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Petrogenesis of Metamorphic Rocks, Notas de estudo de Geologia

Geologia Geral

Tipologia: Notas de estudo

2016

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Baixe Petrogenesis of Metamorphic Rocks e outras Notas de estudo em PDF para Geologia, somente na Docsity! Ta Tea | Rodney Grapes E ns He Petrogenesis Ro e onto fo TO Rocks 8th Edition A Springer Petrogenesis of Metamorphic Rocks Prof.Dr. Kurt Bucher University Freiburg Mineralogy Geochemistry Albertstr. 23 B 79104 Freiburg Germany bucher@uni-freiburg.de Prof. Rodney Grapes Department of Earth and Environmental Sciences Korea University Seoul Korea grapes@korea.ac.kr ISBN 978-3-540-74168-8 e-ISBN 978-3-540-74169-5 DOI 10.1007/978-3-540-74169-5 Springer Heidelberg Dordrecht London New York Library of Congress Control Number: 2011930841 # Springer-Verlag Berlin Heidelberg 2011 This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer. Violations are liable to prosecution under the German Copyright Law. The use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. Cover design: deblik Printed on acid-free paper Springer is part of Springer Science+Business Media (www.springer.com) Preface This new edition of “Petrogenesis of Metamorphic Rocks” has several completely revised chapters and all chapters have updated references and redrawn figures. All chapters of Part II of the book have been rewritten. Also, the chapters “Introduc- tion” and “Grade” have undergone several major changes. The references made to important web sites relating to metamorphic petrology tutorials, software, mail base, etc have been updated and extended. However, it should be noted that some of the links to these sites may fail to work in the future. A large number of new figures showing assemblage stability diagrams have been computed using the Theriak/ Domino software by Ch. de Capitani of the University of Basel. We encourage you to regularly read (or at least glance through) current issues of scientific journals in your library either online or paper copies. In the field of metamorphic petrology, the Journal of Metamorphic Geology is essential reading and some of the other particularly relevant journals include, Journal of Petrology, Contributions to Mineralogy and Petrology, Geofluids, American Mineralogist, European Journal of Mineralogy, Lithos, Chemical Geology and Earth and Plane- tary Science Letters. Freiburg, Germany Kurt Bucher Seoul, South Korea Rodney Grapes November 2010 v . 4.7 Geothermobarometry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 4.7.1 Concept and General Principle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 4.7.2 Assumptions and Precautions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 4.7.3 Exchange Reactions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 153 4.7.4 Net-Transfer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 4.7.5 Miscibility Gaps and Solvus Thermometry . . . . . . . . . . . . . . . . 162 4.7.6 Uncertainties in Thermobarometry . . . . . . . . . . . . . . . . . . . . . . . . . 166 4.7.7 Thermobarometry Using Multi-equilibrium Calculations (MET) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167 4.8 Gibbs Method . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 4.9 Assemblage Stability Diagrams . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 4.10 More P–T Tools . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171 4.10.1 Reactions Involving Fluid Species . . . . . . . . . . . . . . . . . . . . . . . . 171 4.10.2 P–T Tools for Very Low Grade Rocks . . . . . . . . . . . . . . . . . . . 173 References and Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 174 Part II Metamorphism of Specific Rock Types 5 Metamorphism of Ultramafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 191 5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 191 5.2 Ultramafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 191 5.2.1 Rock Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 193 5.2.2 Chemical Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 194 5.3 Metamorphism in the MSH System . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 196 5.3.1 Chemographic Relations in the MSH System . . . . . . . . . . . . . . 196 5.3.2 Progressive Metamorphism of Maximum Hydrated Harzburgite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 198 5.4 Metamorphism in the CMASH System . . . . . . . . . . . . . . . . . . . . . . . . . . . 202 5.4.1 Progressive Metamorphism of Hydrated Al-Bearing Lherzolite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 202 5.4.2 Effects of Rapid Decompression and Uplift Prior to Cooling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 204 5.5 Isograds in Ultramafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 204 5.6 Mineral Assemblages in the Uppermost Mantle . . . . . . . . . . . . . . . . . . 206 5.7 Serpentinization of Peridotite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 208 5.8 Ultramafic Rocks at High Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . 210 5.9 Thermometry and Geobarometry in Ultramafic Rocks . . . . . . . . . . . 211 5.10 Carbonate-Bearing Ultramafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 212 5.10.1 Metamorphism of Ophicarbonate Rocks . . . . . . . . . . . . . . . . . . 212 5.10.2 Soapstone and Sagvandite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 215 5.11 Open System Reactions in Ultramafic Rocks . . . . . . . . . . . . . . . . . . . . . 219 5.12 Potassium-Bearing Peridotites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 221 References and Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 221 Contents ix 6 Metamorphism of Dolomites and Limestones . . . . . . . . . . . . . . . . . . . . . . . . 225 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 225 6.1.1 Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 225 6.1.2 Minerals and Rock Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . 227 6.1.3 Chemographic Relationships . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 228 6.2 Orogenic Metamorphism of Siliceous Dolomite . . . . . . . . . . . . . . . . . . 230 6.2.1 Modal Evolution Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 232 6.3 Orogenic Metamorphism of Limestone . . . . . . . . . . . . . . . . . . . . . . . . . . . 233 6.4 Contact Metamorphism of Dolomites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 235 6.4.1 Modal Evolution Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 238 6.5 Contact Metamorphism of Limestones . . . . . . . . . . . . . . . . . . . . . . . . . . . 239 6.6 Isograds and Zone Boundaries in Marbles . . . . . . . . . . . . . . . . . . . . . . . . 243 6.7 Metamorphic Reactions Along Isothermal Decompression Paths . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 244 6.8 Marbles Beyond the CMS-HC System . . . . . . . . . . . . . . . . . . . . . . . . . . . . 246 6.8.1 Fluorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 247 6.8.2 Aluminium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 248 6.8.3 Potassium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 248 6.8.4 Sodium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 251 6.9 Thermobarometry of Marbles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 251 6.9.1 Calcite–Dolomite Miscibility Gap . . . . . . . . . . . . . . . . . . . . . . . . . 252 6.10 High-Pressure and Ultrahigh-Pressure Metamorphism of Carbonate Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252 References and Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 253 7 Metamorphism of Pelitic Rocks (Metapelites) . . . . . . . . . . . . . . . . . . . . . . . . 257 7.1 Metapelitic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257 7.2 Pelitic Sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257 7.2.1 General . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257 7.2.2 Chemical Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 258 7.2.3 Mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 258 7.3 Pre-metamorphic Changes in Pelitic Sediments . . . . . . . . . . . . . . . . . . . 258 7.4 Intermediate-Pressure Metamorphism of Pelitic Rocks . . . . . . . . . . . . 259 7.4.1 Chemical Composition and Chemography . . . . . . . . . . . . . . . . . . . 259 7.4.2 Mineral Assemblages at the Beginning of Metamorphism . . 260 7.4.3 The ASH System . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 262 7.4.4 Metamorphism in the FASH System . . . . . . . . . . . . . . . . . . . . . . . . . 264 7.4.5 Mica-Involving Reactions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 267 7.4.6 Metamorphism in the KFMASH System (AFM System) . . . . 269 7.5 Low-Pressure Metamorphism of Pelites . . . . . . . . . . . . . . . . . . . . . . . . . . . 278 7.5.1 FASH System . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 278 7.5.2 KFMASH System . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 280 7.6 High-Temperature Metamorphism of Pelites: Metapelitic Granulites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 284 x Contents 7.6.1 Cordierite–Garnet–Opx–Spinel–Olivine Equilibria . . . . . . . . . . 285 7.6.2 The Excess Quartz Condition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 286 7.6.3 Partial Melting and Migmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 288 7.6.4 More About Granulites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 292 7.7 Metamorphism of Mg-rich “Pelites” . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 296 7.8 High Pressure: Low Temperature Metamorphism of Pelites . . . . . . 298 7.9 Additional Components in Metapelites . . . . . . . . . . . . . . . . . . . . . . . . . . . . 302 References and Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 306 8 Metamorphism of Marls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 315 8.1 General . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 315 8.2 Orogenic Metamorphism of Al-Poor Marls . . . . . . . . . . . . . . . . . . . . . . . . 316 8.2.1 Phase Relationships in the KCMAS-HC System . . . . . . . . . . . . 317 8.2.2 Prograde Metamorphism in the KCMAS-HC System at Low XCO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 319 8.3 Orogenic Metamorphism of Al-Rich Marls . . . . . . . . . . . . . . . . . . . . . . . . 321 8.3.1 Phase Relationships in the CAS-HC System . . . . . . . . . . . . . . . . . 323 8.3.2 Phase Relationships in the KNCAS-HC System . . . . . . . . . . . . . 325 8.4 Increasing Complexity of Metamarls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 332 8.5 Low Pressure Metamorphism of Marls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 333 References and Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 336 9 Metamorphism of Mafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 339 9.1 Mafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 339 9.1.1 Hydration of Mafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 340 9.1.2 Chemical and Mineralogical Composition of Mafic Rocks . . 342 9.1.3 Chemographic Relationships and ACF Projection . . . . . . . . . . . 343 9.2 Overview of Metamorphism of Mafic Rocks . . . . . . . . . . . . . . . . . . . . . . 348 9.2.1 Plagioclase in Mafic Rocks, Equilibria in the Labradorite System (NCASH System) . . . . . . . . . . . . . . . . . . . . . . . 352 9.3 Subgreenschist Facies Metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 353 9.3.1 General Aspects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 353 9.3.2 Metamorphism in the CMASH and NCMASH Systems . . . . . 356 9.3.3 Transition to Greenschist Facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 362 9.4 Greenschist Facies Metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 363 9.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 363 9.4.2 Mineralogical Changes Within the Greenschist Facies . . . . . . 363 9.4.3 Greenschist–Amphibolite Facies Transition . . . . . . . . . . . . . . . . . 365 9.5 Amphibolite Facies Metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 367 9.5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 367 9.5.2 Mineralogical Changes Within the Amphibolite Facies . . . . . 367 9.5.3 Low-Pressure Series Amphibolites . . . . . . . . . . . . . . . . . . . . . . . . . . . 368 9.5.4 Amphibolite–Granulite Facies Transition . . . . . . . . . . . . . . . . . . . . 370 9.6 Granulite Facies and Mafic Granulites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 371 Contents xi . Chapter 1 Definition, Conditions and Types of Metamorphism Rock metamorphism is a geological process that changes the mineralogical and chemical composition, as well as the structure of rocks. Metamorphism is typically associated with elevated temperature and pressure, thus it affects rocks within the earth’s crust and mantle. The process is driven by changing physical and/or chemical conditions in response to large-scale geological dynamics. Consequently, it is inherent in the term, that metamorphism always is related to a precursor state where the rocks had other mineralogical and structural attributes. Metamorphism, metamorphic processes and mineral transformations in rocks at elevated tempera- tures and pressures are fundamentally associated with chemical reactions in rocks. Metamorphism does not include, by definition, similar processes that occur near the earth’s surface such as weathering, cementation and diagenesis. The transition to igneous processes is gradual, and metamorphism may include partial melting. The term metasomatism is used if modification of the rocks bulk composition is the dominant metamorphic process. Metamorphic rocks are rocks that have developed their mineralogical and structural characteristics by metamorphic processes. The most typical metamorphism transforms sedimentary rocks to metamorphic rocks by addition of heat during mountain building or by a large volume of magma in the crust. For example, Upper Ordovician shales and nodular limestone (Fig. 1.1) in the Permian Oslo rift were heated to about 420C by syenite plutons. The gray fissile shales composed predominantly of diagenetic clay minerals and quartz were transformed to dark splintery rocks called hornfels, which contain metamorphic minerals such as biotite, cordierite, K-feldspar and sillimanite. The limestone nodules consist of pure CaCO3 in the unmetamorphic sedimentary rock (Fig. 1.1a). At the temperature of metamorphism calcite reacted with the minerals of the shale. The reaction produced concentrically zoned nodules consisting of the Ca-silicates anorthite, wollastonite and diopside, so-called calcsilicate rocks (Fig. 1.1b). All calcite has been used up in the reaction and all CO2 once present has left the rock together with H2O produced by the dehydration of the clay minerals. The example shows all characteristic features of metamorphism: The nodular limestone and shale were the precursor rocks (so-called protolith) of the newly formed metamorphic hornfels. A set of new minerals formed on the expense of previous minerals by chemical reactions in the rocks. The reactions were driven by heat added to the rock. The structure of the original rock was modified as K. Bucher and R. Grapes, Petrogenesis of Metamorphic Rocks, DOI 10.1007/978-3-540-74169-5_1, # Springer-Verlag Berlin Heidelberg 2011 3 expressed by new ultrafine grainsize in the hornfels and by the concentric arrange- ment of chemically distinct reaction zones in the calcsilicate nodules as a result of small scale redistribution of chemical constituents in the rock. Obviously also the chemical composition of the rocks were changed on a local scale because the volume once occupied by carbonate (limestone nodule) was replaced by silicates during the metamorphic process and the rock is devoid of CO2 and H2O. 1.1 Conditions of Metamorphism Large scale geologic events such as global lithospheric plate movements, subduc- tion of oceanic lithosphere, continent–continent collision and ocean floor spreading all have the consequence of moving rocks and transporting heat. Consequently, changes in pressure (depth) and temperature are the most important variables in rock metamorphism. As an example, consider a layer of sediment on the ocean floor that is covered with more sediment layers through geologic time and finally subducted at a destructive plate margin. The mineral, chemical and structural transformations experienced by the sediment can be related to a gradually increas- ing temperature with time. The question may be asked: at what temperature does metamorphism begin? 1.1.1 Low-Temperature Limit of Metamorphism Temperatures at which metamorphism sets in are strongly dependent on the mate- rial under investigation. Transformation of evaporites, of vitreous material and of shale (pelite) limestone nodule calcsilicate hornfels metapelitic hornfels zoned calcsilicate nodule a b Fig. 1.1 Metamorphic transformation of sedimentary rocks (a) to metamorphic hornfels (b) in a contact aureole of Permian plutons in the Oslo rift, Norway: (a) Upper Ordovician shale with fossiliferous limestone nodules. (b) Same rock as (a) but heated to about 430C by a nearby pluton. Shale transformed to hornfels and limestone nodules reacted to zoned calcsilicate rock with anorthite and diopside as main minerals 4 1 Definition, Conditions and Types of Metamorphism metamorphism is not restricted to the Earth’s crust. A given volume of rock in the convecting mantle continuously undergoes metamorphic processes such as recrys- tallization and various phase transformations in the solid state at temperatures in excess of 1,500C. 1.1.3 Low-Pressure Limit of Metamorphism Ascending hot silicate magmas are a typical and globally important occurrence in geologically active areas. The heat released by the cooling magma causes meta- morphism in the surrounding country rocks to produce so-called contact aureoles at shallow depth and pressures of a few megapascal. 1.1.4 High-Pressure Limit of Metamorphism For a long time it was believed that maximum pressures in metamorphic crustal rocks did not exceed 1.0 GPa, which corresponds to the lithostatic pressure at the base of a normal continental crust with a thickness of 30–40 km (pressure units used in this book: Pa – pascal, 1 Pa ¼ 1 N/m2 ¼ 1 kg/(m s2); 105 Pa ¼ 1 bar; 1 GPa ¼ 109 Pa; 1 MPa ¼ 106 Pa; 100 MPa ¼ 1 kbar). As better data on mineral stability relations became available, it was found that mineral assemblages in some meta- morphic mafic rocks often recorded pressures of 1.5–2.0 GPa. It was recognized early that these rocks, called eclogites, represented high density and consequently high pressure equivalents of basalts (Eskola 1921). Spectacular high-pressure rocks of clearly crustal origin were discovered in the Dora-Maira massif of the Alps (Chopin 1984). Gneisses with pure pyrope garnet containing coesite inclusions indicate pressures of at least 3.0 GPa. Similar super-high-pressure rock have been discovered in many other places in the world (e.g. Reinecke 1991; Wain et al. 2000) since. Today, eclogites with coesite inclusions in garnet and even diamond-bearing eclogites have been reported (Smith and Lappin 1989; Okay 1993; Zhang and Liou 1994; Coleman et al. 1995; Schreyer 1995; Ernst 1999; L€u et al. 2009). Some of these rocks require at least 6.0 GPa to form. Rock transformations at such high pressures are called ultra-high-pressure (UHP) metamorphism. It is clear that such super-high pressures must be related to transport of crustal rocks to very great depth of >100 km. Yet, metamorphism as stated above, is not confined to the Earth’s crust or to subducted crustal rocks. Mantle rocks such as garnet peridotites (or garnet–olivine– pyroxene–granofelses) from ophiolite complexes or from xenoliths in kimberlite record pressures of>3–4 GPa (e.g. Yang et al. 1993; Song et al. 2009). Consequently, it is possible to find and study rocks that formed 100–200 km below the earth surface corresponding to pressures >3–6 GPa. 1.1 Conditions of Metamorphism 7 Temperature Pressure Low T-limit Low P-limit 0C for processes in near surface environments, rock–water reactions 0.1 MPa at contact to lava flows at the surface Conventionally, the term metamorphism implies T > 150–200C Upper T-limit High P-limit In crustal rocks: 750–850C (max. recorded T~1,150C) Presently, some rocks collected at the earth surface are known to have once formed at 100–200 km depth, ¼ 3–6 GPaIn many regional scale metamorphic areas T does not exceed ~ 650–700C 1.2 Types of Metamorphism On the basis of geological setting, we distinguish between metamorphism of local and regional extent: Regional extent Local extent Orogenic metamorphism (regional metamorphism) Contact (igneous) metamorphism Subduction metamorphism Cataclastic metamorphism Collision metamorphism Hydrothermal metamorphism Ocean-floor metamorphism Exotic local metamorphism Burial metamorphism Impact metamorphism Lightning metamorphism Combustion metamorphism Such a subdivision is certainly useful, but it should be kept in mind that there commonly exist transitional forms between regional and contact (igneous) cate- gories of metamorphism. 1.2.1 Orogenic Metamorphism Orogenic metamorphism (Miyashiro 1973, p. 24) is caused by mountain building or orogenic process. It acts upon large volumes of rocks and the metamorphic effects of a specific mountain building episode are found over large, regional scale dimensions. The term regional metamorphism can be used as a word with the same meaning. This is the most significan type of metamorphism affecting the rocks of the continental crust on a large scale. Consequently this book deals almost exclusively with orogenic metamorphism supplemented with some aspects of low-P contact metamorphism (Sect. 1.2.4). Metamorphism associated with a specific mountain building episode may show characteristic features. As an example; Alpine metamorphism in the Alps is collec- tively used for all metamorphic processes associated with the formation of the Alps 8 1 Definition, Conditions and Types of Metamorphism in late Cretaceous and Tertiary times. It affected rocks that experienced metamorphic transformations for the first time, i.e. Mesozoic sediments and volcanics. It also once more transformed rocks that already have been metamorphosed during earlier oro- genic periods, such as Variscian gneisses, Caledonian metasediments and even rocks with Precambrian mineral assemblages. Rocks subjected to orogenic metamorphism usually extend over large belts, hundreds or thousands of kilometers long and tens or hundreds of kilometers wide. In many higher temperature parts of orogenic metamorphic belts, syn- or late- tectonic granites are abundant. Orogenic metamorphism and granitic plutons are often intimately associated. In the middle and upper crust, rising granitic magma carries heat and thereby contributes to the increase of temperature on a regional scale leading to typical high-temperature low-pressure terranes. In the lower to middle crust, the granitic magmas were generated by partial melting as a conse- quence of high-grade metamorphism. Orogenic metamorphism is commonly characterized by two fundamentally dif- ferent kinds of regional scale metamorphic transformations that follow each other in time. An early high-pressure low-temperature type metamorphism is related to a subduction zone process, a later younger regional metamorphism following a moderate P–T gradient is related to continental collision. This dualism of orogenic metamorphism is related to major intervals of a Wilson cycle of mountain building (Wilson 1966; Murphy and Nance 1992; Duncan and Turcotte 1994). Typically, metamorphic recrystallization in orogenic belts is accompanied by deformation. Such metamorphic rocks exhibit a penetrative fabric with preferred orientation of mineral grains. Examples are phyllites, schists and gneisses. Oro- genic metamorphism is a long-lasting process of millions or tens of millions of years duration. It includes a number of distinct episodes of crystallization and deformation. Individual deformation phases appear to have definite characteristics, like attitude and direction of schistosity, folds, and lineations. Therefore, several phases of deformation can perhaps be recognized in the field and they often can be put into a time sequence. Texture observations in thin sections under the micro- scope can unravel the relationships between structural features and mineral growth, and establish the time relations of deformation and metamorphism (e.g. Vernon 1976, p. 224). Deformation is intimately associated with most forms of metamorphic recrystal- lization. Almost all metamorphic rocks show distinct features of ductile (Fig. 1.3) or brittle (Fig. 1.4) deformation. Metamorphic rocks are commonly intensely folded like the banded marbles shown in Fig. 1.3a from Engabreen, Nordland, Norway. The marbles represent former calcareous sediments, including limestone, marl and dolomite. The dolomite marble contains tremolite, diopside and phlogopite as characteristic metamorphic minerals. Granites and other unfoliated plutonic rocks can be progressively sheared resulting the transformation to gneiss. The Matorello granite is shown as an example on Fig. 1.3b from the Lagetti locality (Valle Maggia, Ticino, Swiss Alps). It shows the gradual transition of undeformed granite (upper left) to strongly foliated and sheared gneiss (lower right). Deformation is monitored by a regular increasing flattening of spherical mafic inclusions (strain 1.2 Types of Metamorphism 9 T a b le 1 .1 T y p ic al fe at u re s o f im p o rt an t ty p es o f m et am o rp h is m T y p e o f m et am o rp h is m O ro g en ic (s u b d u ct io n ty p e) O ro g en ic (c o ll is io n ty p e) O ce an -fl o o r C o n ta ct G eo lo g ic se tt in g In o ro g en ic b el ts , ex te n d in g fo r se v er al 1 ,0 0 0 k m 2 In o ro g en ic b el ts , ex te n d in g fo r se v er al 1 ,0 0 0 k m 2 In o ce an ic cr u st an d u p p er m an tl e, ex te n d in g fo r se v er al 1 ,0 0 0 k m 2 P ro x im it y to co n ta ct s to sh al lo w le v el ig n eo u s in tr u si o n s; co n ta ct au re o le o f a fe w m u p to a so m e k il o m et er w id th E ar ly p h as e o f o ro g en ic m et am o rp h is m L at e p h as e o f o ro g en ic m et am o rp h is m S ta ti c/ d y n am ic re g im e D y n am ic ; g en er al ly as so ci at ed w it h th ru st in g , sl ic in g D y n am ic , g en er al ly as so ci at ed w it h p o ly p h as e d ef o rm at io n , fo li at io n an d fo ld in g  S ta ti c, ex te n si v e fr ac tu ri n g an d v ei n in g , n o fo li at io n as so ci at ed w it h ex te n si o n an d se a fl o o r sp re ad in g S ta ti c, n o fo li at io n T em p er at u re 1 5 0 – 7 0 0  C 1 5 0 – 8 5 0  C 1 5 0 – 5 0 0  C 1 5 0 – 6 0 0  C (> 7 0 0  C in d ee p su b d u ct io n ) (m ax . T ~ 1 ,0 5 0  C ) (> 5 0 0  C cl o se to m ag m as ) (> 6 0 0  C at g ab b ro co n ta ct s) L it h o st at ic p re ss u re 2 0 0 – 3 ,0 0 0 M P a fo r cr u st al ro ck s 2 0 0 – 1 ,0 0 0 M P a < 3 0 0 M P a F ro m a fe w te n s M P a to 3 ,0 0 0 M P a (i n so m e co ll is io n b el ts u p to 1 4 k b ar s, “d o u b le cr u st ”) T em p er at u re g ra d ie n ts 5 – 1 2  C /k m (v er ti ca l) g ra d ie n t d ep en d s o n su b d u ct io n v el o ci ty 1 2 – 6 0  C /k m (v er ti ca l) g ra d ie n ts d ep en d s o n as so ci at ed ig n eo u s ac ti v it y 5 0 – 5 0 0  C /k m (v er ti ca l o r h o ri zo n ta l) 1 0 0  C /k m o r h ig h er (h o ri zo n ta l) P ro ce ss es A ss o ci at ed w it h su b d u ct io n o f o ce an ic li th o sp h er e (o p h io li te s) an d p ar tl y al so co n ti n en ta l ro ck s C o n ti n en t– co n ti n en t co ll is io n , li th o sp h er ic th ic k en in g , co m p re ss io n an d h ea ti n g H ea t su p p ly b y as ce n d in g as th en o sp h er e an d in tr u d in g m afi c m ag m as at m id -o ce an ri d g es , co m b in ed w it h ci rc u la ti o n o f se a- w at er th ro u g h fr ac tu re d h o t ro ck s in an ex te n si o n al re g im e H ea t su p p ly b y ig n eo u s in tr u si o n s, co m m o n ly al so as so ci at ed w it h ex te n si v e m et as o m at is m ca u se d b y co n v ec ti v e h y d ro th er m al ci rc u la ti o n , T y p ic al m et am o rp h ic ro ck s B lu es ch is t, ec lo g it e, se rp en ti n it e, S la te , p h y ll it e, sc h is t, g n ei ss , m ig m at it e, m ar b le , q u ar tz it e, g re en sc h is t, am p h ib o li te , g ra n u li te M et ab as al t, g re en st o n e, m et ag ab b ro , se rp en ti n it e; p ri m ar y st ru ct u re o ft en w el l p re se rv ed H o rn fe ls , m ar b le , ca lc si li ca te g ra n o fe ls , sk ar n 12 1 Definition, Conditions and Types of Metamorphism Australia (e.g. Smith 1969), Japan (e.g. Seki 1973), and Chile (e.g. Levi et al. 1982). Well-known examples of burial metamorphism in metapelitic and metabasic rocks are also described by Merriman and Frey (1999) and Robinson and Bevins (1999), respectively. Diastathermal metamorphism is a term proposed by Robinson (1987) for burial metamorphism in extensional tectonic settings showing enhanced heat flow. An example of this type of regional metamorphism was described from the Welsh Basin (Bevins and Robinson 1988; see also Merriman and Frey 1999; Robinson and Bevins 1999). 1.2.4 Contact Metamorphism Contact metamorphism usually takes place in rocks in the vicinity of plutonic or extrusive igneous bodies. Metamorphic changes are caused by the heat given off by the cooling igneous body and also by gases and fluids released by the crystallizing magma. The zone of contact metamorphism is termed a contact aureole, the rocks affected are the country rocks of the magma body. The width of contact metamor- phic aureoles typically varies in the range of several meters to a few kilometers. Local deformation associated with emplacement of the igneous mass can be observed in some aureoles. Specifically, the width of aureoles depends on the volume, composition, intru- sion depth of a magmatic body and also the properties of the country rocks, especially their fluid content and permeability. A larger volume of magma carries more heat with it than a smaller one, and the temperature increase in the bordering country rock will last long enough to cause mineral reactions. Rocks adjacent to small dikes, sills or lava flows are barely metamorphosed, whereas larger igneous bodies give rise to a well-defined contact aureole of metamorphic rocks. The temperature of the various types of magma differs widely. Typical solidus temperatures of gabbroic (mafic, basaltic) magmas are well over 1,000C. In contrast, many granitic plutons are formed from water-rich melts at temperatures close to 650–700C. Large granitoid plutons formed from water-deficient magmas, (e.g. charnockite, mangerite) have solidus temperatures close to 900–950C. How- ever, because water-rich granites are the most common plutonic bodies in the continental crust, contact aureoles around granites are the most common examples of contact metamorphism. The intrusion depth of a magmatic body determines the thermal gradient and heat flow across the magma-country rock contact. Exceptionally high thermal gradients are generally confined to the upper 10 km of the Earth’s crust because, at deeper levels, the country rocks are already rather hot and, hence, obvious thermal aureoles are seldom produced. The effects of contact metamorphism are most obvious where non-metamorphic sedimentary rocks, especially shales and limestones, are in contact with large magmatic bodies as detailed in later chapters. On the other hand, country rocks 1.2 Types of Metamorphism 13 that experienced medium- or high-temperature regional metamorphism prior to the intrusion show limited effects of contact metamorphic overprinting because the older mineral assemblages persist at contact metamorphic conditions. Contact metamorphic rocks are generally fine-grained and lack schistosity. The most typical example is called hornfels (see Chap. 2 for nomenclature); but foliated rocks such as spotted slates and schists are occasionally present. Some features of contact metamorphism are stated briefly in Table 1.1 (for more details see Kerrick 1991). Pyrometamorphism is a special kind of contact metamorphism. It shows the effects of particularly high temperatures at the contact of a rock with magma under volcanic or quasi-volcanic conditions, e.g. in xenoliths. Partial melting is common, producing glass-bearing rocks called buchites, and in this respect pyrometamorph- ism may be regarded as being intermediate between metamorphism and igneous processes. A detailed discussion of pyrometamorphism illustrated by numerous examples of rocks of different compositions is given by Grapes (2006) and is not discussed further here. 1.2.5 Cataclastic Metamorphism Cataclastic metamorphism is confined to the vicinity of faults and overthrusts, and involves purely mechanical forces causing crushing and granulation of the rock fabric. Experiments show that cataclastic metamorphism is favored by high strain rates under high shear stress at relatively low temperatures. The resulting cataclas- tic rocks are non-foliated and are known as fault breccia, fault gauge, or pseudo- tachylite. A pseudotachylite consists of an aphinitic groundmass that looks like black basaltic glass (tachylite). Note that mylonites are not regarded as cataclastic rocks, because some grain growth by syntectonic recrystallization and neoblastesis is involved (see e.g. Wise et al. 1984). Dislocation or dynamic metamorphism are sometimes used as synonyms for cataclastic metamorphism; but these terms were initially coined to represent what is now called regional metamorphism. In order to avoid misunderstanding, the name “cataclastic metamorphism” is preferred. 1.2.6 Hydrothermal Metamorphism Hydrothermal metamorphism was originally introduced as a term by Coombs (1961). In hydrothermal metamorphism, hot aqueous solutions or gases flow through fractured rocks causing mineralogical and chemical changes in the rock matrix. The study of hydrothermal processes experienced a dramatic expansion during the last decades. It has been generalized and today the research on water–rock interaction 14 1 Definition, Conditions and Types of Metamorphism vaporization, and extreme chemical reduction, e.g. Essene and Fisher (1986). In soils and sands, the products of lightning metamorphism are typically glassy tubes called fulgurites. Examples of lightning metamorphism are summarized in Grapes (2006). Combustion metamorphism is a type of pyrometamorphism caused by sponta- neous combustion of organic matter, coal, oil and gas at or near the earth’s surface. Temperatures can be extreme, i.e., 1,000–1,500C, and with increasing temperature burnt rocks, clinkers, and slag or paralava are produced. Although contact aureoles caused by combustion metamorphism are generally only a few meters thick, and more rarely may be up to several tens of meters, burnt rocks may crop out over a large area, especially where combusted organic-bearing strata are gently dipping, e.g.>518,000 km2 in the case of the great western coal-bearing Cretaceous–Tertiary basins of the United States extending from Texas in the south to Canada in the north. The interested reader is referred to numerous examples of combustion metamor- phism described in Grapes (2006). References and Further Reading Alt JC (1999) Very low-grade hydrothermal metamorphism of basic igneous rocks. In: Frey M, Robinson D (eds) Low-grade metamorphism. Blackwell Science, Oxford, pp 169–201 Bevins RE, Robinson D (1988) Low grade metamorphism of the Welsh Basin Lower Paleozoic succession: an example of diastathermal metamorphism? J Geol Soc Lond 145:363–366 Boles JR, Coombs DS (1975) Mineral reactions in zeolitic Triassic tuff, Hokonui Hills, New Zealand. Geol Soc Am Bull 86:163–173 Bucher K, Stober I (2010) Fluids in the upper continental crust. Geofluids 10:241–253 Buick IS, Cartwright I (1996) Fluid–rock interaction during low-pressure polymetamorphism of the reynolds range group, Central Australia. J Petrol 37:1097–1124 Burnham CW (1979) Magmas and hydrothermal fluids. In: Barnes HL (ed) Geochemistry of hydrothermal ore deposits. Wiley, New York, pp 71–136 Chopin C (1984) Coesite and pure pyrope in high-grade blueschists of the western Alps: a first record and some consequences. Contrib Miner Petrol 86:107–118 Coleman RG, Wang X (eds) (1995) Ultrahigh pressure metamorphism. Cambridge University Press, Cambridge, 528 pp Coombs DS (1961) Some recent work on the lower grades of metamorphism. Aust J Sci 24:203–215 Dietz RS (1961) Astroblemes. Sci Am 205:50–58 Droop GTR, Al-Filali IY (1996) Interaction of aqueous fluids with calcareous metasediments during high-T, low-P regional metamorphism in the Qadda area, southern Arabian Shield. J Metamorph Geol 14:613–634 Duncan CC, Turcotte DL (1994) On the breakup and coalescence of continents. Geology 22:103–106 Ellis DJ (1980) Osumilite-sapphirine-quartz-granulites from Enderby Land, Antarctica: P-T con- ditions of metamorphism, implications for garnet-cordierite equilibria and the evolution of the deep crust. Contrib Miner Petrol 74:201–210 Emmermann R, Lauterjung J (1997) The German Continental Deep Drilling Program KTB. J Geophys Res 102:18179–18201 Ernst WG (1999) Metamorphism, partial preservation, and exhumation of ultrahigh-pressure belts. I Arc 8:125–153 Eskola P (1921) On the eclogites of Norway. Skr Vidensk Selsk Christiania, Mat-nat Kl I 8:1–118 References and Further Reading 17 Essene EJ, Fisher DC (1986) Lightning strike fusion: extreme reduction and metal-silicate liquid immiscibility. Science 234:189–193 Frey M, Kisch HJ (1987) Scope of subject. In: Frey M (ed) Low temperature metamorphism. Blackie, Glasgow, pp 1–8 Gianelli G, Grassi S (2001) Water-rock interaction in the active geothermal system of Pantelleria, Italy. Chem Geol 181:113–130 Giggenbach WF (1981) Geothermal mineral equilibria. Geochim Cosmochim Acta 45:393–410 Giggenbach WF (1984) Mass transfer in hydrothermal alteration systems – a conceptual approach. Geochim Cosmochim Acta 48:2693–2711 Gillis KM (1995) Controls on hydrothermal alteration in a section of fast-spreading oceanic crust. Earth Planet Sci Lett 134:473–489 Grapes R (2006) Pyrometamorphism. Springer, Berlin, Heidelberg, 275 pp Grieve RAF (1987) Terrestrial impact structures. Ann Rev Earth Planet Sci 15:245–270 Harley SL, Motoyoshi Y (2000) Al zoning in orthopyroxene in a sapphirine quartzite: evidence for >1120C UHT metamorphism in the Napier Complex, Antarctica, and implications for the entropy of sapphirine. Contrib Miner Petrol 138:293–307 Hokada T (2001) Feldspar thermometry in ultrahigh-temperature metamorphic rocks: Evidence of crustal metamorphism attaining 1100C in the Archean Napier Complex, East Antarctica. Am Mineralog 86:932–938 Kawachi DE (1975) Pumpellyite-actinolite and contigous facies metamorphism in the Upper Wakatipu district, southern New Zealand. NZ J Geol Geophys 17:169–208 Kerrick DM (1991) Contact metamorphism., vol 26, Reviews in mineralogy. Mineralogical Society of America, Washington DC, 847 pp Kisch HJ (1987) Correlation between indicators of very-low-grade metamorphism. In: Frey M (ed) Low temperature metamorphism. Blackie, Glasgow, pp 227–300 Kozlovsky YeA (1984) The world’s deepest well. Sci Am 251:106–112 Lamb RC, Smalley PC, Field D (1986) P-T conditions for the Arendal granulites, southern Norway: implications for the roles of P, T and CO2 in deep crustal LILE-depletion. J Metamorph Geol 4:143–160 Levi B, Aguirre L, Nystroem JO (1982) Metamorphic gradients in burial metamorphosed vesicular lavas; comparison of basalt and spilite in Cretaceous basic flows from central Chile. Contrib Mineralog Petrol 80:49–58 L€u Z, Zhang L, Du J, Bucher K (2009) Petrology of coesite-bearing eclogite from Habutengsu Valley, western Tianshan, NW China and its tectonometamorphic implication. J Metamorph Geol 27:773–787 Merriman RJ, Frey M (1999) Patterns of very low-grade metamorphism in metapelitic rocks. In: Frey M, Robinson D (eds) Low-grade metamorphism. Blackwell Science, Oxford, pp 61–107 Miyashiro A (1971) Pressure and temperature conditions and tectonic significance of regional and ocean floor metamorphism. Tectonophysics 13:141–159 Miyashiro A (1973) Metamorphism and metamorphic belts. Allen & Unwin, London, 492 pp M€oller P et al (1997) Paleo- and recent fluids in the upper continental crust – results from the German Continental deep drilling Program (KTB). J Geophys Res 102:18245–18256 Murphy JB, Nance RD (1992) Mountain belts and the supercontinent cycle. Sci Am 266:84–91 Newton RC (1987) Petrologic aspects of Precambrian granulite facies terrains bearing on their origins. In: Kr€oner A (ed) Proterozoic lithospheric evolution, vol 17, Geodynamics series. American Geophysical Union, Washington, DC, pp 11–26 Okay AI (1993) Petrology of a diamond and coesite-bearing metamorphic terrain: Dabie Shan, China. Eur J Miner 5:659–675 Ramsay JG, Allison I (1979) Structural analysis of shear zones in an alpinised Hercynian granite (Maggia Nappen, Pennine zone, central Alps). Schweiz Miner Petrogr Mitt 59:251–279 Reinecke T (1991) Very-high-pressure metamorphism and uplift of coesite-bearing metasedi- ments from the Zermatt-Saas zone, Western Alps. Eur J Miner 3:7–17 18 1 Definition, Conditions and Types of Metamorphism Reyes AG, Grapes R, Clemente VC (2003) Fluid-rock interaction at the magmatic-hydrothermal interface of theMount Cagua geothermal field, Philippines. Soc Econ Geol Spec Publ 10:197–222 Robinson D (1987) Transition from diagenesis to metamorphism in extensional and collision settings. Geology 15:866–869 Robinson D, Bevins RE (1999) Patterns of regional low-grade metamorphism in metabasites. In: Robinson D, Frey M (eds) Low-grade metamorphism. Blackwell Science, Oxford, pp 143–168 Robinson D, Merriman RJ (1999) Low-temperature metamorphism: an overview. In: Frey M, Robinson D (eds) Low-grade metamorphism. Blackwell Science, Oxford, pp 1–9 Ruppert LF, Hower JC, Ryder RT, Trippi MH, Grady WC, Levine JR (2010) Geologic controls on observed thermal maturation patterns in Pennsylvanian coal-bearing rocks in the Appalachian basin. Int J Coal Geol 81:169–181 Sajeev K, Osanai Y (2004) Ultrahigh-temperature metamorphism (1150C, 12 kbar) and multi- stage evolution of Mg-, Al-rich granulites from the Central Highland Complex, Sri Lanka. J Petrol 45:1821–1844 Schertl HP, Schreyer W, Chopin C (1991) The pyrope-coesite rocks and their country rocks at Parigi, Dora Maira Massif, Western Alps: detailed petrography, mineral chemistry and PT- path. Contrib Miner Petrol 108:1–21 Schreyer W (1995) Ultradeep metamorphic rocks: the retrospective viewpoint. J Geophys Res 100:8353–8366 Seki Y (1973) Temperature and pressure scale of low-grade metamorphism. J Geol Soc Jpn 79:735–743 Shatsky VS, Sobolev NV, Vavilow MA (1995) Diamond-bearing metamorphic rocks of the Kokchetav massif (Northern Kazakhstan). In: Coleman RG, Wang X (eds) Ultrahigh pressure metamorphism. Cambridge University Press, Cambridge, pp 427–455 Smith RE (1969) Zones of progressive regional burial metamorphism in part of the Tasman geosyncline, eastern Australia. J Petrol 10:144–163 Smith DC, Lappin MA (1989) Coesite in the Straumen kyanite-eclogite pod, Norway. Terra Nova 1:47–56 Song SG, Niu YL, Zhang LF, Bucher K (2009) The Luliangshan garnet peridotite massif of the North Qaidam UHPM belt, NW China – a review of its origin and metamorphic evolution. J Metamorph Geol 27:621–638 Stern CR,Wyllie PJ (1981) Phase relationships of I-type granite with H2O to 35 kilobars: the Dinkey Lakes biotite-granite from the Sierra Nevada batholith. J Geophys Res 86(B11):10412–10422 Stober I, Bucher K (2007) Hydraulic properties of the crystalline basement. Hydrogeol J 15:213–224 St€offler D, Greive RAF (2007) Impactites. In: Fettes D, Desmonds J (eds) Metamorphic rocks – a classification and glossary of terms; recommendations of the International Union of Geologi- cal Sciences Subcommission on the systematics of metamorphic rocks. Cambridge University Press, New York, pp 82–92 Teichm€uller M (1987) Organic material and very low-grade metamorphism. In: Frey M (ed) Low temperature metamorphism. Blackie, Glasgow, pp 114–161 Vernon RH (1976) Metamorphic processes. Allen & Unwin, London, 247 pp Wain A, Waters D, Jephcoat A, Olijynk H (2000) The high-pressure to ultrahigh-pressure eclogite transition in the Western Gneiss Region, Norway. Eur J Miner 12:667–688 Wilson JT (1966) Did the Atlantic close and then re-open? Nature 211:676–681 Wise DU, Dunn DE, Engelder JT, Geiser PA, Hatcher RD, Kish SA, Odon AL, Schamel S (1984) Fault-related rocks: suggestions for terminology. Geology 12:391–394 Yang J, Godard G, Kienast JR, Lu Y, Sun J (1993) Ultrahigh-pressure (60 kbar) magnesite-bearing garnet peridotites from northeastern Jiangsu, China. J Geol 101:541–554 Zhang R-Y, Liou JG (1994) Coesite-bearing eclogite in Henan Province, central China: detailed petrography, glaucophane stability and PT-path. Eur J Miner 6:217–233 References and Further Reading 19 in the last section of this chapter (Sect. 2.5) on graphical representation of mineral assemblages. However, the quantitative computation of equilibriumcomposition phase diagrams is not discussed in this book. 2.1 Primary Material of Metamorphic Rocks All metamorphic rock-forming processes make rocks from other rocks. The precur- sor rock or protolith determines many attributes of the new metamorphic rock. Metamorphism results from the addition (or removal) of heat and material to discrete volumes of the crust or mantle by tectonic or igneous processes. Metamor- phism, therefore, may affect all possible types of rock present in the Earth’s crust or mantle. Protoliths of metamorphic rocks comprises rocks of all possible chemical compositions and include the entire range of sedimentary, igneous and metamor- phic rocks. Metamorphic processes tend to change the original composition of the protolith. Addition of heat to rocks typically results in the release of volatiles (H2O, CO2, etc.) that are stored in hydrous minerals (e.g. clay, micas, amphiboles), carbonates and other minerals containing volatile components. Therefore, many metamorphic rocks are typically depleted in volatiles relative to their protoliths. Metamorphism that releases only volatiles from the protolith is, somewhat illogically, termed isochemical. On a volatile-free basis, the chemical composition of protolith and product rock is identical in isochemical metamorphism. In truly isochemical meta- morphism, protolith and product rocks are of identical composition including the volatile content. Recrystallization of sedimentary calcite to coarse grained calcite marble is an example of isochemical metamorphism in a strict sense. It affects the structure of the rock, no new minerals are formed. Many, if not most, metamorphic processes also change the cation composition of the protolith. This type of metamorphism is termed allochemical metamorphism or metasomatism. The aqueous fluid released by dehydration reactions during meta- morphism generally contains dissolved solutes. These are then carried away with the fluid and lost by the rock system. It has been found, for example, that many granulite facies gneisses are systematically depleted in alkalis (Na and K) relative to their amphibolite facies precursor rocks. This can be explained by alkali loss during dehydration. Silica saturation is a general feature of nearly all metamorphic fluids. Pervasive or channelled regional scale flow of silica-saturated dehydration fluids may strongly alter silica-deficient rocks (ultramafic rocks, dolomite marbles) that come in contact with these fluids. Unique metamorphic rock compositions may result from metasomatism on a local or regional scale. Efficient diffusion and infiltration metasomatism requires the presence of a fluid phase. Metasomatism is a process of fluid–rock interaction combined with transfer of material between the systems at elevated temperature and pressure. Fluid–rock interaction processes are dominant in chemical weathering, in sedimentary, diagenetic, groundwater, hydro- thermal and other near surface environments, but they play an important role also in 22 2 Metamorphic Rocks the metamorphic domain. In fact, isochemical metamorphism of rocks may be viewed as interaction of rocks with an internally derived fluid that makes dissolu- tion of unstable and precipitation of more stable mineral assemblages possible. Interaction of rocks with externally derived fluids is referred to as allochemical metamorphism. The volatile composition of the fluid may not in equilibrium with the mineral assemblage of the rock and, consequently, the rock may be altered. Some examples: flushing of rocks with pure H2O at high-P–T conditions may initiate partial melting, it may form mica and amphibole in pyroxene-bearing rocks, it may induce wollastonite or periclase formation in marbles. CO2 metasomatism is particularly common in very high-grade rocks. Metasomatism can create rocks of extreme composition which, in turn, may serve as a protolith in subsequent metamorphic cycles. Metasomatic rocks of unusual composition are widespread in regional meta- morphic terrains and contact aureoles. However, the total volume of such types of rocks is negligible. Although interesting petrologically, these exotic rocks will not be discussed in Part II where we present a systematic treatment of prograde metamor- phism of the most important types of bulk rock compositions. 2.1.1 Chemical Composition of Protoliths of Metamorphic Rocks The average composition of crust and mantle is listed in Table 2.1. The mantle constitutes the largest volume of rocks on planet Earth. From geophysical and petrophysical evidence and from mantle fragments exposed at the surface, we know that the mantle consists predominantly of ultramafic rocks of the peridotite family. The bulk of the mantle is in a solid state and experiences continuous recry- stallization as a result of large-scale convective flow in the sub-lithospheric mantle and tectonic processes in the lithospheric mantle. Therefore, nearly all mantle rocks represent metamorphic rocks. The composition of the mantle (Table 2.1) is repre- sentative for the most prominent type of metamorphic rock of this planet. However, mantle rocks can only be transported through the lid of crust to the surface of the Earth by active tectonic or igneous processes. Although outcrops of ultramafic Table 2.1 Composition of the Earth’s crust and mantle (After Carmichael 1989) Peridotite mantle Continental crust Oceanic crust Basalt Tonalite SiO2 45.3 60.2 48.6 47.1 61.52 TiO2 0.2 0.7 1.4 2.3 0.73 Al2O3 3.6 15.2 16.5 14.2 16.48 FeO 7.3 6.3 8.5 11.0 5.6 MgO 41.3 3.1 6.8 12.7 2.8 CaO 1.9 5.5 12.3 9.9 5.42 Na2O 0.2 3.0 2.6 2.2 3.63 K2O 0.1 2.8 0.4 0.4 2.1 H2O <0.1 1.4 1.1 <1.0 1.2 CO2 <0.1 1.4 1.4 <1.0 0.1 2.1 Primary Material of Metamorphic Rocks 23 rocks are common and widespread, particularly in orogenic belts, the total volume of ultramafic rocks exposed on continents is small. Crustal rocks may be divided into rocks from oceanic and continental environ- ments. Characteristic compositions of continental (tonalite) and oceanic crust (basalt) are listed in Table 2.1. It is evident that the average composition of oceanic crust is well represented by an average basalt composition, and the average com- position of continental crust can be described by an average tonalite composition. The composition of the continental crust is more heterogeneous than the oceanic crust which is ~99% basaltic. Table 2.2 lists abundances of types of rocks that make up typical crust and are the predominant protoliths of metamorphic rocks. Igneous rocks of mafic composition (basalts, gabbros of Mid Ocean Ridge Basalt [MORB] affinity) form the oceanic crust which covers much larger areas than continental crust, and constitute an important chemical group of metamorphic rocks (greens- chist, amphibolite, granulite, eclogite). Granite and related rocks such as granodiorite and quartz–diorite (typical granite and tonalite compositions given in Table 2.3) dominate the continental crust. They make up the family of metamorphic rocks termed meta-granitoids (¼ quartzo- feldspathic rocks) and represent 33% of all igneous rocks of the Earth’s crust (Table 2.2). Table 2.2 Abundance of rocks (vol%) in the crust (After Carmichael 1989) Igneous rocks 64.7 Sedimentary rocks 7.9 Metamorphic rocks 27.4 Igneous rocks (64.7) Sedimentary rocks (7.9) Granites 16 Shales 82 Granodiorites/diorites 17 Sandstones, arkoses 12 Syenites 0.6 Limestones 6 Basalts/gabbros 66 Peridotites/dunites 0.3 Table 2.3 Chemical composition of sedimentary and igneous rocks (After Carmichael 1989) Sandstones, graywackes Shales (platforms) Pelites, pelagic clays Carbonates (platforms) Tonalite Granite Basalt MORB SiO2 70.0 50.7 54.9 8.2 61.52 70.11 49.2 TiO2 0.58 0.78 0.78 – 0.73 0.42 2.03 Al2O3 8.2 15.1 16.6 2.2 16.48 14.11 16.09 Fe2O3 0.5 4.4 7.7 1.0 – 1.14 2.72 FeO 1.5 2.1 2.0 0.68 5.6 2.62 7.77 MgO 0.9 3.3 3.4 7.7 2.8 0.24 6.44 CaO 4.3 7.2 0.72 40.5 5.42 1.66 10.46 Na2O 0.58 0.8 1.3 – 3.63 3.03 3.01 K2O 2.1 3.5 2.7 – 2.1 6.02 0.14 H2O 3.0 5.0 9.2 – 1.2 0.23 0.70 CO2 3.9 6.1 – 35.5 0.1 C 0.26 0.67 – 0.23 24 2 Metamorphic Rocks Structure. The arrangement of parts of a rock mass irrespective of scale, including geometric interrelationships between the parts, their shapes and internal features. The terms micro-, meso- and mega- can be used as a prefix to describe the scale of the feature. Micro- is used for a thin-section scale, meso- for hand-specimen and outcrop scale, mega- for larger scales. Fabric. The kind and degree of preferred orientation of parts of a rock mass. The term is used to describe the crystallographic and/or shape orientation of mineral grains or groups of grains, but can also be used to describe meso- and mega-scale features. Layer. One of a sequence of near parallel tabular-shaped rock bodies. The sequence is referred to as being layered (equivalent expressions: bands, banded, laminated). Foliation. Any repetitively occurring or penetrative planar structural feature in a rock body. Some examples: l Regular layering on a cm or smaller scale l Preferred planar orientation of inequant mineral grains l Preferred planar orientation of lenticular (elongate) grain aggregates More than one kind of foliation with more than one orientation may be present in a rock. Foliations may become curved (folded) or distorted. The surfaces to which they are parallel are designated s-surfaces (Fig. 2.1). Schistosity. A type of foliation produced by deformation and/or recrystallization resulting in a preferred orientation of inequant mineral grains. It is common practice in phyllosilicate-rich rocks to use the term slaty cleavage instead of schistosity when individual grains are too small to be seen by the unaided eye (Fig. 2.1). Cleavage. A type of foliation consisting of a regular set of parallel or sub- parallel closely spaced surfaces produced by deformation along which a rock body will usually preferentially split. More than one cleavage may be present in a rock. Slaty cleavage. Perfectly developed foliation independent of bedding resulting from the parallel arrangement of very fine-grained phyllosilicates (Fig. 2.1). Fracture cleavage. A type of cleavage defined by a regular set of closely spaced fractures. Crenulation cleavage. A type of cleavage related to microfolding (crenulation) of a pre-existing foliation. It is commonly associated with varying degrees of meta- morphic segregation. Gneissose structure. A type of foliation on hand-specimen scale, produced by deformation and recrystallization, defined by: l Irregular or poorly defined layering l Augen and/or lenticular aggregates of mineral grains (augen structure, flaser structure) l Inequant mineral grains which are present, however, only in small amounts or which display only a weak preferred orientation, thus defining only a poorly developed schistosity Lineation. Any repetitively occurring or penetrative visible linear feature in a rock body (Fig. 2.1). This may be defined by, for example: 2.2 The Structure of Metamorphic Rocks 27 l Alignment of the long axes of elongate mineral grains (¼ mineral lineation) l Alignment of elongate mineral aggregates l Alignment of elongate objects, bodies (e.g. strongly deformed pebbles in a meta-conglomerate) l Common axis of intersection of tabular mineral grains (or bodies) l Intersection of two foliations (intersection lineation) l Parallelism of hinge lines of small scale folds l Slickenside striations l Striations due to flexural slip More than one kind of lineation, with more than one orientation, may be present in a rock. Lineations may become curved or distorted. The lines to which they are parallel are called L-lines. Reference to a lineation is incomplete without indication of the type concerned. Lineation Schistosity (S1) Schistosity (S2) Intersection lineation (L1) Bedding (S0) D1 D2 D1 D2 Mica Quartz + Feldspar Slaty cleavage Bedding a b c a b c Fig. 2.1 Diagrams to illustrate some structure features in metamorphic rocks. (a) Folded sedi- mentary rocks and slaty cleavage; (b) Schist with alternating mica and quartz–feldspar foliation displaying two schistosity planes (S1 and S2) related to deformation events D1 and D2, respec- tively; The first schistosity, S1, is parallel to sedimentary bedding planes (S0); intersection of S1 and S2 planes produce an intersection lineation (L1); (c) Examples of metamorphic lineation, a ¼ alignment of mineral aggregates; b ¼ alignment of elongate mineral grains (e.g. hornblende); c ¼ alignment of tabular mineral grains (e.g. mica) 28 2 Metamorphic Rocks Joint. A single fracture in a rock with or without a small amount (<1 cm) of either dilatational or shear displacement (joints may be sealed by mineral deposits during or after their formation). Cataclasis. Rock deformation accomplished by some combination of fracturing, rotation, and frictional sliding producing mineral grain and/or rock fragments of various sizes and often of angular shape. Metamorphic differentiation. Redistribution of mineral grains and/or chemical components in a rock as a result of metamorphic processes. Metamorphic process by which mineral grains or chemical components are redistributed in such a way to increase the modal or chemical anisotropy of a rock (or portion of a rock) without changing the overall chemical composition. Textural zones. Regional geological mapping of metamorphic terranes is typically based on criteria such as lithologic associations, metamorphic zones and struc- tural zones. Field and petrographic subdivision of metamorphic rocks has also been made on the basis of textural zones that subdivide rocks in terms of the degree of recrystallization, e.g. foliation, mineral segregation, increasing grain size, with increasing metamorphism. For example, four textural zones have been established for quartzofeldspathic rocks (metagreywacke) that comprise the Otago Schist, New Zealand (Fig. 2.2) in which the boundaries of the 10–50 km wide zones are gradational and are termed isotects (Bishop 1972; Turnbull et al. 2001). The macroscopic/microscopic criteria used to distinguish textural zones are essentially a mapping tool and cannot be used as a mineralogical or isochemical P–T indicator. 2.3 Classification and Names of Metamorphic Rocks The names of metamorphic rocks are usually straightforward and self-explanatory. The number of special terms and cryptic expressions is relatively small. Neverthe- less in order to be able to communicate with other geologists working with metamorphic rocks it is necessary to define commonly used names and expressions and to briefly review currently used classification principles for metamorphic rocks. There is not one sole classification principle used for the description of meta- morphic rocks, which consequently means that all metamorphic rocks may have a series of perfectly correct and accepted names. However, modal mineral compo- sition and mesoscopic structure are the main criteria for naming metamorphic rocks. In addition, the composition and the nature of the protolith (original material) is an important classification criterion. Finally, some well-established special names are used also in metamorphic geology. The names of metamorphic rocks consist of a root name and a series of prefixes. The root of the name may be a special name (e.g. amphibolite) or a name describing 2.3 Classification and Names of Metamorphic Rocks 29 2.3.2 Names for High-Strain Rocks Metamorphism may locally be associated with an extremely high degree of rock deformation. Localized high strain in metamorphic terrains produces rocks with distinctive structures. Some widely used special names for high-strain rocks are defined below: IV III IIB I IIA 100 km Otago Alpine Marlborough N Fig. 2.3 Map showing distribution of textural zones (TZ’s) delineating transition from non- schistose to schistose metagreywacke rocks, Otago, South Island of New Zealand (inset map shows distribution of Otago, Alpine and Marlborough schist) (Redrawn from Turnbull et al. 2001, Fig. 1D). TZI. Non foliated or spaced fracture cleavage. TZIIA. Weak-moderate, anastomosing, penetrative foliation that is stronger in fine-grained rocks. Sedimentary bedding dominant over cleavage. Rocks break into wedge-shaped blocks. Lowest grade semischist (sandstone ¼ phyllite; mudstone¼ slate). TZIIB. Strong, penetrative foliation. Cleavage dominant over bedding which is transposed. No foliation parallel segregation of quartz and mica. Rocks break into parallel-sided slabs. Highest grade semischist (sandstone ¼ phyllite; mudstone ¼ slate). TZIII. Strong, penetra- tive foliation undulating on millimeter-scale. Distributed foliation-parallel quartz segregation lenses <1 mm thick. Psammitic-pelitic contacts sharp on millimeter scale. Lowest grade schist. TZIV. Strong, penetrative foliation undulating on millimeter- to centimeter-scale. Segregation banding (<1 mm thick) of quartz and mica. Psammitic–pelitic contacts blurred on millimeter-scale but distinguishable on centimeter-scale. Higher grade schist 32 2 Metamorphic Rocks Mylonite.A rock produced by mechanical reduction of grain size as a result of ductile, non-cataclastic deformation in localized zones (shear zones, fault zones), resulting in the development of a penetrative fine-scale foliation, and often with an asso- ciated mineral and stretching lineation (Fig. 2.2). Ultramylonite.Amylonite in which most of the megacrysts or lithic fragments have been eliminated (>90% fine-grained matrix). Augen mylonite (blastomylonite). A mylonite containing distinctive large crystals or lithic fragments around which the fine-grained banding is wrapped. Cataclasite. A rock which underwent cataclasis. Fault breccia. Cataclasite with breccia-like structure formed in a fault zone. Pseudotachylite. Ultra-fine-grained vitreous-looking material, flinty in appearance, occurring as thin veins, injection veins, or as a matrix to pseudo-conglomerates or -breccias, which seals dilatancy in host rocks displaying various degrees of fracturing. 2.3.3 Special Terms Some commonly used and approved special terms (names, suffixes, prefixes) include: Mafic minerals. Collective expression for ferro-magnesian minerals. Felsic minerals. Collective term for quartz, feldspar, feldspathoids and scapolite. Mafic rock. Rock mainly consisting of mafic minerals (mainly  modally >50%). Felsic rock. rock mainly consisting of felsic minerals. Meta-. If a sedimentary or igneous origin of a metamorphic rock can be identified, the original igneous or sedimentary rock term preceded by “meta” may be used (e.g. metagabbro, metapelite, metasediment, metasupracrustal). Also used to generally indicate that the rock in question is metamorphic (e.g. metabasite). Ortho- and para-. A prefix indicating, when placed in front of a metamorphic rock name, that the rock is derived from an igneous (ortho) or sedimentary (para) rock respectively (e.g. orthogneiss, paragneiss). Acid, intermediate, basic, ultrabasic. Terms defining the SiO2 content of igneous and metamorphic rocks (respectively, >63, 63–52, 52–45, <45 wt% SiO2). Greenschist and greenstone. Schistose (greenschist ) or non-schistose (greenstone) metamorphic rock whose green color is due to the presence of minerals such as chlorite, actinolite, and epidote (greenschist e.g. epidote-bearing actinolite- chlorite schist; greenstone, e.g. chlorite-epidote granofels). Blueschist. Schistose rock whose bluish color is due to the presence of sodic amphibole (e.g. glaucophane schist). However, the “blue” color of a blueschist will not easily be recognized by a non-geologist (i.e. it is not really blue, although very rare outcrops of really blue glaucophanites do exist). Blueschists are schistose rocks containing amphibole with significant amounts of the M(4) cation position in the amphibole structure occupied by Na (glaucophane, crossite). 2.3 Classification and Names of Metamorphic Rocks 33 Amphibolite. Mafic rock predominantly composed of hornblende (>40%) and plagioclase. Granulite. Metamorphic rock in or from a granulite facies terrain exhibiting characteristic granulite facies mineral assemblages. Anhydrous mafic minerals are modally more abundant than hydrous mafic minerals. Muscovite is absent in such rocks. Characteristic is the occurrence of metamorphic orthopyroxene in both mafic and felsic rocks. The term is not used for marbles and ultramafic rocks in granulite facies terranes. Charnockite, mangerite, jotunite, enderbyite. Terms applied to orthopyroxene- bearing rocks with igneous texture and granitic (charnockite), monzonitic (man- gerite, jotunite), and tonalitic (enderbyite) composition, irrespective of whether the rock is igneous or metamorphic. Eclogite. A plagioclase-free mafic rock mainly composed of omphacite and garnet, both of which are modally adundant. Eclogitic rock. Rocks of any composition containing diagnostic mineral assem- blages of the eclogite facies (e.g. jadeite–kyanite–talc granofels). Marble. A metamorphic rock mainly composed of calcite and/or dolomite (e.g. dolomitic marble). Calc-silicate rock. Metamorphic rock which, besides 0–50% carbonates, is mainly composed of Ca-silicates such as epidote, zoisite, vesuvianite, diopside– hedenbergite, Ca-garnet (grossular–andradite), wollastonite, anorthite, scapolite, Ca-amphibole. Skarn. A metasomatic Ca–Fe–Mg–(Mn)-silicate rock often with sequences of compositional zones and bands, formed by the interaction of a carbonate and a silicate system in mutual contact. Typical skarn minerals include, wollastonite, diopside, grossular, zoisite, anorthite, scapolite, margarite (Ca skarns); heden- bergite, andradite, ilvaite (Ca–Fe skarns); forsterite, humites, spinel, phlogopite, clintonite, fassaite (Mg skarns); rhodonite, tephroite, piemontite (Mn skarns). Blackwall. A chlorite- or biotite-rich rock developed by metasomatic reaction between serpentinised ultramafic rocks and mafic rocks or quartzo-feldspathic rocks, respectively. Rodingite. Calc-silicate rock, poor in alkalis and generally poor in carbonates, generated by metasomatic alteration of mafic igneous rocks enclosed in serpen- tinized ultramafic rocks. The process of rodingitization is associated with oce- anic metamorphism (serpentinization of peridotite, rodingitization of enclosed basic igneous rocks such as gabbroic/basaltic dykes). Metarodingite is a pro- grade metamorphic equivalent of rodingite produced by oceanic metamorphism. Quartzite or metachert. A metamorphic rock containing more than about 80% quartz. Serpentinite. An ultramafic rock composed mainly of minerals of the serpentine group (antigorite, chrysotile, lizardite), e.g. diopside–forsterite–antigorite schist. Hornfels. Is a non-schistose very fine-grained rock mainly composed of silicate  oxide minerals that shows substantial recrystallization due to contact meta- morphism. Hornfelses often retain some features inherited from the original rock such as graded bedding and cross-bedding in hornfelses of sedimentary origin. 34 2 Metamorphic Rocks structures), it is often difficult to determine mineral assemblages that represent a particular stage in the evolutionary history of the rock. One has to keep in mind that mineral assemblages are usually identified in thin sections which represent a two-dimensional section through a volume of rock. In such sections, only two and a maximum of three minerals may be in mutual grain contact. In the three-dimensional rock a maximum of four minerals can be in contact at a point in space, three minerals form contacts along a line, and two minerals contact along surfaces (that show as lines in thin sections). It is therefore strongly recommended to study more than one thin section from the most signifi- cant and interesting samples. As an example, in a series of 20 thin-sections of a single hand specimen of a coarse-grained sapphirine–granulite from the Western Alps, unique assemblages and unique structures were found in nearly all of the sections. The problem is especially acute in coarse-grained samples, where the scale of chemical homogeneity may considerably exceed the size of a thin section. In extremely fine-grained samples the mineral assemblage must be determined by X-ray techniques. In this case it is not possible to maintain the requirement of mutual grain contact in the definition of an assemblage. However, this shortcoming may be overcome by using the high magnification backscatter image mode of an electron probe microanalyser (EPMA). Rocks should always be examined by X-ray techniques in order to identify minerals which are difficult to distinguish under the microscope (e.g. muscovite, paragonite, talc, pyrophyllite, also quartz and untwinned albite). Staining techniques help in distinguishing some important rock-forming miner- als with similar optical properties such as calcite from dolomite. Minerals occurring as inclusions in refractory minerals such as garnet but not in the matrix of the rock do not belong to the main matrix assemblage. In the example of the staurolite–garnet–biotite–kyanite assemblage cited above, garnet may show small composite two-phase inclusions of chlorite and chloritoid. In this case, chlorite–chloritoid–garnet constitutes another, earlier assemblage of the rock. During metamorphism some earlier-formed minerals may become unstable and react chemically to form a new, more stable assemblage. However, metastable relics of the early assemblage may partly survive. Great care must be taken in the study of metamorphic micro-structures in order to avoid mixing up mineral assem- blages. The correct identification of a successive series of mineral assemblages, i.e. the paragenesis of a metamorphic rock, represents the “great art” of metamorphic petrology. It can be learned only by experience. Table 2.4 Practical determination of a mineral assemblage Mineral Staurolite Garnet Biotite Kyanite Staurolite X X X X Garnet X X X Biotite X X Kyanite X A cross in the Grt–St cell means; garnet and staurolite have mutual grain contacts in thin section 2.4 Mineral Assemblages and Mineral Parageneses 37 Figure 2.4 shows some of the aspects related to the recognition of mineral assemblages. Three fictive rocks all contain the minerals, quartz, calcite and wollastonite on the scale of a thin section. The general micro-structure of the three sections (Fig. 2.4) shows the distribution of the three minerals in their respective rocks. Rock A is clearly heterogeneous on the scale of a thin section, the upper part contains the assemblage Qtz + Cal, the lower half of the section is free of carbonate and contains the assemblage Qtz + Wo. The rock does not contain the assemblage Qtz + Cal + Wo. The two parts of the section are different in overall composition. Note, however, on a volatile-free basis, the two domains of the rock may have very similar or identical compositions (same Ca/Si ratio). In rock B, obviously all three minerals can be found in mutual grain contact, and Qtz + Cal + Wo constitutes the mineral assemblage. In rock C, which appears to be composi- tionally homogeneous on thin-section scale, Qtz + Cal and Qtz + Wo form common grain boundaries. However, no Wo + Cal grain boundaries can be observed. Thus, although the three phases Qtz + Cal + Wo do not strictly represent a mineral assemblage, many petrologists would probably approve it being labeled as an equilibrium assemblage. [Comment: after the publication of the first edition of this book Fig. 2.4 (originally Fig. 2.1), inspired a lively discussion among petro- logists on the geo-metamorphism mailbase. A number of comments were made and it was suggested that the figure be modified. We prefer to keep it essentially in its original form and some aspects of the debate will be discussed in Sect. 3.7. We strongly recommend that the reader and anyone interested in rock metamorphism subscribe to the mailbase geo-metamorphism where one sometimes captures a stimulating dispute on questions related to geo-metamorphism. The present address is: geo-metamorphism@jiscmail.ac.uk: you need to send an e-mail with only the following in the body of the message: subscribe geo-metamorphism]. a b c quartz calcite wollastonite Fig. 2.4 (a–c) Mineral assemblages in three different rocks all containing wollastonite, calcite and quartz 38 2 Metamorphic Rocks 2.5 Graphical Representation of Metamorphic Mineral Assemblages Once the mineral assemblage of a rock has been identified, it is convenient or even necessary to represent graphically the chemical composition of the minerals that constitute the assemblage. Such a figure is called a chemograph and represents a composition phase diagram. The geometric arrangement of the phase relation- ships on such a phase diagram is called the topology. Composition phase diagrams can be used only to document the assemblages found in rocks of a given metamor- phic terrain or outcrop. However, such diagrams are an indispensable tool for the analysis of metamorphic characteristics and evolution of a terrain. They can be used to deduce sequences of metamorphic mineral reactions. Finally, composition phase diagrams can also be calculated theoretically from thermodynamic data of minerals which permits the quantitative calibration of field-derived chemographs. Composition phase diagrams display the chemical composition of minerals and the topologic relationships of mineral assemblages. Variables on the diagrams are concentrations or amounts of chemical entities. All other variables that control the nature of the stable mineral assemblage such as pressure and temperature must be constant. Thus, composition phase diagrams are isothermal and isobaric. Also, not more than two composition variables can conveniently be displayed on a two- dimensional xy-diagram (sheet of paper, computer screen). 2.5.1 Mole Numbers, Mole Fractions and the Mole Fraction Line It is useful to change the scale for the compositional variables from wt% to mol%, mole fractions or mole numbers. Most chemographs use mol% or mole fraction as units for the composition variables. The mineral forsterite (Fo), for example, is composed of 42.7 wt% SiO2 and 57.3 wt%MgO. Mole numbers and mole fractions for this mineral are calculated as follows: SiO2: 42.7/60.1 (molecular weight SiO2) ¼ 0.71 (number of moles of SiO2 per 100 g Fo) MgO: 57.3/40.3 (molecular weight MgO) ¼ 1.42 (number of moles of MgO per 100 g Fo) Forsterite has a MgO/SiO2 ratio of 1.42/0.71 ¼ 2 and thus has 2 mol MgO per 1 mol SiO2. The composition is reported as Mg2SiO4 or (2MgO SiO2) or 66.66% MgO + 33.33% SiO2 or 2/3MgO + 1/3SiO2. This is all equivalent. However, the last version has many advantages ) mole fraction basis. The mole fraction is defined as follows: XMgO ¼ Number of moles of MgOð Þ Number of moles of MgOð Þþ Number of moles of SiO2ð Þ 2.5 Graphical Representation of Metamorphic Mineral Assemblages 39 The total number of moles of system components is 8 and the composition of talc can be normalized to a total number of 8 mol of system components. In this case the talc composition will be expressed by ) mole fractions: XH2O ¼ 1/8, XMgO ¼ 3/8, XSiO2 ¼ 4/8, or: XH2O ¼ 0.125, XMgO ¼ 0.375, XSiO2 ¼ 0.50; in mol%: H2O ¼ 12.5%, MgO¼ 37.5%, SiO2¼ 50%. Graphic representation of the talc composition on a mole fraction basis is given by the intersection of the talc phase vector in Fig. 2.6 and the plane S Xi ¼ 1. The mole fraction plane is a regular triangle with the corners Xi ¼ 1. This triangle is called the mole fraction triangle. A representation of the mole fraction triangle is shown in Fig. 2.7. The lines of constant Xi are parallel with the base lines of the triangle. This follows from Fig. 2.6, where it can be seen that the planes Xi ¼ constant and Xj 6¼ i ¼ 0 intersect the mole fraction plane along lines parallel to the base line Xi ¼ 0. Three rulers with Xi¼ 0.1 increments are also shown in Fig. 2.7 for the three components. Triangular coordinate paper is available on the market. A further example of a mineral composition in a three-component system MgO–SiO2–H2O is the chemical formula of chrysotile (one of many serpentine minerals) that can be written as, H4Mg3Si2O9. The sum of moles of system compo- nents is 7; ) mole fractions: XH2O ¼2/7, XMgO ¼ 3/7 , XSiO2 ¼ 2/7, or equivalent: XH2O ¼ 0.29, XMgO ¼ 0.42, XSiO2 ¼ 0.29; expressed in mol%: H2O¼ 29%, MgO¼ 42%, SiO2 ¼ 29%. Because triangular coordinate drawing paper is not always at hand for plotting composition data, the formulae for recalculation into Cartesian coordinates are given in Fig. 2.8 together with the position of the serpentine composition. The value of the scaling factor “f” depends on the desired size of the figure. The scaling factor must be multiplied by 100 when using mole fractions rather than mol%. Talc intersection of talc phase vector with the mole fraction triangle XMgO = 1 XSiO2 = 1XH2O = 1 H2Mg3Si4O12 X S iO 2 = 0 0.38 0.13 0. 50Fig. 2.7 Mole fraction triangle MgO–H2O–SiO2 and intersection coordinates of the talc phase vector with the mole fraction triangle 42 2 Metamorphic Rocks 2.5.3 Projections 2.5.3.1 Simple Projections On a mole fraction triangle two compositional variables of a three-component system can be depicted. Most rocks, however, require more than three components to describe and to understand the phase relationships. Graphical representation of an eight component system requires a seven-dimensional figure. Projection phase diagrams are graphical representations of complex n-component systems that show two composition variables at a time while keeping the other n  3 composition variables constant. The remaining variable is given by the mole fraction equation as outlined above. In a suite of samples of similar bulk composition (e.g. 20 samples of metapelite), one often finds certain mineral species that are present in many of them. This circumstance permits projection of the phase compositions from the composition of one such mineral which is present in excess. For instance, in many metapelitic rocks quartz is modally abundant whereas calcite is the main mineral in most marbles. 0 2 4 6 8 10 2 4 6 8 10 units : mole % y (cm) x (cm) 2.5 4.3 MgO SiO2 H2O 42% M gO 29 % S iO 2 29% H2O 10 0% S iO 2 0% S iO 2 H4Mg3Si2O9 intersection of the chrysotile phase vector with the mole fraction triangle: calculated and displayed in a Cartesian xy-diagram (x = 4.3 cm; y = 2.5 cm). cos (30) = 0.86603; tan (30) = 0.57735 f = 10/100 (scaling factor) y (cm) = f * (mole % of component) * cos (30) x (cm) = {y * tan (30)} + ( f * mole % of component) Fig. 2.8 Mole fraction triangle MgO–H2O–SiO2 in the Cartesian coordinate system 2.5 Graphical Representation of Metamorphic Mineral Assemblages 43 Composition phase diagrams for metapelites can therefore be constructed by projecting through SiO2 onto an appropriate mole fraction triangle, and for marbles by projection the phase compositions from CaCO3. The system MgO–SiO2–H2O (MSH system) can be used to explain the basic principle of making projections. Some phase compositions in the MSH system are given in Table 2.5. This table represents a composition matrix with oxide compo- nents defining unit vectors of the composition space and the mineral compositions used as column vectors. Figure 2.9 shows the chemographic relationships in the ternary system H2O–MgO–SiO2. The corners of the triangle represent Xi ¼ 1 or 100 mol% of the system components. The lines connecting them are the binary subsystems of the three-component system. The compositions of the black circles occur as stable phases in nature (phase components). The ternary system has three binary sub- systems (MgO–SiO2, MgO–H2O, and SiO2–H2O). Some phase compositions in the ternary system fall on straight lines such as Tlc–Ath–En and Brc–Atg–Tlc. This means that the phase compositions along a straight line, e.g. Tlc–Ath–En, are linearly dependent; one of these compositions can be expressed by the other two (4En + Tlc ¼ Ath). Therefore, there are only two components required to describe the compositions of the other phases on the straight line) pseudobinary join. The colinearity is also said to be a compositional degeneracy in the system. Now, one might wish to analyze and discuss phase relationships among the minerals shown in Fig. 2.9 for geologic situations in which an aqueous fluid phase (H2O) is present in all rocks and in equilibrium with the solid phase assemblage. The presence of excess water in all the considered rocks permits projection of the other phase compositions through water onto the MgO–SiO2-binary (in principle on any pseudobinary as well). Imagine that you are standing in the H2O corner of Fig. 2.9. What you will see from there is shown at the bottom of Fig. 2.9. The Table 2.5 Phase compositions in the MSH-system (a) Composition matrix (moles); columns are phase vectors, composition space defined by the system components SiO2, MgO, H2O Fo Brc Tlc En Ath Qtz Per Atg Fl SiO2 1 0 4 2 8 1 0 2 0 MgO 2 1 3 2 7 0 1 3 0 H2O 0 1 1 0 1 0 0 2 1 Sum 3 2 8 4 16 1 1 7 1 (b) Composition matrix (mole fractions); columns are normalized phase vectors, values are coordinates on the molefraction triangle SiO2 0.33 0.00 0.50 0.50 0.50 1.00 0.00 0.29 0.00 MgO 0.67 0.50 0.38 0.50 0.44 0.00 1.00 0.43 0.00 H2O 0.00 0.50 0.13 0.00 0.06 0.00 0.00 0.29 1.00 Sum 1 1 1 1 1 1 1 1 1 (c) Composition matrix (mole fractions); projected through H2O, columns are normalized phase vectors, values are coordinates on the SiO2–MgO-binary (mole fraction line) SiO2 0.33 0.00 0.56 0.50 0.53 1.00 0.00 0.40 0.00 MgO 0.67 1.00 0.43 0.50 0.47 0.00 1.00 0.60 0.00 44 2 Metamorphic Rocks rewritten in terms of a set of new system components. One of these new compo- nents must be the desired new projection composition. For example, we would like to prepare a figure representing the phase composi- tions in the MSH system on a mole fraction triangle with the corners Mg2SiO4 (Fo), Mg(OH)2 (Brc), and Mg3Si4O10(OH)2 (Tlc). Secondly, we would like to study phase relationships in rocks which contain excess forsterite and need a projection from Mg2SiO4 (Fo) onto the Brc–Tlc-binary. The solution to the problem is shown in Table 2.6. In Table 2.6a the phase compositions are expressed in terms of the new system components Mg2SiO4 (Fo), Mg(OH)2 (Brc), and Mg3Si4O10(OH)2 (Tlc). As an example, the composition of enstatite can be expressed by 2Fo  1Brc + 1Tlc, which is equivalent to 6MgO + 6SiO2. Table 2.6b gives the coordinates of the mineral compositions in the mole fraction triangle Fo–Brc–Tlc. The algebraic operation which transforms the com- position space expressed in terms of simple oxide components (Table 2.5a) to the composition space expressed in terms of Mg2SiO4 (Fo), Mg(OH)2 (Brc), and Mg3Si4O10(OH)2 (Tlc) components (Table 2.6a) is also given in Table 2.6. It can be seen that the operation is a pre-multiplication of Table 2.5a by the inverse of the leading 3  3 square matrix in Table 2.5a. The result of the operation is the composition matrix (Table 2.7a) with the mineral compositions expressed by the new set of system components. Today, any standard commercial spreadsheet Table 2.6 Phase compositions in the Fo–Brc–Tlc-system (a) Composition matrix in terms of moles; columns are phase vectors, composition space defined by the system components Fo Brc Tlc En Ath Qtz Per Atg Fl Fo 1.00 0.00 0.00 2.00 4.00 1.00 4.00 0.00 4.00 Brc 0.00 1.00 0.00 1.00 2.00 1.00 1.00 3.00 5.00 Tlc 0.00 0.00 1.00 1.00 5.00 1.00 1.00 1.00 1.00 (b) Composition matrix in terms of mole fractions; columns are normalized phase vectors, values are coordinates on the mole fraction triangle Fo 1.00 0.00 0.00 1.00 0.57 1.00 1.00 0.00 2.00 Brc 0.00 1.00 0.00 0.50 0.29 1.00 0.25 0.75 2.50 Tlc 0.00 0.00 1.00 0.50 0.71 1.00 0.25 0.25 0.50 Initial composition matrix (old system components): [A][BOC] (Table 2.5a) New matrix (new system components) [I][BNC] (Table 2.6a) Matrix operation: [A1][A][BOC] ) [I][BNC] Old basis [A] Inverse of old basis [A1] Identity matrix [I] 1 0 4 0.33 0.67 0.67 1 0 0 2 1 3 0.33 0.17 0.83 0 1 0 0 1 1 0.33 0.17 0.17 0 0 1 (c) Composition matrix in terms of mole fractions; projection through Mg2SiO4 (Fo), columns are normalized phase vectors, values are coordinates on the mole fraction line Fo Brc Tlc En Ath Qtz Per Atg Fl Brc 0.00 1.00 0.00 1.00 0.69 1.00 0.00 0.75 0.83 Tlc 0.00 0.00 1.00 1.00 1.69 1.00 0.00 0.25 0.17 2.5 Graphical Representation of Metamorphic Mineral Assemblages 47 T a b le 2 .7 C o m p o si ti o n m at ri x fo r m in er al s in th e K F M A S H -s y st em (a ) C o m p o si ti o n m at ri x in te rm s o f o x id e co m p o n en ts K y Q tz M s F l F eO M g O A lm P rp A n n P h l F E s E s S t C ld C rd O P X S p l C h l S iO 2 1 1 6 0 0 0 3 3 6 6 5 5 8 2 5 1 .8 0 3 A l 2 O 3 1 0 3 0 0 0 1 1 1 1 2 2 9 2 2 0 .2 1 1 F eO 0 0 0 0 1 0 3 0 6 0 5 0 4 2 0 0 .5 0 0 M g O 0 0 0 0 0 1 0 3 0 6 0 5 0 0 2 1 .3 1 5 K 2 O 0 0 1 0 0 0 0 0 1 1 1 1 0 0 0 0 .0 0 0 H 2 O 0 0 2 1 0 0 0 0 2 2 2 2 2 2 0 0 .0 0 4 (b ) In v er se o f le ad in g 6  6 sq u ar e m at ri x (A -m at ri x ) A 1 ) 0 .0 0 1 .0 0 0 .0 0 0 .0 0 3 .0 0 0 .0 0 1 .0 0 1 .0 0 0 .0 0 0 .0 0 3 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 2 .0 0 1 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .0 0 (c ) C o m p o si ti o n m at ri x in te rm s o f K y , Q tz , M s, H 2 O , F eO an d M g O K y Q tz M s F l F eO M g O A lm P rp A n n P h l F e- E s E s S t C ld C rd O P X S p l C h l K y 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 1 .0 0 2 .0 0 2 .0 0 1 .0 0 1 .0 0 9 .0 0 2 .0 0 2 .0 0 0 .2 0 1 .0 0 1 .0 0 Q tz 0 .0 0 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 2 .0 0 2 .0 0 2 .0 0 2 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 3 .0 0 1 .6 0 1 .0 0 2 .0 0 M s 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 1 .0 0 1 .0 0 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 F l 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 2 .0 0 2 .0 0 0 .0 0 0 .0 0 0 .0 0 4 .0 0 F eO 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 3 .0 0 0 .0 0 6 .0 0 0 .0 0 5 .0 0 0 .0 0 4 .0 0 2 .0 0 0 .0 0 0 .5 0 0 .0 0 0 .0 0 M g O 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 3 .0 0 0 .0 0 6 .0 0 0 .0 0 5 .0 0 0 .0 0 0 .0 0 2 .0 0 1 .3 0 1 .0 0 5 .0 0 (d ) R en o rm al iz ed co m p o si ti o n m at ri x ,c o lu m n v ec to rs ar e co o rd in at es o f m in er al co m p o si ti o n s in th e m o le fr ac ti o n tr ia n g le K y (A ), F eO (F ) an d M g O (M ) () A F M -d ia g ra m ), p ro je ct io n th ro u g h Q tz , M s an d H 2 O K y Q tz M s F l F eO M g O A lm P rp A n n P h l F e- E s E s S t C ld C rd O P X S p l C h l K y 1 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .2 5 0 .2 5 0 .5 0 0 .5 0 0 .2 5 0 .2 5 0 .6 9 0 .5 0 0 .5 0 0 .1 0 0 .5 0 0 .1 7 F eO 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .7 5 0 .0 0 1 .5 0 0 .0 0 1 .2 5 0 .0 0 0 .3 1 0 .5 0 0 .0 0 0 .2 5 0 .0 0 0 .0 0 M g O 0 .0 0 0 .0 0 0 .0 0 0 .0 0 0 .0 0 1 .0 0 0 .0 0 0 .7 5 0 .0 0 1 .5 0 0 .0 0 1 .2 5 0 .0 0 0 .0 0 0 .5 0 0 .6 5 0 .5 0 0 .8 3 48 2 Metamorphic Rocks program running on a PC or MAC will do these algebraic operations for you (e.g. Excel). We are now ready for the construction of the desired forsterite projection. Just as in simple projections, delete the Fo-row in Table 2.6b and renormalize to constant sum. Table 2.6c shows the coordinates of the phase compositions along the Brc–Tlc binary. Some of the compositions project to the negative side of talc, and periclase cannot be projected onto the Brc–Tlc binary at all. This is apparent also from Fig. 2.9, where it can be recognized that, seen from forsterite, periclase projects away from the Brc–Tlc binary. It is also clear from the procedure outlined above, that the compositions which one chooses to project from or which one wants to see at the apexes of the mole fraction triangle must be written as column vectors in the original A-matrix. With the two basic operations, projection and redefinition of system compo- nents, one can construct any thermodynamically valid composition phase diagram for any geologic problem. 2.5.3.4 AFM Projections A classical example of a composite projection is the AFM projection for metapelitic rocks (see also Chap. 7). Many of the phase relationships in metapelitic rocks can be described in the six-component system K2O–FeO–MgO–Al2O3–SiO2–H2O (KFMASH system). A graphical representation of the system requires projection from at least three fixed compositions. Many of the metapelitic rocks contain excess quartz and some aspects of metamorphism can be discussed for water-present conditions. Therefore, a projection from SiO2 and H2O can be easily prepared. However, none of the remaining four components is present as a phase in such rocks. On the other hand, in low- and medium-grade metapelites muscovite is normally present as an excess phase, and in high-grade rocks muscovite is replaced by K-feldspar. Therefore, a useful diagram could be prepared by projecting through muscovite or K-feldspar onto the plane Al2O3–FeO–MgO. Furthermore, under the condition of excess quartz an Al2SiO5 polymorph is always more stable than corundum. This, in turn, requires that the composition matrix for minerals in metapelitic rocks is rewritten in terms of the new system components KAl3- Si3O10(OH)2–Al2SiO5–FeO–MgO–SiO2–H2O as shown in Table 2.7. The coordi- nates of the phase compositions in the AFM mole fraction triangle can then be represented in an AFM diagram (Fig. 2.11). All Mg–Fe-free Al-silicates project to the A apex, pure enstatite, talc and anthophyllite are located at the M apex, ferrosilite and Fe-anthophyllite project to the F apex. Biotites have negative A-coordinates. The iron-magnesium substitution (FeMg1 exchange) in the ferro- magnesian minerals is parallel to the AM binary, the Mg-tschermak substitution (2AlSi1Mg1 exchange) is parallel to the AM binary. Minerals such as staurolite, chloritoid, garnet, cordierite and spinel do not show any tschermak variation, their compositional variation is exclusively parallel to the FeMg1 exchange vector. Other minerals such as chlorites, biotites, orthopyroxenes and orthoamphiboles show both FeMg1 exchange and 2AlSi1Mg1 exchange. The compositions of 2.5 Graphical Representation of Metamorphic Mineral Assemblages 49 assemblage (open squares in Fig. 2.12). All bulk rock compositions which project in the three-phase field Crd + Bt + Sil will be composed of these three minerals of identical composition in modal proportions depending on the com- position of the rock. Three-phase fields are invariant fields at constant P and T, because the mineral compositions do not vary with rock composition. The cordierite coexisting with biotite and sillimanite has the most Fe-rich composi- tion of all cordierite at the P–T conditions of the AFM diagram. Note that, in general, all three-phase fields on an AFM diagram must be separated from each other by two-phase fields. In the three-phase field Grt + Sil + Bt, the mineral compositions are controlled by the assemblage as well (open circles in Fig. 2.12). It follows that biotite in this assemblage must be more Fe-rich than biotite in the assemblage Crd + Sil + Bt. The garnet coexisting with Bt + Sil has the mostMg-rich composition of all garnets at the given pressure and temperature. AFM-type diagrams will be extensively used for discussing metamorphism in metapelitic rocks in Chap. 7. 2.5.3.5 ACF Projections Composition phase diagrams for assemblages in marbles, calcsilicate rocks, calcar- eous metapelites and other rocks with calcic phases such as mafic rocks, can be prepared, for example, for a simple CFMAS system. Calcic phases may include amphiboles, plagioclase, epidote, diopside and carbonate minerals. A graphic representation of the phase relationships can, for instance, be made by projecting from quartz (if present in excess) and from a CO2–H2O fluid phase of constant composition onto the mole fraction triangle Al2O2–CaO–FeO (ACF diagram, Fig. 2.13). The ACF diagram is also a projection parallel to the FeMg1 exchange vector. All minerals of the AFM system can also be represented on ACF diagrams provided that one also projects through muscovite or K-feldspar. The coordinates of mineral compositions can be calculated as explained above for the MSH and KFMASH systems, respectively. Figure 2.14 is a typical ACF diagram, showing phase relationships at a certain pressure and temperature. The tschermak variation is parallel to the AF binary and affects mainly chlorite and amphibole (and pyroxene which is not present at the P–T conditions of the figure). The three- phase fields Pl + Am + Grt and Ky + Grt + Pl are separated by a two-phase field Grt + Pl because of CaMg1-substitution in garnet (grossular component). Both garnet and plagioclase show no compositional variation along the TS exchange vector. Any mineral which can be described by the components K2O–CaO–FeO–MgO–A- l2O3–SiO2–H2O– CO2 can be represented on an ACF diagram such as Fig. 2.13. However, the consequences of FeMg1 substitution in minerals cannot be discussed by means of ACF diagrams. Therefore, any discontinuous reaction relationship deduced from an ACF diagram is in reality continuous and dependent on the Fe–Mg variation (if it involves Fe–Mg minerals). For example, replacement of the Pl–Grt tie line by a more stable tie line between kyanite and amphibole can be related to the 52 2 Metamorphic Rocks reaction: Pl + Grt) Am + Ky Qtz H2O. Equilibrium of the reaction, however, depends not only on pressure and temperature but also on the Fe–Mg variation in garnet and amphibole. The projection coordinates of a selection of mineral compo- sitions is shown in the chemograph above Fig. 2.13. In Fig. 2.14, the same mineral chemography is shown together with two protolith compositions listed in Table 2.3. As the P–T conditions shown on Fig. 2.14 are those of the amphibolite facies (Barrovian kyanite zone) conditions, it can be seen that metamorphosed MOR basalt, would consist of Pl + Am + Grt and would be rich in amphibole. The platform carbonate plots in the Cal + Am + Zo triangle. 2.5.3.6 Other Projections Any other graphic representation of phase relationships on composition phase diagrams can be prepared by the procedure outlined above. The type of graphic representation of assemblages is entirely dictated by the material one is working with and by the problem one wants to solve. The following steps are a guide to the production of adequate phase composition diagrams. ACF-projection Chemography C F A Ca-Fe-Mg garnet plagioclase lawsonite zoisite prehnite am phibole chlorite talc diopside dolomite C px biotite Opx Oam grossular staurolite chloritoid cordierite aluminosilicates pyrophyllite margarite pumpellyite wollastonite calcite + quartz + H2O + CO2 + muscovite Fig. 2.13 Coordinates of phase compositions in the (K2O)–CaO–FeO–MgO–Al2O3–SiO2– H2O–CO2 system projected through (muscovite), quartz, CO2 and H2O and parallel to the MgFe1 exchange vector onto the plane Al2O3, FeO and MgO (ACF projection). Upper coordi- nates of some ACF phase compositions; lower phase relationships among some typical ACF phases at a specified pressure and temperature 2.5 Graphical Representation of Metamorphic Mineral Assemblages 53 1. Group the collected rocks in populations with similar bulk compositions, e.g. “normal” metapelites, metabasalts, etc. 2. Identify minerals which are present in the majority of a given group of rocks (e.g. muscovite and quartz in metapelites). Special assemblages require a special treatment, e.g. in quartz-absent corundum-bearing metapelites one may project through corundum onto an QFM plane. 3. If the excess minerals are not composed of simple oxide components, rewrite the composition matrix in terms of the compositions of the excess phases and the desired compositions at the corners of the mole fraction triangle selected as the new projection plane. 4. Delete rows in the composition matrix containing the excess phases and renor- malize the column vectors. Draw the diagram and keep in mind the proper distribution of one- two- and three-phase fields. Do not forget to write the compositions of the projection phases on the diagram (without this information your figure is meaningless). In Part II we will make extensive use of various kinds of composition phase diagrams. + quartz + CO2 + H2O kyanite garnet plagioclase zoisite calcite dolomite amphibole chlorite talc C F A Ky+Grt+Pl Pl+Am +Grt Grt+Am +Chl Am+Chl+Tlc Cal+Am+Zo Cal+Dol+Am Zo+Pl+Am 3 Fig. 2.14 ACF projection of phase relationships of typical ACF phases at a specified pressure and temperature (same as lower diagram in Fig. 2.13). MORB composition projected at black square, platform carbonate at black diamond (rock compositions listed in Table 2.3) 54 2 Metamorphic Rocks Chapter 3 Metamorphic Processes Rock metamorphism is always associated with processes and changes. Metamor- phism reworks rocks in the Earth’s crust and mantle. Typical effects of rock metamorphism include: l Minerals andmineral assemblages originally not present in a rockmay form, the new mineral assemblages grow at the expense of old ones. Consequently older minerals may disappear (e.g. metapelitic gneiss may originally contain Sil + Grt + Bt. A metamorphic event transforms this rock into one that contains Crd + Grt + Bt in addition to the Qtz and Fsp that also have been previously present in the rock; the old rock contained Sil, the new one Crd). l The relative abundance of minerals in a rock may systematically change and the new rock may have a different modal composition (metamorphism may increase the amount of Crd present in the rock and decrease the volume proportion of Grt + Bt). l Metamorphic minerals may systematically change their composition (e.g. the XFe of Grt and Crd may simultaneously increase during metamorphism). l The structure of rocks in crust and mantle may be modified (e.g. randomly oriented biotite flakes may be parallel aligned after the process). l The composition of the bulk rock may be altered during metamorphism by adding or removing components to, or from the rock from a source/sink outside the volume of the rock considered (e.g. removing K2O, MgO and FeO dissolved in a coexisting aqueous solution from a Grt + Crd + Bt rock may result in the formation of sillimanite). Typical changes in the modal composition of rocks and in the chemical compo- sition of minerals that constitute the rocks are caused by heterogeneous chemical reactions progressing in the rocks. The principles of metamorphism are, therefore, strongly related to the principles of chemical reactions. Mineral- and rock-forming metamorphic processes are mainly controlled by the same parameters that control chemical reactions. Metamorphic petrology studies reaction and transport processes in rocks. Metamorphic processes are caused by transient chemical, thermal and mechanical disequilibrium in confined volumes of the Earth’s crust and mantle. These disequilibrium states ultimately result from large-scale geological processes and the dynamics of the Earth’s planetary system as a whole. Metamorphic processes K. Bucher and R. Grapes, Petrogenesis of Metamorphic Rocks, DOI 10.1007/978-3-540-74169-5_3, # Springer-Verlag Berlin Heidelberg 2011 57 always result from disequilibrium and gradients in parameters that control reaction and transport in rocks; they cease when the rocks reach an equilibrium state. Chemical reaction is always inherent in the term metamorphism. The term metamor- phosis actually means transformation, modification, alteration, conversion and thus is clearly a process-related expression. Metamorphism is a very complex occurrence that involves a large number of chemical and physical processes at various scales. Metamorphic processes can be viewed as a combination of (1) chemical reactions between minerals and between minerals and gasses, liquids and fluids (mainly H2O) and (2) transport and exchange of substances and heat between domains where such reactions take place. The presence of an aqueous fluid phase in rocks undergoing metamorphism is critical to the rates of both chemical reactions and chemical transport. Consequently, an advanced understanding of metamorphism requires a great deal of insight into the quantitative description of chemical reactions and chemical transport processes, especially reversible and irreversible chemical thermodynamics. The term metamorphism as it is related to processes, changes and reactions clearly also includes the aspect of time. Metamorphism occurs episodically and is particularly related to mountain-building or orogenic episodes at convergent plate margins (collision zones) and during subsequent uplift and extension of continental crust, but also during sea-floor spreading and continental rifting. 3.1 Principles of Metamorphic Reactions In the following we briefly explain some basic aspects of metamorphic reactions and introduce some elementary principles that are essential for a basic understand- ing of metamorphism. The treatment is not sufficient for a thorough understanding of metamorphic processes and chemical reactions in rocks. It is therefore recom- mended for the reader who needs to know more, to study textbooks on chemical thermodynamics (e.g. Guggenheim 1986; Lewis and Randall 1961; Moore 1972; Prigogine 1955; Denbigh 1971) or textbooks that deal particularly with the appli- cation of thermodynamics to mineralogy and petrology (Ferry 1982; Fraser 1977; Greenwood 1977; Lasaga and Kirkpatrick 1981; Powell 1978; Saxena and Ganguly 1987; Wood and Fraser 1976). We recommend Chatterjee (1991), Fletcher (1993), Nordstrom and Munoz (1994), Anderson and Crerar (1993), and Ganguly (2008). First, let us consider, for example, a rock that contains the minerals albite and quartz. The Ab–Qtz rock is located at a certain depth in the crust (e.g. at point Th ¼ 1.25 GPa, 550C in Fig. 3.1). At that given pressure and temperature the two minerals (phases) are associated with unique values of molar Gibbs free energy. The Gibbs free energy, usually abbreviated with the symbolG, is a thermodynamic potential with the dimension joules/mole (energy/mole) and it is a function of pressure and temperature. The free energy of minerals and their mixtures are negative quantities (because they refer to the free energy of formation from the elements or oxides rather than absolute energies, e.g. G of albite at 900 K and 58 3 Metamorphic Processes 0.1 MPa is 3,257.489 kJ/mol). The considered Ab + Qtz rock can be formed by mechanically mixing, for example, 1 mol albite and 1 mol quartz. All rocks represent, thermodynamically speaking, mechanical mixtures of phases and are therefore heterogeneous thermodynamic systems.1 The phases, in turn, can be viewed as chemically homogeneous sub-spaces of the considered system, e.g., a volume of rock. Minerals, aqueous fluids, gasses and melts are the thermodynamic phases in rocks. These phases are usually chemical mixtures of a number of phase components (most minerals are solid chemical solutions and show a wide range in composition). In our example, albite and quartz shall be pure NaAlSi3O8 and SiO2 respectively. The total free energy of the rock is the sum of the free energies of its parts, that is in our case, nAbGAb + nQtzGQtz (where ni ¼ number of moles of substance i). The rock is characterized by a unique value of G and, by taking1 mole of each substance, its total composition is NaAlSi4O10. However, the composition of such a mechanical mixture ( rock) can also be obtained by mixing (powders of) jadeite (NaAlSi2O6) and quartz in the appropriate proportions. It is clear also that this rock has a unique Gibbs free energy at the given pressure and temperature and that it corresponds to the sum of GJd + 2GQtz. The free energy of the Ab–Qtz rock may be designated GAQ and that of the Jd–Qtz rock GJQ. The Gibbs free energies of the two rocks at P and T can be calculated from (3.1) and (3.2) provided that the free energy values of the three minerals can be calculated for that P and T: jadeite + quartz albite + quartz Temperature (°C) P re ss ur e (G P a) equilibrium of reaction Jd + Qtz = Ab (ΔGr = 0) isobar equilibrium temperature of reaction 300 400 500 0.5 1.0 1.5 2.0 600200 Te B A ThTl Fig. 3.1 Pressure–temperature diagram showing equilibrium conditions of the reaction jadeite + quartz ¼ albite 1The thermodynamic description of heterogeneous systems has been developed and formulated mainly by W. Gibbs (1878, 1906). Gibbs scientific contributions were fundamental for the development of modern quantitative petrology. 3.1 Principles of Metamorphic Reactions 59 surface conditions relative to the hydrous Al-silicate, kaolinite. In general, at some arbitrary metamorphic P–T condition only one Al-silicate + quartz can be stable, the other two possible two-phase assemblagesmust bemetastable.Metastable persistence of metamorphic minerals and assemblages is common in metamorphic rocks. This is, of course, an extremely fortunate circumstance for metamorphic petrologists. Rocks that formed in the deep crust or mantle with characteristic high-pressure and high- temperature mineral assemblages may be collected at the Earth’s surface. If metasta- ble assemblages were not common in metamorphic rocks we would find only low- P–T rocks at the surface. Metastable assemblages may even survive over geological time scales (hundreds of millions of years). It is obvious from Fig. 3.1 that a rock containing Jd + Qtz is metastable under conditions at the surface of the Earth. However, rocks containing Jd + Qtz are found at surface outcrops and provide evidence that they were formed originally at very high pressures. Some criteria and methods to detect disequilibrium in rocks do exist, however. One may distinguish two main kinds of disequilibrium; structural or textural disequilibrium and chemical disequilibrium. Structural disequilibrium is indicated by distinct shapes and forms of crystals and spatial distribution and arrangement of groups of minerals in heterogeneous systems. Structural disequilibrium may be found in rocks that underwent chemical reactions or successive series of reactions that all ceased before equilibrium structures developed and overall chemical equi- librium was reached. For example, zoned minerals with relic, often resorbed cores, representing an earlier metamorphism overgrown and partly replaced by composi- tionally different rims that formed during later metamorphism. Many different kinds of compositional features of metamorphic rocks and minerals may indicate chemical disequilibrium. As an illustration we can use our Jd + Qtz example again, except that in this case the rock that contains omphacite (Omp) + quartz (sodium- rich clinopyroxene where jadeite is present as a phase component). Independent information suggests that this example rock formed at about 600C. The rock may also contain pure albite in domains, local patches or veins and its composition clearly indicates chemical disequilibrium. The minerals Omp + Ab + Qtz never coexisted in stable or metastable equilibrium because the pyroxene contains a calcic component (diopside) and the feldspar does not (the phase component anorthite is not present in Pl). However, at the inferred temperature of 600C, plagioclase in an omphacite-bearing rock should, in an equilibrium situation (stable or metastable), contain some anorthite component. The presence of albite in such a rock must be related to some metamorphic process that progressed under conditions other than the equilibration of the Omp + Qtz assemblage. It is, on the other hand, important to note that the assemblage Omp + Qtz + Ab represents an overall disequilibrium but the assemblages Omp + Qtz and Ab + Qtz may well be equilibrium assemblages that equilibrated at different times at different P–T conditions. The question of equilibrium is always related to the scale of equilibrium domains. Disequilibrium may exist between large rock bodies in the crust, layers of rocks of different composition in an outcrop, between local domains of a thin section or between two minerals in mutual grain contact. Any chemically zoned mineral also repre- sents disequilibrium. At some scale there is always disequilibrium at any time. 62 3 Metamorphic Processes Chemical properties of coexisting minerals in a rock are often used to help resolve the question of equilibrium. As an example, a series of garnet–biotite pairs have been analyzed from a homogeneous Grt–Bt schist and the data are schemati- cally shown in Fig. 3.2. The data pairs can be arranged into two groups. Group a represents matrix biotite in contact with the rim of garnet; group b represents analyses of biotite grains included in the core of the garnet. Overall chemical equilibrium (stable or metastable) between garnet and biotite requires that both minerals have a uniform composition in the rock. Crossing tie lines are inconsistent with overall equilibrium. However, it is evident that the Grt–Bt pairs from each group are not in conflict with the requirements of chemical equilibrium. They are likely to constitute two different local equilibrium systems. However, the Grt–Bt pair connected with a dashed line in Fig. 3.2 shows crossing tie-line relationships within group a pairs and clearly represents a disequilibrium pair. Note that the absence of crossing tie-line relationships does not necessarily prove that the rock was in a state of overall equilibrium. The lack of obvious disequilibrium phenom- ena can be taken as evidence for, but not proof of, equilibrium. Metastable persistence of minerals and mineral assemblages and also the obvi- ous disequilibrium features in many rocks reflects the controlling factors and circumstances of reaction kinetics. The rate of a mineral reaction may be slower than the rate of, e.g., cooling of a volume of rock. Lacking activation energy for reaction in cooling rocks and nucleation problems of more stable minerals typically affects reaction kinetics. The kinetics of reactions in rocks is extremely sensitive to the presence or absence of H2O. Aqueous fluids serve as both a solvent and a reaction medium for mineral reactions. For example, if kyanite-bearing rocks are brought to pressure and temperature conditions where andalusite is more stable than kyanite, kyanite may be replaced by andalusite in rocks containing a free aqueous fluid in the pore space or along grain boundaries, whilst andalusite may fail to form in fluid-absent or dry rocks. It is commonly reasonable to assume that during progrademetamorphism rocks pass through successive sequences of equilibrium mineral assemblages. These sequences can be viewed as a series of stages, each of them characterized by an equilibrium assemblage and the different stages are connected by mineral reactions. Fe Mggarnet biotiteFe Mg group a group b b a Fig. 3.2 Schematic diagram showing the Fe–Mg distribution between coexisting garnet and biotite pairs in a single rock 3.1 Principles of Metamorphic Reactions 63 This assumption is founded on convincing evidence that prograde metamorphism takes place under episodic or continuous water-present conditions. One would therefore expect to find disequilibrium and metastable assemblages particularly in rocks that were metamorphosed under fluid-absent or fluid-deficient conditions. Aqueous fluids are typically not present in cooling rocks after they have reached maximum metamorphic pressure and temperature conditions. Textural and chemi- cal disequilibrium is also widespread in very high-grade rocks of the granulite facies that have lost their hydrous minerals and aqueous fluids during earlier stages of prograde metamorphism. Microstructures such as reaction rims, symplectites, partial replacement, corrosion and dissolution of earlier minerals are characteristic features of granulite facies rocks. They indicate that, despite relatively high tem- peratures (700–900C), equilibrium domains were small and chemical communi- cation and transport were hampered as a result of dry or H2O-poor conditions. To further illustrate some aspects of disequilibrium, we may consider large bodies of incompatible rock types in the crust. In many orogenic fold belts (e.g. Alps, Caledonides, Tianshan) mantle-derived ultramafic rock fragments were emplaced in the continental crust during the collision phase and stacking of nappes. Mantle fragments (harzburgites, lherzolites, serpentinites) of various dimensions can be found as lenses in granitic crust (Fig. 3.3). Forsterite (in the mantle fragment) + quartz (in the crustal rocks) is metastable relative to enstatite at any geologically accessible pressure and temperature conditions. The presence of Fo-bearing mantle rocks in Qtz-rich crustal rocks represents a large-scale disequilibrium feature. Chemical mass transfer across the contact of the incompatible rock types results in the formation of shells of reaction zones that encapsulate the mantle fragment. The nature of the minerals found in the reaction zones depends on the P–T conditions at the reaction site. In our example, the reaction shells may consist of granite harzburgite Fo + En Qtz + Kfs + Pl ± Bt 100 m Si Mg talc-zone biotite-zone fault Fig. 3.3 Large-scale disequilibrium between a tectonic lens of harzburgite (ultramafic rock fragment from the upper mantle) in granitic (gneissic) crust. The harzburgite lens is enveloped by protective shells of talc and biotite. These shells were produced by chemical reaction between incompatible rocks (rocks of radically different compositions) and their formation required transfer of, e.g. Si from granite to harzburgite and Mg from harzburgite to granite 64 3 Metamorphic Processes Because J/s2 is equivalent to W (watt), it follows that 1 HFU is equal to 0.042 W/m2. (or more convenient 42 mW/m2). The heat flow from the interior of the Earth is on the order of 30 mW/m2. However, heat flowmeasurements at the surface of the Earth vary between about 30 and 120 mW/m2. The total heat flow at the surface is composed of a number of contributions: (1) heat flow from the interior resulting from conductive heat trans- port as described by Fourier’s law, (2) transport of heat by convective mass flow in the mantle (Fig. 3.5), (3) transport of heat generated by decay of radioactive elements. The continental crust consists mostly of granitic rocks that produce about 30 mJ heat per kilogram and year. Oceanic crust, that is generally composed of basaltic rocks, produces about 5 mJ heat per kilogram and year, whereas mantle rocks produce only a small amount of radioactive heat (ca. 0.1 mJ heat per kilogram and year). The extra heat produced in the crust by decay of radioactive elements thus contributes significantly to the observed heat flow at the surface. Heat flow through a specified volume of the crust may occur under the following conditions: (a) Heat flow into the crustal volume is equal to the heat flow out from that volume. In this case the temperature in the volume remains constant. The temperature profile along thermal gradient z in Fig. 3.4 is independent of time (steady-state geotherm). convective mantle mid-ocean ridge cooling oceanic crust continental crust high heat flow very high heat flow young oceanic crust old oceanic crust subduction low heat flow very low heat flow trench lithosphere ascending plume descending plume core Fig. 3.5 Modification of steady-state heat flow resulting from conductive heat transfer by active tectonic processes. Heat flow at the surface varies by a factor of 4 3.2 Pressure and Temperature Changes in Crust and Mantle 67 (b) Heat flow into the crustal volume is greater than the heat flow out from that volume. The excess heat put into the crustal volume will be used in two different ways: (1) to increase the temperature of the rock volume, (2) to drive endothermic chemical reactions in rocks. (c) Heat flow into the crustal volume is lower than the heat flow out of the volume. In this case the heat loss of the volume of rock results in a temperature decrease (the rock cools). In this situation exothermic chemical reactions may produce extra heat to some extent. These reactions have the effect of preventing the rocks from cooling. Different heat flow at the surface also has the consequence that rocks at the same depth in the crust and upper mantle may be at different temperatures leading to lateral heat transport parallel to the earth surface (parallel to xy-surfaces). This is shown in Fig. 3.6. Along a profile from Denmark to southern Norway the observed surface heat flows have been used by Balling (1985) to model the temperature field in the crust and mantle shown in Fig. 3.6. Because flow vectors are always normal to the force field from which the flow results, the heat flow trajectories will roughly look like the flow vectors shown in Fig. 3.6. It is obvious from this two dimensional section that at a given point in the crust the heat flow has a vertical and a horizontal component. Also note that the MOHO under the mid-Paleozoic continental crust of central Europe (Denmark) is at about 700C whereas the MOHO under the Precambrian crust of the Baltic Shield in the north is only at about 350C. It follows that the base of continental crust may be at largely different temperatures surface mantle crust 100 200 300 400 500 600 700 800 DenmarkNorway heat flow trajectories 35 km isotherms Jq z-component x-component T em pe ra tu re ( C ) MOHO Fig. 3.6 Modelled temperature field along a cross section from Denmark to Norway (After Balling 1985) 68 3 Metamorphic Processes depending on the state and evolutionary history of a lithosphere segment and on the thermal regime deeper in the mantle. Observed surface heat flows can be translated into geotherms that may be time-dependent geotherms or steady-state geotherms. Such model geotherms are shown in Fig. 3.7 together with typical associated surface heat flows and temperatures at an average MOHO depth beneath continents of about 35 km (1.0 GPa). The geotherms represent curves that relate temperature with depth. The geothermal gradient (expressed as dT/dz) is the slope of the geotherm at a given depth in the crust or mantle and characterizes the temperature increase per or depth increment. In this book, you will see many P–T-diagrams with T on the x-axis and P on the y-axis. Pressure is increases proportional with depth (for details see Sect. 3.2.4). Thus steep geotherms (or paths of metamorphism) on these figures correspond to low geothermal gradients, flat geotherms represent high geothermal gradients. Geotherms have high gradients near the surface and the gradients become progressively lower with increasing depth. 3.2.2.1 Transient Geotherms Temperature changes in the crust and mantle are caused by adding or removing extra heat to or from the rocks. Heat flow changes may have a number of geological causes. Deep-seated causes include changes in the relative positions of lithosphere plates and their continental rafts relative to mantle convection systems (Fig. 3.5). Collision of lithosphere plates may lead to subduction of oceanic crust and cause Temperature (˚C) P re ss u re ( G P a) 0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 D ep th ( km ) 200 400 600 800 1000 1200 40 10 20 30 0 50 60 86 746140 surface heat flow (mW/m2) MOHO temperatures geotherms 400 700 950 1000 Fig. 3.7 Model geotherms with associated surface heat flow values and MOHO temperatures 3.2 Pressure and Temperature Changes in Crust and Mantle 69 remains constant until one of the two reactant minerals (in this case Jd) is used up. After completion of the reaction, the temperature of the rock will continue to increase. The effect is analogous to heating water to its boiling point. The tempera- ture of the water increases to the boiling temperature as heat is added to the water. During boiling the water temperature remains constant and the heat is used to drive the reaction (phase transition) liquid water ) steam. After all the water is boiled off, further addition of heat will increase the temperature of the steam. If a crustal volume receives less heat than it requires to maintain a steady state geotherm, it will cool and the rocks will be potentially affected by retrograde metamorphism. During cooling, it may become necessary for the rock to change its mineralogical composition in order to maintain a state of minimum Gibbs free energy. In our example, for a rock consisting of Ab + Qtz, the reaction Ab)Qtz + Jd converts all albite into jadeite and quartz as the rock cools to Te (Fig. 3.9) under equilibrium conditions. The heat released by the reaction buffers the temperature to a constant value (Te) until all Ab has been used up and further cooling can take place. The situation is analogous to removing heat from water. The temperature will decrease until water begins to freeze. As long as both reactant (liquid water) and product (ice) are present in the system, the temperature is confined to the freezing temperature. After conversion of all water into ice the temperature will continue to fall if further heat is removed from the ice. 3.2.4 Pressure Changes in Rocks Normally, pressure in the crust and mantle is characterized by isobaric surfaces that are approximately parallel to the Earth’s surface. As outlined above, the Gibbs free energy of minerals and associations of minerals (rocks) is also a function of pressure. Pressure changes in rocks are related to the change of position along the vertical space coordinate (z-axis in Fig. 3.4). The prevailing pressure at a given depth in a crustal profile with a steady-state geotherm is given by the density of the material above the volume of interest. It can be calculated from (dP/dz ¼ gr): P zð Þ ¼ g ðz 0 rðzÞdzþ P z¼0ð Þ (3.10) where g is the acceleration due to gravity (9.81 m/s2), r is the density of the rock at any z (e.g. 2.7 g/cm3¼ 2,700 kg m3) and P(z ¼ 0) is the pressure at the surface (e.g. 105 Pa¼ 1 bar¼ 105 N/m2 ¼ 105 kg m1 s2). The z-axis direction is always taken as negative. With a constant, depth independent density, (3.10) reduces to the simple relationship P ¼ g r z. The pressure, for example at 10 km (10,000 m) depth, is then calculated for an average rock density of 2,700 kg/m3 as 265 MPa or 2.65 kbar. The pressure at the base of continental crust of normal thickness (35–40 km) 72 3 Metamorphic Processes is about 1.0 GPa. The pressure at z results from the weight of the rock column above the volume of interest and is designated as the lithostatic pressure. This pressure is usually nearly isotropic. Non-isotropic pressure (stress) may occur as a result of a number of geologically feasible processes and situations as discussed below. A pressure difference of about 10 MPa at a depth of, e.g. 10 km (P ~ 270 MPa) will occur if a column of nearby rocks has a different average density (e.g. 2,800 instead of 2,700 kg/m3). If the density contrast of two columns of rock is 300 kg/m3 a DP of ~9 MPa at a depth of 30 km is possible. This is nearly a 100 MPa at a total pressure of ~0.8 GPa. However, because crustal material has similar densities, possible pressure differences between columns of rock at the same depth are limited to relatively small fractions of the lithostatic pressure (typically <10%). However, such pressure gradients may cause flow of fluids stored in the interconnected pore space of rocks according to Darcy’s law. The transport of fluids may cause chemical reactions in rocks through which the fluids pass if they are not in equilibrium with the solid assemblage of the rock. Non-isotropic pressure (stress) also results from tectonic forces and from volume changes associated with temperature changes and composition changes of rocks and minerals. Stress is one of the major controlling factors of the structure of metamorphic rocks. It is also a very important force for small-scale migration and redistribution of chemical components in metamorphic rocks. The processes of pressure solution, formation of fabric (such as foliation) and formation of meta- morphic banding from homogeneous protolith rocks are caused by non-isotropic pressure distribution in rocks and associated transport phenomena (Fig. 2.2b). The redistribution of material in stressed rocks attempts to reduce the non-isotropic pressure and to return to isotropic lithostatic pressure conditions. There are two basic regimes how rocks respond by deformation to applied stresses. At high temperatures and lithostatic pressures rocks react by ductile deformation to stress, whereas brittle deformation is dominant at low-T and shallow depth. The two deformation regimes grade into one another at the brittle–ductile transition zone, which is located at about 12 km depth corresponding to about 318 MPa (3.2 kbar) pressure in typical anorogenic continental crust. The temperature at this depth is about 300C. The P–T conditions at the brittle–ductile transition zone correspond to those of the lower greenschist facies (see Chap. 4). The precise depth of the brittle–ductile transition depends on the amount of applied stress, the mechanical properties of the rocks (dependent on the type of rocks present) and the presence or absence of an aqueous fluid. In the brittle deformation regime rocks break and fracture in response to applied stress. The created fractures fill normally with an aqueous fluid (liquid water) because the P–T conditions are in the field of liquid water (below the critical point of water, see Sect. 3.3). The fractures in crustal rocks above the brittle–ductile transition are typically well connected and form a permeable fracture network. The water in the fracture system is under hydrostatic pressure rather than lithostatic pressure conditions as described above. Thus hydro- static pressure operating on water in a extension fracture at 10 km depth is about 100 MPa rather than the lithostatic 270 MPa as on the rock matrix on both sides of 3.2 Pressure and Temperature Changes in Crust and Mantle 73 the fracture. Hydrothermal metamorphism, precipitation of fissure minerals, low-T rock alteration and similar processes occur typically under hydrostatic pressure conditions. Within the ductile deformation regime fluids tend to be close to litho- static pressure or if present in isolated pores even at lithostatic pressure. In general, the lithostatic pressure acting on a volume of rock changes as a result of a change of depth of that rock volume (change of the position along the z-direction). In almost all geological occurrences, depth changes are caused by tectonic processes (that ultimately are also the result of thermal processes, e.g. mantle convection, density changes resulting from temperature changes). A litho- sphere plate can be subducted in a collision zone and be transported to great depth in the mantle before it is resorbed by the mantle. Some of the subducted material may return to the surface before a normal steady-state geotherm is established. Metamorphic rocks of sedimentary or volcanic origin with mineral assemblages that formed at pressures in excess of 3.0 GPa have been reported from various orogenic belts; 3.0 GPa pressure is equivalent to about 100 km subduction depth (z¼P/(gr)). In continental collision zones the continental crust often is thickened to double its normal pre-collision thickness. At the base of the thickened crust the pressure increases from 0.9 to 1.8 GPa The changes in mineralogical composition of the rocks undergoing such dramatic pressure changes depend on the nature and composition of these rocks and will be discussed in Part II. Variations in depth and associated pressure changes also can be related to subsiding sedimentary basins where a given layer a of sediment is successively overlain and buried by new layers b, c . . . of sediments. Layer a experiences a progressive increase in pressure and temperature. If sedimentation and subsidence is slow, layer amay follow a steady state geotherm. Metamorphism experienced by layer a is commonly described under the collective term burial metamorphism as defined in Sect. 1.2.3. Crustal extension or lithosphere thinning is often the cause for the required subsidence creating deep sedimentary basins. Burial metamor- phism is therefore also caused by tectonic processes. 3.3 Gases and Fluids Sedimentary rocks such as shales often contain large modal proportions of hydrous minerals. In fact, sediments deposited in marine environments can be expected to contain, under equilibrium conditions, a mineral association that corresponds to a maximum hydrated state. Adding heat to the hydrous minerals (clays) of a sediment during a metamorphic event will drive reactions of the general form: Hydrous assemblage ) less hydrous or anhydrous assemblageþ H2O (3.11) The general reaction (3.11) describes the dehydration process taking place during prograde metamorphism. The important feature of dehydration reactions is the release of H2O. Steam is, in contrast to solid minerals, a very compressible 74 3 Metamorphic Processes constrained by maximum heat flow differences observed at the surface. The highest heat flows are measured along mid-ocean ridges (120 mW/m2), whilst the lowest are in old continental cratons (30 mW/m2). Consider a crustal layer with a heat flow difference between bottom (75 mW/m2 ) and top (35 mW/m2) of 40 mW m2. This means that every second the rock column receives 0.04 J heat (per m2). If the layer consists of shales (heat capacity¼ 1 kJ kg1 C1) with a volatile content (H2O and CO2) of about 2 mol/kg, the heat received will be used to increase the temperature of the rock and to drive endothermic devolatilization reactions. About 90 kJ heat are required to release 1 mol H2O or CO2. If complete devolatilization occurs in the temperature interval 400–600C, a total of 380 kJ will be consumed by the shale (180 kJ for the reactions and 200 kJ for the temperature increase from 400 to 600C). It takes 9.5  106 s (0.3 years) to supply 1 kg of rock with the necessary energy. Using a density of 2.63 g/cm3, 1 kg of rock occupies a volume of 380 cm3 and it represents a column of 0.38 mm height and 1 m2 ground surface. From this it follows that metamorphism requires about 8 years to advance by 1 cm. Eight million years (Ma) are required to metamorphose a 10 km thick layer of shale. Similarly, if the heat flow difference is only 20 mW/m2, and the shale layer is 20 km thick layer with an initial temperature 200C, it will require about 48 Ma of metamorphism to convert the hydrous 200C shale to an anhydrous 600C meta- pelitic gneiss. This crude calculation shows that typical time spans for regional scale metamor- phic processes are on the order of 10–50 million years. Similar time scales have been derived from radiometric age determinations. 3.5 Pressure–Temperature–Time Paths and Reaction History During tectonic transport, any given volume of rock follows its individual and unique path in space and time. Each volume of rock may experience loss or gain of heat, and changes of its position along the z-coordinate result in changes in lithostatic pressure loaded on the rock. Figure 3.11 shows again a simple model of a destructive plate margin. The situation here depicts a continent–continent collision with the formation of continental crust twice its normal thickness. In A at t1 (0 Ma) there is a volume of rock (a rock unit indicated by an open square) at depth c, the position of which changes depth with time (t1 through to t6 at 30 Ma). In B through E, the position of the rock unit (filled squares) during tectonic transport is shown in terms of P–T space. The time slices are arbitrary and have been chosen in accordance with time scales of the formation of Alpine-type orogenic belts. At t1 the rock unit lies on a stable steady-state geotherm. At t2 (10 Ma) tectonic transport moves the crust together with the rock unit beneath another continental crust of normal 35 km thickness. Increasing depth of the rock unit is accompanied by increasing pressure (3.10). At the same time, the rock unit begins to receive more heat than at its former position at t1. However, because heat transport is a slow process compared with tectonic transport, dP/dT tends to be much steeper 3.5 Pressure–Temperature–Time Paths and Reaction History 77 (Fig. 3.11c) than the corresponding dP/dT slope of the initial steady-state geotherm (Fig. 3.11b). Between t1 and t2, the rock unit has traveled in P–T space along a path that is on the high-pressure side of the initial steady-state geotherm. The rock unit is now on a transient geotherm that changes its shape as time progresses. At t3 the crust is twice its normal thickness (about 70 km), which is about the maximum thickness in continent–continent collision zones. The rocks have reached their maximum depth and consequently their maximum pressure of about 2 GPa (Fig. 3.11c). Continued plate motion does not increase the thickness of the crust and pressure remains constant as long as underthrusting is going on. On the other hand, heat transfer to the rock unit in question increases the temperature as shown in Fig. 3.11d (t4). From this time, a number of feasible mechanisms may control the path of the rock unit. Continued tectonic transport may return slices and fragments of rock to shallower levels in a material counter current (dashed arrow in Fig. 3.11a), or simply, after some period of time, plate convergence stops, e.g. because frictional forces balance the force moving the plate. The thickened crust starts to uplift, and erosion restores the crust to its original thickness. By this mechanism the rock unit may return to its original depth position and, given enough time, the stable steady-state geotherm will be re-established. The path between t4 and t5 is characterized by decompression (transport along the z-axis). If initial uplift rates are slow compared with heat transport rates, the rocks will experience dcb p T t1 t2 t3 t4 steady state geotherm maximum pressure t5 t6 maximum temperature transient geotherms e t1 a t2 t5 t3 t4 t6 a b c time = 30 ma time = 10 ma time = 0 ma time = 20 ma Fig. 3.11 (a–e) Schematic diagram showing the position of a rock unit in the crust as a function of time during continent–continent collision and the corresponding paths followed by the rock unit in P–T space 78 3 Metamorphic Processes a continued temperature increase during uplift as shown in Fig. 3.11e). However, at some stage along the path the rocks must start to lose more heat to the surface than they receive from below, and consequently cooling begins. The point t5 in Fig. 3.11e represents the maximum temperature position of the path traveled by the rock unit. At t6 the rock has returned to its former position on the steady state geotherm (Fig. 3.11a). The consequences for pressure and temperature of the geologic process illu- strated in Fig. 3.11 are summarized in Fig. 3.12. A rock at depth c (P–T–t-path 3) follows a clockwise pressure–temperature loop. Such clockwise P–T–t paths are a characteristic feature of orogenic metamorphism and have been documented from such diverse mountain belts as the Scandinavian Caledonides, western Alps, Appa- lachians, and Himalayas. In detail, clockwise P–T–t loops may show a number of additional complications and local features. Counter-clockwise P–T–t paths have been reported from granulite facies terrains where an event of heating from igneous intrusions precedes crustal thickening. They may also occur in terrains that have experienced an initial phase of crustal extension. Very often “normal” orogenesis, is characterized by the following sequence of P–T–t path sections: isothermal thickening, isobaric heating, isothermal decompression and isobaric cooling. Returning to Fig. 3.12a, it is evident that the maximum temperature point along the P–T–t path followed by a metamorphic rock does not necessarily coincide with the maximum pressure point of the path. This means that maximum pressure and maximum temperature will be generally diachronous. p T dehydration reactions least hydrated state c a M N 1 2 3 steady state geotherm PTt path 3 c b a b piezo- thermal array steady state geotherm PTt path1 PTt path2 PTt path 3 least hydrated state maximum temperature maximum pressure Fig. 3.12 (a, b) (a) P–T–t path of a rock unit at depth c in 0. The mineral assemblage stable at the tangent point of the path with dehydration reaction 2 will, in general, be preserved in metamorphic rocks. The tangent point corresponds to the least hydrated state of the rock. (b) Clockwise P–T–t loops for crustal rock units from different depth levels (0). The mineral assemblage of the individual rock unitscorresponds to the least hydrated state. The P–T points of the least hydrated state from all rocks of a metamorphic terrain define a curve that is known as a piezo-thermal array (~metamorphic field gradient) 3.5 Pressure–Temperature–Time Paths and Reaction History 79 3.6.1 Reactions Among Solid-Phase Components These are often termed “solid–solid” reactions because only phase components of solid phases occur in the reaction equation. Typical “solid–solid” reactions are, for example: 3.6.1.1 Phase Transitions, Polymorphic Reactions Al2SiO5 Ky ¼ And; Ky ¼ Sil; Sil ¼ And CaCO3 Calcite ¼ aragonite C Graphite ¼ diamond SiO2 a-Qtz ¼ b-Qtz; a-Qtz ¼ coesite, ...... KAlSi3O8 Microcline ¼ sanidine 3.6.1.2 Net-Transfer Reactions Such reactions transfer the components of reactant minerals to minerals of the product assemblage. Reactions involving anhydrous phase components only Jd + Qtz = Ab Grs + Qtz = An + 2 Wo 2 Alm + 4 Sil + 5 Qtz = 3 Fe - Crd Volatile-conserving “solid-solid” reactions Tlc + 4 En = Ath Lws + 2 Qtz = Wa 2 Phl + 8 Sil + 7 Qtz = 2 Ms + Crd 3.6.1.3 Exchange Reactions These reactions exchange components between a set of minerals. Reactions involving anhydrous phase components only Fe–Mg exchange between olivine and orthopyroxene: Fo + Fs ¼ Fa + En Fe-Mg exchange between clinopyroxene and garnet: Di + Alm ¼ Hed + Prp Volatile-conserving “solid–solid” reactions Fe–Mg exchange between garnet and biotite: Prp + Ann ¼ Alm + Phl Cl–OH exchange between amphibole and biotite: Cl–Fpa + OH–Ann ¼ OH–Fpa + Cl–Ann 82 3 Metamorphic Processes 3.6.1.4 Exsolution Reactions/Solvus Reactions High-T alkali feldspar ¼ K-feldspar + Na-feldspar Ternary high-T feldspar ¼ meso-perthite + plagioclase Mg-rich calcite ¼ calcite + dolomite High-T Cpx ¼ diopside + enstatite Al-rich Opx ¼ enstatite + garnet The common feature of all “solid–solid” reactions is that equilibrium conditions of the reaction are independent of the fluid phase composition, or more generally speaking, of the chemical potentials of volatile phase components during metamorphism. It is for this reason that all “solid–solid” reactions are potentially useful geologic thermo- meters and barometers. The absence of volatile components in the reaction equation of “solid–solid” reactions should not be confused with general fluid-absent conditions during reaction progress. Metamorphic reactions generally require the presence of an aqueous fluid phase in order to achieve significant reaction progress even in geologi- cal time spans, attainment of chemical equilibrium in larger scale domains and chemical communication over several grain-size dimensions. Although P–T coordi- nates of the simple reconstructive phase transition kyanite¼ sillimanite are indepen- dent of mH2O, the detailed reaction mechanismmay involve dissolution of kyanite in a saline aqueous fluid and precipitation of sillimanite from the fluid at nucleation sites that can be structurally unrelated to the former kyanite (e.g. Carmichael 1968). It is clear that the chlorine-hydroxyl exchange between amphibole and mica represents a process that requires the presence of an saline aqueous fluid. However, the equilibrium of the exchange reaction is independent on the composition of that fluid. Phase transitions, “solid–solid” net transfer and exchange reactions often show straight line equilibrium relationships in P–T space. The slope of the equilibrium line of such reactions can be readily calculated (estimated) from the Clausius– Clapeyron equation (Fig. 3.13). 3.6.2 Reactions Involving Volatiles as Reacting Species 3.6.2.1 Dehydration Reactions Reactions involving H2O are the most important metamorphic reactions. Low- grade metasediments contain modally abundant hydrous minerals. Pelitic and mafic rocks from the subgreenschist facies consist mainly of clay minerals (and zeolites). Serpentine minerals make up lower-grade ultramafic rocks. These sheet silicates contain up to about 12 wt% H2O. The hydrous minerals are successively removed from the rocks by continuous and discontinuous dehydration reactions as heat is added to them as shown in Fig. 3.14 for a shale and basaltic rock. Release of H2O during prograde metamorphism of H2O-rich protoliths ensures that a free H2O fluid is present in the rocks either permanently or periodically during the progress of dehydra- tion reactions. This in turn often permits discussion of metamorphism for conditions 3.6 Chemical Reactions in Metamorphic Rocks 83 where the lithostatic pressure acting on the solids also applies to a free fluid phase. If the fluid is pure H2O (and choosing an appropriate standard state), the condition is equivalent to aH2O¼ 1. The general shape of dehydration equilibria is shown schema- tically in Fig. 3.13. Using the talc breakdown reaction as an example, the general curve shape of dehydration reactions on P–T diagrams can be deduced as follows: Tlc ¼ 3Enþ 1Qtzþ 1H2O ðVÞ (3.13) The dP/dT-slope of a tangent to the equilibrium curve is given at any point along the curve by the Clapeyron equation. All dehydration reactions have a positive DSr (entropy change of reaction) and the curvature is largely controlled by the volume change of the reaction. The volume term can be separated into a contribution from the solids and the volume of H2O, respectively: DVr ¼ DVsolids þ DVH2O (3.14) The volume change of the solids can be calculated from: DVsolids ¼ 3VEn þ 1VQtz  1Vtlc (3.15) DVsolids is a small and negative quantity and varies very little with P and T. Consequently, the curve shape of dehydration reactions largely reflects the volume function of H2O, that is shown in Fig. 3.10. At low pressures, VH2O is large and the Temperature Tlc dehydration by compression dehydration by heating deydration by decompression dT dp = DVr DSr dT dp ~ const dT dp = solid-solid reaction 3 En + Qtz + H2O Clausius- Clapyeron- slope Jd + Q tz Ab dehydration reaction P re ss ur e 8 Fig. 3.13 Schematic equilibrium conditions of solid–solid and dehydration reactions in P–T space and the Clausius–Clapeyron equation (DSr ¼ entropy change and DVr¼ volume change of the reaction) 84 3 Metamorphic Processes
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