, Geologic Society, London, Special Publications, 173. Yılmaz, Y., E. Gökaşan, and A. Y. Erbay (2010) Morphotectonic development of the Marmara Region, Tectonophysics, 488, 51–70. Yılmaz Y., O. Gürpınar, C. Ş. Genç, Ö. Gürer, Z. Karacık, Ş. Altunkaynak, M. Bozcu, and K. Yılmaz (1999b), Milas‐Ören ve Göktepe‐Kale Neojen havzaları ile Bodrum yarımadasının Jeolojisi,TPAO Arama Grubu Rapor No. 4003 (Unpublished 2 reports and the accompaniying 15 suppliments and geology maps), 42. Yılmaz, Y., O. Tüysüz, E. Yiğitbaş, S. C. Genç, and A. M. C. Şengör (1997), Geology and tectonic evolution of the Pontides, in Regional and Petroleum Geology of the Black Sea and Surrounding Region, edited by A. G. Robertson, 83–226, AAPG Memoir, 68. Yılmaz, Y., Ş. C. Genç, E. Yiğitbaş, M. Bozcu, and K. Yılmaz (1995), Geological evolution of the Late Mesozoic continental margin of Northwestern Anatolia, Tectonophysics, 243, 155–171. Yılmaz, Y., Ş. C. Genç, Ö. F. Gürer, Z. Karacık, S. Altunkaynak, M. Bozcu, K. Yılmaz, and A. Elmas (1999a), Ege Denizi ve Ege bölgesinin jeolojisi ve evrimi, in Türkiye Denizleri, Devlet Planlama Teşkilatı, edited by N. Görür, Türkiye Bilimsel ve Teknolojik Araştırma Kurumu yayını, Ankara, 211–337. Yılmaz, Y., Ş. C. Genç, Z. Karacikand S. Altunkaynak (2001), Two contrasting magmatic associations of NW Anatolia and their tectonic significance, J. Geod., 31, 243–271 Yolsal‐Çevikbilen, S., C. B. Biryol, S. Beck, G. Zandt, T. Taymaz, H. E. Adıyaman, and A. A. Ozacar (2012), 3‐D

Major Problems of Western Anatolian Geology  187 crustal structure along the North Anatolian fault zone in north‐central Anatolia revealed by local earthquake tomography, Geophys. J. Int., 188 (3), 819–849; doi:10.1111/ j.1365‐246X.2011.05313.x. Yücel-Öztürk, Y., Helvacı, C., Palmer, M. R., Ersoy, E. Y., & Freslon, N. (2015), Origin and significance of tourmalinites and tourmaline-bearing rocks of Menderes Massif, western Anatolia, Turkey. Lithos, 218, 22–36.

Zanetti, A., M. Mazzucchelli, G. Rivalenti, and R. Vannucci (1999), The Finero phlogopite‐peridotite massif: An example of subduction‐related metasomatism, Contrib. Mineral. Petrol., 134, 107–122. Zhu, L., B. J. Mitchell, N. Akyol, I. Cemen, and K. Kekovalı (2006), Crustal thickness variations in the Aegean region and implications for the extension of continental crust, J. Geophys. Res. Sol. Earth (1978–2012) 111(B), 01301; doi: 10.1029/2005JB00377.

7

The Çataldağ Plutonic Complex in Western Anatolia: Roles of Different Granites on the Crustal Buildup in Connection With the Core Complex Development Ömer Kamacı,1 Alp Ünal,1 Şafak Altunkaynak,1 Stoyan Georgiev,2 and Zeki M. Billor3

ABSTRACT This study documents the geology, structure and age of the Çataldağ Plutonic Complex (ÇPC), the rock association within the footwall of the Çataldağ detachment fault zone (ÇDFZ). ÇPC consists of two contrasting granitic bodies, an older granite‑gneiss‑migmatite complex (GGMC) and a younger I‑type granodioritic body (ÇG: Çataldağ granodiorite). GGMC is a heterogeneous body consisting of migmatite, gneiss, and two‑mica granite, and represents a deep‑seated pluton. By contrast, ÇG represents a discordant, shallow level intrusive body. New U‑Pb zircon (LA‑ICP‑MS) and monazite ages of GGMC yielded magmatic ages of 33.8 and 30.1 Ma (Latest Eocene–Early Oligocene). 40Ar/39Ar muscovite, biotite, and K‑feldspar from the GGMC yielded the deforma­ tion age span 21.38 ± 0.05 Ma and 20.81 ± 0.04 Ma, which is also the emplacement age (20.84 ± 0.13 Ma and 21.6 ± 0.04 Ma) of ÇG. The ÇG‑GGMC duo represents a core complex, which was exhumed in the Early Miocene as a dome structure in the footwall of a ring‑shaped low‑angle detachment surface. A number of micro­ scale and mesoscale shear sense indicators display evidence that the rocks underwent a ductile deformation in the earlier stage of the elevation, which was superimposed later by a semibrittle and brittle deformation. They indicate top‑to‑north and top‑to‑northeast sense of movement. The exhumation process was partly contempo­ raneous with the development of the major core complexes of the region (e.g., the Menderes Massif and the Kazdağ Massif) as a result of combined effects of thermal weakening and rollback of the Aegean subducted slab during the Oligocene–Early Miocene. 7.1. INTRODUCTION Studies on the extensional tectonics and the low‐angle normal faults facilitated our understanding of the connection between granite ascent and emplacement mechanisms. Such insights have contributed greatly to “space problem” or “room problems,” which have been 1 ̇  Department of Geological Engineering, Istanbul Technical ̇ University, Istanbul, Turkey 2  Department of Geochemistry and Petrology, Geological Institute, Bulgarian Academy of Sciences, Sofia, Bulgaria 3  Department of Geology and Geography, Auburn University, Auburn, Alabama, USA

discussed since 1835 by Charles Lyell [Hutton et al., 1990]. These processes are now believed to operate in many places including the Basin and Range [Gans et  al., 1989], Sierra Nevada [Tobish et al., 1993], the Cyclades [Gautier et al., 1993], Papua New Guinea [Hill et al., 1995], Yemen [Geoffrey et al., 1998], the Tyrrhenian Sea [Jolivet et al., 1998], Green­ land [Strachan et  al., 2001], and the Menderes Massif [Çemen et al., 2006; Dilek et al., 2009; Erkül, 2010; Erkül and Erkül, 2012; Işık et al., 2004; Catlos et al., 2012; Okay and Satır, 2000], where granite intrusions into extensional shear zones have been conclusively demonstrated. The Aegean Province is a well‐known region of exten­ sion [McKenzie, 1978; Lister et al., 1984; Taymaz et al., 1991; Jackson, 1994; Gautier and Brun, 1994; Dinter, 1998;

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 189

190  ACTIVE GLOBAL SEISMOLOGY

McClusky et al., 2000; Jolivet and Faccenna, 2000; Jolivet et al., 2013]. As a part of the Aegean extensional prov­ ince, western Anatolia experienced intense extensional deformation and magmatism in the Late Oligocene–Early Miocene, as manifested by development of extensional basins, metamorphic core complexes, and widespread magmatism. Exhumation of lower to middle crustal rocks and the development of extensional metamorphic domes [Menderes and Kazdağ metamorphic core com­ plexes: Bozkurt and Park, 1994; Hetzel et al., 1995; Işık and Tekeli, 2001; Gessner et al., 2001; Ring et al., 2003; Bozkurt, 2007; van Hinsberg, 2010; Okay and Satir, 2000;  Bonev et  al., 2009; Cycladic core complex: Buick, 1991; Gautier and Brun, 1994; Vandenberg and Lister, 1996; Rhodope core complex: Dinter and Royden, 1993; Crete core complex: Jolivet et  al., 1994; Kilias et  al., 1994] occurred during the extension in western Anatolia and the adjacent regions. Granitic plutonism and extension are broadly contem­ poraneous in many core complexes. There are many studies that display the detachment faults and the related normal faults played an important role in the exhumation of these granites in an extensional regime. Examples from the Menderes core complex include the Salihli, Turgutlu,

Egrigöz, Koyunoba, Baklan, and Alasehir granites [Işık  et  al., 2003; Çemen et  al., 2006; Dilek et  al., 2009; Catlos et  al., 2010; Erkül and Erkül 2012; Altunkaynak et al., 2012a] and from the Kazdağ core complex include the Evciler and Eybek plutons [Okay and Satir, 2000; Beccaletto et al., 2005; Bonev et al., 2009]. However, there are some other studies on the granites of western Anatolia, suggesting that granite generation and emplacement of the coeval intrusions are not linked directly to any core complex development (the Eğrigöz and Koyunoba granites; Hasözbek et  al. [2010]). Other granitic bodies straddling on both sides of the Izmir‐Ankara‐Erzincan suture zone (IAESZ) and around the core complexes were uplifted also, when the overlying rocks were exhumed in the same period. Therefore, western Anatolia is among the best places to study interplay between extension and granite emplacement. We have studied the Çataldağ Plutonic Complex (ÇPC) that is exposed to the immediate north of the IAESZ (Fig. 7.1). The Çataldağ Plutonic Complex shows close spatial and temporal relationship to a detachment fault zone (Çataldağ detachment fault zone: ÇDFZ) and consists of two contrasting granitic bodies displaying different textural, structural, and lithological features.

ontid

Intrap

rs

a rd Va

Thrace basin

Rhodope

ag

e ur

l Pe

ut

r Saka

n

ia

on

6

An at Ta olide u pla ride tfo rm

4 3

Me

nd e

res

e

n zo

Kazdag˘

2

1

s

n ian

c Ly

ic arc

te

He

lle

nic

pe

ap

de

Cre

e

n ya zo

IAE suture

5

Cy cla

e sutur

lcan Recent vo

Oligo-Miocene plutons Eocene plutons Metamorphic massifs Study area Detachment fault

N

Thrust fault

Tre n

ch

Movement direction of hanging-wall 200 km

Figure 7.1  Simplified tectonic map of the Aegean region showing major tectonic units, metamorphic massifs, Eocene and Oligo‐Miocene granitoids [modified from Bonev el al., 2006; Cavazza et  al., 2009; Erkül, 2010]. Red box shows the study area. Extension‐related Early Miocene plutons: 1–6, 1‐Salihli, 2‐Turgutlu, 3‐Eğrigöz, 4‐Koyunoba, 5‐Yenice, 6‐Eybek granite.

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  191

Our geochronological data suggests these granitic bodies have also different emplacement ages corresponding to the Eo‐Oligocene and the Early Miocene and thus repre­ sent magmatic phases shortly preceding and during the extension. This is the first detailed structural study that was conducted on Çataldağ Plutonic Complex and the ÇDFZ. Therefore, main purposes of this paper are (1) to introduce structural (microstructure and macrostructures) and geochronological (LA‐ICP‐MS ages on zircon and monazite and 40Ar/39Ar ages) data sets obtained from these two diverse granitic body, (2) to compare and cor­ relate these data with Çataldağ detachment fault zone (ÇDFZ) and the related deformations. The broader goal is to better define emplacement mechanisms of the Çataldağ Plutonic Complex and its role in the Cenozoic crustal buildup of western Anatolia, which can serve as a model for other regions of the world. 7.2. ÇATALDAG PLUTONIC COMPLEX (ÇPC) An older granite‐gneiss‐migmatite complex (GGMC) constitutes the eastern half of the ÇPC, whereas a younger granodioritic body (ÇG: Çataldağ granodiorite) and associated sills and dikes form in the western half. 7.2.1. The Granite‐Gneiss‐Migmatite Complex (GGMC) The GGMC of the ÇPC is located around Turfaldag Mountain (Fig. 7.2a). The GGMC is an internally heter­ ogeneous igneous body. It consists mainly of two‐mica granite, quartz‐feldspar gneiss, and migmatite. All of the members gradually pass into one another. Two‐mica granite is common in the central zone of GGMC. It is variably deformed but has preserved magmatic textures. Locally preserved magmatic textures indicate that they are equigranular and weakly porphyritic medium‐ grained granite. Two‐mica granite contains quartz, ortho­ clase, oligoclase, muscovite, and reddish‐brown biotite with occasional garnet (Fig.  7.3a, b). Apatite, titanite, zircon, and monazite are common accessories. Muscovite is seen as both primary and secondary phases. The latter, as the alteration product, formed particularly along cleavage planes of feldspars. Cataclastic deformation is a common feature in the GGMC, which overprints the magmatic textures. In deformed two‐mica granites, quartz is surrounded by recrystallized small subgrains. Ellipsoidal K‐feldspar and plagioclase are aligned along the main foliation. Garnet is generally present as shattered porphyroclasts in the rock. Gneiss is a quartzofeldspatic, leucocratic rock repre­ sented by an assemblage that consist mainly of quartz, plagioclase, K‐feldspar, and muscovite (Fig. 7.3c, d). It also has garnet, biotite, and zircon. Gneiss gradually passes into two‐mica granite, which also contains quartz,

plagioclase, K‐feldspar, and muscovite as the major minerals suggesting its granitic origin. Foliation and line­ ation are defined by muscovite, biotite, and quartz. In the field, gneiss and migmatite are observed to be intricately intermixed (Fig. 7.3e). Common mineral composition of the migmatite is bio­ tite + quartz + K‐feldspar + plagioclase + garnet ± cordierite ± sillimanite. This mineral paragenesis indicates that the migmatite underwent amphibolite facies (upper amphibo­ lite?) metamorphism. Leucosome and melanosome (restite) rich migmatites [Sawyer, 2008] can be distinguished within the GGMC (Fig.  7.2a). Melanosome rich migmatite is a common rock particularly in the northwestern and the eastern borders of the GGMC and toward the center, pass­ ing into leucosome‐rich migmatite. The mineral assemblage of the leucosome is similar to that of gneiss and granite. Boundaries between leucosome and melanosome are marked by an anomalously biotite‐rich zone. It shows schlieren and fold‐type structures, where the migmatization reached an advanced stage (Fig. 7.3e). Mafic synplutonic dikes were identified within the migmatite close to northern border of GGMC (Fig. 7.4). Structures and contact relations of mafic dikes and the host rock indicate that the host was totally molten or mush, when the dike was intruded. It is gabbroic diorite in composition and consists of plagioclase, hornblende, clinopyroxene, biotite, and quartz (as rounded xenocryst). Microgranular texture is predominant in the rock. During the continuous ductile deformation, the dikes were pinched out to form a number of separate small dikes aligned along the same direction surrounded by migma­ tite (Fig. 7.4). Along the northern and eastern borders, the GGMC is bounded by a low‐angle normal fault zone (ÇDFZ) that juxtaposes granite‐migmatite complex with the green­ schist facies metamorphic rock of the basement rocks [Karakaya complex and Bornova flysch: Okay, 1984; Okay et al., 2001] (see Structural Data, Section 7.3, for details on the ÇDFZ). Haplogranitic sheets and pegma­ titic dikes of ÇG were emplaced through this fault and cut the GGMC along its western contact. 7.2.2. The Çataldağ Granodiorite (ÇG) The Çataldağ granodiorite consists mainly of homoge­ neous granodiorite with haplogranitic peripheral rocks. The main granodiorite body is dominated by biotite gran­ odiorite and subordinate hornblende‐epidote granodiorite in the south. From north to south they gradationally pass into one another. The granodiorite is easily identified in the field by its distinctive porphyritic texture consisting of large megacrysts of K‐feldspar reaching up to 5 cm (Fig. 7.5a, b). It is made up of plagioclase (oligoclase‐andesine), +  perthite), and biotite. quartz,  K‐feldspar (orthoclase 

192  ACTIVE GLOBAL SEISMOLOGY

620000

615000

610000

605000

(a) N

Pa¸salar Karapür¸cek

C′

A

442400

8

20,7

Sünlük

20,8

B′

60

21,1 20,8 21,2

441400

Tirnova

ÇG

21.1 C

Serçeören

33,6 30.1

21,3 B Ar/Ar ages

Eo-Oligocene granite-gneissmigmatite complex (GGMC)

Quaternary alluvium Supradetachment sediments Neogene lacustrine sediments Early Miocene Çataldagˇ Granodiorite (ÇG) Haplogranite and pegmatite dykes

440000

21,1

33,8

440400

Cover rocks

A′

32,2

?

?

GGMC

6

Susurluk

Biotite

Gneiss/Two-mica granite

Muscovite

Migmatite

K-feldspar

Basement rocks

U/Pb ages Monazite

Bomova flysch (Upper Cretaceous) Karakaya Complex (Permo-Triassic)

Zircon

Porphyritic granodiorite

?

?

4 km

Possible detachment surface

Undefined fault

Normal-dextral fault

Detachment fault

A

Direction of cross-section

(b) SE

NW

A

(c)

S

A′

N

2 km B

(d)

B′

S

N

C

C′

Figure 7.2  (a) Geological map of the Çataldağ Plutonic Complex; (b) northwest‐southeast cross section of CPC (A‐A′); (c) north‐south cross section showing the GGMC (B‐B′); (d) north‐south cross section of the ÇG (C‐C′). Map coordinates: Universal Transverse Mercator (UTM) projection, zone 35.

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  193 (a)

(b) Ms

Q

Bi

(c)

Pl

(d)

Ms

Gr

Kf

(e)

Figure 7.3  GGMC units: (a, b) Two‐mica granite; (c, d) gneiss; (e) Schollen‐type migmatites [after Mehnert, 1968], including melanosome and leucosomes (UTM coordinate 35S 612620/4421440). Ms = muscovite; Kf = K‐feldspar; Bi = biotite; Q = quartz; Pl = plagioclase; Gr = garnet. Magnification for (b) and (d) is 4X, crossed plane.

Minor hornblende and epidote are observed where the granite is in contact with calc‐silicate rocks along south­ ern contact zone. Titanite, zircon, apatite, and opaques occur as accessory minerals. Chlorite, sericite, and epi­ dote form the secondary minerals. Haplogranitic sheets crop out along the western and eastern borders of the main granodiorite body. It is a leucocratic and fine‐grained rock displaying aplitic or graphic‐granopyhric texture. It consists mainly of quartz, K‐feldspar, plagioclase (oligoclase), and minor biotite and garnet. The haplogranite contact with the main body of the granodiorite is sharp. It displays a chilled margin against the metapelitic‐metabasic basement rocks of the Karakaya complex and Bornova flysch [Okay, 1984, Okay et al., 2001] in the west.

The ÇG intruded into metapelitic‐metabasic basement rocks of the Karakaya complex (Fig.  7.5c). A primary igneous contact zone is observed along the southern border. The ÇG was intruded into calc‐silicate country rocks and formed a well‐developed contact aureole where mineral paragenesis of calcite + diopside + garnet + wollastonite ± vesuvianite indicate that the contact meta­ morphism reached pyroxene‐hornfels facies in the inner aureole. Within the host granodiorite, euhedral horn­ blende and epidote were also locally developed as a result of Ca metasomatism. Aplite sills and dikes, pegmatite dikes and enclaves of the country rocks are common along the southern margin of the ÇG. The ÇG is bounded along the northern and eastern con­ tact by a low‐angle normal fault (the Çataldağ fault zone).

194  ACTIVE GLOBAL SEISMOLOGY

In the northern part of the ÇG, the granite laterally passes into augen‐gneiss (Fig.  7.6). Syngenetical hap­ logranite and pegmatite dikes of ÇG were intruded in this fault zone and formed a narrow dike belt between the ÇG and GGMC (Fig.  7.2a, b). Younger (oblique) normal‐dextral and normal faults cut and concealed the primary contact of the ÇG in the western margin. 7.3. STRUCTURAL DATA Synplutonic dyke

7.3.1. Çataldağ Detachment Fault Zone (ÇDFZ)

GGMC

Figure  7.4  Mafic synplutonic dyke within the GGMC (UTM coordinate 35S 618185/4417585).

GGMC and ÇG are bounded along northern and eastern contacts by a low‐angle normal fault we named herein the Çataldağ detachment fault zone (Fig. 7.2). All the structural features (e.g., slickenlines, undulation, tur­ tleback) that are observed within the fault zone and fault plane collectively display normal fault character of ÇDFZ (Fig. 7.7a, b). ÇDFZ separates GGMC and ÇG, which represents the footwall, from the basement rocks (Karakaya complex and Bornova flysch of Okay [1984]; Okay et  al. [2001]), and the sedimentary cover rocks located on the hanging wall. Attitude of the fault plane of the detachment fault is N85E 8NW along the northern border (Fig.  7.7a). It turns to N5W 6NE toward the (b)

(a)

Kf

4 cm

1 mm

(c)

ÇG

ÇG

Country rock

Figure  7.5  (a, b) Typical K‐feldspar megacrystals forming porphyritic texture of ÇG (UTM coordinate 35S 607975/4414830); (c) intrusive contact between ÇG and metamorphic country rock in western boundary (UTM coordinate 35S 601249/4403663). Kf = K‐feldspar. Magnification for (b) is 4X, crossed plane.

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  195 N

Step-over

S

C 2 cm

Figure  7.6  Augen‐gneiss in the ÇDFZ with top to the north sense of feldspar‐fish and step‐over structures in northern border of ÇG. C = shear plane; S = schistosity plane.

(a)

N ck en

lin

e

(b)

rik eo

undulation

f fa

sli

St

ult

pla

ne

Figure 7.7  The fault plane structures of the ÇDFZ: (a) low‐angle fault plane and undulation; (b) slickenline and rake (UTM coordinate 35S 616580/4422555).

eastern border (Fig. 7.8a). Slickenlines are clearly observed around the Sunluk area (northern border of GGMC) and trend N5W (Fig. 7.7b). This is also the same as the dip of the fault plane. Rake measurement for fault plane is calcu­ lated as (‐90), which indicates a normal fault characteristic (Fig.  7.7b). The mylonitic texture exhibits progressive changes from ductile to brittle deformation in the footwall rocks. A top‐to‐the north sense of motion is determined from the structures of the mylonitic shear zone (Fig. 7.8b, c, d). On the fault plane, some undulations, subparallel to direction of the shear foliation, are observed (Fig. 7.7a). The fault planes undulate and, thus, one of the typical indicators of low‐angle normal faults, local turtleback structures are observed in the eastern border of ÇG, as exemplified from the east of the Serçeören area (Fig. 7.9).

Within the ÇG, the fault and related deformation are not clearly seen everywhere due to weathering and erosion. In some places, the fault plane is commonly eroded and/or truncated, and therefore displaced (hidden) by younger (oblique) dextral normal and normal faults. Attitude of the dextral normal fault plane is N20W 60SW. To show the original shape and geometry of this dome and the associated detachment fault, the ~3.5–4 km right‐slip displacement along the northwest‐southeast‐oriented dextral normal fault that cuts across the dome has been restored. Further retrodeforming the dome by restoring the right‐oblique displacements along the east‐west‐striking faults in its southern half, produces an east‐west‐oriented elliptical body of crystalline rocks bounded in its northern, eastern, and southern parts by ÇDFZ (Fig. 7.10).

(a)

(b) W

E

S

N

Marble (K arakaya) GGMC

(c)

(d) N

Figure 7.8  (a) Hanging wall and footwall of the ÇDFZ in the eastern border of GGMC (UTM coordinate 35S 629500/4418835); (b) mylonites (UTM coordinate 35S 618730/4419565); (c) microphotograph of top to the north sense muscovite‐fish from the outer zone of the GGMC, crossed plane, 10X magnification; (d) boudinage of the leucosomes in migmatite of the GGMC (UTM coordinate 35S 612345/4419795).

(a)

N

ÇG GGMC

250 m

(b)

Figure 7.9  (a) Google Earth image of the low‐angle fault plane of the ÇDFZ and (b) turtle‐back structure in south of the ÇG (UTM coordinate 35S 612574/4409832).

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  197 N

(a)

(b)

Pas¸alar

Karapürçek

Sünlük GGMC Susurluk

Tirnova

ÇG

ÇG

Supradetachment sediments

GGMC

Serçeören

Çataldag˘ granodiorite Granite-gneiss-migmatite complex Detachment fault

Dextral normal fault

Possible detachment surface

Trace of detachment fault

Figure 7.10  Structural reconstruction of ÇG, GGMC and ÇDFZ: (a) the initial shape of the ÇDFZ before oblique dextralnormal faulting; (b) recent shape of ÇDFZ after the oblique dextral‐normal faulting.

Supradetachment sediments are clearly seen above the northern contact of the ÇDFZ. These sedimentary rocks are represented by poorly lithified conglomerate in immediate contact, and composed of clasts of grano­ diorite and granite‐gneiss derived from the ÇG and GGMC, and onlap the detachment surface. Bedding planes tilted into the detachment surface, and bedding thickness, increase toward the detachment surface. This indicates synextensional deposition of these rocks, which is reminiscent of supradetachment basin sequences [Oner and Dilek, 2011, 2013). 7.3.2. Deformation of the Granite‐Gneiss‐Migmatite Complex (GGMC) Structures indicate that the GGMC underwent both ductile and brittle deformation. A total of 315 foliations were measured from the GGMC (Fig.  7.11). The lower hemisphere steographic plots of the foliations and line­ ations are shown in Figure  7.12. Deformation style is different in the inner zone (IZ on Fig. 7.11) and the outer zone (OZ in Fig. 7.11) of the GGMC, which are described separately below. In the inner zone, the rock was ductily deformed, and the foliation of the rock is defined by the orientation of elongated feldspar and mica minerals, whereas in the outer zone, the GGMC commonly displays semibrittle to brittle deformation. Steographic plots of the foliations of the inner zone are shown in Figure 7.12a. Dips of the foliations are clustered mainly northwest and southeast in direction. Strike and dips of foliations demonstrate two northeast‐southwest

trending smaller scale subdomes within a larger dome recognized on the whole body of GGMC (Fig. 7.11). In the inner zone, microscale deformation is charac­ terized by an incipient mylonitization that is evidenced by the beginning of grain‐size reduction and softening reactions generated by quartz recrystallization pro­ cesses, such as grain boundary migration and subgrain rotations. Feldspars are commonly represented by por­ phyroclasts. K‐feldspar frequently displays semiductile features such as bulging recrystallization. Micas are in plastic forms commonly around feldspars. The whole microstructural data suggest that the temperature ranged approximately from 450 °C to 250 °C [Passchier and Trouw, 1996] in the inner zone during the synkinematic deformation. The outer zone is represented by a mylonitic zone that has been variably overprinted by late brittle defor­ mation along the borders of GGMC. Mylonitization increases toward the outer zone indicating the ÇDFZ is responsible from the deformation (Fig. 7.11). Trend of foliations is parallel‐subparallel to the trend of the ÇDFZ (Fig. 7.11). The plunges of lineations of GGMC are roughly parallel with the dip directions of ÇDFZ (Figs.  7.11 and 7.12e) regardless of the attitude of the foliation measured all over the region. However, movement direc­ tion of the hanging wall is top‐to‐north and top‐to‐ northeast (Fig. 7.13). Overprinting brittle deformation is represented by microstructures such as microfaults and cracks that are seen in quartz and feldspar. The brittle deformational

198  ACTIVE GLOBAL SEISMOLOGY

12 10

14 12

10

15

16

12 10 11 30 13 12 11 15 20

31 28

20

30

10 15

15

14 20 23 18 IZ

16

16

8

13 22

21 24 16 18 18

ÇG 27

15

25 27

12

Inner zone (IZ)

12

32

12

26 22

26

14 20 40 20

29

21

23 32

25 35 20 16

20

16

15 22

25 19

30

15

23

26 12 14

GGCM

14 40

Outer zone (OZ)

12

16 20 12 20

24

22 34 28 14

16 20 15

20

2116

18

40

12

35 26

OZ 15

17

16

15

27 19 23

26 10 13 25 32

20

17

21 12 20 21

20

8

Strike and dip of foliation

18 Plunge and azimuth of stretching lineation

Figure 7.11  Strike and dip of foliations and plunge and azimuth of stretching lineations in structural map of ÇPC.

behavior of quartz is stated to be below 250 °C by Passchier and Trouw [1996]. Shear direction is determined from the small‐scale faults. All of the microfaults display dip slip displace­ ment and the motion is “top to north and northeast” from north to east, respectively. The same shear sense can also be determined from mica‐fish texture (Fig.  7.8c), which suggests that the rocks underwent different stages of a progressive deformation. They passed through ductile deformation at an early stage. Later, a brittle deformation superimposed on the ductile deformation. In summary, GGMC displays effects of a progressive deformation. The inner zone of GGMC (Fig.  7.11) displays ductile deformation, whereas the outer zone exhibits both ductile and brittle deformation. The data collectively suggest that GGMC underwent continuous deformation starting from about a temperature interval

of 450 °C–250 °C and it continued in the lower tempera­ tures [Passchier and Trouw, 1996]. 7.3.3. Deformation of the Çataldağ Granodiorite (ÇG) Similar to the outer zone of the GGMC, the north‐ northwest border of ÇG displays evidence for both ductile and brittle deformation in every scale. Along the north boundary of ÇG, the structural evidence indicates a regional top‐to‐north shear sense of semiplastic mylo­ nitic deformation. This is reflected by the orientation and elongation of biotite and feldspar defining foliation that are locally observed (Fig.  7.6). Their dips commonly trend to the north‐northwest (Fig. 7.12d). The grain boundary migrations (GBR), subgrain rota­ tions (SGR), and feldspar‐fish structures are common in the northern border (Fig.  7.6). According to Passchier

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  199

(a)

N

148 data (c)

68 data (d)

N

94 data

N

(b)

N

N

24 data

(e)

66 data Figure  7.12  Lower hemisphere equal‐area projections of foliation (a, b, c, d) and lineation (e) from the ÇPC: (a) foliations for inner zone of the GGMC; (b) foliation for the eastern outer zone of the GGMC; (c) foliations for the northern outer zone of the GGMC; (d) foliations for the northern border of the ÇG; (e) lineation for the GGMC.

and Trouw [1996], the structures such as feldspar‐fish and mica‐fish are formed at the temperatures of 600 °C and 300 °C, respectively. Toward the northern border of the ÇG, the CDFZ separates the basement rocks that are exposed on the hanging wall from the ÇG on the footwall. Cataclasis increases toward the fault plane. The shear sense indica­ tors, for example, feldspar‐fish structures (Fig.  7.6), in the northern part of the ÇG consistently define top‐to‐ the north movement. The microfaults observed in the fault zone share the same top‐to‐the north and north­ west sense of shear. In contrast to northern and western borders, other con­ tacts of ÇG with the metamorphic country rock are pri­ marily intrusive in character. The structures that resulted in the exhumation of the plutonic bodies do not cut the southern border.

7.4. LA‐ICP‐MS GEOCHRONOLOGY 7.4.1. Analytical Methods Cathodoluminescence (CL) and back‐scattered (BSE) images were collected prior to zircon analyses to identify inherited cores, cracks, and inclusions using the scanning electron microscope JSM‐ 6610 LV at the University of Belgrade. U‐Pb isotope analyses were carried out using a New Wave Research (NWR) Excimer 193 nm laser‐ablation system UP‐193FX attached to a Perkin‐Elmer ELAN DRC‐e inductively coupled plasma mass spectrometer (LA‐ICP‐MS) at the Geological Institute, Bulgarian Academy of Science in Sofia, Bulgaria. Spatial resolution was about 35 µm (rarely 20–25 µm) for the thinner zones and frequency of 8 Hz. The well characterized zircon

200  ACTIVE GLOBAL SEISMOLOGY

OZ

IZ

ÇG

GGCM

OZ

Outer zone (OZ) Movement direction of hanging-wall Inner zone (IZ)

Figure 7.13  Movement direction of hanging wall.

standard (GEMOC GJ‐1; 608 Ma with near‐concordant Pb) was used as external standard. Ablation of the stand­ ard and samples in He was employed to increase sample transport efficiency, decrease ablation product deposi­ tion, and stabilize and increase the signals. During the analyses, the SuperCell of NWR allowed the use of low‐ carrier gas flow. The measurement protocol included the isotopes 206Pb, 207Pb, 208Pb, 204Pb, 232Th, 238U, and 201Hg. In nearly all cases, 204Pb content was below the level of detec­ tion and no correction was applied. Operating conditions for the laser were kept stable to ensure constant U/Pb fractionation. Measurement procedure involved calibra­ tion against the standard zircon GJ‐1 at the beginning of the “block” with a run of two analyses, followed by eight zircon analyses, one standard GJ‐1, the next seven unknowns, and finishing the block with two further analyses of the GJ‐1 standard. This technique allows a suitable correction for instrumental drift along with the minimization of elemental fractionation effects. Raw data were processed using GLITTER, a data reduction program of the GEMOC, Macquarie University, Australia (http://www.gemoc.mq.edu.au/). 207Pb/206Pb, 208Pb/232Th, 206 Pb/238U, and 207Pb/235U ratios were calculated and the

time‐resolved ratios for each analysis were then carefully examined. Optimal signal intervals for the background and ablation data were selected for each sample and automatically matched with the standard zircon analyses, thus correcting for the effects of ablation/transport‐related U/Pb fractionation and the mass bias of the mass spec­ trometer. Net background‐corrected count rates for each isotope were used for calculation of sample ages. U‐Pb Concordia ages are calculated and plotted using ISOPLOT [Ludwig, 2003]. Rho value is taken as 0.5 for the U‐Pb Concordia diagrams. 7.4.2. Results LA‐ICP‐MS ages were achieved from zircon and mon­ azite from GGMC (Table 7.1). Most of the zircon grains in the samples of GGMC reveal the complex internal structure and inherited cores and autocrystic magmatic rims (Fig.  7.14). Most of the cores have endured mag­ matic corrosion, which is represented by rounding, and some of the newly grown zircon zones penetrate in them. Some of the margins between the cores and the rims are cracked, probably due to thermal effect. The newly grown

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  201 Table 7.1  U/Pb LAICPMS Age Data of the CPC LA–ICP–MS Zircon Radiometric Age Determination of Sample SA1 Spot

206

Pb/238U

1 Sigma

1r 1c 2r 2c 3r 3c 4 5r 5c 6r 7r 7c 8c 8r 9r 10r 11r 11c 12c 13r 14r 14c 15r 15c 16r 17r 18c 18r 19c 19r 21r1 21r2 21c 20r 23r 24r 25c 25r 26r 26c 28r 29r 29c 30c 30r

0.00499 0.15479 0.00499 0.04815 0.005 0.08722 0.00508 0.00496 0.00507 0.00502 0.00512 0.04966 0.07113 0.00506 0.0051 0.00799 0.00486 0.04475 0.08576 0.0234 0.00597 0.01295 0.00494 0.00505 0.00495 0.00473 0.30158 0.00583 0.08576 0.00511 0.00518 0.11821 0.14077 0.00504 0.00523 0.00497 0.12022 0.00782 0.00503 0.00499 0.00507 0.00493 0.00503 0.11992 0.0072

0.0001 0.00188 0.0001 0.00081 0.00006 0.00103 0.00006 0.00006 0.00011 0.00007 0.00009 0.00111 0.00112 0.00007 0.00008 0.00012 0.00007 0.00052 0.00116 0.00027 0.00009 0.00029 0.00007 0.00013 0.00007 0.00008 0.00356 0.0001 0.00121 0.00009 0.00007 0.00169 0.00179 0.00006 0.00008 0.00007 0.00199 0.00015 0.00009 0.00007 0.00008 0.00008 0.00009 0.00207 0.00013

Pb/235U

1 Sigma

0.0322 1.50848 0.04281 0.39383 0.03231 0.72829 0.04182 0.03224 0.04281 0.03883 0.03497 0.3611 0.53793 0.03842 0.04661 0.05664 0.03822 0.37014 0.89753 0.15997 0.16259 0.6599 0.03271 0.04572 0.0329 0.03483 6.72922 0.05297 0.93112 0.04529 0.03326 1.15861 1.4072 0.03675 0.03892 0.03281 1.05509 0.05993 0.03251 0.03291 0.03257 0.03153 0.05197 1.07056 0.04652

0.00289 0.04061 0.00327 0.02362 0.00115 0.01872 0.00134 0.00124 0.00359 0.0014 0.00216 0.03515 0.02918 0.00142 0.00205 0.00271 0.00182 0.00946 0.03088 0.00435 0.00529 0.02793 0.00144 0.00471 0.00154 0.00228 0.17055 0.0033 0.03483 0.00263 0.00132 0.04526 0.04414 0.00106 0.00182 0.00153 0.05513 0.00432 0.00227 0.00132 0.00169 0.00174 0.00329 0.06089 0.00333

207

Pb/232Th

208

0.00214 0.0469 0.02299 0.02249 0.00235 0.02945 0.00426 0.00207 0.0023 0.01297 0.00294 0.02781 0.02986 0.00742 0.02196 0.00533 0.01018 0.01839 0.05189 0.01139 0.04987 0.01999 0.00171 0.00449 0.00839 0.00914 0.09812 0.00365 0.03967 0.00573 0.00221 0.04989 0.04836 0.00578 0.0148 0.00462 0.04347 0.00067 0.00698 0.00318 0.00194 0.0021 0.00494 0.0414 0.00126

1 Sigma 0.00034 0.00283 0.00332 0.00176 0.00024 0.00194 0.00033 0.00021 0.00027 0.00127 0.00041 0.00302 0.00296 0.00078 0.00252 0.00062 0.00092 0.00137 0.00367 0.00119 0.00363 0.00149 0.00033 0.00067 0.00103 0.00142 0.00823 0.00061 0.00403 0.0007 0.00025 0.00536 0.00479 0.00048 0.00145 0.00067 0.00309 0.002 0.0016 0.00033 0.0002 0.00085 0.00048 0.00368 0.00061

Pb/206Pb

207

0.04677 0.07068 0.06216 0.05932 0.04686 0.06055 0.05971 0.04716 0.06125 0.05612 0.04955 0.05273 0.05485 0.05511 0.06623 0.05144 0.05704 0.05999 0.07591 0.04959 0.19757 0.36956 0.04802 0.06561 0.04821 0.05338 0.16184 0.06584 0.07875 0.06432 0.04655 0.07109 0.07251 0.05288 0.05399 0.04785 0.06365 0.0556 0.04686 0.04788 0.04657 0.04643 0.07489 0.06475 0.04689

1 Sigma 0.00427 0.00194 0.00488 0.00364 0.0017 0.00158 0.00196 0.00185 0.00526 0.00206 0.00313 0.00523 0.00303 0.00208 0.00299 0.00251 0.00277 0.00156 0.00268 0.00137 0.00682 0.01725 0.00216 0.00694 0.00229 0.00358 0.00416 0.00421 0.00301 0.00384 0.00188 0.00283 0.0023 0.00156 0.00258 0.00227 0.00341 0.0041 0.00334 0.00197 0.00246 0.00262 0.00488 0.00378 0.00342

Age, Ma Pb/238U

1 Sigma

32.1 927.8 32.1 303.1 32.2 539.1 32.7 31.9 32.6 32.3 32.9 312.4 443 32.5 32.8 51.3 31.3 282.2 530.4 149.1 38.4 83 31.8 32.5 31.8 30.4 1699.1 37.5 530.4 32.8 33.3 720.2 849 32.4 33.6 32 731.8 50.2 32.4 32.1 32.6 31.7 32.4 730.1 46.2

0.66 10.51 0.66 4.99 0.4 6.09 0.41 0.41 0.72 0.42 0.56 6.84 6.76 0.43 0.49 0.76 0.46 3.23 6.87 1.73 0.6 1.86 0.44 0.86 0.45 0.54 17.63 0.67 7.18 0.57 0.45 9.71 10.09 0.39 0.49 0.45 11.44 0.93 0.57 0.43 0.48 0.48 0.6 11.92 0.81

Age, Ma Pb/238U

1 Sigma

206

LA–ICP–MS Zircon Radiometric Age Determination of Sample SA2 Spot 1r 2r 2c 3

Pb/238U

1 Sigma

0.00538 0.06663 0.0053 0.0055

0.00006 0.00086 0.00007 0.00012

206

Pb/235U

1 Sigma

0.03654 0.59767 0.03773 0.03554

0.00107 0.01985 0.00129 0.00332

207

Pb/232Th

208

0.00206 0.02428 0.00281 0.00156

1 Sigma 0.00013 0.0013 0.0004 0.00062

Pb/206Pb

207

0.04924 0.06506 0.0516 0.04684

1 Sigma 0.00147 0.00221 0.0018 0.00446

206

34.6 415.8 34.1 35.4

0.41 5.21 0.42 0.75 (continued)

202  ACTIVE GLOBAL SEISMOLOGY Table 7.1 (Continued) LA–ICP–MS Zircon Radiometric Age Determination of Sample SA2 Spot 4r 4r 4c 5c 6r 6c 7r 7c 8 9c 9r 10 11r 11c 13r 13c 14r 14c 17r 17c 16r 16c 15 18r 18c 19r 19c 19c2 20c 21r 21c 22c 23r 23c 24r 24c 25c 26c 27c 27r 25r

Pb/238U

1 Sigma

0.00549 0.00531 0.24849 0.07338 0.00535 0.00636 0.00535 0.07421 0.04877 0.31196 0.00536 0.0054 0.00498 0.00529 0.00506 0.03814 0.00515 0.01641 0.00518 0.00521 0.00522 0.0069 0.04989 0.00514 0.01711 0.00518 0.39834 0.00733 0.00664 0.00523 0.00515 0.09822 0.00512 0.02898 0.00519 0.00533 0.01285 0.00514 0.11215 0.00802 0.0053

0.00013 0.00006 0.00281 0.00085 0.00007 0.00015 0.00008 0.00098 0.00066 0.00397 0.00008 0.00014 0.00008 0.00009 0.00006 0.00064 0.00008 0.00041 0.00008 0.0001 0.00009 0.0002 0.00068 0.00007 0.00033 0.00008 0.0046 0.00025 0.00019 0.00008 0.0001 0.00161 0.00009 0.0004 0.00007 0.00016 0.00024 0.00013 0.00154 0.00012 0.00008

206

Pb/235U

1 Sigma

0.03603 0.06376 5.67676 0.5959 0.0422 0.05865 0.03927 0.5976 0.39024 5.48444 0.03469 0.03622 0.03387 0.03433 0.03446 0.29061 0.03772 0.10631 0.03373 0.03349 0.03995 0.04507 0.45205 0.03337 0.13989 0.04499 8.99838 0.07591 0.06211 0.05561 0.03326 0.81209 0.03352 0.24435 0.0349 0.03569 0.12391 0.03305 1.01843 0.06377 0.03473

0.00433 0.00167 0.1167 0.01463 0.00168 0.00562 0.00192 0.02142 0.015 0.15315 0.00199 0.00518 0.00176 0.00272 0.00107 0.01724 0.0018 0.0134 0.00179 0.00314 0.00242 0.00624 0.01615 0.00153 0.01018 0.0022 0.20809 0.01031 0.00792 0.00253 0.00277 0.04568 0.00225 0.00989 0.00158 0.0052 0.0079 0.00431 0.03738 0.00298 0.00186

207

Pb/232Th

208

0.00231 0.009 0.07505 0.02339 0.00699 0.00298 0.00888 0.02337 0.0237 0.08643 0.0036 0.00165 0.0043 0.00184 0.00259 0.0214 0.00517 0.00057 0.00209 0.00138 0.01045 0.00154 0.02239 0.00244 0.01656 0.0048 0.10398 0.00257 0.00381 0.01686 0.00149 0.03081 0.00197 0.01276 0.00229 0.00147 0.02622 0.0016 0.0383 0.00567 0.00304

1 Sigma 0.00033 0.0005 0.00418 0.00161 0.00058 0.00155 0.00101 0.00175 0.00263 0.00649 0.00059 0.00038 0.00053 0.00022 0.00023 0.00168 0.00074 0.00086 0.00039 0.00051 0.00128 0.00067 0.00128 0.00039 0.00138 0.00058 0.00713 0.00418 0.00138 0.00161 0.00017 0.00279 0.00064 0.00124 0.00035 0.00048 0.00239 0.00036 0.0033 0.00062 0.00068

Pb/206Pb

207

0.04756 0.08705 0.16567 0.05889 0.05724 0.06682 0.05322 0.0584 0.05802 0.12749 0.04694 0.04867 0.0493 0.04707 0.04938 0.05526 0.05314 0.04699 0.04721 0.04666 0.05554 0.04737 0.06571 0.0471 0.05928 0.06293 0.16382 0.07515 0.06787 0.07705 0.04685 0.05996 0.04751 0.06116 0.0488 0.04855 0.06996 0.04665 0.06585 0.05764 0.0475

1 Sigma 0.0058 0.00233 0.00346 0.00147 0.00233 0.00656 0.00265 0.00213 0.00227 0.00361 0.00275 0.00706 0.00262 0.00378 0.00156 0.00336 0.00261 0.00602 0.00255 0.00445 0.00345 0.00668 0.0024 0.0022 0.00441 0.00316 0.00385 0.01048 0.00883 0.00362 0.00396 0.00344 0.00327 0.00254 0.00225 0.0072 0.00458 0.00618 0.00246 0.00275 0.00259

Age, Ma Pb/238U

206

1 Sigma

35.3 34.2 1430.7 456.5 34.4 40.9 34.4 461.5 307 1750.3 34.5 34.7 32 34 32.5 241.3 33.1 104.9 33.3 33.5 33.5 44.3 313.9 33 109.4 33.3 2161.4 47.1 42.6 33.7 33.1 604 32.9 184.1 33.3 34.3 82.3 33 685.2 51.5 34.1

0.85 0.41 14.51 5.11 0.47 0.98 0.51 5.88 4.08 19.5 0.54 0.93 0.49 0.6 0.4 3.99 0.49 2.59 0.51 0.66 0.58 1.27 4.16 0.46 2.09 0.52 21.21 1.6 1.2 0.53 0.62 9.45 0.58 2.53 0.47 1 1.51 0.84 8.9 0.77 0.52

Age, Ma Pb/238U

1 Sigma

25.7 30 64.7 33.7 28.3 32 56.3

0.46 0.4 1.07 0.79 0.37 0.52 15.23

LA–ICP–MS Zircon Radiometric Age Determination of Sample SA3 Spot 1r 2r 2c 3 4r 5 6

Pb/238U

1 Sigma

0.00399 0.00467 0.01009 0.00524 0.0044 0.00498 0.00877

0.00007 0.00006 0.00017 0.00012 0.00006 0.00008 0.00238

206

Pb/235U

1 Sigma

0.02804 0.03745 0.06978 0.03759 0.02945 0.0351 1.71385

0.00195 0.00133 0.00447 0.00433 0.00109 0.00191 0.3177

207

Pb/232Th

208

0.00753 0.00603 0.00495 0.00149 0.00334 0.00172 0.05345

1 Sigma 0.00171 0.00045 0.00116 0.00037 0.00065 0.00014 0.01196

207

Pb/206Pb

0.05099 0.05817 0.05016 0.05203 0.04855 0.05111 1.41638

1 Sigma 0.00363 0.00211 0.00327 0.00609 0.00182 0.00284 0.46489

206

(continued)

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  203 Table 7.1 (Continued) LA–ICP–MS Zircon Radiometric Age Determination of Sample SA3 Spot 7r 7c 8c 9r 9c 10r 10c 11r 11c1 11c2 12r 13r 14r 13c 14c

Pb/238U

1 Sigma

0.00523 0.0052 0.0053 0.00524 0.08745 0.00532 0.04819 0.00505 0.00546 0.00535 0.0052 0.00524 0.00522 0.38923 0.069

0.00007 0.00009 0.00014 0.00007 0.00113 0.00009 0.00087 0.00007 0.0001 0.00009 0.00007 0.00011 0.00009 0.00482 0.00124

206

Pb/235U

1 Sigma

0.03384 0.03694 0.03353 0.03909 0.70524 0.03475 0.39332 0.03256 0.03534 0.0426 0.0688 0.04296 0.11605 8.91671 0.57502

0.00152 0.00236 0.00417 0.00129 0.02266 0.00223 0.02527 0.00129 0.00282 0.00253 0.00245 0.00334 0.00469 0.24045 0.03834

207

Pb/232Th

1 Sigma

208

0.00197 0.00177 0.00157 0.00528 0.02735 0.00634 0.01945 0.00223 0.00291 0.00222 0.02809 0.01023 0.12272 0.10498 0.03841

0.00022 0.00014 0.00029 0.00041 0.0016 0.00104 0.0013 0.00037 0.00037 0.00034 0.00209 0.00159 0.0097 0.0079 0.00326

Pb/206Pb

207

0.04692 0.0515 0.04585 0.0541 0.05846 0.04735 0.05917 0.04673 0.04692 0.05771 0.09597 0.05944 0.16111 0.16612 0.06044

1 Sigma 0.00214 0.00336 0.0058 0.00183 0.00192 0.0031 0.00389 0.00189 0.00382 0.00352 0.00356 0.00474 0.00692 0.00467 0.00413

Age, Ma Pb/238U

206

1 Sigma

33.6 33.4 34.1 33.7 540.5 34.2 303.4 32.5 35.1 34.4 33.4 33.7 33.6 2119.2 430.1

0.47 0.59 0.87 0.43 6.69 0.56 5.36 0.43 0.66 0.59 0.47 0.68 0.58 22.35 7.51

1 Sigma

Age, Ma 206 Pb/238U

1 Sigma

0.0016 0.00324 0.0015 0.00207 0.0017 0.00144 0.00429 0.0019 0.00243 0.0018 0.00176

30 30.2 30.2 30.3 30 30.2 30.4 30.1 30.1 30 30.1

0.36 0.44 0.37 0.39 0.38 0.37 0.49 0.41 0.39 0.38 0.38

LA–ICP–MS Monazite Radiometric Age Determination of Sample SA3 Spot mz1 mz2 mz3 mz4 mz6 mz7 mz11 mz12 mz13 mz16 mz18

Pb/238U

1 Sigma

0.00467 0.0047 0.0047 0.00471 0.00467 0.0047 0.00473 0.00468 0.00468 0.00467 0.00468

0.00006 0.00007 0.00006 0.00006 0.00006 0.00006 0.00008 0.00006 0.00006 0.00006 0.00006

206

Pb/235U

1 Sigma

0.0387 0.05514 0.03528 0.04193 0.03588 0.03383 0.06512 0.03234 0.04881 0.0362 0.03616

0.00101 0.00204 0.00096 0.00132 0.00108 0.00092 0.0027 0.00121 0.00154 0.00115 0.00112

207

208

Pb/232Th

1 Sigma

0.00119 0.00114 0.00123 0.00126 0.00125 0.00126 0.00123 0.00127 0.00124 0.00126 0.00118

magmatic zircons are presented most often as outer zones with well‐developed oscillatory zoning with inclusion oriented along them (Fig. 7.14). ••Gneiss: A total of 45 spot analyses from 27 grains of zoned zircons were done (Table 7.1). Forty‐five spot analyses of distinct zircon zones of 27 grains are made (Table 7.1). The Concordia age of 33.80 ± 0.14 Ma repre­ sents the magmatic crystallization age. LA‐ICP‐MS age data for the xenocrystic zircons and cores define several clusters of the inherited component (Fig. 7.14). ••Garnet‐bearing two‐mica granite: Forty‐five spot analyses of distinct zircon zones of 30 grains were made (Table 7.1). The Concordia age of 32.15 ± 0.13 Ma is accepted as magmatic crystallization age. LA‐ICP‐MS age data for the xenocrystic zircons and cores define several clusters of the inherited component (Fig. 7.14). ••Two‐mica Granite: Twenty‐two spot analyses of distinct zircon zones of 14 grains are made (Table 7.1). The Concordia age of 33.36 ± 0.69 Ma represents the

0.00006 0.00006 0.00008 0.00008 0.00008 0.00009 0.00009 0.00009 0.0001 0.0001 0.0001

Pb/206Pb

207

0.06015 0.08514 0.05449 0.0646 0.05571 0.05222 0.09992 0.05014 0.07566 0.05624 0.05607

magmatic crystallization age. LA‐ICP‐MS age data for the xenocrystic zircons and cores define several clusters of the inherited component. Additionally, in the sample are dated monazite crystals revealing 206Pb/238U weighted average age of 30.13 ± 0.23 that can be interpreted as the age of hydrothermal alteration or reheating by new magmatic impulse (Fig.  7.14). It can be preferred as magmatic age because of being compatible with zircon ages. Some zircons 206Pb/238U give younger age, which is due to Pb loss. 7.5. AR/AR GEOCHRONOLOGY 7.5.1. Analytical Methods Ar/39Ar geochronology was carried out on biotite, muscovite, and K‐feldspar samples from the ÇPC at the Auburn University Noble Isotope Mass Analysis Laboratory (ANIMAL). The rock samples were selected 40

206Pb/238U

204  ACTIVE GLOBAL SEISMOLOGY

0.0054

0.0052

0.0050 Two-mica granite SA1/Zircon

Concordia Age = 32.15 ± 0.13 Ma

0.0048

MSWD (of concordance) = 1.4,

207Pb/235U

0.0046 0.022

0.030

0.034

0.038

0.042

0.046

0.050

206Pb/238U

0.0059

0.026

Probability (of concordance) = 0.23

0.0057 0.0055 0.0053

Gneiss

0.0051

SA2/Zircon

Concordia Age = 33.80 ± 0.14 Ma MSWD (of concordance) = 3.6,

0.0049

Probability (of concordance) = 0.058

207Pb/235U

0.0047 0.02

0.0055

0.04

0.05

206Pb/238U

0.0057

0.03

0.0053 0.0051

Two-mica granite SA3/Zircon

Concordia Age = 33.36 ± 0.69 Ma

0.0049 0.0047

MSWD (of concordance) = 10.6, Probability (of concordance) = 0.001

207Pb/235U

0.020

0.024

0.028

0.032

0.036

0.040

0.044

0.048

0.052

31.0 30.8 Pb206/U238

30.6 30.4 30.2 30.0 29.8 29.6 29.4 29.2

Two-mica granite SA3/Monazite

Mean = 30.13 ± 0.23 Ma [0.77%] MSWD = 0.097, probability = 1.000

29.0

Figure 7.14  Cathodoluminescence (CL), back‐scattered (BSE) images and U/Pb LA–ICP–MS Zircon and monazite ages of GGMC.

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  205

after petrographic analysis of thin sections from each sample. The selected samples were crushed and sieved for whole rock and biotite grains. The individual biotite (850–600 µm), muscovite, and K‐feldspar grains were handpicked under a binocular microscope to be generally free from visible inclusion of other phases and free from visible alterations. The selected samples were washed with deionized water in ultrasonic cleaner. Samples were located in aluminum disk with monitor FC‐2 (age = 28.02 Ma, Renne et  al. [1998]) along with CaF2 flux monitor. All samples and standards were irradiated in the USGS TRIGA reactor located at the Denver Federal Center, USA. The ANIMAL analytical lab is equipped with a low‐volume, high‐sensitivity 10 cm radius sector mass spectrometer and automated sample extraction system (50 W Syndrad CO2 laser). The analyses of these samples were conducted by laser incremental heating analysis (LIH) of approximately 100 grains of whole rock. The biotite, muscovite, and K‐feldspar samples were dated from single crystal total fusion (SCTF) and single crystal incremental heating method (SCIH). All statistical 40Ar/39Ar ages in this study (weighted means, plateau, or isochron) are quoted at the standard deviation in precision of measurement, whereas errors in individual measurements are quoted at one standard deviation. Data reduction was calculated through use of custom Microsoft Excel application. The plateau and correlation ages were calculated by isoplot [Ludwig, 2003]. Plateau in this study was defined as at least three or more contiguous increments containing more than 50% of the 39ArK in three or more contiguous steps with no resolvable slope among ages. 7.5.2. Results Single grain plateau ages were obtained by laser con­ trolled incremental heating to provide more detail on the outgassing behaviors of the samples. SCTF and SCIH ages were achieved from biotite, muscovite, and K‐feldspar samples of ÇG and GGMC. The results are given in Table 7.2, and Figures 7.15 and 7.16. ••Çataldağ granodiorite (ÇG): SCTF and SCIH bio­ tite ages of biotite‐granodiorite samples (OS‐55 and OS‐72) of ÇG yielded well‐developed plateau ages of 20.84 ± 0.13 Ma (MSWD = 0.4) and 21.16 ± 0.04 Ma (MSWD = 0.26), respectively; by total of 36 biotite grains (Fig. 7.15). ••Granite‐gneiss‐migmatite complex (GGMC): Muscovite of two‐mica granite (OS‐75 and OS‐105), gneiss (OS‐150), mylonitic granite (OS‐77), and biotite in migmatite (OS‐111) of GGMC have been dated. SCIH and SCTF muscovite age of two‐mica granite (OS‐75) has a plateau age of 21.38 ± 0.05 Ma (MSWD = 0.84). SCTF

muscovite age of other two‐mica granite (OS‐105) yields 20.81 ± 0.04 Ma with the highest MSWD of 1.4 (Fig. 7.16). Gneiss (OS‐150) has been dated on biotite and K‐feldspar. SCIH biotite and K‐feldspar ages of 20.80 ± 0.08 Ma (MSWD = 0.35) and 21.12 ± 0.06 Ma (MSWD = 0.60), respectively. Mylonitic granite (OS‐77) of ÇDFZ has a well‐­developed plateau age of 21.28 ± 0.09 Ma (MSWD = 0.34), which is very close to the age of two‐mica granite (OS‐75). Migmatite, sample OS‐111, has SCIH and SCTF biotite ages of 21.18 ± 0.07 Ma (MSWD = 0.58) and 21.11 ± 0.21 Ma, respectively. In summary, 40Ar/39Ar mica ages of ÇG and GGMC are similar and range from 21.3 to 20.8 Ma. 7.6. DISCUSSION 7.6.1. Emplacement of the Granite‐Gneiss‐ Migmatite Complex (GGMC) The granite‐gneiss‐migmatite complex exhibits a domal structure; at the core of dome, granitic gneiss and S‐type two‐mica granites are dominant and both are concordant with migmatite of amphibolite facies. Within the migma­ titic zones, discrete parts of melanosome, rich in biotite, sillimanite, cordierite, and garnet represent restites from partial melting. These data suggest that high‐grade metamorphic rocks (amphibolite to upper amphibolite P/T conditions) of metasediments of pelitic‐psammitic composition may be viewed as precursor of GGMC. The igneous textures and some microscale features, presence of primary garnet and muscovite, and also inherited zircon cores are the supporting evidence for an anatectic melt of crustal origin (Fig. 7.14). Granite with a typical igneous texture passing into the migmatite of amphibolite facies toward the peripheral zone and absence of contact metamorphism around the GGMC are regarded as supporting evidence for the granitic rocks of deep‐seated origin. The granitic magma that formed GGMC was possibly emplaced in middle crustal depths. The intrusion age of the S‐type, two‐mica granites, and the precursor of the migmatite, dates the migmatization and partial melting processes beneath western Anatolia. The zircon and monazite ages for the GGMC are 33.8–32.1 Ma and 30.1 Ma, respectively. The age obtained from the S‐type, two‐mica granite, gneiss, and migmatite falls into the same range, indicating that all the rock groups of GGMC were formed during the Eo‐Oligocene. This period may thus be evaluated as the age of the partial melting, migmatization, and deformation. Along the northern, eastern, and southeastern borders, the GGMC is bounded by a low‐angle normal fault zone

1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8

a b c d e f g h i j k l m n o p q r

15 15 15 15 15 15 15 15 15 15 15 15 15 15 15 15 15 15

t

2.90667 1.81656 2.14506 1.02731 2.53285 1.55998 1.65332 2.55158 1.72717 1.20647 1.80762 1.21186 0.80610 2.50317 1.97377 0.60162 5.17647 3.61627

± 0.00176 ± 0.00106 ± 0.00206 ± 0.00108 ± 0.00203 ± 0.00190 ± 0.00139 ± 0.00170 ± 0.00269 ± 0.00121 ± 0.00184 ± 0.00133 ± 0.00084 ± 0.00163 ± 0.00211 ± 0.00122 ± 0.00208 ± 0.00451

40 V 1.52446 0.95329 1.09413 0.53694 1.32176 0.81464 0.86929 1.34672 0.90697 0.62030 0.92486 0.63848 0.40325 1.31881 1.02312 0.30880 2.71954 1.80887

± 0.00148 ± 0.00105 ± 0.00152 ± 0.00053 ± 0.00143 ± 0.00093 ± 0.00148 ± 0.00099 ± 0.00143 ± 0.00083 ± 0.00083 ± 0.00077 ± 0.00101 ± 0.00111 ± 0.00101 ± 0.00046 ± 0.00139 ± 0.00227

39 V 0.02027 0.01234 0.01432 0.00680 0.01710 0.01064 0.01133 0.01740 0.01181 0.00778 0.01199 0.00841 0.00514 0.01700 0.01323 0.00392 0.03543 0.02394

± 0.00022 ± 0.00014 ± 0.00016 ± 0.00009 ± 0.00021 ± 0.00013 ± 0.00013 ± 0.00018 ± 0.00016 ± 0.00008 ± 0.00015 ± 0.00014 ± 0.00005 ± 0.00010 ± 0.00015 ± 0.00010 ± 0.00017 ± 0.00018

38 V

0.48 0.52 0.56 0.6 0.65 0.72 0.8 0.9 1 1.1 1.2

P

20 15 15 15 15 15 15 15 15 15 15

t

0.14441 0.15981 0.09554 0.79669 0.61497 0.59505 0.70009 0.72678 0.63063 0.61463 0.09427

± 0.00048 ± 0.00047 ± 0.00055 ± 0.00099 ± 0.00072 ± 0.00058 ± 0.00117 ± 0.00113 ± 0.00073 ± 0.00104 ± 0.00039

40 V 0.01581 0.03405 0.03535 0.39370 0.33555 0.31872 0.38371 0.39742 0.34690 0.33671 0.05017

± 0.00015 ± 0.00016 ± 0.00014 ± 0.00043 ± 0.00080 ± 0.00075 ± 0.00115 ± 0.00066 ± 0.00092 ± 0.00035 ± 0.00015

39 V 0.00028 0.00046 0.00044 0.00517 0.00408 0.00401 0.00488 0.00522 0.00437 0.00419 0.00059

± 0.00005 ± 0.00004 ± 0.00006 ± 0.00010 ± 0.00005 ± 0.00006 ± 0.00008 ± 0.00014 ± 0.00006 ± 0.00005 ± 0.00004

38 V 0.00057 0.00106 0.00037 0.00799 0.02148 0.09705 0.03498 0.00784 0.00948 0.03509 0.03005

± 0.00016 ± 0.00017 ± 0.00020 ± 0.00017 ± 0.00043 ± 0.00092 ± 0.00028 ± 0.00026 ± 0.00011 ± 0.00061 ± 0.00048

37 V

± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002

0.000363 0.000343 0.000109 0.000334 0.000070 0.000106 0.000038 0.000043 0.000022 0.000054 0.000015

± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00001 ± 0.00002 ± 0.00002 ± 0.00002

36 V

0.000616 0.000395 0.000611 0.000225 0.000597 0.000364 0.000360 0.000444 0.000350 0.000320 0.000551 0.000272 0.000296 0.000457 0.000484 0.000164 0.001052 0.001250

36 V

1.01E‐15 1.12E‐15 6.69E‐16 5.58E‐15 4.31E‐15 4.17E‐15 4.90E‐15 5.09E‐15 4.42E‐15 4.30E‐15 6.60E‐16

Moles 40Ar*

2.04E‐14 1.27E‐14 1.50E‐14 7.19E‐15 1.77E‐14 1.09E‐14 1.16E‐14 1.79E‐14 1.21E‐14 8.45E‐15 1.27E‐14 8.49E‐15 5.65E‐15 1.75E‐14 1.38E‐14 4.21E‐15 3.63E‐14 2.53E‐14

Moles 40Ar*

25.85% 36.54% 66.29% 87.70% 96.94% 96.08% 98.79% 98.34% 99.08% 97.89% 97.99%

%Rad

93.84% 93.65% 91.74% 93.59% 93.18% 93.17% 93.63% 95.01% 94.22% 92.21% 91.07% 93.42% 89.30% 94.71% 92.77% 91.98% 94.03% 90.07%

%Rad

2.3617 1.7151 1.7916 1.7746 1.7768 1.7943 1.8026 1.7984 1.8011 1.7870 1.8420

R

1.7892 1.7847 1.7987 1.7907 1.7856 1.7842 1.7807 1.8001 1.7943 1.7935 1.7800 1.7731 1.7853 1.7977 1.7896 1.7919 1.7899 1.8008

R

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R = 40Ar*/39Ar (Ar*: radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

1 2 3 4 5 6 7 8 9 10 11

Steps

± 0.00026 ± 0.00031 ± 0.00055 ± 0.00016 ± 0.00045 ± 0.00019 ± 0.00032 ± 0.00030 ± 0.00077 ± 0.00018 ± 0.00024 ± 0.00020 ± 0.00025 ± 0.00023 ± 0.00014 ± 0.00011 ± 0.00029 ± 0.00135

37 V 0.03442 0.01813 0.04295 0.00589 0.04361 0.01106 0.01223 0.04497 0.04430 0.00566 0.01750 0.00590 0.01672 0.03304 0.00450 0.00295 0.02401 0.12579

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐55 of ÇG (Mineral: Biotite)

P

Sample

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐55 of ÇG (Mineral: Biotite; N: 18)

Table 7.2  40Ar/39Ar Ages Obtained from the CPC

± 0.05 ± 0.07 ± 0.06 ± 0.13 ± 0.06 ± 0.08 ± 0.09 ± 0.05 ± 0.09 ± 0.11 ± 0.08 ± 0.10 ± 0.16 ± 0.06 ± 0.08 ± 0.34 ± 0.03 ± 0.06

27.43 19.96 20.85 20.65 20.67 20.88 20.97 20.92 20.96 20.79 21.43

%‐sd

0.22% 0.34% 0.29% 0.62% 0.30% 0.37% 0.46% 0.26% 0.43% 0.52% 0.36% 0.49% 0.79% 0.29% 0.40% 1.63% 0.14% 0.28%

%‐sd

± 3.86 14.06% ± 1.96 9.81% ± 1.78 8.53% ± 0.15 0.71% ± 0.18 0.88% ± 0.20 0.97% ± 0.16 0.74% ± 0.13 0.64% ± 0.20 0.95% ± 0.19 0.90% ± 1.20 5.61%

Age (Ma)

20.82 20.77 20.93 20.84 20.78 20.76 20.72 20.94 20.88 20.87 20.71 20.63 20.77 20.92 20.82 20.85 20.83 20.95

Age (Ma)

1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6

a b c d e f g h i j k l m n o p q r

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

6.94779 3.67676 6.12546 3.26473 6.21500 3.93429 7.06005 5.72659 3.34128 6.75694 3.45974 5.48701 6.27494 5.60479 2.62769 8.17385 5.82363 6.35603

± 0.00416 ± 0.00178 ± 0.00465 ± 0.00366 ± 0.00400 ± 0.00372 ± 0.00593 ± 0.00219 ± 0.00415 ± 0.00247 ± 0.00211 ± 0.00437 ± 0.00597 ± 0.00537 ± 0.00294 0.00429 0.00319 0.00536

40 V 3.50444 1.82032 2.99501 1.63483 3.07517 1.94234 3.55612 2.86471 1.66687 3.34976 1.67867 2.70378 3.16015 2.82942 1.32577 4.10257 2.90998 3.13932

± 0.00288 ± 0.00161 ± 0.00306 ± 0.00278 ± 0.00241 ± 0.00266 ± 0.00373 ± 0.00195 ± 0.00196 ± 0.00213 ± 0.00291 ± 0.01173 ± 0.00178 ± 0.00223 ± 0.00107 ± 0.00464 ± 0.00191 ± 0.00296

39 V 0.05058 0.02675 0.04404 0.02354 0.04318 0.02770 0.05115 0.04135 0.02403 0.04912 0.02480 0.03996 0.04495 0.04173 0.01852 0.05929 0.04269 0.04467

± 0.00021 ± 0.00016 ± 0.00027 ± 0.00015 ± 0.00023 ± 0.00021 ± 0.00037 ± 0.00029 ± 0.00012 ± 0.00015 ± 0.00010 ± 0.00026 ± 0.00029 ± 0.00031 ± 0.00015 ± 0.00041 ± 0.00034 ± 0.00018

38 V

0.35 0.4 0.44 0.48 0.52 0.56 0.6 0.65 0.72 0.8 0.95 1.1 1.3 1.6 1.8

P

20 15 10 10 10 10 10 10 10 10 10 10 10 10 10

t

0.29405 0.29914 0.24816 0.37289 0.39346 0.55585 0.76335 1.05005 1.31896 1.84699 2.77645 2.03776 1.32459 0.01196 0.03097

± 0.00061 ± 0.00036 ± 0.00044 ± 0.00066 ± 0.00048 ± 0.00095 ± 0.00140 ± 0.00221 ± 0.00164 ± 0.00168 ± 0.00263 ± 0.00243 ± 0.00199 ± 0.00024 ± 0.00027

40 V 0.11661 0.13302 0.11946 0.18477 0.19704 0.28166 0.38544 0.52992 0.66952 0.93553 1.40855 1.02962 0.66832 0.00554 0.01461

± 0.00046 ± 0.00035 ± 0.00046 ± 0.00073 ± 0.00048 ± 0.00069 ± 0.00080 ± 0.00073 ± 0.00091 ± 0.00160 ± 0.00195 ± 0.00164 ± 0.00121 ± 0.00022 ± 0.00013

39 V 0.00173 0.00197 0.00183 0.00274 0.00283 0.00406 0.00544 0.00767 0.00955 0.01346 0.02070 0.01521 0.00969 0.00005 0.00017

± 0.00005 ± 0.00008 ± 0.00005 ± 0.00005 ± 0.00007 ± 0.00007 ± 0.00009 ± 0.00007 ± 0.00010 ± 0.00008 ± 0.00020 ± 0.00020 ± 0.00008 ± 0.00006 ± 0.00005

38 V 0.00332 0.00253 0.00195 0.00193 0.00195 0.00264 0.00361 0.00517 0.00867 0.02261 0.08949 0.11136 0.10050 0.00165 0.00712

± 0.00017 ± 0.00015 ± 0.00023 ± 0.00017 ± 0.00015 ± 0.00012 ± 0.00016 ± 0.00027 ± 0.00015 ± 0.00039 ± 0.00039 ± 0.00121 ± 0.00057 ± 0.00021 ± 0.00023

37 V

0.000189 0.000132 0.000021 0.000032 0.000026 –0.000001 0.000025 0.000032 0.000023 0.000040 0.000080 0.000118 0.000095 –0.000039 0.000003

Moles 40Ar*

4.87E‐14 2.57E‐14 4.29E‐14 2.29E‐14 4.35E‐14 2.76E‐14 4.94E‐14 4.01E‐14 2.34E‐14 4.73E‐14 2.42E‐14 3.84E‐14 4.39E‐14 3.93E‐14 1.84E‐14 5.72E‐14 4.08E‐14 4.45E‐14

Moles 40Ar*

%Rad

98.54% 96.14% 95.73% 97.67% 96.45% 96.60% 98.03% 97.67% 97.54% 96.87% 94.56% 95.96% 98.46% 98.49% 98.83% 97.67% 97.43% 96.37%

%Rad

R

1.9536 1.9420 1.9580 1.9506 1.9495 1.9567 1.9462 1.9525 1.9553 1.9540 1.9491 1.9475 1.9551 1.9511 1.9589 1.9460 1.9500 1.9512

R

± 0.03 ± 0.05 ± 0.04 ± 0.07 ± 0.04 ± 0.06 ± 0.04 ± 0.03 ± 0.06 ± 0.03 ± 0.07 ± 0.10 ± 0.03 ± 0.04 ± 0.06 ± 0.03 ± 0.04 ± 0.04

Age (Ma)

21.09 20.97 21.14 21.06 21.05 21.12 21.01 21.08 21.11 21.09 21.04 21.02 21.11 21.06 21.15 21.01 21.05 21.06

Age (Ma)

%‐sd

0.16% 0.22% 0.20% 0.35% 0.18% 0.28% 0.18% 0.17% 0.28% 0.16% 0.31% 0.49% 0.16% 0.19% 0.30% 0.16% 0.18% 0.19%

%‐sd

(continued)

2.06E‐15 81.12% 2.0455 22.08 ± 0.75 3.40% 2.09E‐15 87.01% 1.9567 21.12 ± 0.63 3.00% 1.74E‐15 97.62% 2.0279 21.89 ± 0.76 3.47% 2.61E‐15 97.55% 1.9687 21.25 ± 0.45 2.11% 2.76E‐15 98.09% 1.9588 21.15 ± 0.39 1.84% 3.89E‐15 100.10% 1.9743 21.31 ± 0.32 1.49% 5.35E‐15 99.08% 1.9623 21.18 ± 0.21 1.01% 7.35E‐15 99.13% 1.9643 21.20 ± 0.15 0.73% 9.24E‐15 99.54% 1.9609 21.17 ± 0.12 0.59% 1.29E‐14 99.47% 1.9639 21.20 ± 0.11 0.51% 1.94E‐14 99.43% 1.9600 21.16 ± 0.07 0.35% 1.43E‐14 98.78% 1.9551 21.11 ± 0.09 0.45% 9.28E‐15 98.55% 1.9534 21.09 ± 0.15 0.70% 8.38E‐17 198.05% 2.1860 23.58 ± 14.07 59.66% 2.17E‐16 98.88% 2.0968 22.63 ± 5.17 22.86%

± 0.00003 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003

± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00002

36 V

0.000368 0.000509 0.000943 0.000282 0.000793 0.000473 0.000531 0.000493 0.000293 0.000748 0.000681 0.000794 0.000375 0.000339 0.000110 0.000701 0.000543 0.000810

36 V

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R  = 40Ar*/39Ar (Ar* = radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15

Steps

± 0.00092 ± 0.00061 ± 0.00061 ± 0.00054 ± 0.00104 ± 0.00046 ± 0.00084 ± 0.00087 ± 0.00021 ± 0.00062 ± 0.00068 ± 0.00145 ± 0.00135 ± 0.00133 ± 0.00022 ± 0.00150 ± 0.00093 ± 0.00082

37 V 0.08009 0.09655 0.19326 0.08350 0.15627 0.06673 0.19785 0.13743 0.04989 0.10587 0.14964 0.14656 0.15755 0.17391 0.02145 0.18749 0.12330 0.09767

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐72 of ÇG (Mineral: Biotite)

P

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐72 of ÇG (Mineral: Biotite; N: 17)

1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6

a b c d e f g h i j k l m n o p q r

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

9.01388 5.45546 5.10721 3.84697 3.10267 5.06933 5.72735 3.56234 2.46505 3.69409 2.82005 3.50115 2.96882 4.03798 3.56585 3.77374 6.09154 6.56079

± 0.00595 ± 0.00537 ± 0.00548 ± 0.00211 ± 0.00347 ± 0.00326 ± 0.00310 ± 0.00454 ± 0.00311 ± 0.00361 ± 0.00253 ± 0.00304 ± 0.00160 ± 0.00303 ± 0.00316 0.00540 0.00972 0.00718

40 V 4.29703 2.52139 2.47217 1.80233 1.46497 2.29341 2.57725 1.66191 1.08977 1.78427 1.32331 1.68858 1.43433 1.95944 1.53130 1.72616 2.69042 3.15511

± 0.00617 ± 0.00297 ± 0.00252 ± 0.00274 ± 0.00117 ± 0.00216 ± 0.00232 ± 0.00232 ± 0.00145 ± 0.00164 ± 0.00184 ± 0.00106 ± 0.00154 ± 0.00248 ± 0.00256 ± 0.00112 ± 0.00263 ± 0.00264

39 V 0.06203 0.03684 0.03540 0.02640 0.02088 0.03359 0.03789 0.02399 0.01569 0.02495 0.01859 0.02385 0.02000 0.02802 0.02178 0.02466 0.03809 0.04518

± 0.00047 ± 0.00036 ± 0.00023 ± 0.00033 ± 0.00011 ± 0.00020 ± 0.00027 ± 0.00017 ± 0.00012 ± 0.00017 ± 0.00011 ± 0.00011 ± 0.00013 ± 0.00029 ± 0.00009 ± 0.00015 ± 0.00020 ± 0.00034

38 V 0.04541 0.11439 0.03973 0.06079 0.05599 0.02200 0.01539 0.03326 0.02244 0.02889 0.00179 0.02131 0.01469 0.01055 0.04261 0.01608 0.05360 0.03909

0.35 0.4 0.44 0.48 0.52 0.56 0.6 0.72 0.8 0.9 1 1.3 1.6

P

20 15 10 10 10 10 10 10 10 10 10 10 10

t

1.49656 1.30725 0.77988 0.75657 0.91699 1.14168 0.55016 0.04283 0.04435 0.04379 0.08244 0.01108 0.00217

± 0.00133 ± 0.00104 ± 0.00160 ± 0.00108 ± 0.00161 ± 0.00075 ± 0.00098 ± 0.00050 ± 0.00051 ± 0.00030 ± 0.00041 ± 0.00026 ± 0.00021

40 V 0.55581 0.66178 0.39620 0.37913 0.45987 0.57354 0.27486 0.02214 0.02104 0.02283 0.04035 0.00440 0.00066

± 0.00008 ± 0.00009 ± 0.00006 ± 0.00009 ± 0.00009 ± 0.00014 ± 0.00009 ± 0.00005 ± 0.00004 ± 0.00006 ± 0.00005 ± 0.00005 ± 0.00004

38 V

± 0.00134 0.00806 ± 0.00110 0.00931 ± 0.00067 0.00561 ± 0.00081 0.00537 ± 0.00068 0.00656 ± 0.00085 0.00841 ± 0.00045 0.00383 ± 0.00010 0.00035 ± 0.00015 0.00032 ± 0.00011 0.00029 ± 0.00031 0.00057 ± 0.00014 0.00002 ± 0.00012 –0.00012

39 V 0.00086 0.00027 0.00020 0.00005 0.00094 0.00492 0.02122 0.01003 0.01094 0.00572 0.02165 0.00514 0.00104

± 0.00016 ± 0.00020 ± 0.00012 ± 0.00021 ± 0.00020 ± 0.00026 ± 0.00030 ± 0.00024 ± 0.00034 ± 0.00010 ± 0.00022 ± 0.00017 ± 0.00022

37 V

± 0.00003 ± 0.00004 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00004 ± 0.00005 ± 0.00002

36 V

1.05E‐14 9.16E‐15 5.46E‐15 5.30E‐15 6.42E‐15 8.00E‐15 3.85E‐15 3.00E‐16 3.11E‐16 3.07E‐16 5.77E‐16 7.76E‐17 1.52E‐17

Moles 40Ar*

72.56% 99.57% 98.06% 98.80% 97.98% 99.02% 99.01% 104.30% 85.74% 139.20% 126.86% 425.47% 998.47%

R

%‐sd

0.20% 0.25% 0.23% 0.29% 0.34% 0.26% 0.20% 0.30% 0.44% 0.26% 0.36% 0.28% 0.35% 0.23% 0.39% 0.33% 0.28% 0.21%

%‐sd

± 0.19 0.92% ± 0.21 0.99% ± 0.24 1.16% ± 0.22 1.05% ± 0.18 0.88% ± 0.16 0.75% ± 0.26 1.22% ± 3.62 17.30% ± 3.48 18.13% ± 3.44 16.70% ± 3.37 15.20% ± 33.12 119.31% ± 125.04 347.01%

Age (Ma)

± 0.04 ± 0.05 ± 0.05 ± 0.06 ± 0.07 ± 0.05 ± 0.04 ± 0.06 ± 0.09 ± 0.05 ± 0.08 ± 0.06 ± 0.07 ± 0.05 ± 0.08 ± 0.07 ± 0.06 ± 0.04

Age (Ma) 20.72 20.63 20.82 20.73 20.67 20.76 20.78 20.85 20.74 20.76 20.65 20.61 20.78 20.75 20.74 20.78 20.75 20.76

20.72 20.86 20.47 20.91 20.72 20.90 21.02 20.95 19.19 20.58 22.17 27.76 36.04

1.9537 1.9452 1.9632 1.9551 1.9493 1.9577 1.9593 1.9659 1.9554 1.9573 1.9474 1.9434 1.9592 1.9563 1.9558 1.9596 1.9565 1.9578

R

1.9537 1.9669 1.9301 1.9716 1.9538 1.9711 1.9819 1.9755 1.8086 1.9406 2.0915 2.6226 3.4126

93.13% 89.90% 95.03% 91.59% 92.04% 88.57% 88.17% 91.71% 86.45% 94.54% 91.38% 93.73% 94.65% 94.93% 83.99% 89.63% 86.41% 94.15%

%Rad

%Rad

6.31E‐14 3.82E‐14 3.58E‐14 2.69E‐14 2.17E‐14 3.55E‐14 4.01E‐14 2.49E‐14 1.73E‐14 2.59E‐14 1.97E‐14 2.45E‐14 2.08E‐14 2.83E‐14 2.50E‐14 2.64E‐14 4.27E‐14 4.59E‐14

Moles 40Ar*

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad: % radiogenic argon; R = 40Ar*/39Ar (Ar*: radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

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

Steps

± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003

36 V 0.002109 0.001899 0.000871 0.001113 0.000853 0.001968 0.002298 0.001009 0.001137 0.000692 0.000823 0.000750 0.000542 0.000696 0.001945 0.001329 0.002818 0.001310

0.001390 0.000019 0.000051 0.000031 0.000063 0.000039 0.000025 –0.000003 0.000025 –0.000056 –0.000068 –0.000120 –0.000066

± 0.00056 ± 0.00075 ± 0.00041 ± 0.00095 ± 0.00099 ± 0.00025 ± 0.00049 ± 0.00029 ± 0.00020 ± 0.00032 ± 0.00023 ± 0.00053 ± 0.00029 ± 0.00020 ± 0.00027 ± 0.00025 ± 0.00038 ± 0.00044

37 V

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐150 of GGMC (Mineral: Biotite)

P

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐150 of GGMC (Mineral: Biotite; N: 18)

Table 7.2 (Continued)

1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6

P

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

20.44546 9.14790 17.07355 19.97864 5.32071 12.91263 13.87396 13.61879 14.12572 13.77383 14.52831 12.59796 10.40971 12.20579 3.71236

± 0.03561 ± 0.00570 ± 0.01162 ± 0.01854 ± 0.00482 ± 0.00943 ± 0.00832 ± 0.01158 ± 0.01164 ± 0.00695 ± 0.00881 ± 0.00706 ± 0.00570 ± 0.01333 ± 0.00276

40 V 9.29526 4.46746 7.87247 8.28891 2.41077 6.10857 6.32197 5.39708 6.85951 6.11066 6.98691 5.84731 4.94959 5.92447 1.53345

± 0.02260 ± 0.00684 ± 0.00842 ± 0.00784 ± 0.00263 ± 0.00492 ± 0.00718 ± 0.00414 ± 0.00622 ± 0.00403 ± 0.00664 ± 0.00619 ± 0.00426 ± 0.00389 ± 0.00197

39 V 0.12567 0.06008 0.10578 0.11244 0.03265 0.08110 0.08575 0.07212 0.09143 0.08188 0.09400 0.07892 0.06703 0.07954 0.02077

± 0.00057 ± 0.00025 ± 0.00045 ± 0.00058 ± 0.00021 ± 0.00033 ± 0.00051 ± 0.00022 ± 0.00035 ± 0.00019 ± 0.00069 ± 0.00029 ± 0.00056 ± 0.00037 ± 0.00023

38 V 0.04966 0.05580 0.03420 0.08412 0.40278 0.04161 0.08086 0.13657 0.02996 0.37241 0.04211 0.04478 0.03162 0.03953 0.26772

± 0.00028 ± 0.00027 ± 0.00020 ± 0.00036 ± 0.00142 ± 0.00031 ± 0.00030 ± 0.00032 ± 0.00011 ± 0.00103 ± 0.00016 ± 0.00027 ± 0.00025 ± 0.00047 ± 0.00138

37 V

20 20 20 20 20 20 20 20 20 20 20 20

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

0.39534 1.18624 3.23475 4.36492 5.25007 0.73171 1.77145 3.06491 1.28474 0.44647 0.06505 0.03996

± 0.00052 ± 0.00164 ± 0.00213 ± 0.00295 ± 0.00287 ± 0.00142 ± 0.00164 ± 0.00184 ± 0.00151 ± 0.00068 ± 0.00039 ± 0.00039

40 V 0.10752 0.45980 1.58336 2.11065 2.54508 0.36420 0.88791 1.53843 0.64426 0.22517 0.03127 0.02016

± 0.00037 ± 0.00085 ± 0.00185 ± 0.00163 ± 0.00240 ± 0.00064 ± 0.00166 ± 0.00188 ± 0.00178 ± 0.00064 ± 0.00014 ± 0.00017

39 V 0.00156 0.00612 0.02166 0.02794 0.03416 0.00497 0.01188 0.02007 0.00839 0.00295 0.00031 0.00017

37 V

36 V

Moles 40Ar*

2.15E‐14 9.00E‐15 3.13E‐15 4.56E‐16 2.80E‐16

%Rad

93.03% 96.92% 92.69% 88.56% 91.11% 97.23% 91.86% 82.54% 95.60% 90.24% 97.00% 93.02% 95.49% 96.65% 91.34%

%Rad

67.64% 84.91% 96.67% 96.99% 96.89% 98.56% 98.50% 98.76% 98.79% 101.77% 106.07% 126.56%

1.43E‐13 6.41E‐14 1.20E‐13 1.40E‐13 3.73E‐14 9.04E‐14 9.72E‐14 9.54E‐14 9.89E‐14 9.65E‐14 1.02E‐13 8.82E‐14 7.29E‐14 8.55E‐14 2.60E‐14

Moles 40Ar*

2.77E‐15 8.31E‐15 2.27E‐14 3.06E‐14 3.68E‐14 5.12E‐15

± 0.00016 ± 0.00007 ± 0.00016 ± 0.00018 ± 0.00012 ± 0.00014 ± 0.00016 ± 0.00038 ± 0.00027 ± 0.00021 ± 0.00010 ± 0.00018 ± 0.00012 ± 0.00010 ± 0.00010

± 0.00007 0.00115 ± 0.00014 0.000433 ± 0.00008 ± 0.00008 0.00310 ± 0.00017 0.000607 ± 0.00008 ± 0.00020 0.00617 ± 0.00014 0.000367 ± 0.00008 ± 0.00015 0.00648 ± 0.00020 0.000446 ± 0.00009 ± 0.00035 0.00554 ± 0.00015 0.000554 ± 0.00007 ± 0.00008 0.00050 ± 0.00008 0.000036 ± 0.00006 ± 0.00010 0.00092 ± 0.00019 0.000090 ± 0.00007 ± 0.00011 0.00219 ± 0.00011 0.000129 ± 0.00018 ± 0.00007 0.00074 ± 0.00007 0.000053 ± 0.00006 ± 0.00007 –0.00017 ± 0.00017 ‐0.000027 ± 0.00008 ± 0.00004 –0.00062 ± 0.00013 ‐0.000013 ± 0.00009 ± 0.00005 –0.00041 ± 0.00011 ‐0.000036 ± 0.00008

38 V

0.004836 0.000970 0.004237 0.007757 0.001722 0.001224 0.003845 0.008089 0.002112 0.004662 0.001485 0.002990 0.001600 0.001394 0.001168

36 V

2.4873 2.1906 1.9749 2.0059 1.9987 1.9802 1.9651 1.9676 1.9700 1.9828 2.0784 1.9797

R

2.0463 1.9846 2.0101 2.1347 2.0110 2.0553 2.0160 2.0827 1.9687 2.0341 2.0171 2.0041 2.0082 1.9913 2.2117

R

26.34 23.21 20.94 21.27 21.19 21.00 20.84 20.86 20.89 21.02 22.03 20.99

%‐sd

0.41% 0.28% 0.32% 0.33% 0.72% 0.35% 0.39% 1.01% 0.61% 0.50% 0.24% 0.48% 0.38% 0.29% 0.87%

%‐sd

(continued)

± 2.45 9.31% ± 0.56 2.40% ± 0.16 0.77% ± 0.14 0.64% ± 0.09 0.42% ± 0.55 2.63% ± 0.25 1.19% ± 0.37 1.78% ± 0.32 1.54% ± 1.09 5.20% ± 9.27 42.08% ± 11.82 56.28%

Age (Ma)

± 0.09 ± 0.06 ± 0.07 ± 0.07 ± 0.15 ± 0.08 ± 0.08 ± 0.22 ± 0.13 ± 0.11 ± 0.05 ± 0.10 ± 0.08 ± 0.06 ± 0.20

Age (Ma) 21.69 21.04 21.31 22.63 21.32 21.79 21.37 22.08 20.88 21.57 21.39 21.25 21.29 21.11 23.44

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R = 40Ar*/39Ar (Ar* = radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

0.35 0.4 0.45 0.475 0.5 0.55 0.6 0.7 0.8 0.9 1.1 1.3

t

Steps P

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐150 of GGMC (Mineral: K‐Feldspar)

* Rejected data.

a b c d* e f g h i j k l n o p*

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐150 of GMMC (Mineral: K‐Feldspar; N: 15)

1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6

a b c d e f g h i j k l m n o p q r s

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

4.28314 1.10437 2.49371 1.00582 1.57782 2.94448 2.75529 1.48780 2.09295 1.27339 1.19081 1.07738 1.88780 1.94377 2.71035 1.52531 1.24562 0.86556 0.99449

± 0.00569 ± 0.00093 ± 0.00251 ± 0.00112 ± 0.00266 ± 0.00282 ± 0.00227 ± 0.00084 ± 0.00245 ± 0.00103 ± 0.00151 ± 0.00153 ± 0.00305 ± 0.00148 ± 0.00216 0.00156 0.00171 0.00125 0.00170

40 V 2.05257 0.52747 1.16384 0.47521 0.66001 1.26576 1.34073 0.69696 0.99631 0.59956 0.55713 0.51543 0.87428 0.90075 1.23907 0.66761 0.56796 0.39027 0.43800

± 0.00327 ± 0.00063 ± 0.00132 ± 0.00113 ± 0.00129 ± 0.00159 ± 0.00189 ± 0.00095 ± 0.00134 ± 0.00104 ± 0.00122 ± 0.00091 ± 0.00093 ± 0.00156 ± 0.00128 ± 0.00144 ± 0.00085 ± 0.00103 ± 0.00108

39 V 0.02700 0.00680 0.01525 0.00617 0.00880 0.01679 0.01821 0.00900 0.01298 0.00827 0.00740 0.00678 0.01188 0.01183 0.01636 0.00914 0.00748 0.00527 0.00581

± 0.00017 ± 0.00011 ± 0.00012 ± 0.00004 ± 0.00008 ± 0.00009 ± 0.00018 ± 0.00009 ± 0.00011 ± 0.00011 ± 0.00006 ± 0.00003 ± 0.00017 ± 0.00007 ± 0.00010 ± 0.00014 ± 0.00006 ± 0.00008 ± 0.00004

38 V 0.00239 0.00097 0.00132 0.00066 0.00341 0.00390 0.00286 0.00043 0.00249 0.00070 0.00054 0.00097 0.00665 0.00175 0.00169 0.00061 0.00025 0.00110 0.00035

0.35 0.4 0.44 0.48 0.52 0.56 0.6 0.72 0.8 0.95 1.1 1.3

P

15 15 10 10 10 10 10 10 10 10 10 10

t

0.03703 0.28766 0.25558 0.61380 0.92411 1.03457 1.26843 0.39980 0.07567 0.02278 ‐0.00212 0.00295

± 0.00038 ± 0.00035 ± 0.00049 ± 0.00109 ± 0.00064 ± 0.00217 ± 0.00212 ± 0.00053 ± 0.00064 ± 0.00036 ± 0.00026 ± 0.00026

40 V 0.00530 0.05182 0.11840 0.30229 0.45548 0.51641 0.62600 0.20355 0.03826 0.01214 0.00014 0.00065

38 V

± 0.00013 0.00006 ± 0.00004 ± 0.00029 0.00073 ± 0.00005 ± 0.00048 0.00145 ± 0.00005 ± 0.00055 0.00394 ± 0.00008 ± 0.00068 0.00592 ± 0.00008 ± 0.00096 0.00667 ± 0.00010 ± 0.00120 0.00814 ± 0.00007 ± 0.00064 0.00259 ± 0.00007 ± 0.00014 0.00049 ± 0.00004 ± 0.00019 0.00010 ± 0.00004 ± 0.00009 –0.00007 ± 0.00005 ± 0.00010 –0.00007 ± 0.00005

39 V 0.00448 0.00346 0.00091 0.00140 0.00170 0.00132 0.00209 0.00033 0.00084 0.00020 0.00007 0.00020

± 0.00022 ± 0.00021 ± 0.00014 ± 0.00016 ± 0.00019 ± 0.00016 ± 0.00017 ± 0.00019 ± 0.00032 ± 0.00014 ± 0.00016 ± 0.00013

37 V

Moles 40Ar*

%Rad

R

95.04% 95.25% 93.09% 94.09% 83.35% 85.45% 96.61% 93.28% 94.84% 94.82% 93.41% 94.53% 92.32% 92.16% 90.08% 88.06% 91.17% 88.20% 88.85%

%Rad 1.9832 1.9942 1.9945 1.9915 1.9926 1.9878 1.9853 1.9912 1.9922 2.0138 1.9966 1.9760 1.9935 1.9887 1.9704 2.0120 1.9995 1.9561 2.0173

R

± 0.06 ± 0.17 ± 0.08 ± 0.17 ± 0.13 ± 0.08 ± 0.07 ± 0.11 ± 0.08 ± 0.14 ± 0.16 ± 0.16 ± 0.11 ± 0.10 ± 0.08 ± 0.14 ± 0.13 ± 0.19 ± 0.19

Age (Ma)

21.03 21.14 21.15 21.12 21.13 21.08 21.05 21.11 21.12 21.35 21.17 20.95 21.14 21.09 20.89 21.33 21.20 20.74 21.39

Age (Ma)

± 0.00003 2.59E‐16 64.64% 4.5176 47.55 ± 16.57 ± 0.00003 2.01E‐15 45.77% 2.5408 26.90 ± 1.81 ± 0.00003 1.79E‐15 98.98% 2.1365 22.64 ± 0.67 ± 0.00003 4.30E‐15 99.56% 2.0215 21.43 ± 0.29 ± 0.00003 6.47E‐15 99.83% 2.0254 21.47 ± 0.20 ± 0.00003 7.25E‐15 101.16% 2.0036 21.24 ± 0.18 ± 0.00003 8.88E‐15 98.74% 2.0007 21.21 ± 0.14 ± 0.00003 2.80E‐15 102.93% 1.9643 20.83 ± 0.42 ± 0.00003 5.30E‐16 113.60% 1.9796 20.99 ± 2.40 ± 0.00003 1.60E‐16 160.15% 1.8778 19.92 ± 6.77 ± 0.00003 –1.49E‐17 –620.22% –14.8078 –165.28 – ± 958.69 ± 0.00004 2.07E‐17 1321.87% 4.5901 48.30 ± 220.06

36 V

3.00E‐14 7.73E‐15 1.75E‐14 7.04E‐15 1.10E‐14 2.06E‐14 1.93E‐14 1.04E‐14 1.47E‐14 8.92E‐15 8.34E‐15 7.55E‐15 1.32E‐14 1.36E‐14 1.90E‐14 1.07E‐14 8.72E‐15 6.06E‐15 6.96E‐15

Moles 40Ar*

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R = 40Ar*/39Ar (Ar* = radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

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

Steps

± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00003

36 V 0.000720 0.000178 0.000584 0.000201 0.000890 0.001451 0.000317 0.000339 0.000366 0.000224 0.000266 0.000200 0.000493 0.000516 0.000910 0.000616 0.000372 0.000346 0.000375

0.000046 0.000529 0.000009 0.000010 0.000006 –0.000040 0.000055 –0.000040 –0.000035 –0.000046 –0.000052 –0.000122

± 0.00029 ± 0.00020 ± 0.00023 ± 0.00015 ± 0.00012 ± 0.00017 ± 0.00018 ± 0.00012 ± 0.00015 ± 0.00015 ± 0.00015 ± 0.00013 ± 0.00012 ± 0.00015 ± 0.00015 ± 0.00019 ± 0.00022 ± 0.00010 ± 0.00014

37 V

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐77 of GGMC (Mineral: Muscovite)

P

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐77 of GGMC (Mineral: Muscovite; N: 18)

Table 7.2 (Continued)

34.84% 6.74% 2.96% 1.36% 0.94% 0.82% 0.67% 2.03% 11.44% 33.98% 580.03% 455.57%

%‐sd

0.30% 0.79% 0.40% 0.81% 0.64% 0.39% 0.32% 0.53% 0.39% 0.64% 0.77% 0.76% 0.50% 0.49% 0.38% 0.64% 0.63% 0.93% 0.91%

%‐sd

1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6

a b c d e f g h i j k l m n o p q r

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

1.39700 2.68141 1.99662 2.58414 1.92219 1.63477 1.88743 2.63330 2.62018 3.02661 1.62109 1.04785 1.88345 2.14930 2.38225 0.79850 2.14607 4.87437

± 0.00184 ± 0.00488 ± 0.00226 ± 0.00282 ± 0.00269 ± 0.00189 ± 0.00145 ± 0.00716 ± 0.00339 ± 0.00288 ± 0.00135 ± 0.00181 ± 0.00173 ± 0.00097 ± 0.00154 0.00069 0.00311 0.00350

40 V 0.67380 1.19844 0.91092 1.26832 0.88905 0.73270 0.92142 1.12259 1.05856 1.31435 0.78684 0.48546 0.92440 0.87555 1.11774 0.38185 1.02023 2.45866

± 0.00165 ± 0.00279 ± 0.00152 ± 0.00175 ± 0.00120 ± 0.00121 ± 0.00194 ± 0.00296 ± 0.00128 ± 0.00134 ± 0.00132 ± 0.00067 ± 0.00140 ± 0.00189 ± 0.00148 ± 0.00109 ± 0.00121 ± 0.00204

39 V 0.00910 0.01646 0.01233 0.01751 0.01202 0.01001 0.01226 0.01422 0.01441 0.01807 0.01050 0.00658 0.01253 0.01186 0.01497 0.00527 0.01361 0.03286

± 0.00007 ± 0.00013 ± 0.00011 ± 0.00015 ± 0.00009 ± 0.00010 ± 0.00010 ± 0.00134 ± 0.00011 ± 0.00014 ± 0.00010 ± 0.00009 ± 0.00010 ± 0.00008 ± 0.00007 ± 0.00010 ± 0.00010 ± 0.00020

38 V 0.00036 0.00101 0.00051 0.00110 0.00179 0.00146 0.00036 0.03146 0.00132 0.00113 0.00151 0.00039 0.00080 0.00050 0.00143 0.00045 0.00139 0.00685

0.35 0.4 0.44 0.48 0.52 0.56 0.6 0.65 0.72 0.8 0.9 1 1.15 1.3 1.6 1.7

P

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

0.00785 0.15851 0.36400 0.58544 0.82012 0.93893 1.16519 1.54642 1.65264 0.57857 0.51578 0.63578 0.21951 0.16413 0.08899 0.00171

± 0.00032 ± 0.00066 ± 0.00070 ± 0.00075 ± 0.00138 ± 0.00160 ± 0.00167 ± 0.00190 ± 0.00246 ± 0.00089 ± 0.00101 ± 0.00113 ± 0.00038 ± 0.00073 ± 0.00036 ± 0.00026

40 V 0.00302 0.06023 0.17094 0.28659 0.40449 0.46728 0.58021 0.76686 0.82278 0.29084 0.26133 0.32433 0.11193 0.08312 0.04591 0.00292

± 0.00009 ± 0.00032 ± 0.00075 ± 0.00052 ± 0.00079 ± 0.00102 ± 0.00112 ± 0.00165 ± 0.00129 ± 0.00049 ± 0.00090 ± 0.00051 ± 0.00055 ± 0.00033 ± 0.00019 ± 0.00015

39 V 0.00014 0.00091 0.00226 0.00385 0.00528 0.00586 0.00767 0.01019 0.01132 0.00377 0.00363 0.00427 0.00163 0.00113 0.00056 0.00001

± 0.00005 ± 0.00004 ± 0.00004 ± 0.00005 ± 0.00004 ± 0.00011 ± 0.00008 ± 0.00009 ± 0.00014 ± 0.00008 ± 0.00008 ± 0.00009 ± 0.00008 ± 0.00006 ± 0.00005 ± 0.00004

38 V 0.00036 0.00045 0.00039 0.00047 0.00031 –0.00008 0.00009 0.00007 0.00107 0.00018 –0.00027 –0.00023 –0.00032 –0.00026 –0.00008 –0.00011

± 0.00009 ± 0.00010 ± 0.00015 ± 0.00009 ± 0.00011 ± 0.00016 ± 0.00010 ± 0.00007 ± 0.00011 ± 0.00009 ± 0.00011 ± 0.00012 ± 0.00012 ± 0.00011 ± 0.00007 ± 0.00012

37 V

± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002

36 V

5.49E‐17 1.11E‐15 2.55E‐15 4.10E‐15 5.74E‐15 6.58E‐15 8.16E‐15 1.08E‐14 1.16E‐14 4.05E‐15 3.61E‐15 4.45E‐15 1.54E‐15 1.15E‐15 6.23E‐16 1.20E‐17

Moles 40Ar*

126.40% 83.03% 95.19% 97.08% 99.00% 100.12% 98.21% 97.59% 98.60% 100.16% 100.65% 101.06% 102.57% 105.08% 104.96% 60.70%

R

± 0.11 ± 0.10 ± 0.09 ± 0.07 ± 0.09 ± 0.11 ± 0.09 ± 0.13 ± 0.09 ± 0.07 ± 0.09 ± 0.15 ± 0.09 ± 0.11 ± 0.08 ± 0.18 ± 0.08 ± 0.04

%‐sd

0.52% 0.48% 0.43% 0.33% 0.43% 0.54% 0.43% 0.60% 0.45% 0.34% 0.44% 0.70% 0.42% 0.53% 0.36% 0.86% 0.36% 0.20%

%‐sd

(continued)

± 23.19 82.61% ± 1.20 5.07% ± 0.45 2.06% ± 0.28 1.33% ± 0.20 0.93% ± 0.16 0.75% ± 0.14 0.67% ± 0.10 0.48% ± 0.09 0.43% ± 0.38 1.75% ± 0.30 1.40% ± 0.22 1.03% ± 0.66 3.13% ± 0.86 4.06% ± 1.44 6.90% ± 22.61 591.92%

Age (Ma)

21.36 21.22 21.35 21.21 21.41 21.22 21.36 21.34 21.25 21.31 21.21 21.32 21.25 21.24 21.20 21.40 21.32 21.25

Age (Ma)

28.08 23.57 21.88 21.41 21.67 21.69 21.29 21.24 21.38 21.47 21.30 21.16 21.17 21.31 20.93  3.82

1.9784 1.9661 1.9775 1.9646 1.9834 1.9657 1.9783 1.9765 1.9682 1.9739 1.9646 1.9749 1.9681 1.9675 1.9638 1.9828 1.9754 1.9688

R

2.6059 2.1851 2.0268 1.9832 2.0072 2.0093 1.9722 1.9680 1.9804 1.9894 1.9736 1.9602 1.9610 1.9744 1.9383 0.3521

95.42% 87.88% 90.22% 96.42% 91.73% 88.10% 96.58% 84.26% 79.52% 85.72% 95.36% 91.49% 96.60% 80.15% 92.14% 94.82% 93.91% 99.31%

%Rad

%Rad

9.78E‐15 1.88E‐14 1.40E‐14 1.81E‐14 1.35E‐14 1.14E‐14 1.32E‐14 1.84E‐14 1.83E‐14 2.12E‐14 1.14E‐14 7.34E‐15 1.32E‐14 1.51E‐14 1.67E‐14 5.59E‐15 1.50E‐14 3.41E‐14

Moles 40Ar*

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R = 40Ar*/39Ar (Ar* = radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16

Steps

± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003

36 V 0.000217 0.001101 0.000661 0.000313 0.000538 0.000659 0.000219 0.001412 0.001817 0.001463 0.000255 0.000302 0.000217 0.001444 0.000634 0.000140 0.000443 0.000116

–0.000007 0.000091 0.000059 0.000058 0.000028 –0.000004 0.000071 0.000126 0.000079 –0.000003 –0.000011 –0.000023 –0.000019 –0.000028 –0.000015 0.000002

± 0.00011 ± 0.00015 ± 0.00008 ± 0.00007 ± 0.00009 ± 0.00011 ± 0.00010 ± 0.00104 ± 0.00011 ± 0.00011 ± 0.00007 ± 0.00010 ± 0.00013 ± 0.00010 ± 0.00012 ± 0.00005 ± 0.00010 ± 0.00019

37 V

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐75 of GMMC (Mineral: Muscovite)

P

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐75 of GGMC (Mineral: Muscovite; N: 18)

1.6 1.6 1.6 1.6 1.6 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8 1.8

a b c d e f g h i j k l m n o p q

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

0.77817 2.37181 1.47555 1.08964 2.10305 1.61812 1.11227 0.98192 3.31411 3.98042 2.11764 0.79270 1.76059 1.56674 1.40520 1.23994 2.23900

± 0.00126 ± 0.00274 ± 0.00173 ± 0.00122 ± 0.00228 ± 0.00118 ± 0.00170 ± 0.00157 ± 0.00349 ± 0.00273 ± 0.00213 ± 0.00119 ± 0.00219 ± 0.00260 ± 0.00148 0.00168 0.00119

40 V 0.38414 1.16539 0.72863 0.48868 1.00874 0.78138 0.51308 0.40356 1.59111 1.79549 0.97317 0.39654 0.77213 0.46403 0.64077 0.53025 1.11734

± 0.00092 ± 0.00185 ± 0.00081 ± 0.00113 ± 0.00176 ± 0.00142 ± 0.00103 ± 0.00076 ± 0.00099 ± 0.00092 ± 0.00163 ± 0.00121 ± 0.00070 ± 0.00083 ± 0.00109 ± 0.00095 ± 0.00163

39 V 0.00561 0.01685 0.01019 0.00714 0.01464 0.01131 0.00767 0.00577 0.02235 0.02544 0.01369 0.00548 0.01153 0.00692 0.00915 0.00762 0.01601

± 0.00009 ± 0.00015 ± 0.00009 ± 0.00012 ± 0.00015 ± 0.00017 ± 0.00014 ± 0.00003 ± 0.00012 ± 0.00014 ± 0.00009 ± 0.00007 ± 0.00016 ± 0.00007 ± 0.00008 ± 0.00007 ± 0.00017

38 V 0.00062 0.00064 0.00041 0.00103 0.00107 0.00074 0.00072 0.00050 0.00126 0.00419 0.00079 0.00071 0.00211 0.00115 0.00072 0.00059 0.00094

± 0.00010 ± 0.00009 ± 0.00008 ± 0.00013 ± 0.00013 ± 0.00014 ± 0.00011 ± 0.00010 ± 0.00016 ± 0.00007 ± 0.00013 ± 0.00011 ± 0.00013 ± 0.00015 ± 0.00010 ± 0.00012 ± 0.00014

37 V

0.35 0.4 0.44 0.48 0.52 0.56 0.6 0.65 0.72 0.8 0.9 1 1.2 1.5

P

15 15 10 10 10 10 10 10 10 10 10 10 10 10

t

0.05417 0.54507 0.57040 0.59619 0.57585 0.84419 1.04142 1.20038 1.37646 0.41756 0.03847 0.02603 0.05224 0.27016

± 0.00036 ± 0.00071 ± 0.00085 ± 0.00100 ± 0.00124 ± 0.00106 ± 0.00143 ± 0.00127 ± 0.00193 ± 0.00085 ± 0.00030 ± 0.00031 ± 0.00027 ± 0.00034

40 V 0.02377 0.26020 0.27779 0.29128 0.28416 0.41948 0.52101 0.61280 0.70168 0.21239 0.01938 0.01293 0.02618 0.13710

± 0.00014 ± 0.00061 ± 0.00061 ± 0.00068 ± 0.00057 ± 0.00098 ± 0.00053 ± 0.00069 ± 0.00132 ± 0.00057 ± 0.00014 ± 0.00012 ± 0.00013 ± 0.00048

39 V 0.00032 0.00361 0.00386 0.00416 0.00391 0.00583 0.00714 0.00843 0.00967 0.00321 0.00031 0.00021 0.00034 0.00186

37 V

± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002

± 0.00002 ± 0.00002 ± 0.00004 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00003 ± 0.00003 ± 0.00002 ± 0.00003 ± 0.00002 ± 0.00002 ± 0.00002 ± 0.00002

36 V

± 0.00005 0.00083 ± 0.00047 0.000044 ± 0.00005 0.00102 ± 0.00025 0.000152 ± 0.00006 0.00095 ± 0.00022 0.000059 ± 0.00008 0.00099 ± 0.00015 0.000057 ± 0.00006 0.00097 ± 0.00032 0.000057 ± 0.00009 0.00100 ± 0.00018 0.000049 ± 0.00011 0.00104 ± 0.00021 0.000071 ± 0.00006 0.00046 ± 0.00032 0.000023 ± 0.00013 0.00047 ± 0.00017 –0.000034 ± 0.00008 0.00029 ± 0.00026 –0.000024 ± 0.00005 ‐0.00051 ± 0.00016 –0.000035 ± 0.00004 ‐0.00057 ± 0.00031 –0.000038 ± 0.00005 0.00159 ± 0.00021 –0.000033 ± 0.00004 0.00010 ± 0.00023 –0.000028

38 V

0.000088 0.000318 0.000220 0.000336 0.000489 0.000369 0.000200 0.000744 0.000769 0.001566 0.000719 0.000109 0.000572 0.002204 0.000560 0.000424 0.000305

36 V

3.79E‐16 3.82E‐15 3.99E‐15 4.18E‐15 4.03E‐15 5.91E‐15 7.29E‐15 8.41E‐15 9.64E‐15 2.92E‐15 2.69E‐16 1.82E‐16 3.66E‐16 1.89E‐15

Moles 40Ar*

5.45E‐15 1.66E‐14 1.03E‐14 7.63E‐15 1.47E‐14 1.13E‐14 7.79E‐15 6.88E‐15 2.32E‐14 2.79E‐14 1.48E‐14 5.55E‐15 1.23E‐14 1.10E‐14 9.84E‐15 8.68E‐15 1.57E‐14

Moles 40Ar*

76.24% 91.79% 96.97% 97.18% 97.07% 98.29% 97.99% 99.44% 100.73% 101.71% 126.99% 143.60% 119.14% 103.06%

%Rad

96.66% 96.04% 95.61% 90.90% 93.14% 93.27% 94.69% 77.63% 93.15% 88.38% 89.97% 95.96% 90.41% 58.43% 88.22% 89.91% 95.98%

%Rad

1.7373 1.9227 1.9910 1.9890 1.9670 1.9780 1.9587 1.9479 1.9617 1.9662 1.9821 2.0082 2.0010 1.9705

R

1.9582 1.9546 1.9361 2.0268 1.9418 1.9316 2.0528 1.8887 1.9402 1.9593 1.9578 1.9183 2.0614 1.9727 1.9347 2.1025 1.9233

R

Note: P: % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R = 40Ar*/39Ar (Ar* = radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

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

Steps

Single Crystal Incremental Heating Ar/Ar Dating of the Sample OS‐111 of GMMC (Mineral: Biotite)

P

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐111 of GGMC (Mineral: Biotite; N: 17)

Table 7.2 (Continued)

± 0.19 ± 0.07 ± 0.10 ± 0.15 ± 0.09 ± 0.10 ± 0.13 ± 0.17 ± 0.05 ± 0.04 ± 0.09 ± 0.18 ± 0.09 ± 0.23 ± 0.10 ± 0.14 ± 0.07

18.77 20.76 21.49 21.47 21.23 21.35 21.14 21.03 21.18 21.22 21.40 21.68 21.60 21.27

± 3.12 ± 0.27 ± 0.42 ± 0.23 ± 0.26 ± 0.18 ± 0.21 ± 0.14 ± 0.12 ± 0.40 ± 4.08 ± 5.42 ± 2.67 ± 0.54

Age (Ma)

21.14 21.10 20.90 21.88 20.96 20.85 22.15 20.39 20.95 21.15 21.13 20.71 22.25 21.30 20.89 22.69 20.76

Age (Ma)

16.62% 1.28% 1.93% 1.05% 1.21% 0.86% 0.99% 0.66% 0.57% 1.87% 19.07% 25.02% 12.38% 2.54%

%‐sd

0.91% 0.35% 0.47% 0.71% 0.42% 0.49% 0.61% 0.85% 0.24% 0.21% 0.44% 0.89% 0.42% 1.09% 0.49% 0.64% 0.33%

%‐sd

1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6 1.6

a b c d e f g h i j k l m n o p q r s

10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10 10

t

6.94384 4.51736 4.95417 1.67869 2.52739 2.55738 1.94237 2.46576 2.26855 2.51609 2.31747 2.45882 1.71005 1.70450 2.49780 3.00921 2.43836 1.73189 1.69883

± 0.00170 ± 0.00404 ± 0.00359 ± 0.00109 ± 0.00220 ± 0.00162 ± 0.00239 ± 0.00166 ± 0.00122 ± 0.00201 ± 0.00211 ± 0.00195 ± 0.00191 ± 0.00175 ± 0.00143 0.00227 0.00230 0.00134 0.00164

40 V 3.24699 2.04868 2.31687 0.80894 1.16634 1.18976 0.88519 1.23066 1.08333 1.11044 1.05460 1.18506 0.79203 0.68568 1.11444 1.45712 1.19322 0.77806 0.81447

± 0.00220 ± 0.00183 ± 0.00235 ± 0.00076 ± 0.00184 ± 0.00158 ± 0.00147 ± 0.00172 ± 0.00125 ± 0.00122 ± 0.00106 ± 0.00158 ± 0.00156 ± 0.00106 ± 0.00080 ± 0.00198 ± 0.00156 ± 0.00100 ± 0.00173

39 V 0.04407 0.02803 0.03159 0.01067 0.01534 0.01606 0.01164 0.01617 0.01423 0.01452 0.01421 0.01548 0.01043 0.00949 0.01466 0.01899 0.01536 0.01036 0.01073

± 0.00024 ± 0.00022 ± 0.00027 ± 0.00010 ± 0.00008 ± 0.00015 ± 0.00010 ± 0.00010 ± 0.00007 ± 0.00010 ± 0.00020 ± 0.00013 ± 0.00011 ± 0.00012 ± 0.00011 ± 0.00011 ± 0.00014 ± 0.00008 ± 0.00008

38 V 0.00836 0.00353 0.01929 0.00005 0.00174 0.00089 0.00062 0.00102 0.00033 0.00268 0.00090 0.00048 0.00040 0.00096 0.00103 0.00124 0.00178 0.00063 0.00036

± 0.00019 ± 0.00027 ± 0.00027 ± 0.00012 ± 0.00018 ± 0.00024 ± 0.00017 ± 0.00017 ± 0.00019 ± 0.00012 ± 0.00019 ± 0.00016 ± 0.00016 ± 0.00018 ± 0.00012 ± 0.00016 ± 0.00020 ± 0.00016 ± 0.00015

37 V 0.001924 0.001587 0.001382 0.000358 0.000748 0.000751 0.000712 0.000251 0.000485 0.001151 0.000844 0.000471 0.000576 0.001205 0.000997 0.000520 0.000353 0.000732 0.000330

± 0.00003 ± 0.00005 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00003 ± 0.00002

36 V 4.86E‐14 3.16E‐14 3.47E‐14 1.18E‐14 1.77E‐14 1.79E‐14 1.36E‐14 1.73E‐14 1.59E‐14 1.76E‐14 1.62E‐14 1.72E‐14 1.20E‐14 1.19E‐14 1.75E‐14 2.11E‐14 1.71E‐14 1.21E‐14 1.19E‐14

Moles 40Ar* 91.82% 89.62% 91.79% 93.69% 91.25% 91.32% 89.16% 97.00% 93.68% 86.50% 89.24% 94.34% 90.05% 79.11% 88.21% 94.90% 95.73% 87.52% 94.26%

%Rad 1.9637 1.9762 1.9628 1.9443 1.9774 1.9630 1.9565 1.9435 1.9617 1.9599 1.9610 1.9575 1.9443 1.9665 1.9770 1.9599 1.9562 1.9481 1.9660

R

Note: P = % power of 60 W (60 W *[P/10] Synrad CO2 laser); t = laser duration time; V = volt; %Rad = % radiogenic argon; R = 40Ar*/39Ar (Ar* = radiogenic argon); J value = 0.0064870 ± 0.000011 (1σ).

P

Mineral

Single Crystal Total Fusion Ar/Ar Dating of the Sample OS‐105 of GGMC (Mineral: Muscovite; N: 19)

20.82 20.96 20.81 20.62 20.97 20.82 20.75 20.61 20.80 20.78 20.80 20.76 20.62 20.85 20.96 20.78 20.74 20.66 20.85

± 0.03 ± 0.08 ± 0.05 ± 0.11 ± 0.09 ± 0.09 ± 0.11 ± 0.08 ± 0.09 ± 0.10 ± 0.10 ± 0.08 ± 0.13 ± 0.15 ± 0.08 ± 0.07 ± 0.08 ± 0.13 ± 0.11

Age (Ma)

0.16% 0.38% 0.24% 0.51% 0.41% 0.41% 0.53% 0.38% 0.42% 0.50% 0.49% 0.41% 0.63% 0.73% 0.39% 0.34% 0.41% 0.64% 0.51%

%‐sd

214  ACTIVE GLOBAL SEISMOLOGY Plateau steps are filled, rejected steps are open

16

box heights are 2σ

ÇG – biotite granodiorite Sample OS-55 Biotite incremental heating

36 32

14

Single Grain incremental heating Plateau age = 20.84±0.13 Ma

12

24

8 6

16

4

Age (Ma)

20

8 0.0

2 Cumulative 39Ar Fraction 0.2

0.4

0.6

Plateau steps are filled, rejected steps are open

0.8

0 1.0 19.0

Age (Ma) 19.5

20.0

20.5

21.0

21.5

22.0

22.5

23.0

box heights are 1σ

ÇG – biotite granodiorite Sample OS-72 Biotite incremental heating

30

Relative probability

10

Plateau age = 20.84±0.13 Ma MSWD = 0.74, probability = 0.68 Includes 100% of the 39Ar

Number

28

12

Total fusion Single grain Ages Mean = 20.838±0.038 [0.18%] 95% conf. N = 18

10

Total fusion Single grain ages Mean = 21.08±0.02 [0.093%] 95% conf. N = 18

Single Grain incremental heating Plateau age = 21.16±0.04 Ma

Age (Ma)

10

6 Plateau age = 21.16±0.042 Ma MSWD = 0.26, probability = 0.997 Includes 100% of the 39Ar

0 0.0

Number

20

Relative probability

8

4

2 Cumulative 39Ar Fraction 0.2

0.4

0.6

0.8

0 1.0 20.6

Age (Ma) 20.8

21.0

21.2

21.4

21.6

Figure 7.15  Results of 40Ar/39Ar geochronology of porphyritic granodiorite of ÇG.

(ÇDFZ) that juxtaposes deep‐seated granite‐migmatite complex with the greenschist facies metamorphic rocks of Karakaya complex (Figs.  7.2a, b, c and 7.17). Along the western border, it was intruded by ÇG and associated haplogranite and pegmatite dikes of Early Miocene. These two plutons forming the ÇPC were uplifted together along ÇDFZ during the Early Miocene (see Section 7.6.3). 7.6.2. Emplacement of Çataldağ Granodiorite (ÇG) In contrast to GGMC, ÇG is an I‐type granodiorite [Kamacı and Altunkaynak, 2011a, 2011b; Kamacı et  al., 2013], and represents a discordant, shallow‐level intrusive body that intruded into the neighboring GGMC and metamorphic basement rocks of Sakarya continent. It is formed mainly from K‐feldspar megacryst biotite‐grano­ diorites, haplogranitic peripheral zone, and associated dikes and sills. It commonly displays porphyritic texture. Granonophyric‐graphic textures are also rarely observed.

These textural features and discordant nature of the ÇG indicate that the granitic magma reached shallow levels in the crust, most probably in the epizone. Along the northern and eastern contacts, the ÇG is also bounded by the ÇDFZ. The Cataldağ granodiorite is relatively undeformed in its center, but shows discrete mylonitic bands and cataclastic textures in the footwall of the detachment fault in its northern edge (Figs.  7.2 and 7.17). The reason for not showing very extensive deformation fabric is possibly the significant amount of exhumation and uplift that had already taken place. Hot granodioritic melt was synkinematically emplaced into an already partly exhumed GGMC located in the footwall of the detachment. High temperature ductile deformation products (e.g., grain boundary migration and chess‐board twins of quartz) have occurred in rap­ idly cooling ÇG as synkinematic textures [Kamaci and Altunkaynak, in preparation]. Cooling and ductile defor­ mation appears to be coeval. The T‐time history of some

THE ÇATALDAĞ PLUTONIC COMPLEX IN WESTERN ANATOLIA  215 Plateau steps are filled, rejected steps are open

box heights are 1σ

GGMC – two-mica granite Sample OS-75 Muscovite incremental heating

40

10

Single Grain incremental heating 21.38±0.05 Ma

Total fusion Single grain ages Mean = 21.27±0.04 [0.18%] 95% conf N = 18

Relative probability

8 Total Gas age = 21.38±0.5 Ma

6 20

Age (Ma)

4

2 Age (Ma)

Cumulative 39Ar Fraction 0.4

0.6

Plateau steps are filled, rejected steps are open

0.8

box heights are 1σ

GGMC – migmatite melanosome Sample OS-111 Biotite incremental heating

40

30

8

5 20

4

21.6

22.0 22.2

21.8

Single Grain incremental heating Age = 21.18±0.07 Ma

Total fusion Single grain ages Mean = 21.11±0.21 [0.99%] 95% conf. N = 16

Age (Ma)

2 1 0.2

0.4

0.6

Plateau steps are filled, rejected steps are open

0.8 box heights are 1σ

GGMC – gneiss Sample OS-150 Biotite incremental heating

0 1.0 19.0

Age (Ma) 21.0

20.0

22.0

14 12

30

23.0

20 6

Age (Ma)

Plateau age = 20.80±0.08 Ma MSWD = 0.35, probability = 0.98 Includes 100% of the 39Ar

24.0

Single Grain incremental heating Age = 20.80±0.08 Ma

Total fusion Single grain ages Mean = 20.74±0.03 [0.14%] 95% conf. N = 18

10 8

10

21.4

Relative probability

40

21.2

3

Cumulative 39Ar Fraction

0 0.0

21.0

7 6

10

20.8

10 9

Plateau age = 21.18±0.07 Ma MSWD = 0.58, probability = 0.87 Includes 100% of the 39Ar

20.6

Relative probability

0.2

0 1.0 20.4

Number

0 0.0

Number

10

Plateau age = 21.38±0.05 Ma MSWD = 0.84, probability = 0.62 Includes 98.6% of the 39Ar

Number

30

4 2

0 0.0

Cumulative 39Ar Fraction 0.2

0.4

0.6

0.8

0 1.0 20.1

Age (Ma) 20.3

20.5

Figure 7.16  Results of Ar/ Ar geochronology of gneiss and two‐mica granite of GGMC. 40

39

20.7

20.9

21.1

21.3

Fig.5a

Fig.8a

Fig.7a

Fig.3e

Fig.6

216  ACTIVE GLOBAL SEISMOLOGY

ÇG

GGCM

Marble (Karakaya) GGMC

Fig.9b

Fig.3c

Fig.5c

ÇG

Country rock Turtle-back

Figure 7.17  Graphical abstract showing contact relationships and deformational structures in ÇPC.

detachment zones reveals rapid cooling (>50 °C/my) of the footwall rocks, followed by a more protracted cooling history (2 km), the two fault types are expected to merge into one fault zone [modified from McNeill et al., 2004].

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  247

Figure 9.7  The Central Aegean trough (CAT) fault system according to GreDaSS. The inset at the bottom right shows the aftershock distribution and the focal mechanism of the 26 July 2001 (Mw 6.4) Skyros earthquake [Ganas et al., 2005].

South of NAB and NAT, a similar but smaller‐scale tectonic structure has been formed; that is the central Aegean trough (CAT, Figs. 9.4 and 9.7), which is consid‑ ered as a southern branch of the NAF [e.g., Şengör, 1979; Barka and Kadinsky‐Cade, 1988; Barka, 1992; Şaroğlu et  al., 1992; Herece, 1990; Barka and Reilinger, 1997]. This branch roughly separates Biga Peninsula to the north from Edremit Gulf to the south, enters the Aegean Sea [Martin and Mascle, 1989; Mascle and Martin, 1990; Boztepe‐Güney et al., 2001; Işler et al., 2008; Yaltırak et al., 2012], and ends up at Skyros basin [Martin and Mascle, 1989; Mascle and Martin, 1990; Papanikolaou et  al., 2015]. The shape of Skyros basin (Fig. 9.7) is similar to the shape of NAB with its margins being tectonically delimited by normal faults [Papanikolaou et  al., 2015]. According to GreDaSS, three marginal normal dip‐slip fault segments (Fig.  9.7) have produced three strong events [Jackson et  al., 1982; Papazachos et  al., 1984; Kiratzi et al., 1985; Ekström and England, 1989; Barakou et al., 2001; Kiratzi and Louvari, 2003; Papazachos et al., 1991; Taymaz et al., 1991]: (1) on 4 March 1967 (MW 6.2) by the central Aegean trough west ISS845; (2) on 19 December 1981 (MW 6.8) by the central Aegean trough seg. A (ISS840); and on 27 December 1981 (MW 6.3) by the central Aegean trough seg. B (ISS841). The two latter events occurred on two adjacent segments, ISS840 and ISS841, located at the northeast‐southwest–trending

southern scarp of the basin, while the former occurred on a perpendicular northwest‐southeast–trending segment at the western scarp of the basin (Fig.  9.7). Although Skyros basin exhibits normal faulting, the focal mecha‑ nisms of the 26 July 2001 (MW 6.5) Skyros earthquake, which occurred offshore and northwest of Syros Island, west of Skyros basin, revealed strike‐slip faulting [Benetatos et  al., 2002; Ganas et  al., 2005; Roumelioti et  al., 2003] represented by the Skyrosfault (ISS835) in GreDaSS. 9.1.2.4. Seismicity of North Greece Seismicity in North Greece is not homogeneously distributed, either in frequency or density (Fig.  9.8). Consequently, historical information is not uniform all over the region and even the most complete and updated earthquake catalogues [Ambraseys and Jackson, 1990; 1998; Guidoboni et al., 1994; Papazachos and Papazachou, 1997; 2003; Guidoboni and Comastri, 2005; Ambraseys, 2001; 2009] are in some cases contradicting, or even doubtful, in describing past events, especially the older ones. In other words, the quantitative exploitation of data and the association of the events to any recognized fault, sometimes become hard [Caputo et  al., 2008]. The Aegean region (Fig. 9.8) and specifically northern Greece, show a sparse and infrequent seismicity in east‑ ern Macedonia and Thrace regions, where no significant seismicity is instrumentally recorded. On the contrary,

248  ACTIVE GLOBAL SEISMOLOGY

Figure 9.8  Historical and instrumental seismicity of the Aegean Region [Sboras, 2011], taken from the online catalogue of the Seismological Station of AUTh [http://geophysics.geo.auth.gr/ss/station_index_en.html; Papazachos et al., 2000, 2009].

central Macedonia and Chalkidiki Peninsula show intense activity including several strong instrumentally recorded events (e.g., the 1905 Athos [MS 7.5], the 1932 Ierissos [MS 6.9], and the 1902 [MS 6.6] and 1978 [MW 6.4] Thessaloniki earthquakes) and plenty of records of pre‑ instrumental historic events. West Macedonia is consid‑ ered a low seismicity area [Voidomatis, 1989; Papazachos, 1990], which however produced the strong (Mw 6.5) 15 May 1995 earthquake, which occurred in the vicinity of Kozani‐Grevena towns. The pattern then changed back to sparse and infrequent seismicity with few instrumen‑ tally recorded moderate events in the area of Epirus and, generally, broader northwestern Greece. A possible explanation for this difference is observational bias: the historical record is obviously richer in densely populated areas, while mountainous regions like Epirus and western Macedonia were always sparsely populated. Nevertheless, instrumental seismicity seems to verify this difference, although the time window is too narrow. Finally, the North Aegean Sea shows the most intense and frequent instrumentally recorded seismicity. The highly active structures of NAT and NAB have produced numerous strong to major events, many of which have been instru‑ mentally recorded. However, historical events are hard to associate to specific seismogenic sources, given that the extent of the sea prevents intensity distribution doc‑ umentation and any credible observation.

9.1.3. Seismogenic Sources and Active Faults 9.1.3.1. Definitions Generally, the common terminology for defining active faults and capable faults varies in the sense of the time window used. In this chapter, the term active fault is preferred instead of capable fault, since the former is con‑ sidered also as capable of producing a future earthquake. Machette [2000], after revising various definitions given by several authors, organizations, and committees, con‑ cludes that the time window used for characterizing faults should depend on the regional tectonic setting and the range of completeness of the regional seismic hazards. Thus, the term active fault is applied when a fault shows evidence of recent reactivation and/or is capable of being reactivated in the future. The definition of recent is quite relative and can be very subjective, but it always is comparable to the seismotectonic regime of the Aegean region. Moreover, activity can be implied and constrained by evidence. Therefore, the criteria of Pavlides et  al. [2007], which are suitably adapted for the faulting charac‑ ter of the Aegean, can be assigned to the active faults. Active faults in continental Greece show a large variety at all scales. However, small individual faults are insignifi‑ cant for the aims of seismic hazard assessment (SHA). For this reason, the structures studied in this paper are the so‐called seismogenic sources. Based on the definition of

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  249

Kastelic et al. [2008] for the Italian Database of Individual Seismogenic Sources (DISS), seismogenic sources are active faults capable of generating Mw > 5.5 earthquakes. The latter definition is also suitable for the case of the study area, given that Pavlides and Caputo [2004] suggest that a magnitude larger than 5.5 is needed in order to have linear morphogenic earthquakes sensu [Caputo, 2005] in the Aegean area. Morphogenic earthquakes are those seismic events capable of generating, or modifying, the surface morphology instantaneously and perma‑ nently. Information on past events is obtained following two distinct methodological approaches: the geological one (morphotectonic investigations and palaeoseismo‑ logical excavations) and the historical one (contempora‑ neous descriptions and surveys of coseismic ruptures) [Caputo, 2005]. 9.1.4. Fault Inventories: An Historical Review 9.1.4.1. Worldwide National Concepts Having foreseen the importance to hold and manage all collected data from active fault investigations, many research institutes worldwide have built databases at a national scale. For example, the Geological National Survey in New Zealand (GNS Science; http://data.gns.cri.nz/af/), the Institute of Advanced Industrial Science and Tech­ nology in Japan (AIST; https://gbank.gsj.jp/activefault/ index_e_gmap.html), the United States Geological Survey (USGS; http://earthquake.usgs.gov/hazards/ qfaults/), and the Istituto Nazionale di Geofisica e Volcanologia in Italy (INGV; http://diss.rm.ingv.it/diss/) have certainly the most developed databases of active faults: the Database of Individual Seismogenic Sources [DISS; Valensise and Pantosti, 2001; Basili et  al., 2008; 2009] for Italy and the European Database of Seismogenic Faults [EDSF; Basili et al., 2013] for Europe. Concerning the Aegean region, the early efforts were focused on faults that were related to either historically or instrumentally recorded earthquakes. Ambraseys and Jackson [1998] associated historical and instrumental events to faulting for the eastern Mediterranean area, while Papazachos et al. [1999; 2001] compiled a map of “rupture zones” representing faults responsible for recent events, also suggesting their geometric (strike, dip) and kinematic (rake) parameters. Nevertheless, their work was almost exclusively based on historical and instru‑ mental seismological data. At the same time, in the frame of the European project FAUST (Faults as a Seismologist’s Tool; Mucciarelli and FAUST Working Group [2000]), the first, though preliminary, version of a database with fully parameterized active faults was compiled for the Aegean region, including around 50 sources. Although this was the first database complete of all principal seismotectonic parameters, most seismogenic sources were associated

with recently reactivated faults, with few exceptions where geological information was also partially considered. From a geographic point of view, a more enriched and detailed study was the map of capable faults in Greece and the broader Aegean region compiled by Pavlides et al. [2007]. A main difference from the previous compi‑ lations is the fact that this map delineates in detail scarps of all active faults with a clear morphological expression that meet one or more of the criteria concerning the iden‑ tification of active faults, whether they are associated with a known earthquake or not. The map, however, gives no other parametric information except the geographical one. More recently, Karakaisis et  al. [2010] reassessed and enriched the previous seismologically based compilations of Papazachos [1999; 2001], while Mountrakis et  al. [2006], using geological and seismological evidence, presented a review of the active faults of a limited sector of northern Greece. Like other compilations, this is rather descriptive without containing any quantitative parametric information. It is obvious that in all previous works crucial informa‑ tion was lacking, for example, maps do not contain any other parametric data except the geographic location, while catalogues and compilations provide (and not always) only partial information, exclusively concerning some geometric and kinematic parameters. 9.1.4.2. The Greek Database of Seismogenic Sources (GreDaSS) The GreDaSS (http://eqgeogr.weebly.com/; http:// gredass.unife.it/) is a continuously updatable open‐access database based on GIS software environment (MapInfo). The overall database structure and informatics derive from the well‐tested, time‐proven, and worldwide acknowledged DISS (Database of Individual Seismogenic Sources) data‑ base, proposed by INGV, which represents the result of almost 20 yr research experience of its Working Group [e.g., Valensise and Pantosti, 2001; Basili et al., 2008; 2009]. In the first stage, only shallow (crustal) tectonic structures have been considered, leaving the deeper structures of the Hellenic subduction zone open for future investigations. Shallow structures are more important in terms of seismic hazard assessment (SHA), since they are distributed all over the Aegean region, close to and sometimes even directly affecting inhabited areas. Not only does GreDaSS contain and combine all kinds of information that previ‑ ous studies dealt with, but also it provides data of vari‑ ous levels, which can be handled either independently or combined. GreDaSS has contributed to the European Database of Seismogenic Faults [EDSF; Basili et  al., 2013], a part of the FP7‐funded SHARE project (Seismic Hazard Harmonization in Europe; http://www. share‐eu.org/, wih documentation hosted in the EFEHR - European Facilities for Earthquake Hazard & Risk; http://www.efehr.org:8080/jetspeed/portal/).

250  ACTIVE GLOBAL SEISMOLOGY

Although GreDaSS is a multilayered informational database with several classes of seismogenic sources, two are currently the most important and advanced ones: the Individual Seismogenic Sources (ISSs) and the Composite Seismogenic Sources (CSSs). Both of those types are described in detail by Basili et al. [2008; 2009] for DISS and by Sboras [2011] and Caputo et al. [2015] specifically for GreDaSS. 9.1.5. North Aegean Broader Region: The Seismogenic Sources of North Greece 9.1.5.1. Introduction The main active faults of northern Greece characterized as possible seismogenic sources are presented in detail in the GreDaSS database [Pavlides et al., 2010; Sboras, 2011; Caputo et al., 2012; Caputo et al., 2015]. After observing the geometry and kinematics of the sources in this area, Sboras [2011] and Caputo et al. [2012] roughly divided the area into five sectors (Fig. 9.9): (1) an east‐west–trending fault belt in Thrace and eastern Macedonia; (2) a complex system east(southeast)‐west(northwest)–trending fault ­ affecting Chalkidiki Peninsula and possibly associated with low‐angle detachments; (3) the northeast‐southwest–

trending Hellenides fault system in western Macedonia and Epirus (northwest Greece); (4) the east(southeast)‐ west(northwest)–trending Thessalian fault system (central Greece); and (5) the (east)northeast‐(west)southwest– trending North Aegean Sea fault system, which is almost exclusively located offshore. Based on this discrimina‑ tion, only seismogenic sources belonging to the North Aegean Sea fault system (5) will be synthetically described and discussed in the following sections, in terms of their seismotectonic parameters, their relations to historical and instrumental earthquakes, and their seismotectonic behavior wherever it differentiates. In the following descriptions, seismogenic sources are presented grouped when a CSS spans one or more ISSs. The discussion in such cases is cumulative and, like in individual cases, it includes reviews and critical analyses of data and nonparametric information that might concern faults’ behavior, interaction, earthquake asso‑ ciations, and so on. Furthermore, some minor fault segments might be discussed even if they do not qualify to be included in the database individually due to their small size and their low potential magnitude capability (below M 5.5). Discussion, however, might contain useful information about their behavior (interaction, linkage, etc.) inside the fault zone.

Figure  9.9  Map of the seismogenic sources (CSSs and ISSs) completed for North Greece within the frame of GreDaSS project [Sboras, 2011; Caputo et al., 2012]. Sources that are in progress in the time of publication, or out of the study area are not shown. The green dashed lines divide the area into five sectors (green numbers) as described in the text. Light colored CSSs represent ambiguous sources.

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  251

9.1.5.2. The North Aegean Sea Fault System A strong limitation in investigating offshore faults is the complete lack of direct observations; in terms of morphotectonic and geological approaches, only detailed bathymetric maps [e.g., Maley and Johnson, 1971; IOC, 1981; Papanikolaou et  al., 2002] and seismic reflection investigations [e.g., Brooks and Ferentinos, 1980; Mascle and Martin, 1990; Roussos and Lyssimachou, 1991; Papanikolaou et al., 2006; Sakellariou et al., in press] are available. Accordingly, most of the information con‑ cerning offshore faults is provided by seismic data like focal mechanisms and microseismic distribution [e.g., Dziewonski et al., 1983; 1984; Rocca et al., 1985; Kiratzi, 1991; Taymaz et al., 1991; Koukouvelas and Aydin 2002; update by GreDaSS]. As aforementioned, the dominating regional tectonic structures of the North Aegean Sea are the North Aegean basin (NAB), and the North Aegean trough (NAT; Figs. 9.4, 9.9, and 9.10). Along strike, the NAB is bounded by two major fault zones (CSSs), which are characterized mainly by oblique‐ slip kinematics: the South Chalkidiki offshore and the North Aegean basin CSSs (Fig. 9.10). Minor structures affect the interposed region, but the lack of sufficient

data, their likely limited dimensions (hence limited maximum expected magnitude), and their location far from the Greek coastlines make at present their recog‑ nition more problematic and precise seismotectonic definition less urgent for SHA. The NAT, represented by a crustal‐scale negative flower structure (Fig. 9.6b), is a transtensional shear zone mainly consisting of two CSSs: the North NAT and the South NAT (Fig. 9.10). These two antithetic and roughly parallel CSSs progressively converge eastward probably merging into a unique structure near the entrance of Saros Gulf (Figs. 9.6, 9.10) [Çagatay et al., 1998; Yaltirak et al., 1998; Coskun, 2000; Kurt et al., 2000; Yaltirak and Alpar, 2002]. Other active tectonic structures also occur in the North Aegean region, and are synthetic or kinematically associ‑ ated with those of the NAB and NAT; those are the Aghios Efstratios and Samothraki ISSs and the Lemnos CSS. 9.1.5.3.  South Chalkidiki Offshore CSS280 (Athos ISS282) The South Chalkidiki offshore CSS (Fig. 9.10) repre‑ sents the northern boundary of the NAB. Its trace runs immediately south of Sithonia and Athos peninsulas, forming a northeast‐southwest–trending steep submarine

Figure 9.10  The seismogenic sources of the North Aegean Sea fault system (GreDaSS) [Sboras, 2011; Caputo et al., 2012] (as seen in Fig. 9.9). Orange sources and numbers: CSSs and their codes; red sources and numbers: ISSs and their numbers; blue sources: DSSs. The epicenters (stars) have been taken from the catalogue of Papazachos et al. [2000, 2009; http://geophysics.geo.auth.gr/ss/station_index_en.html].

252  ACTIVE GLOBAL SEISMOLOGY

slope for a total length of around 90 km, as documented by bathymetric surveys [Maley and Johnson, 1971; Papanikolaou et al., 2002] and seismic profiles [Ferentinos et al., 1981; Roussos and Lyssimachou, 1991; Papanikolaou et  al., 2006]. Based on slight strike variations and especially on the occurrence of a large earthquake likely associated with this fault zone, the Athos ISS is considered as a separate segment. The event occurred on 8 November 1905 and the macroseismicmagnitude was 7.5 [Papazachos and Papazachou, 1997; 2003] or Ms 7.3 [Ambraseys, 2001]. This fault segment is approximately 54 km long [Papanikolaou et  al., 2006]. The maximum expected magnitude for the entire CSS is between 6.9 and possibly 7.2 as the worst case scenario. 9.1.5.4. North Aegean Basin CSS810 (NAB Segment A ISS810 and Segment B ISS811) The North Aegean basin CSS (Fig. 9.10) is one of the longest structures of the region, striking northeast‐ southwest and dipping to the northwest. It represents the southern boundary of the NAB, running from north of Lemnos Islands, up to the Sporades Islands. The fault zone is well expressed morphologically, as it forms a deep and steep northwest‐dipping slope, clearly imaged in detailed bathymetric maps [Maley and Johnson, 1971; IOC, 1981; Papanikolaou et  al., 2002]. The cumulative downthrow of the seafloor is about 1300 m. The fault zone is also clearly imaged in the seismic reflection pro‑ files indicating significant cumulative displacements and especially the deformation of the seafloor sediments [Brooks and Ferentinos, 1980; Ferentinos et  al., 1981; Mascle and Martin, 1990; Roussos and Lyssimachou, 1991; Papanikolaou et  al., 2006; Sakellariou et  al., in press]. Along this zone, two strong earthquakes, likely produced by two adjacent fault segments, are well docu‑ mented. Accordingly, two ISSs are distinguished: the NAB Segments A and B (eastern and central, respec‑ tively; Fig. 9.10). The NAB Segment A ISS generated the 6 August 1983 (Ms 6.8 after Kiratzi et al., [1991]) earthquake that reacti‑ vated the northeastern part of the NAB CSS. Based on the aftershock spatial distribution [Rocca et  al., 1985; Taymaz et al., 1991] and the magnitude of the mainshock, a length of 44 km is inferred. The NAB Segment B ISS is associated with the 18 January 1982 (Mw = 6.6 after Dziewonski et al. [1983]) earthquake. Again, a length of 33 km is inferred from the aftershock spatial distribution [Taymaz et  al., 1991] and the magnitude of the main‑ shock. There are many other focal mechanisms proposed for both the 1982 and 1983 seismic crises [Dziewonski et  al., 1984; Papazachos et  al., 1984; Rocca et  al., 1985; Ekström et al., 1987; Ekström and England, 1989; Kiratzi et  al., 1991; Taymaz et  al., 1991; Jackson et  al., 1992; Vannucci and Gasperini, 2003; 2004], all suggesting very

steep (northwest‐dipping) to subvertical nodal planes and an almost pure right‐lateral strike‐slip kinematics. The proposed segmentation is in agreement with the seis‑ mogenic source segments proposed by Koukouvelas and Aydin [2002] and Papanikolaou and Papanikolaou [2007]. The 1982 (Mw 6.6) and 1983 (Mw 6.8) events likely rep‑ resent the maximum expected magnitudes for the two segments, while in case of a unique rupture event (worst‐ case scenario) the expected magnitude is 7.6 [Papanikolaou and Papanikolaou, 2007]. Nevertheless, many older strong events (1 June 1366; 12 November 1456; 1471; 12 August 1564; 12 April 1572; 28 June 1585; 5 December 1776; 3 February 1779; 4 June 1947; Fig. 9.11) are reported in the vicinity of the NAB CSS, especially close to its central and northeastern parts (Fig.  9.11) [Papazachos and Papazachou, 1997; 2003; Ambraseys, 2009], but they are either ambiguous or loosely located in order to be assigned to a specific ISS or CSS. In order to highlight this uncertainty, two examples of historical events are given, one of 1 June 1366 (M 6.6) and the other of 1471 (M 7.0). Both events are absent from Ambraseys’s [2009] catalogue, while only the second one is included in the catalogue of Guidoboni and Comastri [2005] as an event that occurred sometime between June 1470 and 1472, in the area north of Lemnos Island. However, a conflict might also be possible regarding the 8 November 1905 Athos earthquake. According to the descriptions, it is preferred to be assigned to the Athos ISS (see the previous South Chalkidiki offshore CSS). 9.1.5.5. South NAT CSS800 The South NAT CSS (Fig. 9.10) is probably the longest active structure in the Aegean Sea. It is oriented east‐ northeast–west‐southwest, running north of Lemnos and Gökçeada (Imbros) islands, controlling the Saros Gulf and crossing the Gelibolu Peninsula [Fig.  9.6; Stanley and Perissoratis, 1977; Lybéris, 1984; Ambraseys and Finkel, 1987; Mercier et al., 1989; Kiratzi, 1991; Tüysüz et al., 1998; Yaltirak et  al., 1998; Armijo et  al., 1999; Saatçilar et  al., 1999; Papadimitriou and Sykes, 2001; Rockwell et al., 2001; Ambraseys, 2002; Koukouvelas and Aydin, 2002; Yaltirak and Alpar, 2002; Altunel et al., 2004; McNeill et al., 2004; Kaya et  al., 2004; Janssen et  al., 2009; Koral et  al., 2009; Chatzipetros et al., 2013]. This structure is hard to separate from its mechanical continuity, the North Anatolian fault. Ganos fault segment that ruptured the Gelibolu Peninsula (August 9, 1912; Mw 7.4; Gaziköy) from Saros Gulf to Marmara Sea [Ambraseys and Finkel, 1987; Tüysüzet  al., 1998; Yaltirak et al., 1998; Rockwell et al., 2001; Ambraseys, 2002; Yaltirak and Alpar, 2002; Altunel et al., 2004; Kaya et  al., 2004; Janssen et  al., 2009] will not be further dis‑ cussed, as it is far from the study area. The fault zone geometry is relatively well constrained based on bathymetric data, seismic reflection profiles,

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  253

Figure 9.11  Map showing strong earthquakes that are associated, discussed, or referenced in GreDaSS [Sboras, 2011]. The epicenters have been taken from the catalogue of Papazachos et al. [2000; 2009; http://geophysics. geo.auth.gr/ss/station_index_en.html].

microseismic distributions, and focal mechanisms of moderate and strong events [Maley and Johnson, 1971; Mascle and Martin, 1990; Kiratzi et al., 1991; Papazachos et  al., 1991; Taymaz et  al., 1991; Çagatay et  al., 1998; Saatçilar et  al., 1999; Coskun, 2000; Kurt et  al., 2000; McNeill et  al., 2004; Karabulut et  al., 2006; Ustaömer et  al., 2008]. Focal mechanisms document a prevailing strike‐slip motion with some dip‐slip component, while the distribution of microseismicity suggests a steep (to sub‑ vertical) fault plane. Total length of the entire fault zone probably exceeds 200 km. Based on the overall geometry, several segments certainly exist, but lack of specific data does not allow the determination of segment boundaries and especially their nature. Accordingly, the maximum expected magnitude could range between 6.5 to probably more than 7.5. 9.1.5.6.  North NAT CSS290 (Saros Gulf ISS290 and Samothraki SE ISS291) Mechanically associated with the South NAT CSS (see previous CSS) is the North NAT CSS (Fig. 9.10), which runs subparallel and is antithetic (south‐southeast–dip‑ ping) to the former. Those two fault zones progressively converge eastward probably merging into a unique crus‑ tal‐scale flower structure at the entrance in the Saros Gulf (Fig. 9.6) [Çagatay et al., 1998; Yaltirak et al., 1998; Coskun, 2000; Kurt et al., 2000; Yaltirak and Alpar, 2002; McNeill et al., 2004]. The total length is probably more than 120 km demonstrating oblique‐slip (to strike‐slip)

kinematics, according to available focal mechanisms of moderate to large earthquakes. Based on the slightly modular geometry, at least two segments have been recognized and included in GreDaSS: the 26 km long Saros Gulf ISS and the 24 km long Samothraki SE ISS. Those faults ruptured during two Mw 6.6 and Mw 5.7 earthquakes, occurring on 27 March 1975 and 6 July 2003, respectively, north of Gelibolu Peninsula and east of Samothraki Island [Kiratzi et al., 1985; Taymaz et al., 1991; Papazachos et  al., 1991; Jackson et  al., 1992; Papazachos and Papazachou, 1997; 2003; Vannucci and Gasperini, 2003; 2004; Karabulut et  al., 2006]. Focal mechanisms suggest an oblique‐slip motion (normal and dextral components) on a moderate southeast‐dipping fault plane for the Saros Gulf ISS [Kiratzi et  al., 1985; Taymaz et  al., 1991; Papazachos et  al., 1991; Jackson et  al., 1992] and an almost pure dextral strike‐slip motion on a steeply southeast‐dipping fault plane for the Samothraki SE ISS [Karabulut et al., 2006]. The investiga‑ tion of the 2003 sequence [Karabulut et al., 2006] indicates a thick seismogenic layer of about 19 km and a structure of more than 20 km long. However, according to the dimensions, the potential magnitude of the Samothraki SE ISS should be approximately 6.6, although the greatest shock of the 2003 sequence was only Mw 5.7. Two historical large events are also reported in this area [Papazachos and Papazachou, 1997; 2003]: the first on 23 July 1719 (Me 6.7) and the second on 6 August 1860 (Me = 6.2). Although both events cannot be accurately

254  ACTIVE GLOBAL SEISMOLOGY

Figure 9.12  (a) Morphotectonic map of Samothraki Island [after Pavlides et al., 2005], showing the fault traces (red lines) of the North Samothraki ISS (GRIS288), which define the northern coast of the island and part of the North NAT CSS (GRCS290), which defines the southeastern coast of the island. The margins of the hanging valley of Giali (ca. 600 m a.s.l.) are marked with yellow dashed line. (b) 3D terrain model of Samothraki Island observed from the east showing the morphological change along the fault traces (red lines) that delimit the mountain front of Mt. Saos. The vertical exaggeration is 2×. (c) 3D terrain model of the southeastern side of Samothraki Island observed roughly from the south showing in more detail the hanging valley of Giali (black dashed line) and the triangular facets of the southeastern coast of the island that are formed due to the footwall uplift produced by the activity of the offshore North NAT fault zone (GRCS290).

located, at least the former could be associated to the Samothraki SE fault segment. Reactivation of the entire length of the North NAT CSS seems impossible. The active western part of the fault zone, where the 24 May 2014 (Mw 6.9) earthquake occurred, is also evident on land, at the southeastern coast of Samothraki Island (Fig.  9.12). The island lies on the footwall next to the North NAT CSS and has undergone uplift that formed triangular facets and a hanging valley (Giali) at about 600 m a.s.l. (Fig. 9.12) [Pavlides et al., 2005]. Based on the fault zone’s geometry and the empirical relationships of

Wells and Coppersmith [1994], a maximum magnitude of 7.1 could be the result of the worst‐case scenario. 9.1.5.7. North Samothraki ISS288 The fault crosses the northern coast of Samothraki Island and continues unilaterally offshore (Fig.  9.10). This fault can be interpreted as a normal dip‐slip secondary structure of the NAT. According to morphotectonic inves‑ tigations [Pavlides et al., 2005], the fault is characterized by discrete scarps, most of them aligned in an east‐south‑ east–west‐northwest direction and dipping northward.

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  255

The scarps form a steep morphology that controls the drainage pattern and causes deposition of massive colluvial and alluvial deposits. The North Samothraki ISS is probably the causative structure for the homony‑ mous earthquake that occurred on 9 February 1893 [Papazachos and Papazachos, 1997; 2003; Ambraseys, 2009; Pavlides et al., 2005]. Based on the macroseismic magnitude of 6.8 [Papazachos and Papazachos, 1997; 2003], a minimum length of at least 22 km can be esti‑ mated (Fig. 9.12). 9.1.5.8. Lemnos CSS825 Lemnos CSS (Fig. 9.10) is a northeast‐southwest–striking, subvertical strike‐slip right‐lateral structure running across the homonymous island [Koukouvelas and Aydin, 2002; Pavlides et al., 2009; Chatzipetros et al., 2013]. It consists of several smaller segments, some of them controlling the coastline of the northeastern part of the island. The loca‑ tion, geometry, and kinematics of this structure imply a probable connection with the South NAT CSS to the north (Figs.  9.10, 9.13). Even though regional instru‑ mental seismicity is rather low, a strong (M 7.0) historical (197 BC) earthquake is reported in the catalogue of Papazachos and Papazachou [1997; 2003] based on scripts of Pausanias, who refers to the sinking of a small island (Chrysi) northeast of Lemnos. According to empirical relationships [Wells and Coppersmith, 1994], the maximum expected magnitude of a total rupture of the fault zone is estimated to be Mw 6.8. Lemnos Island also bears smaller active faults [Pavlides et al., 1990, 2009; Koukouvelas and Aydin, 2002; Tranos, 2009; Chatzipetros et al., 2013] as seen in Fig. 9.13. At the northern part of the island, the most common fault strikes are southwest‐northeast to west‐southwest–east‐northeast, characterized by dextral strike‐slip component, which in cases also demonstrates a strong normal component [Koukouvelas and Aydin, 2002; Pavlides et  al., 2009; Tranos, 2009; Chatzipetros et al., 2013]. At the southern part of the island, faults mostly show an east‐northeast– west‐southwest strike and normal dip‐slip kinematics [Koukouvelas and Aydin, 2002; Pavlides et  al., 2009; Tranos, 2009; Chatzipetros et al., 2013]. Only at the west‑ ern part of the island (e.g., Kaspakas fault, Fig.  9.13) faulting shows west‐northwest–east‐southeast direction and oblique‐slip kinematics [Koukouvelas and Aydin, 2002; Pavlides et al., 1990; 2009; Tranos, 2009; Chatzipetros et al., 2013]. 9.1.5.9. Aghios Efstratios CSS/ISS831 The Aghios Efstratios ISS is a northeast‐southwest– striking, strike‐slip tectonic structure, running south of Lemnos Island and crossing through the small Aghios Efstratios Island (Figs. 9.10, 9.13) [Pavlides et al., 1990; Pavlides and Tranos, 1991]. Although it is parallel to the

NAB CSS, its motion is mainly transcurrent based on the available focal mechanism of the 1968 event, which ruptured most of the structure [McKenzie, 1972; Kiratzi et al., 1991; Taymaz et al., 1991; Vannucci and Gasperini, 2003, 2004]. Field mapping indicates a northwest‐dipping steep to subvertical fault with significant right‐lateral motion [Pavlides and Tranos, 1991]. Fault length is based on the aftershock distribution [Drakopoulos and Economides, 1972; North, 1977], while width and average displacement are based on North’s [1977] estimations over the aftershock distribution and the Ms 7.1 magnitude of the 1968 earthquake, respectively. However, more recent studies [Nalbant et  al., 1998; Papadimitriou and Sykes, 2001] suggest different dimensions and displacements (see discussion in Pavlides et al. [2009]). 9.2. THE 24 MAY 2014 NORTH AEGEAN EARTHQUAKE: STRESS CHANGE PATTERNS 9.2.1. Introduction The 24 May 2014 event occurred at 09:25:03 (UTC) in the offshore area between Lemnos and Samothraki islands (Fig. 9.14) and caused light damages and injuries in the surrounding area [Sboras et  al., 2015]; however, damages and injuries were more intense and widespread in Turkey (Gökçeada, Marmara Sea, Adapazarı). The given magnitude varies from Mw 6.3 (AUTh) up to Mw 7.0 [Görgün and Görgün, 2015]; nevertheless, the majority of institutes suggested a magnitude of Mw 6.9 (Table 9.1). The epicentral area lies in the North Aegean trough, which is part of the larger tectonic system of the North Anatolian fault, connected with the Marmara Sea fault system to the east and the North Aegean basin to the west (Figs. 9.4, 9.5,. 9.6, 9.10, and 9.14). Considering the magnitude and epicenter of the 2014 earthquake, the causative fault is only a segment of the North Aegean trough, located near Samothraki Island (Fig.  9.15). The aftershock sequence clearly shows three main clusters (Fig.  9.14), the central one along the fault plane and the other two situated bilat‑ erally, east and west of the 2014 seismogenic source, along the trough. Görgün and Görgün [2015] have taken into account in their model a much longer fault (length > 80 km), including the entire aftershock sequence, that is, all the aforementioned three main clusters. Such a case would imply an extremely large tectonic structure with much higher potential magnitude than the 2014 event. The proximity of the 2014 seismogenic source to other fault segments to the east, the North Aegean basin to the west, and some other neighboring faults in the surround‑ ing area, raises possible scenarios for stress loading and consequently triggering effects.

256  ACTIVE GLOBAL SEISMOLOGY

Figure  9.13  (a) Morphotectonic map of Lemnos and Aghios Efstratios islands [Chatzipetros et  al., 2013]. The main fault zones (LN01  =  Mourtzouflos fault; LN02  =  Kaspakas fault; LN03  =  Kontias‐Kotsinas fault zone; LN04 = Moudros fault; LN05 = Fanos‐Aghia Sophia fault zone; AE01 = Aghios Efstratios seismic fault; AE02 = Aghios Efstratios conjugate fault) and their sense of displacement are shown. White squares show the locations where rejuvenated drainage has been observed. (b) View of Kaspakas (LN02a) fault scarp at its westernmost edge. Kaspakas segment is one of the most prominent faults on Lemnos Island, since it forms a well‐defined fault scarp, mostly on the volcanic basement. Dashed and dotted lines delineate the lower and the upper limits of the free face, respectively, showing a clear sign of relatively recent reactivation. (c) The Fanos segment (LN05a) is exposed as a series of low scarps, which occasionally show fresh fault surfaces. (d) Detail of the polished fault surface, indicative of predominantly normal displacement. (e) Lower hemisphere stereographic projection of a bend in LN05 trace, acting as a deformation transfer zone. (f) Outcrop of LN03d dextral strike‐slip fault. A positive flower structure is formed in this restraining step‐over of the fault zone, causing a local reverse deformation zone and bending. This structure is accommodated by small‐scale reverse shear zones on both walls of the fault zone.

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  257

Figure 9.14  Map (top) and profile (bottom) of the aftershock sequence (from 24 May until 6 September) of the 24 May 2014 main shock (from the GeIn‐NOA; http://bbnet.gein.noa.gr/HL/databases/database). The events shown in the profile belong to a 0.5° wide buffer zone at both sides of the profile line. The yellow star on the profile indicates the main shock.

Table 9.1  Moment Tensor Solutions for the 24 May 2014 Main Event in the North Aegean Sea Location Lat 1 2 3 4 5 6 7 8 9 10 11

40.28 40.29 40.27 40.28 40.30 40.305 40.30 40.211 40.30 40.312 40.311

Long

M0 [Nm]

25.40 25.40 25.36 25.38 25.67 25.453 25.50 25.307 25.50 25.452 25.452

4.15E + 18 1.693E + 19 2.5E + 19 2.1E + 19 2.47E + 19 2.56E + 19 – – – 1.952E + 18 4.6E + 19

Plane 1

Plane 2

MW

Depth [km]

Strike

Dip

Rake

Strike

Dip

Rake

Reference

6.3 6.8 6.9 6.8 6.9 6.9 6.9 6.5 6.9 6.7 7.0

15.0 14.0 20.5 20.0 12.0 11.5 29.0 10.0 24.0 22.0 15.0

245 70 72 343 73 165 341 167 219 65 82

72 85 73 76 85 79 88 87 88 59 77

171 −167 −167 −12 −177 13 −8 9 173 −146 149

338 338 338 76 343 72 71 76 309 317 180

81 77 77 77 87 77 82 81 83 52 60

18 −5 −18 −164 −5 168 −178 177 2 −47 15

AUTh NOA INGV GFZ Harvard USGS IPGP* ERD KOERI* KOERI1 Görgün and Görgün [2015]

Source: * Data taken from EMSC 1   Earthquake Report. Note: See also Figure 9.16. AUTh = Aristotle University of Thessaloniki; NOA = National Observatory of Athens; INGV = Istituto Nazionale di Geofisica e Vulcanologia; GFZ = GeoForschungsZentrum; USGS = U.S. Geological Survey; IPGP = Institut de Physique du Globe de Paris; ERD = Disaster and Emergency Management Presidency, Earthquake Department; KOERI = Kandilli Observatory and Earthquake Research Institute.

258  ACTIVE GLOBAL SEISMOLOGY

Figure 9.15  Map of the 45 km long fault model of the 2014 earthquake and the nearby seismogenic sources of the North Aegean Sea. The selected focal mechanism of NOA is also shown. Table 9.2  Receiver Fault Parameters taken from GreDaSS that are Used for the Calculation of Stress Change Code GRIS290 GRIS291 GRIS288 GRCS800 GRCS825 GRIS282 GRIS810 GRIS811

Name

Strike (°)

Dip (°)

Rake (°)

Saros Gulf Samothraki SE North Samothraki South NAT Lemnos Athos NAB segment A NAB segment B

 68  83 286 255 223  55 227 222

65 75 60 78 83 77 79 70

−145 −175 −110 −152 −162 −155 180 −172

9.2.2. Prior Seismicity and Faulting The North Aegean Sea exhibits a rich historical and instrumental record of earthquakes, in terms of both fre‑ quency and magnitude. Nevertheless, historical events are difficult to associate to specific faults due to often large errors in the determination of epicentral locations, especially of the offshore earthquakes and the large num‑ ber of active faults in the broader area. Moreover, the recent earthquake showed that rupture directivity diversely affected urban areas situated on both extensions of the fault tips, a fact that could give wrong estimations on macroseismic epicenters for historical events. Thus, there are few reliable events that can be directly associated with specific faults. In the frame of the GreDaSS project, all available seismotectonic data were collected

Latest earthquake 27 March 1975, MW6.6 06 July 2003, MW5.7 09 February 1893, MW6.8 – – 08 November 1905, M7.5 06 August 1983, MW6.6 18 January 1982, MW6.6

and combined to reconstruct the active fault regime of North Greece, including the North Aegean Sea [Sboras, 2011; Caputo et  al., 2012] (Fig.  9.10).Thus, seismogenic sources associated with specific and reliable enough earth‑ quakes can be found in GreDaSS as Individual Seismogenic Sources (ISSs). In particular, there are six ISSs in the study area that are related to events prior to 2014 (Table  9.2). From east to west, these are (Fig. 9.10): ••Saros Gulf ISS (GRIS290), which is associated with the 27 March 1975 (Mw 6.6) earthquake ••Samothraki SE ISS (GRIS291), which is associated with the 6 July 2003 (Mw 5.7) earthquake sequence ••North Samothraki ISS (GRIS288), which is associ‑ ated with the 9 February 1893 (Mw 6.8) earthquake ••NAB segment A ISS (GRIS810), which is associated with the 6 August 1983 (Mw 6.6) earthquake

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  259

••NAB segment B ISS (GRIS811), which is associated with the 18 January 1982 (Mw 6.6) earthquake ••Athos ISS (GRIS282), which is associated with the 8 November 1905 (M 7.5) earthquake Further information on the seismogenic sources and reliability of earthquake associations, the parametric data, and references for the aforementioned ISSs are available at the GreDaSS project website (http://eqgeogr. weebly.com/; http://gredass.unife.it) The most recent event in the broader region occurred on 8 January 2013 with Mw 5.7. Its epicenter was located about 25 km SE of Lemnos Island and about 25 km north‐northeast from the central Aegean basin [Ganas et  al., 2014; Kiratzi and Svigkas, 2014]. Coulomb static stress change calculations by Ganas et  al. [2014] and Karakostas et al. [2014] show that static stress changes lie far from the epicentral area of the 2014 event. Static stress change modeling of older events for the broader Aegean Sea has been carried out by a few researchers [Nalbant et  al., 1998; Papadimitriou and Sykes, 2001; Paradisopoulou et al., 2010; Rhoades et al., 2010] for various time spans and events. Rhoades et  al. [2010] apply the Coulomb failure criterion for a series of strong events from 1964 to 2001, showing the individual stress change after each event, while Paradisopoulou et al. [2010] make the same calculations, but for the period between 1904 and 1999, and show the cumulative stress change after every event for the next one. The complexity of the North Aegean fault pattern, the rapid crustal deformation of the region, the low quality of seismological data for old events (perhaps older than two‐three decades ago), and the occurrence of plenty of small events, which can cumulatively affect the stress status in the region in contrast with the absence of any strong earthquake, all of these issues make any stress estimation before the rup‑ ture of the 2014 event inaccurate. For this reason, we only calculate the static stress change that was created after the 2014 rupture and how this change can affect the neighboring major faults. Thus, we considered the stress status prior to this event equal to zero. 9.2.3. The 24 May 2014 Earthquake (Mw 6.9) The proposed focal mechanisms of the 24 May 2014 earthquake (Table  9.1; Fig.  9.16) show slight differences, except of No. 10 (KOERI1), which suggests a more oblique slip motion with significant normal dip‐slip component. Based on the tectonic setting of the area and the epicentral location of the event, only (pure) strike‐slip, south‐south‑ east–dipping nodal planes could be considered as most appropriate (focal mechanisms 2, 3, 4, 6, 7, and 8; Table 9.1). The aftershock spatial distribution (Fig. 9.14), accord‑ ing to GeIn‐NOA (http://bbnet.gein.noa.gr/HL/database), lies along a west‐southwest–east‐northeast–trending axis

Figure  9.16  The proposed focal mechanisms as they are referred in Table 9.1; P and T axes are also shown.

and can be divided into three major clusters: (1) the mid‑ dle one is located at the hypocenter of the mainshock, (2) the western one is located immediately south of Athos Peninsula (ca. 85 km west‐southwest from the mainshock and ca. 50 km west‐southwest from the middle cluster), and (3) the eastern cluster is almost next to the middle one (ca. 15 km north of Gökçeada) without any significant gap. This pattern implies stress accumulation bilaterally from the two tips of the causative fault. The vertical dis‑ tribution (Fig. 9.14) shows that seismicity is concentrated approximately between 5 and 35 km depth. 9.2.4. Coulomb Stress Change When an earthquake occurs, the average stress on the activated fault is reduced, while stress is increased at the tips of its plane and at sites around it [e.g., Reasenbeg and Simpson, 1992; Harris et al., 1995]. An immediate result of this stress transfer is the generation of aftershocks [e.g., Toda et al., 2002].

260  ACTIVE GLOBAL SEISMOLOGY

The accumulation and release of stress on a fault are controlled not only by the regional stress field and rock properties, but also by its surrounding faults. Earthquakes will cause stress to increase or decrease at other faults, and thereby alter the timing of earthquake occurrence on them; this effect is known as earthquake triggering and delaying [e.g., King et al., 1994; Hodgkinson et al., 1996; King and Cocco, 2001; Zhang et al., 2003]. In the present work, we study the static fault interaction during the 2014 earthquake sequence in the surrounding area of NAT, using the Coulomb failure criterion. In order for a fault to slip, Coulomb stress change Δσf should exceed a threshold value on its plane:

f

s

n

(9.1)

where Δτs is the shear stress change on the failure plane, μ′ is the friction coefficient, and Δσn is the normal stress change. In order to visualize the areas where stress is increased or decreased for certain types of faulting, we used the Coulomb v3.3 application [Toda et  al., 2005], which resolves the shear and normal components of the stress change on a grid, or on specified receiver fault planes, in a homogeneous, elastic, and isotropic half‐space. Accor­ ding to Toda et al. [2011], source faults are the faults that have slipped and receiver faults are planes with a specified strike, dip, and rake, on which the stresses are imparted by the source faults. Thus, shear stress change is dependent on the position, geometry, and slip of the source fault and on the position and geometry of the receiver fault, including its rake, while normal stress change (clamping or unclamp‑ ing) is independent of the receiver fault rake. The friction coefficient may vary from less than 0.4 for highly lubricated faults to ≈ 0.8 for jagged, anastomosing faults [Chan and Stein, 2009]. However, many authors have supported that the results are not very sensitive to the changes of μ′ [e.g., Deng and Sykes, 1997a; 1997b; King et al., 1994; Stein et al., 1997]. Therefore, a common value of μ′ = 0.4 was selected [Nalbant et  al., 1998; Papadimitriou, 2002]. 9.2.5. Fault Models The first step for building the fault model is to define the constant source fault, which is the causative fault of the 24 May 2014 earthquake. Based on the earthquake epicenter and the seafloor morphology, the fault must be located at the northern scarp of the trough (i.e., the North NAT). This means that a south‐southeast–dipping fault plane should be expected. The local geodynamic regime implies the occurrence of an almost pure right‐lateral strike‐slip fault. Among the available focal mechanisms, only seven of them comply with (almost pure) strike‐slip and south‐ southeast–dipping nodal planes. The differences between these focal mechanisms practically do not significantly affect the stress change and displacement pattern results;

Table 9.3  Source Parameters Used for the Fault Model of the 2014 Earthquake Source fault Strike: Dip: Rake: Length: Width: Min. depth: Max. depth: Slip: M0: MW:

70° 85° −167° 45.0 km 12.0 km 3 km 15 km 1.04 m 1.693 × 1019 Nm 6.8

therefore, for strike, dip, and rake in our fault model, we used the focal mechanism of NOA (strike = 70°, dip = 85°, and rake = −167°). The calculated average slip, after apply‑ ing the relationship of Aki [1966] for the given M0 from NOA, is 1 m. After using the relationship of Aki [1966] for seismic moment and the empirical relationships of Wells and Coppersmith [1994] of magnitude versus rupture length and rupture width for strike‐slip faults, we estimated the fault to be approximately 45 km in length (L) and 12 km in width (W). Based on (1) the fault dimensions, (2) the proposed epicentral location and focal depth of the main‑ shock, (3) the aftershock spatial distribution (Fig. 9.14), (4) the seafloor morphology, and (5) the position of the adjacent ISSs, the fault can be placed immediately south from Samothraki Island and west from Samothraki SE (GRIS291) ISS (Fig. 9.15). For most of the aforementioned reasons (dimensions and hypocentral depths of mainshock and aftershocks), it is also suggested that the rupture of the specific event reached the sea bottom at the distinctive fault scarp formed south of Samothraki Island and westward. The source fault model is mapped in Figure 9.15 (colored in red) and its parameters are listed in Table 9.3. The receiver faults have been modeled according to GreDaSS (Table 9.2 and Fig. 9.15) and include both ISSs and CSSs. ISSs are modeled in such a way that they can directly provide the parameters needed for the stress transfer models. CSSs, however, use a range of values of many parameters. For this study, the average values have been chosen for the CSSs. Parameters of both ISSs and CSSs are shown in Table 9.2. 9.2.6. Stress Change Pattern The static stress change is calculated for each fault in Table 9.2 at the depth of 12 km (Fig. 9.17), which is near the hypocentral depth of the 2014 earthquake. The results show that the remotest faults, Athos (GRIS282) and NAB Segment B (GRIS811), are not affected by the 24 May 2014 rupture. Saros Gulf (GRIS290) and NAB Segment A (GRIS810) ISSs are slightly and partially affected with stress increase. The adjacent Samothraki

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  261

Figure 9.17  Coulomb static stress change pattern after the 24 May 2014 main shock (source fault colored in red) calculated for each receiver fault (colored in black and labeled) shown in Table 9.3 and Figure 9.15. Faults not participating in the calculations are shown in light gray.

262  ACTIVE GLOBAL SEISMOLOGY

SE ISS (GRIS291) shows high values of stress rise. The North Samothraki ISS (GRIS288) to the north and the Lemnos CSS (GRCS825) to the south show significant stress drop. The long South NAT CSS (GRCS800) to the south lies mostly in the stress drop areas except a small part near the Gökçeada (Imbros) Island where a moderate stress rise occurs. The bilateral stress accumulation beyond the tips of the source fault (for receiver faults with similar attributes, such as Saros Gulf or Samothraki SE) is probably the reason why the aftershock distribution was constrained along the 2014 fault strike axis (Fig. 9.14). 9.2.7. Displacement Pattern We calculated the displacement pattern of the sequence, not only to assess the ground deformation and verify the InSAR data, but also to verify our fault model. The displacement pattern of the 24 May 2014 earthquake at zero depth (surface) is also modeled using the Coulomb v3.3 application [Toda et al., 2005], which is based on the Okada [1992] dislocation solution formulae in a homoge‑

Figure 9.18  Horizontal (up) and vertical (down) displacements caused by the 24 May 2014 main shock, as modeled using the Okada [1992] formulae.

neous, elastic, and isotropic half‐space (Fig. 9.18). As a source fault, the same fault model used in the Coulomb stress change calculations has been taken into account. The 24 May 2014 rupture was recorded by GPS stations [Sboras et al., 2015] in the surrounding area. The horizontal and vertical components of the recordings are compared with the modeled displacement pattern, resolved into its horizontal and vertical components at each GPS station. Preliminary results show that there is a quite good agreement between the modeled and the actual rupture components (Fig. 9.18). The most significant displacements are observed on the three surrounding islands (Samothraki, Gökçeada [Imbros] and Lemnos). In fact, the horizontal displacement pattern indicates that the highest displacement took place on Samothraki Island (ca. 0.16 m toward northeast). 9.3. CONCLUSIONS The complex Aegean geodynamics have been discussed in this paper, while the state‐of‐the art of active fault studies in northern Greece and the North Aegean sea have also been presented. Emphasis is given on the active fault geometry of the North Aegean trough. NAT is the northern branch of the North Anatolia fault, extending westward from the Sea of Marmara and the Gulf of Saros into the Aegean Sea. Fault strike changes from west‐southwest–east‐northeast at the Saros Gulf and Samothraki Island, to southwest‐northeast south of Chalkidiki Peninsula and to the coast of the Greek main‑ land (Thessaly), where it terminates (North Aegean basin). The fault is of almost pure strike‐slip character within Turkish territory, while it shows oblique‐slip to normal sense of movement in the North Aegean Sea (transtensional tectonics). The causative fault of the 24 May strong earthquake is a segment (45 km long and 12 km wide) of the North Aegean trough (NAT), part of the western North Anatolian fault (NAF) extension sys‑ tem, located offshore Samothraki and between Samothraki and Lemnos islands on the northern branch of the NAT fault zone. This is supported by the earth‑ quake epicenter, the aftershock distribution and the sea‑ floor morphology. An east‐northeast–west‐southwest striking right‐lateral strike‐slip, south‐southeast–dipping fault plane has been modeled. The receiver faults have been modeled according to GreDaSS (Fig. 9.17; Table 9.2) and include both ISSs and CSSs. The static stress change after the 2014 mainshock on nearby faults (Fig. 9.17) shows that only the immediately eastern segment of the North NAT CSS (CSS290), that is, the Samothraki SE ISS (ISS291), bears significant stress rise. This can explain the eastern aftershock cluster that lies along its fault plane. Static stress rises on the Samothraki SE ISS and triggering effect could be expected. Although this source was reactivated during

NORTH AEGEAN ACTIVE FAULT PATTERN AND THE 24 MAY 2014, MW 6.9 EARTHQUAKE  263

the 1975 earthquake, the rapidly deforming crust in this region and the effects of other earthquakes since then, either strong or weak, leave the triggering issue open. Moreover, it is not clear how the 2014 aftershock eastern cluster affects the stress state on this fault. The normal dip‐slip North Samothraki ISS (ISS288) is situated in the stress drop area, as well as the entire Samothraki Island (for faults of similar geometry and kinematics). Stress relief on the Lemnos CSS (CSS825) implies delay effect. The last fault that is significantly affected by the 2014 rupture is the South NAT CSS (CSS800), which is almost entirely situated in the stress drop area, while a small part of it (toward its northeastern tip) is lying in a moderately stress rise area. The 2014 earthquake fault plane rupture was not enough for the static stress change to reach substantially more distant faults, like the Saros Gulf (ISS290) and the NAB Segment A (ISS810). Both Athos (ISS282) and NAB Segment B (ISS811) faults are indeed beyond the reach of any stress change. Thus, the western aftershock cluster located near the Athos ISS (Fig. 9.14) cannot be explained by static stress change calculations. The deformation pattern after the 24 May 2014 main‑ shock (Fig.  9.18) is in agreement with the geodynamic regime and the GPS motions. Samothraki Island has nota‑ bly moved northeastward (ca. 16 cm), Gökçeada (Imbros) Island toward west‐northwest (ca. 8 cm), and Lemnos Island toward southwest (ca. 6 cm). The most sizeable ver‑ tical movements are observed on Samothraki Island (uplift), matching the geological evidences prior to the 2014 event (Fig.  9.12). Much lower vertical movements (1 Hz) since they involve intense computational power and high‐resolution earth properties that are not always available. Stochastic ­models, on the other hand, not only provide effective solutions up to higher frequencies of engineering interest but also include the inherent randomness regarding the high frequencies. In summary, stochastic models require much less computational effort at the expense of losing the full wave propagation effects. Previously, several of these alternative approaches have been used to simulate past earthquakes that occurred in Turkey as well as

scenario events [e.g., Durukal, 2002; Pulido et al., 2004; Yalcinkaya, 2005; Sørensen et  al., 2007; Ugurhan and Askan, 2010; Tanırcan, 2012]. In this study, we present synthetic intensity maps ­prepared for the eastern section of the North Anatolian fault zone for a selected set of scenario earthquakes. Since we aim to practically simulate ground‐motion levels up to the frequencies of engineering interest, ­ which correlate closely with the observed seismic damage, we use stochastic ground‐motion simulation method herein. In the next two sections, we introduce detailed information on the study region and the simulation methodology. Then, relationships between ground‐ motion parameters and felt‐intensity values are presented. Finally, the applications are demonstrated followed by the conclusions. 10.2. STUDY REGION The North Anatolian fault zone is one of the most active fault zones in the world and the one that causes the most destructive events in Turkey. The sequence of major events along this fault zone started with the 1939 Erzincan (Ms ~8.0) event and moved westward through the last century [e.g., Ketin, 1969; Ambraseys, 1970; Sengor, 1979; Toksoz et  al., 1979; Sengor and Canitez, 1982; Barka, 1992, 1996; Stein et al., 1997; Armijo et al., 1999; Armijo et al., 2002; Sengor et al., 2005]. The final two major events were the 1999 Kocaeli (Mw = 7.4) and 1999 Düzce (Mw = 7.2) earthquakes that occurred close to the Istanbul region (Fig.  10.1a). Since the western parts of the North Anatolian fault zone include the industrial facilities coupled with dense population, these sections have been studied in detail. However, despite the intense seismic activity, Erzincan region in the east is not as much investigated as the western sec‑ tions of NAFZ. Erzincan city center is built on Erzincan basin, which is an alluvium pull‐apart basin [e.g., Tatar et al., 2013; Sarp, 2015] of size 50 km x 15 km. Ovacik fault and the North East Anatolian fault (both left‐lateral strike‐slip faults), intersect the right‐lateral North Anatolian fault at the southern and northern edge of the Erzincan basin, respec‑ tively [e.g., Bernard et  al., 1997; Avsar et  al., 2013] as shown in Figure  10.1a and b. This region, close to the Karliova triple junction in the east, is one of the most hazardous regions of the world. Historical records evi‑ dence around 20 large earthquakes in the close vicinity of Erzincan within the past 1000 yr [Barka, 1993]. In the last century, other than the 1939 event, Erzincan experienced another destructive earthquake in 1992 (Mw = 6.6, accord‑ ing to USGS and Daniell et al., 2011] that caused major structural loss leading to a large number of fatalities.

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  275

Figure  10.1  (a) Major tectonic structures around the Anatolian Plate and major earthquakes on the North Anatolian fault zone in the last century (adapted from Akyuz et al. [2002]); (b) Seismotectonics in the Erzincan region with the fault systems and the epicenters of the 1939 and 1992 events. (b) The light gray region is the Erzincan county and the dark gray area is the Erzincan basin where the city center is located. The rectangle shows the area of computation in this study.

276  ACTIVE GLOBAL SEISMOLOGY

10.3. GROUND‐MOTION SIMULATIONS In this section, initially the methodology for strong ground‐motion simulations will be presented briefly. Then, applications of the simulation method for scenario earthquakes in Erzincan are demonstrated along with the results. 10.3.1. Methodology: Stochastic Ground‐Motion Simulations The objective of all strong ground‐motion simulation methods is to estimate reliable synthetic records by mod‑ eling the earthquake source mechanisms and regional wave propagation. Alternative simulation methods provide different levels of accuracy and computational cost for modeling ground‐motion records. Deterministic methods are numerical solutions of the wave equation for full wave propagation purposes. These methods require highly resolved velocity models in order to generate broadband records with significant accuracy despite the high com‑ putational cost [e.g., Frankel, 1993; Olsen et  al., 1996]. Stochastic methods, on the other hand, construct the ground motions as a combination of the far field S‐wave spectrum and random phase angles [Boore, 1983; Beresnev and Atkinson, 1997]. It has inherent limitations due to lack of full wave propagation and simpler source models, however, it models the higher frequencies simply yet effectively. Stochastic method is employed successfully to model earthquakes at different regions of the world includ‑ ing Turkish events [e.g., Castro et al., 2001; Motazedian and Moinfar, 2006; Ugurhan and Askan, 2010; Ugurhan et al., 2012]. Despite studies concerning detailed local tomography of the upper crust in the Erzincan region [e.g., Aktar et al., 2004; Gokalp 2007], there are no high‐ resolution velocity models for the shallow soil layers. Thus, fully deterministic models that aim higher fre‑ quencies of engineering interest (up to 10–15 Hz) are not implemented here. Stochastic method is used in this study, which is particularly preferred since it produces broadband frequencies that are of interest to engineering structures. We use a recent form of stochastic finite‐fault modeling based on a dynamic corner frequency approach introduced by Motazedian and Atkinson [2005]. Stochastic finite‐fault method generates ground motions radiating from a rectangular finite‐fault discretized into subfaults each of which is modeled as a stochastic point source with an ω−2 spectrum. In this method, the hypo‑ center is located on one of the subfaults and the rupture is assumed to start propagating radially from the hypo‑ center with a constant rupture velocity. Each subfault is triggered when the rupture reaches the center of that subfault. The contribution of all subfaults is summed with appropriate time delays in order to obtain the entire

fault plane’s contribution to the seismic field at any obser‑ vation point. In the dynamic corner frequency concept, the total energy radiated from the fault is conserved regardless of the selected subfault size. In this study, we use the dynamic corner frequency approach as imple‑ mented in the computer program EXSIM [Motazedian and Atkinson, 2005]. In the dynamic corner frequency approach, the accelera‑ tion spectrum Aij( f )of the ijth subfault is expressed in terms of source, path and site effects as follows: Aij f

2 f

CM 0ij Hij 1



f fcij

fRij

2 2

e

Q f

G Rij D f e

f

(10.1) · 2 is a scaling factor, ℜθ φ is the radia‑ 3 4 tion pattern, ρ is the density, β is the shear‐wave velocity, M 0Sij M 0ij is the seismic moment, Sij is the relative nl nw S kl k 1 l 1 where C

slip weight, and fcij (t ) is the dynamic corner frequency of 1/ 3

ij

th

subfault where fcij (t )

N R (t )

1/ 3

4.9 10

6

M 0ave

.

Here Δσ is the stress drop, NR(t) is the cumulative number of ruptured subfaults at time t, and M 0ave M 0 /N is the average seismic moment of subfaults. Rij is the distance from the observation point, Q( f )is the quality factor, G(Rij ) is the geometric spreading factor, D( f ) is the site amplification term, and e f is a high‐cut filter to include the spectral decay at high frequencies described with the κ factor of soils [Anderson and Hough, 1984]. Hij is a scaling factor introduced to conserve the high‐frequency spectral level of the subfaults. 10.3.2. Simulations Along Eastern Segments of NAFZ In this study, we perform ground‐motion simulations for scenario earthquakes of size Mw = 6.0, 6.5, and 7.0 on the same fault where the 1992 Erzincan event (Mw = 6.6) occurred (Fig. 10.1b). We also simulated the 1992 event. Since the epicenter of the 1992 earthquake was critical in terms of its close distance to the city center, the epicenter of the earthquake is kept the same for all scenario events. A total of 225 nodes is selected in the Erzincan region where full waveforms are simulated. Among these nodes, 125 are located within the city center. Each node in the city center represents a small box of size 1 km by 1 km while a coarser mesh of 10 km by 10 km is employed for the nodes that remain outside the city center. The source, path, and site parameters are selected according to a

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  277

previous study by the authors of this chapter [Askan et  al., 2013], where validations are performed in terms of simulated and observed ground‐motion time histo‑ ries of the 1992 Erzincan earthquake. The simulation parameters are displayed in Table 10.1. When the source parameters, such as fault dimensions and stress drop, are concerned, in this work for different scenarios, Mw is scaled with the length and width of the fault using the empirical relationships by Wells and Coppersmith [1994]. We also employed an empirical model, which assumes a varying stress drop as a function of fault width by Mohammadioun and Serva [2001].

Figures  10.2–10.5 display the spatial distribution of Peak Ground Acceleration (PGA) and Peak Ground Velocity (PGV) in Erzincan region for the 1992 Erzincan earthquake (Mw = 6.6) as well as scenario events. The shear‐wave velocity profiles at selected nodes were obtained by a microtremor array method as explained in detail in Sisman et al. [2013]. It is observed that at Mw = 6.0, the maximum PGA in the region is around 0.45 g with relatively smaller ampli‑ tudes (~0.3 g) in the Erzincan city center. For Mw = 6.5, the city center is subjected to PGA values around 0.6 g with larger amplitudes close to the fault plane (~0.7 g). For the

Table 10.1  Parameters Used in the Simulation of Scenario Earthquakes Parameter

Value

Hypocenter location Hypocenter depth Depth to the top of the fault plane Fault orientation Fault dimensions Crustal shear wave velocity Rupture velocity Crustal density Stress drop Quality factor Geometrical spreading

39°42.3 N, 39°35.2 E 9 km 2 km Strike: 125°, Dip: 90° Wells and Coppersmith [1994] 3700 m/sec 3000 m/sec 2800 kg/m3 Mohammadioun and Serva [2001] Q 122f 0.68 R 1.1, R 25 km R 0.5 , R 25 km T T0 0.05R Saragoni‐Hart Regional kappa model (κ0 = 0.066) Local model at each station [Sisman et al., 2013]

Duration model Windowing function Kappa factor Site amplification factors

Figure 10.2  Spatial distribution of the simulated (a) PGA and (b) PGV values of the 1992 Erzincan earthquake in Erzincan region. The rectangle shows the Erzincan city center.

278  ACTIVE GLOBAL SEISMOLOGY

Figure 10.3  Spatial distribution of the simulated (a) PGA and (b) PGV values of the scenario earthquake with Mw = 6.0 in Erzincan region. The rectangle shows the Erzincan city center.

Figure 10.4  Spatial distribution of the simulated (a) PGA and (b) PGV values of the scenario earthquake with Mw = 6.5 in Erzincan region. The rectangle shows the Erzincan city center.

largest scenario earthquake with Mw = 7.0, Erzincan region experiences very high ground‐motion levels with PGA values exceeding 1 g at several locations. In the same sce‑ nario event, significant ground‐motion levels are observed in the city center with PGA values that reach 1 g. These ground‐motion amplitudes can be explained by very close distances from the fault plane and the soft‐soil conditions. As mentioned before, the city center is located on a basin of deep soft‐soil deposits. The soil media is relatively stiffer in the north regions, however, the fault plane is close

to those nodes located in the north resulting in overall higher amplitudes. The PGV values are also observed to be consistently higher for larger scenario events. When the anticipated ground motions of the 1992 Erzincan earthquake are studied, we observe that the PGA values reach almost 1 g around the city center. The corresponding PGV values reach 90 cm/sec. These amplitudes explain the widespread damage observed in the residential buildings in the city center during the earthquake despite the moderate magnitude.

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  279

Figure 10.5  Spatial distribution of the simulated (a) PGA and (b) PGV values of the scenario earthquake with Mw = 7.0 in Erzincan region. The rectangle shows the Erzincan city center.

10.4. RELATIONSHIPS BETWEEN PEAK GROUND‐MOTION PARAMETERS AND FELT‐INTENSITY VALUES Correlations between measured (instrumental) ground‐ motion parameters and felt‐intensity values are crucial for several purposes. A major use of such relationships is the preparation of intensity maps for real or scenario earthquakes in order to assess the effects of the earth‑ quake and indirectly the spatial distribution of potential seismic damage. Until recently, there was no such study performed for earthquakes that occurred in Turkey. It is possible to use the correlations from studies performed elsewhere, however, since the local properties of the build‑ ings and observed ground‐motion characteristics affect such relationships, local correlations should be derived. In a previous study by the authors of this chapter [Bilal and Askan, 2014], local correlations between MMI and PGA as well as PGV are derived for Turkey. Regional ground‐motion and intensity datasets from previous earthquakes along the North Anatolian fault zone were employed. A total of 92 data pairs from 14 earthquakes of 5.7 ≤ Mw ≤ 7.4 were used in standard linear regression analyses. To check the validity of linear regressions, the residuals were shown to be normally distributed through probability plots [Bilal and Askan, 2014]. The simple linear relationships between MMI and PGA as well as PGV are given as follows:

MMI

0.132 3.884 * log PGA (10.2)



MMI

2.673 4.340 * log PGV (10.3)

In Bilal and Askan [2014], conversion equations (10.2) and (10.3) were compared to other similar relationships from other regions in the world. It was observed that the correlations from other regions clearly underesti‑ mate the observed MMI levels in Turkey particularly for higher ground‐motion amplitudes. This observation aids to conclude that intensity maps must be prepared with local equations. Finally, damage patterns monitored in the past events revealed that PGA and PGV correlate well with stiffer structures and flexible structures, respec‑ tively [Erberik, 2008a, 2008b]. Since the governing struc‑ tural types in Turkey are masonry and reinforced concrete buildings, both of the correlations given in equations (10.2) and (10.3) are necessary depending on the region of interest. 10.5. APPLICATIONS IN THE STUDY REGION We employ equations (10.2) and (10.3) to convert the PGA and PGV values in Figures  10.2–10.5 into MMI. Since both PGA and PGV are of interest to the existing structural types in Turkey, we present two intensity maps (one converted from PGA and the other from PGV) for each scenario earthquake. We select the same region as in Figures 10.2–10.5 to present the computed intensity maps. Before simulating the scenario earthquakes, we first compare the observed versus computed intensity map of the 1992 Erzincan earthquake. Figure 10.6.a and b are the intensity maps prepared in this study, respectively, from PGA and PGV. Figure  10.6c displays the ShakeMap prepared by the United States Geological Survey (USGS) for the same event. The observed intensity

280  ACTIVE GLOBAL SEISMOLOGY

Figure 10.6  Seismic intensity map of the 1992 Erzincan earthquake in terms of MMI scale (a) obtained using the MMI‐PGA correlation given in equation (10.2); (b) obtained using the correlation given in equation (10.3); (c)  prepared by the USGS ShakeMap software [adapted from http://earthquake.usgs.gov/earthquakes/shakemap/ atlas/shake/199203131718/]; (d) prepared in the field by Turkish Ministry of Construction [Unal et al., 1993]. The rectangle in (a) and (b) shows the Erzincan city center.

map of the earthquake prepared by the Turkish Ministry of Construction in the field after the event is presented in Figure 10.6d [Unal et al., 1993]. Through the comparison of these figures within a region enclosed by 39°N–40°N and 39°E–40°E, it is observed that the spatial distribution of the MMI values computed in this study (Fig.  10.6a and b) are in the same range as the USGS ShakeMap

(Fig.  10.6c) with MMI values of X and above in the northeastern parts of the study region. In particular, the intensity map obtained from MMI‐PGV correlation yields very close estimates to the intensity values on the USGS map. Main reason of this observation is that most of the damaged buildings are reinforced concrete struc‑ tures whose damage patterns correlate well with PGV

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  281

Figure 10.6  (Continued)

282  ACTIVE GLOBAL SEISMOLOGY

[Erberik, 2008b]. The observed intensity map prepared by the Turkish Ministry of Construction reports the maxi‑ mum MMI value to be IX in the city center. Most of the observed intensity values in the study region are one unit less than the USGS map and the maps prepared in this study. This is most probably due to the inherent subjectivity involved with assigning intensity values to observation points in the field. However, Figures 10.6a–c cannot also be accepted as “reference” maps due to the modeling

errors in both simulations and regressions. Finally, despite the mentioned discrepancies, the overall spatial distri‑ butions of the intensity values from Figures 10.6a–d are found to be consistent with each other. Next, we study the intensity maps for the scenario events of Mw = 6.0, 6.5, and 7.0. For the event of Mw = 6.0, we observe that intensity maps from PGA and PGV yield MMI values apart from each other approximately by one unit (Fig. 10.7a and b). The city center is exposed to

Figure 10.7  Synthetic intensity map of the scenario event with Mw = 6.0 in Erzincan region using (a) equation (10.2) and (b) equation (10.3). The rectangle shows the Erzincan city center.

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  283

Figure 10.8  Synthetic intensity map of the scenario event with Mw = 6.5 in Erzincan region using (a) equation (10.2) and (b) equation (10.3). The rectangle shows the Erzincan city center.

MMI values of VII and VIII for this event, which has a moderate magnitude. Figure 10.8a shows that the MMI values of the scenario event with Mw = 6.5 are close to those anticipated in the 1992 earthquake (Fig.  10.6a) when PGA is used as the main parameter. However, when the MMI‐PGV correlation is used, the intensity map for the Mw = 6.5 scenario (Fig. 10.8b) shows values that are one unit less than those in Fig. 10.8a. We note that the MMI‐PGV relationship has a larger standard deviation than that of MMI‐PGA relationship [Bilal and Askan,

2014], which could possibly explain the discrepancies between Fig. 8a and b. The largest scenario with Mw = 7.0 yields extremely severe shaking intensities of X and above in the city center (Fig.  10.9). For all of the cases, it is observed that the northern part of Erzincan is exposed to higher MMI values compared with the southern part since the fault lies in the north. Finally, we note that as the magnitude increases, the MMI values obtained from PGA and PGV converge to each other. One possible reason for better performance of the equations for large

284  ACTIVE GLOBAL SEISMOLOGY

Figure 10.9  Synthetic intensity map of the scenario event with Mw = 7.0 in Erzincan region using (a) equation (10.2) and (b) equation (10.3). The rectangle shows the Erzincan city center.

events is that the original relationships between MMI and PGA as well as PGV are derived using data from mostly large events [Bilal and Askan, 2014]. 10.6. CONCLUSIONS In this study, synthetic intensity maps are presented for the eastern section of the North Anatolian fault zone for a selected set of scenario earthquakes. These parts of the NAFZ are sparsely monitored and less investigated

compared with the western parts, which include the industrial facilities of the country and the most popu‑ lated city Istanbul. However, the Erzincan region in the east, located at the intersection of three fault systems, has always been one of the most hazardous regions in the world and experienced one of the largest events of the last century (Ms ~8.0). Thus, in this study, we aimed to study the seismicity in the region through ground‐motion simu‑ lations performed for potential scenario earthquakes of various magnitudes. Even though the peak ground‐motion

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  285

parameters are important for earth scientists and earth‑ quake engineers, felt‐intensity values are globally employed for rapid response and disaster management purposes. Thus, the shaking intensities in terms of MMI scale were also computed from correlations between MMI and PGA as well as PGV. The spatial distribution of peak ground‐motion param‑ eters around the city center reveals that Erzincan is under significant seismic threat due to its close distance from the fault system in the north. In addition, Erzincan city center is in a pull‐apart basin underlain by soft sediments, which significantly amplify the ground‐motion levels. When the shaking‐intensity values are assessed in the synthetic MMI maps, we observed that even for a potential event with moderate magnitude (Mw = 6.0), MMI values of VIII are expected in some parts of the city center. Expectedly, larger events have the potential of MMI values of X and above. The spatial distribution of both peak ground‐motions parameters and MMI values demonstrate that the soft‐soil sites as well near‐fault sites have higher damage potential. It must be noted that the results presented herein contain inherent uncertainties arising from various sources. First, modeling errors are involved with the ground‐motion simulations in terms of selected input model parameters as well as the inherent limitations of the method employed [Motazedian and Atkinson, 2005]. Second, the conversion process of PGA and PGV to MMI values involves standard errors of regression. Another major source of uncertainty comes from the collection of intensity data in the field. Thus, even though the results presented here are obtained from physical and realistic models, they should be evalu‑ ated carefully in the existence of the mentioned inherent model and data uncertainties. Finally, based on the results presented here, we recom‑ mend evaluation of the existing buildings in the city for their seismic capacities before another large event occurs, in order to avoid future economic and structural losses as well as fatalities. On a larger scale, seismic design and construction practices should be improved as much as possible in the near future to eliminate the typical damage patterns that have been observed in the past recent earth‑ quakes in Turkey. This study and similar studies are important for future applications in rapid response, disaster mitigation, and estimation of the insurance premiums in earthquake‐ prone regions all over the world. As the ground‐motion and felt‐intensity databases expand, these studies should be updated and improved. ACKNOWLEDGMENTS This study is partially funded by Turkish National Geodesy and Geophysics Union through the project with grant number: TUJJB‐UDP‐01‐12.

REFERENCES Aktar, M., C. Dorbath, and E. Arpat (2004), The seismic velocity and fault structure of the Erzincan basin, Turkey, using local earthquake tomography, Geophys. J. Int., 156 (3), 497–505. Akyuz, H. S., R. Hartleb, A. Barka, E. Altunel, G. Sunal, B. Meyer, and R. Armijo (2002), Surface rupture and slip distri‑ bution of the 12 November 1999 Düzce earthquake (M 7.1), North Anatolian fault, Bolu, Turkey, Bull Seism. Soc. Amer., 92, 61–66. Ambraseys, N. N. (1970), Some characteristic features of the North Anatolian fault zone, Tectonophysics, 9, 143–165. Anderson, J. G., and S.E. Hough (1984), A model for the shape of the Fourier amplitude spectrum of acceleration at high frequencies, Bull Seism. Soc. Amer., 74, 1969–1993. Armijo, R., B. Meyer, A. A. Barka, and A. Hubert (1999), Propagation of the North Anatolian fault into the Northern Aegean: timing and kinematics, Geology, 27, 267–270. Armijo, R., B. Meyer, S. Navarro, G. King, and A. Barka (2002), Asymmetric slip partitioning in the Sea of Marmara pull‐ apart: A clue to propagation processes of the North Anatolian fault, Terra Nova, 14(2), 80–86. Askan, A., F. N. Sisman, and B. Ugurhan, (2013), Stochastic strong ground motion simulations in sparsely‐monitored regions: A validation and sensitivity study on the 13 March 1992 Erzincan (Turkey) earthquake, Soil Dynam. Earth. Eng., 55, 170–181. Avsar, U., E., Turkoglu, M. Unsworth, I. Caglar, and B. Kaypak, (2013), Geophysical images of the North Anatolian fault zone in the Erzincan basin, eastern Turkey, and their tectonic implications, Pure Appl. Geophys., 170 (3), 409–431. Barka, A. (1992), The North Anatolian fault zone, Ann. Tectonicae, 6, 164–195. Barka, A. (1993), The tectonics of Erzincan basin and 13 March 1992 Erzincan earthquake, Proceedings of the 2nd Turkish National Earthquake Engineering Conference, 259–270, Istanbul Technical University Structures and Earthquake Applications‐Research Center (in Turkish). Barka, A. (1996), Slip distribution along the North Anatolian fault associated with the large earthquakes of the period 1939 to 1967, Bull Seism. Soc. Amer., 86, 1238–1254. Beresnev, I., and G. M. Atkinson, (1997), Modeling finite‐fault radiation from the ωn spectrum, Bull. Seism. Soc. Amer., 87, 67–84. Beresnev, I. A., and G. M. Atkinson (1998), FINSIM: a FORTRAN program for simulating stochastic accelera‑ tion time histories from finite faults, Seism. Res. Lett., 69, 27–32. Bernard, P., J. C. Gariel, and L. Dorbath, (1997), Fault location and rupture kinematics of the magnitude 6.8, 1992 Erzincan earthquake, Turkey, from strong ground motion and regional records, Bull. Seism. Soc. Amer., 87, 1230–1243. Bilal, M., and A. Askan (2014), Relationships between felt ıntensity and recorded ground‐motion parameters for Turkey, Bull Seism. Soc. Amer., 104, 484–496. Boore, D. M. (1983), Stochastic simulation of high‐frequency ground motions based on seismological models of the radiated spectra, Bull. Seism. Soc. Amer., 73, 1865–1894.

286  ACTIVE GLOBAL SEISMOLOGY Bouchon, M. (1981), A simple method to calculate Green’s functions for elastic layered media, Bull. Seism. Soc. Amer., 71, 959– 971. Castro, R. R., A. Rovelli, M. Cocco, M. Di Bona, and F. Pacor (2001), Stochastic simulation of strong‐motion records from the 26 September 1997 (Mw 6), Umbria–Marche (central Italy) earthquake, Bull. Seism. Soc. Amer., 91, 27–39. Daniell, J. E., B. Khazai, F. Wenzel, and A. Vervaeck (2011), The CATDAT damaging earthquakes database, Nat. Hazards Earth Syst. Sci., 11, 2235–2251. Durukal, E. (2002), Critical evaluation of strong motion in Kocaeli and Duzce (Turkey) earthquakes, Soil Dynam. Earth. Eng., 22, 589–609. Erberik, M. A. (2008a), Generation of fragility curves for Turkish masonry buildings considering in‐plane failure modes, Earth. Eng. Struct. Dyn., 37, 387–405. Erberik, M. A. (2008b), Fragility‐based assessment of typical mid‐rise and low‐rise RC buildings in Turkey, Eng. Struct., 30, 1360–1374. Faenza, L., and A. Michelini (2010), Regression analysis of MCS intensity and ground motion parameters in Italy and its application in ShakeMap, Geophys. J. Int., 180, 1138–1152. Frankel, A. (1993), Three‐dimensional simulations of the ground motions in the San Bernardino valley, California, for hypothetical earthquakes on the San Andreas fault, Bull. Seism. Soc. Amer., 83, 1020–1041. Gallovic, F., and J. Brokesova (2007), Hybrid k‐squared Source Model for Strong Ground Motion Simulations: Introduction, Phys. Earth Planet. Inter., 160, 34–50. Gokalp, H. (2007), Local earthquake tomography of the Erzincan basin and the surrounding area in Turkey, Ann. Geophys., 50 (6), 707–724. Heaton, T. H., and S. H. Hartzell (1989), Estimation of strong ground motions from hypothetical earthquakes on the Cascadia subduction zone Pacific Northwest, Pageoph, 129, 131–201. Hisada, Y. (2008), Broadband strong motion simulation in layered half‐space using stochastic Green’s function technique, J. Seis., 12, 265–279. Irikura, K. (1983), Semi‐empirical estimation of strong ground motions during large earthquakes, Bull. Dis. Prev. Res. Inst., 33, 63–104. Irikura, K., and K. Kamae, (1994), Estimation of strong ground motion in broad‐frequency band based on a seismic source scaling model and an empirical Green’s function technique, Ann. Geofisica, 37, 1721–1743. Kaka, S. L., and G. M. Atkinson, (2004), Relationships between instrumental ground‐motion parameters and modified Mercalli intensity in eastern North America, Bull. Seism. Soc. Amer., 94, 1728–1736. Kamae, K., K., Irikura, and A. Pitarka, (1998), A technique for simulating strong ground motion using Hybrid Green’s func‑ tion, Bull. Seism. Soc. Amer. 88, 357–367. Ketin, I. (1969), Uber die nordanatolische Horizontalverschiebung, Bull. Mineral Res. Explor. Inst. Turkey, 72, 1–28, 1969. Mohammadioun, B., and L. Serva (2001), Stress drop, slip type, earthquake magnitude, and seismic hazard, Bull. Seism. Soc. Am., 91 (4), 694-707.

Motazedian, D., and A. Moinfar (2006), Hybrid stochastic finite fault modeling of 2003, M 6.5, Bam, earthquake (Iran), J. Seism., 10, 91–103. Motazedian, D., and G. M. Atkinson (2005), Stochastic finite‐ fault modeling based on a dynamic corner frequency, Bull. Seism. Soc. Amer., 95, 995–1010. Olsen, K. B., R. J., Archuleta, and J. R. Matarese, (1996), Three‐ dimensional simulation of a magnitude 7.75 earthquake on the San Andreas fault, Science, 270, 1628–1632. Pulido N, A, Ojeda, K. Atakan, and T. Kubo (2004), Strong ground motion estimation in the Sea of Marmara region (Turkey) based on a scenario earthquake, Tectonophysics, 391, 357–74. Sarp, G. (2015), Tectonic controls of the North Anatolian fault System (NAFS) on the geomorphic evolution of the alluvial fans and fan catchments in Erzincan pull‐apart basin; Turkey, J. Asian Earth Sci., 98, 116–125. Sengor, A. M. C. (1979), The North Anatolian transform fault: its age, offset and tectonic significance, J. Geological Society London, 136, 269–282. Sengor, A. M. C., and N. Canitez (1982), The North Anatolian fault, in Alpine and Mediterranean Geodynamics, edited by H. Berckhemer and K. Hsu, 205–216, Geodynamical Series, American Geophysical Union, 7. Sengor, A. M. C., O. Tuysuz, C. Imren, M. Sakinc, H. Eyidogan, N. Gorur, X. Le Pichon, and C. Rangin (2005), The North Anatolian fault: A new look, Ann. Rev. Earth Planet. Sci., 33, 37–112. Sisman, F. N, M. Asten, and A. Askan (2013), Determination of the site characterization properties in eastern Segment of the North Anatolian fault zone in Turkey based on the MMSPAC method, ASEG Extended Abstracts 2013, 23rd Geophysical Conference, 1–4. Sørensen B. M., N. Pulido, and K. Atakan (2007), Sensitivity of ground‐motion simulations to earthquake source para­ meters: A case study for Istanbul, Turkey. Bull. Seism. Soc. Amer., 97(3), 881–900. Stein, R. S., A. Barka, and J. H. Dieterich (1997), Progressive failure on the North Anatolian fault since 1939 by earth‑ quake stress triggering, Geophys. J. Int., 128, 594–604. Tanırcan, G. (2012), Ground Motion Simulation for Istanbul with Three‐dimensional Velocity Model, J. Fac. Eng. Arch. Gazi Univ., 27 (1), 27–35 (in Turkish). Tatar, O., Z. Akpinar, H. Gursoy, J. D. A. Piper, F. Kocbulut, B. L. Mesci, A. Polat, and A. P. Roberts (2013), Palaeomagnetic evidence for the neotectonic evolution of the Erzincan basin North Anatolian fault zone, Turkey, J. Geod., 65, 244–258. Toksoz, M. N., A. F. Shakal, and J. Michael (1979), Space‐time migration of earthquakes along the North Anatolian fault zone and seismic gaps, Pageoph, 117, 1258–1270. Tselentis, G., and L. Danciu (2008), Empirical relationships between modified Mercalli intensity and engineering ground‐motion parameters in Greece, Bull. Seism. Soc. Amer., 98, 1863–1875. Ugurhan, B., and A. Askan, (2010), Stochastic strong ground motion simulation of the 12 November 1999 Düzce (Turkey) earthquake using a dynamic corner frequency approach, Bull Seism. Soc. Amer., 100, 1498–1512. Ugurhan, B., A. Askan, A. Akinci, and L. Malagnini (2012), Strong‐ground‐motion simulation of the 6 April 2009

SEISMIC INTENSITY MAPS FOR THE EASTERN PART OF THE NORTH ANATOLIAN FAULT ZONE  287 L’Aquila, Italy, earthquake, Bull. Seism. Soc. Amer., 102, 1429–1445. Unal, E., H. Cindemir, O. Ergunay, M. Yildiz, H. Kozaci, S. Tekeli, R. Ceylan, and E. Aytac (1993), 13 March 1992 Erzincan earthquake, Turkish Ministry of Construction (Technical report in Turkish). Wald, D. J., V. Quitoriano, L. Dengler, and J. Dewey (1999a), Utilization of the Internet for rapid community intensity maps, Seism. Res. Lett., 70, 680–697. Wald, D. J., V. Quitoriano, T. H. Heaton, and H. Kanamori (1999b), Relationships between peak ground acceleration,

peak ground velocity and modified Mercalli intensity in California, Earth. Spectra, 15, 557–564. Wells, D. L., and K. J. Coppersmith (1994), New Empirical Relationships among Magnitude, Rupture Length, Rupture Width, Rupture Area and Surface Displacement, Bull. Seism. Soc. Am., 84 (4), 974–1002. Yalcinkaya, E. (2005), Stochastic finite‐fault modeling of ground motions from the June 27, 1998, Adana‐Ceyhan earthquake, Earth Planet Space, 57, 107–115; http:// earthquake.usgs.gov/earthquakes/shakemap/atlas/shake/ 199203131718/.

INDEX Accretionary prisms, 121–25, 123f Advanced Industrial Science and Technology, Japan (AIST), 249 Aegean region. See also Çataldağ Plutonic Complex; Western Anatolia boarder region of, 250–55, 250f–251f, 253f–254f, 256f CAT in, 245f, 247, 247f, 259 Corinth Rift in, 241, 242f–243f, 245f fault inventories in, 249–50 geodynamic setting of, 242–44, 242f–243f North Aegean Sea in, 244–47, 245f–247f Northern Greece in, 244, 245f, 247–48, 248f geology of, 141–45, 142f, 154, 169–70 geotectonic setting of, 239–42, 240f, 242f gravity modeling of, 107f, 109f–110f, 111–15, 113f–115f seismogenic sources of, 248–49 strike‐slip fault zones in, 244–47, 246f, 252–55, 256f, 259–62 tectonic map of, 190, 190f 2014 earthquake in conclusions on, 262–63 Coulomb Stress change in, 259–62, 261f displacement pattern of, 262, 262f fault models of, 258f, 258t, 260 focal mechanisms of, 257f, 257t, 259, 259f introduction to, 255, 257f, 257t prior seismicity of, 258–59, 258t African Plate, 1, 12f, 19, 107, 116, 242–44 Aghios Efstratios island, 255, 256f AIST. See Advanced Industrial Science and Technology Aki, K., 260 Aksoy, E., 56 Aladağ Mountain, 39f, 40 Alaşehir graben, 105–8, 106f, 116 Alaska subduction, 123, 123f Alboran Sea–Betics, 121, 122f Alborz Mountains, Iran, 99 Alpine fault, New Zealand, 101 Altiplano‐Puna plateau, 121, 122f Altunkaynak, S., 216 Amanos Mountains, 15, 15f–17f, 20–28, 21f–22f, 25f, 55 Ambraseys, N. N., 252 Amik depression, 15f, 21f–22f, 23–26, 24f, 55f Anatolia African Plate in, 12f, 19 Amanos Mountains of, 15, 15f–17f, 20–28, 21f–22f, 25f, 55

Amik depression of, 15f, 21f–22f, 23–26, 24f, 55f Arabian Plate in, 11–19, 18f, 23, 29, 36f, 38, 42, 74–75 Bitlis‐Zagros suture mountains in, 16f–17f, 29f, 33f, 36f, 42, 55f Black Sea in, 14, 36f, 42–46, 45f, 53, 56–62, 57f, 64f–65f Bosporus Channel in, 58f, 61–62, 65f CAFZ in, 39–40, 39f central plateau of, 37–42, 38f–39f Deliçay fault zone of, 16f, 18f, 20f, 27–28 DSF of, 12f, 15f, 17f–18f, 19–28, 21f–22f, 24f, 55f east high plateau of, 29–37, 29f–37f EATF in, 55–56, 55f–56f Eskişehir fault in, 39f, 40–41, 75–76 Gediz graben in, 12f, 66f, 70–71, 70f, 73f gravity modeling of, 107f, 109f–110f, 111–15, 113f–115f introduction to, 11–14, 12f–13f Karasu depression of, 15f, 23–28, 55–56 Keldağ horst in, 15f, 21f, 24f, 26–27 Kırşehir‐Niğde Massif of, 37–40, 38f, 75 Maraş triple junction in, 15f–18f, 23, 27–28, 55, 55f Marmara region in, 56–66, 57f–58f, 60f–65f Misis‐Andırın fault zone in, 12f, 17f, 23, 47–48 morphotectonic map of, 12f NATF in, 49–55, 50f–52f, 54f–55f NeoTethyan Ocean in, 16, 29, 36f, 42, 74–75 Oltu fault zone of, 32, 37f peripheral mountains of, 42–49, 43f–45f, 47f–48f Söke graben, 66f, 70f, 71–74, 73f Solhan volcano of, 33f, 35 southeast orogenic belt of, 14–29, 15f–17f, 20f–22f, 24f–25f, 29f summary on, 74–76 Taurus Mountains in, 17f, 38f–39f, 41–42, 46–49, 47f, 48f, 66f Türkoğlu‐Haruniye fault zone in, 18f, 20f, 27–28 Uludağ Mountain of, 41, 57f–58f, 60f, 76 western region of, 66–74, 66f–67f, 69f–73f Anatolian Plate. See Western Anatolia Anatolian Scholle, 100–101 Andırın‐Misis fault zone, 12f, 17f, 23, 47–48 ANIMAL. See Auburn University Noble Isotope Mass Analysis Laboratory Antithetic approach, 97, 97f Antithetic pull‐apart basins, 99–100, 99f–100f

Active Global Seismology: Neotectonics and Earthquake Potential of the Eastern Mediterranean Region, Geophysical Monograph 225, First Edition. İbrahim Çemen and Yücel Yılmaz © 2017 American Geophysical Union. Published 2017 by John Wiley & Sons, Inc. 289

290 INDEX Apennines, 121, 122f Arabian Plate in geodynamical modeling, 122, 124 morphotectonic development of, 11–19, 18f, 23, 29, 36f, 38, 42, 74–75 in NATF, 49 in western Anatolian geology, 169 Ar/Ar geochronology analytical methods of, 203–5 results of, 205, 206t–213t, 214f–215f Asi River, 24f, 26–27 Askan, A., 279 Asthenospheric material, 4, 105–16, 110f, 111t, 115f Ates, A., 114 Atkinson, G. M., 276 Auburn University Noble Isotope Mass Analysis Laboratory (ANIMAL), 203–5 Augen‐gneiss, 194, 195f Back‐arc spreading, 106, 169–170 Back‐scattered images (BSE), 199, 204f Bayrakçı, G., 233 Bergama graben, 163f–164f, 167 Beyhan, A., 40 Bilal, M., 285 Bitlis Poturge Massif (BPM), 33 Bitlis‐Zagros orogenic belt, 13, 18f, 74, 125 Bitlis‐Zagros suture mountains, 16f–17f, 29f, 33f, 36f, 42, 55f Black Sea, 14, 36f, 42–46, 45f, 53, 56–62, 57f, 64f–65f Block rotation, 227 Bodrum caldera, 156, 157f Boğaziçi University, 112 Bornova flysch, 191–94 Bosporus Channel, 58f, 61–62, 65f Bouchon, M., 226 Bouguer gravity anomalies, 108–9, 109f Boulton, S. J., 26 Bozdağ horst, 148–53, 162–67 Bozkurt, E., 167–69 Boztuğ, D., 46, 216–17 BPM. See Bitlis Poturge Massif BSE. See Back‐scattered images Buck, W. R., 171 Bulgaria, 141 Bulut, M., 227 Büyük Menderes grabens, 12f, 70–71, 73f gravity modeling of, 105–8, 106f, 111, 116 CAFZ. See Central Anatolian fault zone Canitez, N., 49–50 Caputo, R., 249, 250, 250f Carpathians, 121, 122f CAT. See Central Aegean trough Çataldağ detachment fault zone (ÇDFZ) introduction to, 190–91, 190f structural data of, 194–97, 195f–197f Çataldağ granodiorite (ÇG)

Ar/Ar geochronology of, 191–94, 194f–195f, 203–5, 206t–213t, 214f–215f deformation of, 198–99, 199f emplacement of, 214–16, 216f introduction to, 5 Çataldağ Plutonic Complex (ÇPC). See also Western Anatolia abstract on, 189 ÇDFZ in, 190–91, 190f, 194–97, 195f–197f Cenozoic crustal buildup of, 217–18 ÇG in, 191–94, 194f–195f Ar/Ar geochronology of, 203–5, 206t–213t, 214f–215f deformation of, 198–99, 199f emplacement of, 214–16, 216f conclusion on, 218 GGMC in, 191, 192f–194f deformation of, 197–98, 198f–200f emplacement of, 204f, 205–14 LA‐ICP‐MS age data on, 199–203, 201t–203t, 204f introduction to, 5, 189–91, 190f K‐feldspar in, 191–93, 192f–194f, 203–5, 209t, 214 LA‐ICP‐MS age data on, 199–203, 201t–203t, 204f monazite radiometric age data on, 200–203, 201t–203t, 204f SCIH in, 205–14, 206t–212t, 214f–215f SCTF in, 205–14, 206t–213t, 214f–215f timing of uplift of, 216–17 Zircon radiometric age data, 200–203, 201t–203t, 204f Cathodoluminescence (CL), 199, 204f ÇDFZ. See Çataldağ detachment fault zone Cenozoic sediments, 113f–114f, 114–18 Central Aegean trough (CAT), 245f, 247, 247f, 259 Central Anatolian fault zone (CAFZ), 39–40, 39f ÇG. See Çataldağ granodiorite Christensen, N., 111 Çiftçi, N. B., 168 Çine horst, 152, 152f Circum‐Rhodope belt, 19, 241 CL. See Cathodoluminescence Colorado plateau, 121, 122f Comastri, A., 252 Composite Seismogenic Sources (CSSs), 6, 250, 250f Aghios Efstratios, 255 Lemnos, 6, 251, 255, 262–63 North Aegean Basin, 251f, 252 North NAT, 253–54, 254f South Chalkidiki Offshore, 251–52, 251f South NAT, 251f, 252–53 Continental delamination. See Lithospheric delamination Convergence velocity, 126–36, 126f–127f, 133f–134f Copley, A., 136 Coppersmith, J. K., 254, 260 Corinth Rift, 241, 242f–243f, 245f ÇPC. See Çataldağ Plutonic Complex CSS. See Composite Seismogenic Sources Cyclades, 189 Database of Individual Seismogenic Sources (DISS), 249–50 Dead Sea transform fault zone (DSF), 12f, 15f, 17f–18f, 19–28, 21f–22f, 24f, 55f

INDEX  291 Delamination. See Lithospheric delamination Deliçay fault zone, 16f, 18f, 20f, 27–28 Dhont, D., 47 Dilek, T., 169 Dilek nappe, 145, 149f, 151, 151f Dirik, K., 39 DISS. See Database of Individual Seismogenic Sources Dry Olivine Mantle, 125t DSF. See Dead Sea transform fault zone Duman, T. Y., 23 Düzce earthquake (1999), 274, 275f EAAC. See East Anatolian accretionary complex EAFZ. See East Anatolian fault zone Earthquakes Düzce (1999), 274, 275f Erzincan (1939), 273–74, 275f Erzincan (1992), 273–85, 275f, 277f–284f, 277t Ganos (1912), 231 Izmit and Duzce (1999), 225–27, 232, 233f Kocaeli (1999), 274, 275f North Aegean (2014) conclusions on, 262–63 Coulomb Stress change in, 259–62, 261f displacement pattern of, 262, 262f fault models of, 258f, 258t, 260 focal mechanisms of, 257f, 257t, 259, 259f introduction to, 255, 257f, 257t prior seismicity of, 258–59, 258t Saros (2013), 231 Skyros (2001), 247, 247f East Anatolian accretionary complex (EAAC), 128–29, 130f East Anatolian fault zone (EAFZ) fluvial geomorphology of, 93–101, 94f–100f gravity modeling of, 106 East Anatolian high plateau, 29–37, 29f–37f East Anatolian transform fault (EATF), 12, 55–56, 55f–56f East Anatolian uplift abstract on, 121 accretionary prisms, 121–25, 123f conclusions on, 135–36 convergence velocity in, 126–36, 126f–127f, 133f–134f introduction to, 121–23, 122f–123f LABMODEL of, 131–36, 131t, 133f–134f laboratory experiments for, 130–35, 131f, 131t, 133f–134f numerical experiments for, 124–30, 124f, 125t, 126f–130f NUMMODEL experiments in, 124–28, 124f, 125t, 126f–127f, 130f Eastern Mediterranean. See also Marmara Sea; North Aegean Trough; North Anatolian fault zone elevation map of, 2f overview of, 1–2 EATF. See East Anatolian transform fault EDSF. See European Database of Seismogenic Faults EIGEN. See European Improved Gravity model of Earth Elkins‐Tanton, L. T., 125 Emre, Ö., 23 Erinç, S., 94

Erol, O., 26, 39 Erzincan region earthquake of, 273–85, 275f, 277f–284f, 277t introduction to, 6 pull‐apart basin of, 93, 94f Erzurum‐Kars plateau, 129, 130f Eskişehir fault, 39f, 40–41, 75–76 Eurasian Plate, fluvial geomorphology of, 93–101, 94f–100f European Database of Seismogenic Faults (EDSF), 249 European Improved Gravity model of Earth (EIGEN), 108, 109f EXSIM program, 276 External Hellenides, 241, 244 Faccenna, C., 109 Fault plane solutions (FPS), 227, 233–36, 233f, 235f–236f Faults as Seismologist’s Tool (FAUST), 249 FBFZ. See Fethiye‐Burdur strike‐slip fault zone Feldspar‐fish, 195f, 198–99, 217–18 Felt‐intensity values, 279 Fernandez‐Blanco, D., 41 Fethiye‐Burdur strike‐slip fault zone (FBFZ), 106f, 108–11 Filyos, 95f Flake tectonics, 135 Fluvial geomorphology, 93–101, 94f–100f FPS. See Fault plane solutions Ganas, A., 259 Ganos earthquake (1912), 231 GBR. See Grain boundary migrations Gediz graben, 143f, 146f, 162–67, 165f morphotectonic development of, 12f, 66f, 70–71, 70f, 73f Geological National Survey, New Zealand (GNS), 249 GGMC. See Granite‐Gneiss‐Migmatite Complex GLITTER program, 200 Gneiss, 191–97, 192f–193f, 195f, 203–5, 215f. See also Granite‐ Gneiss‐Migmatite Complex GNS. See Geological National Survey GOCE. See Gravity Field and Steady‐State Ocean Circulation Explorer Gögüş, O. H., 121, 135–36 Gökçeada Island, 244, 252, 255, 262–63 Gölova pull‐apart basin, 95f Görgün, B., 255 Görgün, E., 255 GRACE. See Gravity Recovery and Climate Experiment Grain boundary migrations (GBR), 197–99, 214 Granite‐Gneiss‐Migmatite Complex (GGMC), 191, 192f–194f. See also Çataldağ Plutonic Complex deformation of, 197–98, 198f–200f emplacement of, 204f, 205–14 LA‐ICP‐MS age data on, 199–203, 201t–203t, 204f Gravity data modeling anomalies and analysis in, 108–11, 110f conclusions on, 115–16 databases in, 108, 109f initial construction of, 111–12, 111t introduction to, 105–8, 106f–107f

292 INDEX Gravity data modeling (cont’d ) of NAFZ, 106, 112 P‐wave velocity in, 111, 111t results and discussion on, 112–15, 113f–115f Gravity Field and Steady‐State Ocean Circulation Explorer (GOCE), 108 Gravity Recovery and Climate Experiment (GRACE), 108 GreDaSS. See Greek Database of Seismogenic Sources Gerede, Turkey, 94 Greek Database of Seismogenic Sources (GreDaSS), 6, 247–53, 250f–251f, 253f, 258, 258t, 260, 262 Greenland, 189 Ground‐motion simulation, 276–78, 277f–279f, 277t Guidoboni, E., 252 Hazro anticline structure, 15f, 19 Hellenic‐Cyprus trench conclusions on, 115–16 gravity data modeling of, 105–12, 106f–107f, 109f–110f, 111t results and discussion on, 112–15, 113f–115f Hellenides, 240f, 241–44, 250 HypoDD algorithm, 228, 231f, 234, 234f IA. See Isparta angle IAESZ. See Iżmir‐Ankara‐Erzincan suture zone Impend, C., 66 Individual Seismogenic Sources (ISSs) Aghios Efstratios, 255, 256f Athos, 251–52, 251f introduction to, 6 NAB Segment, 252 North Samothraki, 6, 254–55 Samothraki SE, 6, 253–54 Saros Gulf, 253–54, 254f INGV. See Istituto Nazionale di Geofisica e Volcanologia Internal Hellenides, 241 Iskenderun Bay, 27 ISOPLOT software, 200 Isparta angle (IA), 12f, 39f, 41, 47–49, 47f, 106f, 108, 158 ISSs. See Individual Seismogenic Sources Istituto Nazionale di Geofisica e Volcanologia, Italy (INGV), 249 Istranca Massif, 153–54 Iżmir‐Ankara‐Erzincan suture zone (IAESZ), 190, 190f, 217 Iżmir‐Ankara suture zone gravity modeling of, 105–8, 106f in lithoshpere delamination, 129f in Western Anatolian geology, 141–45, 143f–145f, 171 Izmit and Duzce earthquake (1999), 225–27, 232, 233f Izmit‐Sapanca zone, 225–26. See also Marmara Sea Jackson, J., 136 Jull, M, 125 KAIVF. See Kirka‐Afyon‐Isparta volcanic field Kandilli Observatory and Earthquake Research Institute (KOERI), 226, 228, 229f, 231 Karabulut, H., 226

Karakaisis, G. F., 249 Karakaya complex, 191–94, 192f, 214 Karakostas, V., 259 Karasu depression, 15f, 23–28, 55–56 Karasu river, 93, 94f Karliova triple junction, 274, 275f Kastelic, V., 249 Kazdağ Massif, 199, 216–18 Keirogens discussion on, 100–101 gentle bending rivers in, 95–99, 96f–98f introduction to, 3, 93–94, 93f–96f pull‐apart basins in, 99–100, 99f–100f Keldağ horst, 15f, 21f, 24f, 26–27 Kelemen, P. B., 125 Keskin, M., 46, 47 K‐feldspar, 191–93, 192f–194f, 203–5, 209t, 214 Kidd, W. S. F., 31 Kirka‐Afyon‐Isparta volcanic field (KAIVF), 107f, 109–11 Kırşehir‐Niğde Massif, 37–40, 38f, 75 Kocaeli earthquake (1999), 274, 275f Koçyigit, A., 40 KOERI. See Kandilli Observatory and Earthquake Research Institute Komut, T., 170 Kozak horst, 156, 156f, 163f–164f, 168 Kula volcanic field (KVF), 107f, 109–11, 114 Kumburgaz basin, 235, 236f, 237 KVF. See Kula volcanic field Kyrenia‐Misis fault zone, 16f, 23 LABMODEL, of lithospheric delamination, 131–36, 131t, 133f–134f LAGEOS. See Laser Geodynamics Satellites Lagrangian‐Eulerian finite element technique, 124 Laser‐ablation inductively coupled plasma mass spectrometer (LA‐ICP‐MS), 199–203, 201t–203t, 204f Laser Geodynamics Satellites (LAGEOS), 108 Laser incremental heating analysis (LIH), 205 Latacia‐Antakya fault zone, 25 Lemnos composite seismogenic source, 6, 251, 255, 262–63 Lemnos island, 244, 252, 255, 256f, 259, 262 Le Pichon, X., 227, 229f, 233, 233f, 235 LIH. See Laser incremental heating analysis Lithospheric delamination abstract on, 121 accretionary prisms in, 121–25, 123f conclusions on, 135–36 convergence velocity in, 126–36, 126f–127f, 133f–134f introduction to, 4, 121–23, 122f–123f LABMODEL of, 131–36, 131t, 133f–134f laboratory experiments for, 130–35, 131f, 131t, 133f–134f numerical experiments for, 124–30, 124f, 125t, 126f–130f NUMMODEL experiments in, 124–28, 124f, 125t, 126f–127f, 130f Lycian nappes, 105, 150–54 Lyell, Charles, 189

INDEX  293 Macedonia, 244, 245f, 247–48, 248f, 250, 250f Mafic synplutonic dyke, 194f Makran subduction, 123, 123f Maraş triple junction, 15f–18f, 23, 27–28, 55, 55f Marmara region morphotectonic development of, 56–66, 57f–58f, 60f–65f in western Anatolian geology, 143–45, 143f, 153–57, 168–72 Marmara Sea, 101, 143f abstract to, 225 conclusions to, 236–37 fault zone complexities in, 231–36, 232f–236f FPS in, 227, 233–36, 233f, 235f–236f introduction to, 225–26 microseismic processes of, 226–28 nonseismic processes of, 228 seismological data of, 228–31, 229f–231f subparallel faults in, 227, 235 Maxwell relaxation time, 131 McClusky, S., 56 McKenzie, D., 49–50 Menderes Massif, 189, 218. See also Çataldağ Plutonic Complex conclusions on, 171–73 discussion of, 150–55, 153f–154f geology map of, 146f introduction to, 141–45, 142f–145f major problems of, 145–55, 146f–155f magmatic associations in, 155–60, 155f–158f Neogene cover rocks in, 160–69, 160f–166f north‐south extensional regime in, 169–71 ophiolite nappes in, 145, 147–51, 149f, 151f Sakarya continent in, 141–45, 143f–144f, 149f, 170–71 stratigraphic section of, 147f, 150f–151f tectonic evolution of, 148f–149f Menderes metamorphic core complex (MMCC), 105–8, 106f Microseismic processes, 226–28 Misis‐Andırın fault zone, 12f, 17f, 23, 47–48 MMCC. See Menderes metamorphic core complex Modified Mercalli Intensity (MMI), 6, 279–85, 280f Moho boundaries, 111 Moho depth map, 115 Moho variation, 128–30 Mohr‐Coulomb failure law, 124 Monazite radiometric age data, 200–203, 201t–203t, 204f Mooney, W., 111 Moores, E. M., 241 Morelli, A., 47 Morphotectonic development African Plate in, 1, 12f, 19 of Amanos Mountains, 15, 15f–17f, 20–28, 21f–22f, 25f, 55 of Amik depression, 15f, 21f–22f, 23–26, 24f, 55f of Anatolia peripheral mountains, 42–49, 43f–45f, 47f–48f of Arabian Plate, 11–19, 18f, 23, 29, 36f, 38, 42, 74–75 of Bitlis‐Zagros suture mountains, 16f–17f, 29f, 33f, 36f, 42, 55f of Black Sea, 14, 36f, 42–46, 45f, 53, 56–62, 57f, 64f–65f Bosporus Channel in, 58f, 61–62, 65f CAFZ in, 39–40, 39f

of central Anatolian plateau, 37–42, 38f–39f of Deliçay fault zone, 16f, 18f, 20f, 27–28 of DSF, 12f, 15f, 17f–18f, 19–28, 21f–22f, 24f, 55f of east Anatolian high plateau, 29–37, 29f–37f of EATF, 55–56, 55f–56f Eskişehir fault in, 39f, 40–41, 75–76 Gediz graben in, 12f, 66f, 70–71, 70f, 73f introduction to, 11–14, 12f–13f Karasu depression in, 15f, 23–28, 55–56 of Keldağ horst, 15f, 21f, 24f, 26–27 of Kırşehir‐Niğde Massif, 37–40, 38f, 75 in Maraş triple junction, 15f–18f, 23, 27–28, 55, 55f of Marmara region, 56–66, 57f–58f, 60f–65f Misis‐Andırın fault zone in, 12f, 17f, 23, 47–48 of NATF, 49–55, 50f–52f, 54f–55f of NeoTethyan Ocean, 16, 29, 36f, 42, 74–75 Oltu fault zone in, 32, 37f Söke graben in, 66f, 70f, 71–74, 73f Solhan volcano in, 33f, 35 of southeast Anatolia, 14–29, 15f–17f, 20f–22f, 24f–25f, 29f summary on, 74–76 Taurus Mountains in, 17f, 38f–39f, 41–42, 46–49, 47f, 48f, 66f of Türkoğlu‐Haruniye fault zone, 18f, 20f, 27–28 Uludağ Mountain in, 41, 57f–58f, 60f, 76 of western Anatolia, 66–74, 66f–67f, 69f–73f Motazedian, D., 276 Mountrakis, D., 249 Musadağ horst, 21f, 23, 26 NAF. See North Anatolian fault Nafe and Drake relation, 111 NAFZ. See North Anatolian fault zone NAT. See North Aegean Trough NATF. See North Anatolian transform fault National Geospatial‐Intelligence Agency (NGA), 108 Neogenic cover rocks, 160–69, 160f–166f Neotectonics. See also Morphotectonic development introduction to, 1–3 use in literature of, 13 NeoTethyan Ocean morphotectonic development in, 16, 29, 36f, 42, 74–75 in western Anatolian geology, 144f, 147–51, 149f, 170 New Zealand, 101 NGA. See National Geospatial‐Intelligence Agency Niğde‐Kırşehir Massif, 37–40, 38f, 75 North, R. G., 255 North Aegean Trough (NAT) boarder region in, 250–55, 250f–251f, 253f–254f, 256f CAT in, 245f, 247, 247f, 259 conclusions on, 262–63 Corinth Rift in, 241, 242f–243f, 245f fault inventories in, 249–50 geodynamic setting of Aegean region in, 242–44, 242f–243f North Aegean Sea in, 244–47, 245f–247f Northern Greece in, 244, 245f, 247–48, 248f geotectonic setting of, 239–42, 240f, 242f

294 INDEX North Aegean Trough (NAT) (cont’d ) introduction to, 6 seismogenic sources of, 248–49 strike‐slip fault zones in, 244–47, 246f, 252–55, 256f, 259–62 2014 earthquake in conclusions on, 262–63 Coulomb Stress change in, 259–62, 261f displacement pattern of, 262, 262f fault models of, 258f, 258t, 260 focal mechanisms of, 257f, 257t, 259, 259f introduction to, 255, 257f, 257t prior seismicity of, 258–59, 258t North Anatolian fault (NAF) in Marmara Sea seismology, 225–37, 229f–236f in 2014 earthquake, 242–44, 243f, 252, 255, 262 North Anatolian fault zone (NAFZ) gravity modeling in, 106, 112 introduction to, 5–6 river courses in, 93–101, 94f–100f seismic intensity maps of, 273–85, 276f–281f synthetic intensity map of, 282–84, 282f–284f North Anatolian transform fault (NATF), 12, 49–55, 50f–52f, 54f–55f North Greece, seismology of, 247–48, 248f North Samothraki, 6, 245–55 NUMMODEL experiments, 124–28, 124f, 125t, 126f–127f, 130f Okada, Y., 262 Okay, A. I., 151 Oltu fault zone, 32, 37f Ophiolite nappes, 142–45, 144f, 147–51, 149f, 151f Orogenic collapse model, 107 Ovacik fault, 6 seismic intensity maps of, 273–85, 275f, 277f–281f synthetic intensity maps of, 282–85, 282f–285f Ova regime, 38 Oxburgh, E. R., 135 Papazachos, B. C., 249 Papua New Guinea, 189 Paradisopoulou, P. M., 259 Parallel strike‐slip fault, 97–98, 97f Passchier, C. W., 197–99 Pavlides, S. B., 248–49 Peak Ground Acceleration (PGA), 6 in seismic intensity mapping, 277–85, 277f–280f Peak Ground Velocity (PGV), 6 in seismic intensity mapping, 277–85, 277f–280f Peyret, M., 44 PGA. See Peak Ground Acceleration PGV. See Peak Ground Velocity Piromallo, C., 47 Pontide mountains, 42–49, 43f–45f, 47f–48f Porphyritic granodiorite, 192f, 214f. See also Çataldağ granodiorite Pull‐apart basins, 93, 94f–96f, 99–100, 99f–100f

Purvis, M., 167 P‐wave velocity, 111, 111t Pysklywec, R. N., 121, 135–36 Rayleigh‐Taylor‐type instability, 121–22 Reilinger, R., 56, 135 Rhoades, D. A., 259 Rhodope massif, 241, 244 Riedel structures, 5, 227–28, 235–37, 246f River courses discussion on, 100–101 false displacement in, 100, 100f of gently bending rivers, 95–99, 96f–98f introduction to, 3, 93–94, 94f–96f Pull‐apart basins in, 99–100, 99f–100f Robertson, A. H. F., 167, 239 Rodgers, D. A., 99 Rojay, B., 167–69 Rolling hinge model, 171 Sakarya continent, 143f–144f, 145, 149f, 170–71 Samothraki SE, 6, 253–54 San Andreas fault system, 93, 101 Sandvol, E., 169 Sapanca Lake, 54f, 56–58, 57f Sarıkaya, M. A., 40 Şaroğlu, F., 31 Saros earthquake (2013), 231 Sboras, S., 250, 250f Schildgen, T. F., 41, 47 Schmeling, H., 132 Schollen‐type migmatites, 193f SCIH. See Single crystal incremental heating Scott, B., 163–64, 167 SCTF. See Single crystal total fusion Sea of Marmara. See Marmara Sea Seismic hazard assessment (SHA), 248–51 Seismic Hazard Harmonization in Europe (SHARE), 249 Seismic intensity maps conclusions on, 284–85 of Erzincan region, 273–85, 275f, 277f–284f felt‐intensity values in, 279 ground‐motion simulations in, 276–78, 277f–279f, 277t introduction to, 273–74 NAFZ study area of, 274, 275f Şengör, A. M. C., 13, 31, 49–50, 52, 66, 99–101, 146, 233 Şengül, E., 41 Serbomacedonian massif, 241 Serçeören area, 192f, 195, 196f Seyitoglu, G., 163–64, 167 Seyrek, A., 26, 27 SGR. See Subgrain rotations SHA. See Seismic hazard assessment ShakeMap software, 273–74, 279–80, 280f SHARE. See Seismic Hazard Harmonization in Europe Sierra Nevada, U.S., 121, 122f, 125, 189 Single crystal incremental heating (SCIH), 205–14, 206t–212t, 214f–215f

INDEX  295 Single crystal total fusion (SCTF), 205–14, 206t–213t, 214f–215f Sisman, F. N., 277 Skyros earthquake (2001), 247, 247f Slab‐tear effects conclusions on, 115–16 database in, 108 discussion of, 112–15, 113f–115f gravity anomalies in, 108–11, 109f–110f gravity models in, 111–12, 111t introduction to, 4, 105–8, 106f 2.5‐D gravity model in, 106f–107f, 109f, 111, 113f Smith, A. G., 241 Software EXSIM, 276 GLITTER, 200 ISOPLOT, 200 ShakeMap, 273–74, 279–80, 280f SOPALE, 124 Söke graben, 66f, 70f, 71–74, 73f, 143f Solhan volcano, 33f, 35 SOPALE software, 124 South Chalkidiki Offshore composite, 251–52, 251f Southeast Anatolia, 14–29, 15f–17f, 20f–22f, 24f–25f, 29f Stochastic ground‐motion simulation, 276. See also Ground‐ motion simulation Strike‐slip fault zones discussion on, 100–101 FBFZ of, 106f, 108–11 gentle bending rivers in, 95–99, 96f–98f introduction to, 3, 93–94, 94f–96f in NAT, 244–47, 246f, 252–55, 256f, 259–62 pull‐apart basins in, 99–100, 99f–100f Subgrain rotations (SGR), 198–99 Subparallel faults, 24f, 26–27, 227, 235 Suess, E., 47 Swiss Alps, 135–36 Synthetic approach, 97 Synthetic intensity map, 282–84, 282f–284f Synthetic pull‐apart basin, 99, 99f Taşova‐Erbaa pull‐apart basin, 96f Tauride Mountains morphotectonic development of, 42–49, 43f–45f, 47f–48f in western Anatolian geology, 141–51, 143f–144f, 148f–149f, 171 Taurus Mountains, 17f, 38f–39f, 41–42, 46–49, 47f, 48f, 66f Tectonic escape model, 106 Tekirdag basin, 233 Tethyan Ocean, 148, 170, 239. See also NeoTethyan Ocean; North Aegean Trough Three‐stage continuous extension model, 107 Tibetan plateau, 121, 122f, 124, 127 Tietze, E., 99 Toda, S., 260 TPAO. See Turkish Petroleum Corporation Troodos Mountain, 15f, 23

Trouw, R. A. J., 197–99 Tunoğlu, C., 42 Turfaldag Mountain, 191, 192f Turkish Ministry of Construction, 279–82, 280f–281f Turkish Petroleum Corporation (TPAO), 19, 29, 68, 116 Türkoğlu‐Haruniye fault zone, 18f, 20f, 27–28 Tüysüz, O., 42 Tuz Gölü basin, 12f, 38–39, 75 2.5‐D gravity model, 106f–107f, 109f, 111, 113f Two‐mica granite, 191, 192f–193f, 215f Tyrrhenian Sea, 189 Uludağ Mountain, 41, 57f–58f, 60f, 76 United States Geological Survey (USGS), 249, 279–82, 280f–281f Universal Transverse Mercator (UTM) projection, 192f–196f USGS. See United States Geological Survey UTM. See Universal Transverse Mercator Wavelength filtering, 108–9 Weijermars, R., 132 Wells, D. L., 254, 260 Westaway, R., 47 Western Anatolia. See also Çataldağ Plutonic Complex Aegean region in, 141–45, 142f, 154, 169–70 Bozdağ horst in, 148–53, 162–67 compositional variation in, 158, 158f conclusions on, 171–73 Dilek nappe in, 145, 149f, 151, 151f geology map of, 143f, 155f–157f, 166f introduction to, 4–5, 141–45, 142f–145f Iżmir‐Ankara suture in, 141–45, 143f–145f, 171 K‐feldspar in, 191–93, 192f–194f, 203–5, 209t, 214 major problems of magmatic associations in, 155–60, 155f–158f Menderes Massif in, 145–55, 146f–155f Neogene cover rocks in, 160–69, 160f–166f north‐south extensional regime in, 169–71 major suture zones of, 155f morphotectonic development of, 66–74, 66f–67f, 69f–73f ophiolite nappes in, 142–45, 144f, 147–51, 149f, 151f Sakarya continent in, 141–45, 143f–144f, 149f, 170–71 Taurides in, 141–51, 143f–144f, 148f–149f, 171 tectonic evolution of, 144f–145f tectonostratigraphic section of, 160f–161f Wet Quartzite Crust, 125t Yemen, 189 Yeşilırmak river, 96f Yilmaz, Y., 42, 53, 56, 146 Yıldırım, M., 43, 46 Yılmaz, Y., 31 Zero offset shear, 95–99, 96f–97f Zircon radiometric age data, 200–203, 201t–203t, 204f Zircon SHRIMP age, 216–17. See also Ar/Ar geochronology