Contact Metamorphism 9781501509612, 9780939950317

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CONTACT METAMORPHISM DERRILL M . K

, Editor

The authors, by affiliation: Mark D. Barton Robert P. Ilchik David A. Johnson Mark A. Marikos John-Mark Staude Dept. of Geosciences University of Arizona Tucson, Arizona 85721 George W. Bergantz Dept. of Geological Sciences University of Washington, AJ-20 Seattle, Washington 98195 James R. Bowers Kevin P. Furlong Derrill M. Kerrick Stuart P. Raeburn Dept. of Geosciences The Pennsylvania State University University Park, Pennsylvania 16802 James M. Brenan Geophysical Laboratory 5251 Broad Branch Road, N.W. Washington, D.C. 20015 John M. Ferry Dept. Earth & Planetary Sciences The Johns Hopkins University Baltimore, Maryland 21218 T. Kenneth Fowler, Jr. Scott R. Paterson Dept. of Geological Sciences University of Southern California Los Angeles, California 90089 B. Ronald Frost Dept. of Geology & Geophysics University of Wyoming Laramie, Wyoming 82071 R. Brooks Hanson Science, 1333 H Street, N.W. Washington, D.C. 20005 and

The Series Editor:

R. Brooks Hanson Dept. of Mineral Sciences Smithsonian Institution Washington, D.C. 20560 Raymond L. Joesten Dept. of Geology & Geophysics University of Connecticut Storrs, Connecticut 06269 Theodore C. Labotka Dept. of Geological Sciences University of Tennessee Knoxville, Tennessee 37996 Antonio C. Lasaga Geological Laboratory Yale University New Haven, Connecticut 06520 Peter I. Nabelek Dept. of Geological Sciences University of Missouri Columbia, Missouri 65211 David R. M. Pattison Dept. of Geology & Geophysics University of Calgary Calgary, Alberta, Canada T2N 1N4 Eleanour A. Snow Dept. of Geology University of South Florida Tampa, Florida 33620 Robert J. Tracy Dept. of Geological Sciences Virginia Polytechnic Institute and State University Blacksburg, Virginia 24061 Ron H. Vernon School of Earth Sciences MacQuarie University Sydney, N.S.W. 2109, Australia

Paul H. Ribbe Department of Geological Sciences Virginia Polytechnic Institute and State University Blacksburg, Virginia 24061

Copyright 1991 MINERALOGICAL SOCIETY of AMERICA Printed by BookCrafters, Inc., Chelsea, Michigan 48118

REVIEWS

in

MINERALOGY

Formerly: SHORT COURSE NOTES

ISSN 0275-0279 V o l u m e 26: CONTACT

METAMORPHISM

ISBN 0-939950-31-6 ADDITIONAL COPIES of this volume and those listed below may be obtained from the MINERALOGICAL SOCIETY OF AMERICA

1130 Seventeenth Street, N.W., Suite 330, Washington, D.C. 20036 U.S.A. Vol.

Year

Pages

Editorfsl

Title

1

1974

284

P. H. Ribbe

SULFIDE MINERALOGY

2

1983

362

P. H. Ribbe

FELDSPAR MINERALOGY ( 2 n d e d i t i o n )

3

out of

print

OXIDE MINERALS

4

1977

232

F. A. Mumpton

MINERALOGY AND GEOLOGY OF NATURAL ZEOLITES

5

1982

450

P. H. Ribbe

ORTHOSIUCATES (2nd edition)

6

1979

R. G. Bums

MARINE MINERALS

7

1980

380 525

C. T. Prewitt

PYROXENES

8

1981

398

A. C. Lasaga R. J. Kirkpatrick

KINETICS OF GEOCHEMICAL PROCESSES

9A

1981

372

D. R. Veblen

AMPHIBOLES AND OTHER HYDROUS PYRIBOLES— MINERALOGY

9B

1982

390

D. R. Veblen P. H. Ribbe

AMPHIBOLES: PETROLOGY AND EXPERIMENTAL PHASE RELATIONS

10

1982

397

J. M. Ferry

CHARACTERIZATION OF METAMORPHISM THROUGH MINERAL EQUILIBRIA

11 12

1983

394

R. J. Reeder

CARBONATES: MINERALOGY AND CHEMISTRY

1983

644

E. Roedder

FLUID INCLUSIONS ( M o n o g r a p h )

13

1984

584

S. W. Bailey

MICAS

14

1985

428

S. W. Kieffer A. Navrotsky

MICROSCOPIC TO MACROSCOPIC : ATOMIC ENVIRONMENTS TO MINERAL THERMODYNAMICS

15

1990

406

M. B. Boisen, Jr. G. V. Gibbs

MATHEMATICAL CRYSTALLOGRAPHY ( R e v i s e d )

16

1986

570

J. W. Valley H. P. Taylor, Jr. J. R. CWeil

STABLE ISOTOPES IN HIGH TEMPERATURE GEOLOGICAL PROCESSES

17

1987

500

THERMODYNAMIC MODELLING OF GEOLOGICAL MATERIALS: MINERALS, FLUIDS, MELTS

18

1988

698

H. P. Eugster I. S. E. Carmichael F. C. Hawthorne

SPECTROSCOPIC METHODS IN MINERALOGY AND GEOLOGY

19

1988

698

S. W. Bailey

HYDROUS PHYLLOSILICATES (EXCLUSIVE OF MICAS)

20

1989

369

D. L. Bish J. E. Post

MODERN POWDER DIFFRACTION

21

1989

348

B. R. Lipin G. A. McKay

GEOCHEMISTRY AND MINERALOGY OF RARE EARTH ELEMENTS

22

1990

406

D. M. Kerrick

THE Al 2 Si0 5 POLYMORPHS (Monograph)

23

1990

603

M. F. Hochella, Jr. A. F. White

MINERAL-WATER INTERFACE GEOCHEMISTRY

24

1990

314

MODERN METHODS OF IGNEOUS PETROLOGY— UNDERSTANDING MAGMATIC PROCESSES

25

1991

509

J. Nicholls J. K. Russell D. H. Lindsley

OXIDE MINERALS: PETROLOGIC AND MAGNETIC SIGNIFICANCE

CONTACT

METAMORPHISM FOREWORD

The Mineralogical Society of America sponsored a short course on Contact Metamorphism, October 17-19, 1991, at the Pala Mesa Resort, Fallbrook, California, prior to its annual meeting with the Geological Society of America. Derrill Kerrick convened the course and edited this volume of Reviews in Mineralogy, the largest ever published in the 18-year history of this series. One might have guessed that petrologists would hold the record. As series editor I am definitely not encouraging a competition! Other volumes in MSA's series are described on the facing page. I thank Derrill Kerrick for a splendid job of editing chapters and harrassing authors to finish them on time—unfortunately he did not succeed on the latter account. Michael Alter and Madelyn Smith did most of the paste-up for the camera-ready copy; Debra Thomson helped with typing. At headquarters, Susan Myers was very helpful, as always. PaulH. Ribbe Blacksburg, VA September 11, 1991 PREFACE AND ACKNOWLEDGMENTS As reviewed in Chapter 1, contact aureoles have unique attributes for elucidating the processes and controls of metamorphism. Within the last two decades there has been considerable evolution in our knowledge of metamorphism. This evolution spans a wide range of scales from submicroscopic analysis of grain boundaries through to regional scale analysis of contact metamorphism associated with batholith terrains. Geological sciences is becoming increasingly multidisciplinary in nature. Traditionally, contact aureoles were primarily studied by metamorphic petrologists. Their mapping of isograds and mineral zones in aureoles, coupled with microscopic analysis of the prograde metamorphic evolution of textures, structures and mineralogy, has provided an excellent framework for our understanding of contact metamorphism. However, complete understanding of the processes and controls of contact metamorphism requires a multidisciplinary analysis from a wide range of geological subdisciplines. This volume provides a multidisciplinary review of our current knowledge of contact metamorphism. As in any field of endeavor, we are provided with new questions, thereby dictating future directions of study. Hopefully, this volume will provide inspiration and direction for future research on contact metamorphism. Following journal-style reviewer guidelines, several individuals served as "formal" reviewers for selected chapters. Accordingly, L.P. Baumgartner, A.R. Cruden, J.J. DeYoreo, J.M. Ferry, R.B. Hanson, T.D. Hoisch, M.J. Holdaway, S.J. Mackwell, and J.M. Rice, kindly provided prompt, thorough reviews. As acknowledged in individual chapters, several other reviewers interacted directly with the authors. I am very grateful to all of the reviewers for their help in improving the quality of this volume. L.M. Miller of The Pennsylvania State University provided invaluable, friendly, efficient help with this volume and with organization of the short course. Susan Myers of the Mineralogical Society of America has been responsible for the budgeting, organization and logistics of this and many previous M.S.A. short courses. The Mineralogical Society of America owes her a large debt of gratitude for the overall success of the M.S.A short course series. For his incredible efforts as editor of the Reviews in Mineralogy series, the Society should consider elevating Paul Ribbe to sainthood. Derrill M. Kerrick University Park, PA September 3,1991 iii

TABLE Page ii iii

OF

CONTENTS

Copyright; List of additional volumes of Reviews in Mineralogy Foreword; Preface and Acknowledgments Derrill M. Kerrick

Chapter 1

OVERVIEW OF CONTACT METAMORPHISM 1 1 2 4 7

8 8 9 9 10 10 12

INTRODUCTION GEOLOGIC SETTING INTRUSIVES COMPARISON WITH REGIONAL METAMORPHISM METAMORPHIC PROCESSES

Coarsening Neocrystallization Metasomatism Anatexis Deformation

FLUIDS REFERENCES

Chapter 2

George W. Bergantz

PHYSICAL AND CHEMICAL CHARACTERIZATION OF PLUTONS 13

INTRODUCTION

16

INTENSIVE VARIABLES

24

EXTENSIVE VARIABLES

30

DYNAMIC STATE: COMBINATION OF INTENSIVE AND EXTENSIVE VARIABLES

34 42

REFERENCES ACKNOWLEDGMENTS

13 16 18 20 21 24 26 30 31

Contact metamorphism as a conjugate system

Sequence of crystallization Volatiles Estimating temperature and pressure Physical properties

Size and shape of plutons Open-system behavior

Crystallization Conduction and convection models

Chapter 3

Theodore C. Labotka

CHEMICAL AND PHYSICAL PROPERTIES OF FLUIDS 43 48

49 52 57

62

THE EXISTENCE OF METAMORPHIC FLUIDS THERMOCHEMICAL PROPERTIES OF CONTACT METAMORPHIC FLUIDS

H20 CO2 CO2-H2O mixtures C0 2 -H 2 0-NaCl

iv

67

MINERAL-FLUID EQUILIBRIA

92

TRANSPORT PROPERTIES OF AQUEOUS FLUIDS

96 97

ACKNOWLEDGMENTS REFERENCES

67 69 70 73 79 81 86 88 92 92 94 95 96

Quartz-H20 Calcite-H20 Graphite-H20 Graphitic pelitic homfelses Graphitic limestones Mixed volatile equilibria in CO2-H2O systems Assemblages in NaCl-CC>2-H20 fluids The compositions of metamorphic fluids

Viscosity Thermal conductivity Self-diffusion coefficient Comparison of the transport properties of H2O Summary

D.R.M. Pattison & R.J. Tracy

Chapter 4

PHASE EQUILIBRIA AND THERMOBAROMETRY OF M E T A P E L I T E S 105 105

PROLOGUE INTRODUCTION

110

DESCRIPTIVE ASPECTS OF CONTACT METAMORPHOSED PELITES

115

CHEMOGRAPHIC ANALYSIS OF ASSEMBLAGES

120 120

CONTACT AUREOLES IN PELITES CONTACT METAMORPHIC FACIES SERIES

105 107 109 109 110 110 118

124 124 124 124 124 124 125 125 125 125 126 126 129 129 129 129 130 130 130 130 132 132

Historical background Distinction between contact and regional metamorphism of pelites The importance of the petrogenetic grid Approach used in this chapter Facies of contact metamorphism Textures of contact metamorphic pelites Metamorphism vs. metasomatism

Facies series 1 Facies series 1, Type la Diagnostic features The Comrie aureole, Scotland Other aureoles Model reaction sequence Mineral compositions Facies series 1, Type lb Diagnostic features The Ballachulish aureole, Scotland Other aureoles from sub-facies series Type lb Regional metamorphic examples Model reaction sequence Mineral compositions Facies series 1, Type lc Diagnostic features The McGerrigle aureole, Quebec The Tono aureole, Japan Other aureoles from sub-facies series Type lc Regional examples of facies series Type lc Model reaction sequence Mineral compositions v

132 133 133 133 134 135 135 135 138 138 138 140 141 141 141 142 142 143 143 145 146 147 147 147 147 148 148 148 148 148 148 149 150 151 151 154 154 155 156 156 156 156 156 157 157 159 160 160 164 165 165 169 170 170 173 173 173

Facies series 2 Facies series 2, Type 2a

The Cupsuptic aureole, Maine The Kiglapait aureole, Labrador Other aureoles of sub-fades series Type 2a Regional examples Model reaction sequence Mineral compositions

Facies series 2, Type 2b

Diagnostic features The Ardara aureole, Donegal The Ronda aureole, Spain Other aureoles from facies series Type 2b Further subdivision ofsub-facies series Type 2b Regional examples Model reaction sequence Mineral compositions

Facies series 3 Facies series Type 4 Summary of facies series types Northwestern Maine—an example of regional-scale contact metamorphism Notable mineral associations in contact metamorphosed pelites

Cordierite-garnet-muscovite Staurolite-cordierite-muscovite Kyanite Quartz-absent assemblages Corundum Spinel Hypersthene

LOW- AND HIGH-P ULTRAMETAMORPHISM: BUCHITES AND EMERIES Buchites Emeries

Cortlandt Complex emeries

PROPOSED PETROGENETIC GRID

Restrictions

Garnet stability—the effect of Ca and Mn Staurolite-cordierite-muscovite K-feldspar + chlorite K-feldspar + staurolite Hypersthene

Univariant and divariant reactions Orientation of the grid Reactions at the onset of anatexis Facies series and bathozones

ANATEXIS

Comparison with the bathozone scheme of Carmichael (1978) High-grade assemblages in the anatectic zone Stability of kyanite + biotite in andalusite-sillimanite sequences

Significance of anatexis to phase equilibria

The muscovite-melt bathograd

CALIBRATION OF THE PETROGENETIC GRID IN P - T SPACE

Accuracy of the calibrated grid C- and S-bearing fluid species and variations in aH20 Pressures of facies series types

THERMOBAROMETRY

Low to intermediate grade metapelites

Thermometry

vi

175 175 175 111 177 111 178 179 179

Barometry Multi-equilibrium calculations The zoned-garnet Gibbs' method High-grade metapelites Thermometry Barometry Multi-equilibrium calculations Pressure estimates using stratigraphie arguments Comparison of exchange thermobarometry with petrogenetic grid constraints

181 181

EPILOGUE SUGGESTIONS FOR FUTURE RESEARCH

182 182

ACKNOWLEDGMENTS BIBLIOGRAPHY AND REFERENCES

181 181 181 181 182 182

Use of modal mineralogy and textures to document reaction processes Testing techniques for P-T path determination Detailed investigation of anatexis in contact metamorphism Experimental refinement of key metapelitic equilibria Resetting of geothermobarometers Testing of multi-equilibrium techniques

Chapter 5

R.J. Tracy & B.R. Frost

PHASE EQUILIBRIA AND THERMOBAROMETRY OF CALCAREOUS, ULTRAMAFIC AND MAFIC ROCKS, AND IRON FORMATIONS 207

207 207 208 208

INTRODUCTION

Historical background Bowen's decarbonation series Pressure series of contact metamorphism Purpose of this chapter

208

CONTACT METAMORPHISM OF CALCAREOUS ROCKS

248

CONTACT METAMORPHISM OF ULTRAMAFIC ROCKS

208 209 212 216 218 223 224 225 226 226 229 229 231 233 237 239 240 241 246 246 249 249

Phase equilibria Fluid composition buffering in mixed-volatile equilibria CMS (H2O-CO2) equilibria CMAS(H20-CC>2) equilibria KCMAS(H20-C02) equilibria Other systems Kinetic control of contact phase equilibria Bulk compositional control ofphase equilibria in calcareous rocks Thermobarometry in contact metamorphosed calcareous rocks Well studied contact metamorphic calcareous rocks Adirondacks Alta aureole Ballachulish aureole Boulder and Marysville aureoles Christmas Mountains and Marble Canyon aureoles Crestmore Elkhorn aureole Notch Peak aureole Beinn an Dubhaich aureole, Skye Carbonate rock xenoliths in magmas

Phase equilibria The system CMS(H20)

vii

250 252 252 252 255 255 257

The system CFMS(H20) The system CMS(H20-C02) The system FMAS(H20) The system CFMAS(H20) Well studied contact metamorphic ultramafic rocks Ber geli aureole Paddy-Go-Easy Pass

251

CONTACT METAMORPHISM OF MARC ROCKS

269

CONTACT METAMORPHISM OF IRON FORMATIONS

211 279 280 280

SUMMARY DIRECTIONS FOR FUTURE RESEARCH ACKNOWLEDGMENTS BIBLIOGRAPHY AND REFERENCES

257 259 259 263 264 264 265 267 269 269 269 273 273 275 211

Difficulties in studying contact metamorphosed metabasites Phase equilibria Amphibole reactions Effects of oxygen fugacity Well studied contact metamorphic mafic rocks Skye and Skaergaard The Karmutsen Morton Pass

Phase equilibria The system FS(H20) The system CFMS(H20) The system FMS(H20-C02-02) Well-studied contact metamorphic iron formations Metamorphism of the Gunflint Formation Metamorphism of the Biwabik Formation

Chapter 6

James Brenan

DEVELOPMENT AND MAINTENANCE OF METAMORPHIC PERMEABILITY: IMPLICATIONS FOR FLUID TRANSPORT 291 291

INTRODUCTION OVERVIEW OF LIKELY FLUID TRANSPORT MECHANISMS

293

CRACK PROPAGATION AND THE FATE OF CRACK-INDUCED PERMEABILITY

291 291 292 292 292 293 293 294 294 294 295 297 298 298 298 300 301 302

Hydrofracture Porous flow Crack networks Inter granularfluid Permeability estimation Surface tension-driven infiltration

Basics of fracture mechanics Processes that affect the onset of unstable crack extension Chemical environment Effect ofP, T and microstructure Stress corrosion Hydrofracturing Processes controlling the longevity of crack-induced permeability Crack sealing Crack healing Effects ofP.T and fluid composition on crack healing Other effects on crack healing Effect of deformation

viii

302

INTERGRANULAR DISTRIBUTION OF FLUID

315 315 315

FINAL COMMENTS ACKNOWLEDGMENTS REFERENCES

302 302 303 304 304 309 310 311 311 312

Overview Introduction Basic principles of fluid distribution theory Results of dihedral angle measurements Experiments with quartz Experiments with calcite Results with other minerals Absence of grain-boundary films Diffusion in fluid-bearing systems Observations of natural wetting features

Chapter 7

M.D. Barton, R.P. Ilchik & M.A. Marikos METASOMATISM

321 321 321 323 324 325 325 325 329 330 331 331 331 333 333 334 334 335 336 336 336 341 345 345

INTRODUCTION CHARACTERIZATION

Modal changes Chemical changes Reactions

TYPES

Igneous and clastic host rocks Na-Mg-Ca-K types Hydrogen-ion metasomatism Volatile addition and silication Fenitiziation Carbonate and ultramafic host rocks Carbonate rocks Ultramafic rocks

ENVIRONMENTS AND DISTRIBUTION

Mafic intrusive suites Intermediate intrusive suites Felsic intrusive suites Strongly alkaline suites

DLSCUSSSION

Intrusion compositions Distibution and zoning

ACKNOWLEDGMENTS REFERENCES

Chapter 8

John M. Ferry

DEHYDRATION AND DECARBONATION REACTIONS AS A RECORD OF FLUID INFILTRATION 351 351 352

352 352 352 352

INTRODUCTION HISTORICAL PERSPECTIVE THE PETROLOGIC RECORD OF FLUID INFILTRATION DURING CONTACT METAMORPHISM: SELECTED EXAMPLES

Five aspects of the petrologic record Sequences of isograds Mineral assemblages Reaction progress

ix

356 356 356 357 358 360 360 362 363 363 365 367 367 370 370 370 370 371 371 372 373 373 373 374 374 374 377 377 379 380 380 380 382 382 383 384 387 387 388 388 388 390 390 390 390 391 391

Spacing of isograds Direction ofmetamorphic grade Scope of review INFILTRATION THEORY

Box models Models for coupled fluid flow and chemical reaction General theory - systems with local mineral-fluid equilibrium General theory - input of disequilibrium fluid at the inlet of the flow system

APPLICATIONS TO SIMPLE SYSTEMS

Effect of fluid flow along a temperature gradient Model siliceous limestone Model pelite Effect of fluid flow along a pressure gradient Model siliceous limestone Model pelite Summary and generalization of applications to simple systems Infiltration-driven contact metamorphism Time-integrated fluid fluxes of infiltration-driven contact metamorphism Equilibrium vs. disequilibrium fluids Relative importance of increasing temperature vs. decreasing pressure during equilibrium flow Carbonate rocks vs. pelites as petrologic records of fluid-rock infiltration

FLUID FLOW IN CONTACT AUREOLES : IDENTIFICATION AND MEASUREMENT

Model rock compositions Phase equilibria and conditions of contact metamorphism Prograde contact metamorphism—zero vs. infinite time-integrated fluid flux Predicted mineralogy: zero time-integrated flux Predicted mineralogy: infinite time-integrated flux Mineralogical criteria for distinguishing zero from infinite fluid flux Quantitative constraints on time-integrated fluid flux from observed mineral assemblages Lower bounds on qm from divariant assemblages Upper bounds on qmfrom invariant assemblages Bounds on qmfrom selected univariant assemblages and the invariant assemblage dolomite-calcite-tremolite-forsteritediopside Limits on time-integrated fluid fluxes during contact metamorphism Fluid flow or large porosity? Quantitative constraints on time-integrated fluid flux from measured reaction progress Quantitative constraints on time-integated fluid flux from measured spacing of isograds

DISCUSSION AND OVERVIEW

Role of infiltration in the mineralogical evolution of contact aureoles Time-integrated fluid flux Lithology and fluid flow Inner and outer aureoles Late hydration Fluid flow direction and hydrologic models for contact metamoiphic events Petrologic evidence of flow direction Hydrologic models of contact metamorphism ACKNOWLEDGMENTS REFERENCES

x

Chapter 9

Peter I. Nabelek STABLE ISOTOPE MONITORS

395 396 397

INTRODUCTION LIST OF SYMBOLS NOTATION AND FRACTIONATION FACTORS

400

ISOTOPIC CHANGES IN CLOSED SYSTEMS

404

OPEN SYSTEMS: L o s s OF VOLATILES

397 397 401 402

The 3-value and standards The fractionation factor

Temperature effects on stable isotope compositions of minerals Isotopic changes in minerals during net-transfer reactions Combined 3180-3 13 C depletions in calc-silicate metasediments

407 408

OPEN SYSTEMS: FLUID-ROCK INTERACTION

415 418

ISOTOPE SYSTEMATICS OF OPEN CONTACT AUREOLES ISOTOPE SYSTEMATICS OF LARGELY CLOSED AUREOLES

424 428

STABLE ISOTOPE THERMOMETRY IN CONTACT AUREOLES KINETIC CONSIDERATIONS AND RETROGRADE EXCHANGE

430 430 430

CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

408 410 413 418 421 422 428 428

Zero-dimensional (one-box models) and mixing One-dimensional flow; numerical multibox models One-dimensional flow; chromatographic exchange

The roles of permeability and structure The origin of fluids in largely-closed contact aureoles Skarns

Kinetic considerations Retrograde exchange

Chapter 10

K.P. Furlong, R.B. Hanson & J.R. Bowers MODELING THERMAL REGIMES

437 438 439 440 443 444 444 445 447 448 448 448 451 452 452 456 456 458 458 462 463 463

INTRODUCTION LIST OF SYMBOLS MODELING HEAT TRANSFER MODELING COOLING INTRUSIVES

Temperature dependence of thermal diffusivity Conduction as rate controlling process Shortcomings of analytical models

CRUSTAL THERMAL REGIMES

Transient geotherms Crustal heterogeneities

NUMERICAL SOLUTIONS

Finite difference approximation Boundary conditions

CONDUCTION-DOMINATED SYSTEMS

Classifying systems Intrusive geometry One-dimensional models Two-dimensional models Three-dimensional models Latent heat Endothermic reactions Host rock thermal gradients

xi

465 467

Multiple intrusive events Aureole deformation

467 468

LOW-PRESSURE METAMORPHIC BELTS MODELING HYDROTHERMAL EFFECTS

497 497 498

CONCLUSION ACKNOWLEDGMENTS REFERENCES

469 470 470 471 472 472 473 474 475 475 478 481 481 482 484 484 484 487 487 491 492 495 496

General equations and principles Conservation of mass (or continuity) equation Darcy's law Conservation of energy equation Solution methods and basic assumptions Models with A • u = 0 Fluid equation of state Permeability arid porosity First-order controls on fluid flow: effects of geometry and permeability General features Cooling times Bulk fluid! rock ratios Evolution of flow patterns Flow directions and flow paths Two-phase flow Comples geometries and anistropy Topography Pore pressure increases and permeability changes Enhancement of permeability Reduction of permeability Fluid production Overview, summary, and application of models Applications

Chapter 11

Raymond L. Joesten

KINETICS OF COARSENING AND D I F F U S I O N - C O N T R O L L E D MINERAL GROWTH 507 508 509 511

511 512 513 514 515 516 516 517 518 520 521 523 524

WHY STUDY KINETICS IN CONTACT AUREOLES ? INTRODUCTION TO THE KINETICS OF COARSENING EMPIRICAL DESCRIPTION OF COARSENING KINETICS THEORETICAL DESCRIPTION OF COARSENING KINETICS

Surface energy reduction-the driving force for coarsening Surface energy and particle coarsening Surface energy and matrix coarsening Coarsening of dispersed particles - porphyroblast coarsening Particle growth controlled by matrix volume diffusion Particle growth controlled by matrix grain-boundary diffusion Particle growth controlled by dissolution kinetics at the particle/matrix interface Generalfeatures ofparticle coarsening rate Integrated particle coarsening equations Matrix coarsening - coarsening of monomineralic aggregates Grain size distributions for particle and matrix coarsening Effect of volume fraction on particle coarsening Application of theoretical coarsening models to the analysis of coarsening in laboratory and natural systems

xii

525

EXPERIMENTAL COARSENING OF ROCK-FORMING MINERALS

536

MATHEMATICAL MODELING OF THE KINETICS OF THERMALLY-ACTIVATED PROCESSES IN CONTACT AUREOLES

525 530 533 533 533 534 535 536

536 538 539

539 543 543 546 551 553 554 558 558

558 561 562 562 564

564 566 566 567 569 569 569 569

569 570 571 573 575 578 579

Calcite Quartz Wollastonite Extrapolation of experimental coarsening results to geologic times Calcite Quartz Does grain diameter stabilize at long times? Compensation equation for retrieval of Arrhenius coefficients for normal grain growth from measurements at a single temperature

Kinetic description of non-isothermal coarsening in contact aureoles Crank-Nicolson finite difference model for the crystallization and cooling of a dike using the enthalpy method

RECRYSTALLIZATION AND COARSENING KINETICS IN CONTACT AUREOLES

Calcite coarsening in the contact aureoles of basalt dikes Quartz coarsening Retrieval of coarsening parameters using a thermal history model for the Christmas Mountains contact aureole A field test of coarsening models in the Ballachulish contact aureole Traversella contact aureole Low temperature quartz coarsening in the Ouachita orogen Wollastonite recrystallization and coarsening in the Christmas Mountains contact aureole

DIFFUSION IN CONTACT METAMORPHISM KINETICS OF DIFFUSION IN POLYCRYSTALLINE AGGREGATES

Grain-boundary diffusion kinetics Scale of diffusion penetration in metamorphic rocks Retrieval of oxygen grain-boundary diffusion coefficients from coarsening coefficients Diffusion and the kinetics of heterogeneous mineral reactions

CHERT NODULE REACTION RIMS IN CARBONATE ROCKS IN CONTACT AUREOLES

Chert nodule reaction rims in limestone Chert nodule reaction rims in dolostone Alta contact aureole Beinn an Dubhaich contact aureole Kaizuki-yama granit, Japan Santa Oliala tonalite, Spain Aankit basaltic feeder dike, U.S.S.R.

STEADY STATE MODELING OF THE DIFFUSION-CONTROLLED GROWTH OF CHERT NODULE REACTION RIMS

Constrained mass balance modeling of mineral assemblage zoning Model chert limestone reaction rims Model chert dolostone reaction rims

EXPERIMENTAL MODELING OF CHERT NODULE REACTION RIMS KINETICS OF DIFFUSION-CONTROLLED REACTION RIM GROWTH IN THE CHRISTMAS MOUNTAINS CONTACT AUREOLE ACKNOWLEDGMENT REFERENCES

xiii

D.M. Kerrick, A.C. Lasaga & S.P. Raeburn

Chapter 12

KINETICS OF HETEROGENEOUS REACTIONS 583 584 586

INTRODUCTION LIST OF SYMBOLS KINETICS OF OVERALL REACTIONS: THEORY

601 601

KINETICS OF OVERALL REACTIONS: EXPERIMENTAL STUDIES Methodology

586 587 587 596 603 603 609 610 610 613 613 613 616 618 619 619 620 621

Elementary versus overall reactions Steady state and metastability Heterogeneous reaction rate laws A simple flow and reaction model

Experimental studies of selected reactions Andalusite osillimanite Calcite + quartz wollastonite + CO2 Tremolite + dolomite =>forsterite + calcite + CO2 + H2O Tremolite + calcite + quartz =>diopside + CO2 + H2O Dolomite + quartz =>diopside + CO2 Muscovite + quartz andalusite + K-feldspar + H2O Brucite periclase + H2O Critique of experimental studies on heterogeneous metamorphic reaction kinetics Application of experimental studies to contact metamorphism Calcite + quartz wollastonite + CO2 Tremolite + dolomite =>forsterite + calcite + CO2 + H2O Tremolite + calcite + quartz =>diopside + CO2 + H2O Critique of the application of experimental kinetic data to contact metamorphism

622 622

KINETICS OF OVERALL HETEROGENEOUS REACTIONS: FIELD STUDIES Overstepping and kinetic isograds

639

KINETICS OF NUCLEATION AND GROWTH

634 634 636 637 639 639 641 644 644 645 645 652 652 653 656 658

Kinetic analysis of selected heterogeneous reactions Andalusite =>sillimanite reaction Reactions in metamorphosed siliceous dolomites Retrograde heterogeneous reaction kinetics in contact metamorphism

Theories of nucleation and growth in metamorphism Nucleation rates Growth rates Overall transformation rates Textural information relevant to crystallization kinetics Spatial distribution of crystals Crystal size distributions Compositional information relevant to crystallization kinetics Theories of compositional zoning in garnet Conversion of compositional data to growth rate information Significance of non-bell-shaped MnO profiles to crystallization kinetics Discussion of published metamorphic crystallization kinetic studies

662

SUGGESTIONS FOR FUTURE RESEARCH

666 666

ACKNOWLEDGMENTS REFERENCES

662 662

Experimental studies Field studies

xiv

Chapter 13

S.R. Paterson, R.H. Vernon & T.K. Fowler, Jr. AUREOLE TECTONICS

673 673 675

WHY STUDY STRUCTURES IN AUREOLES? HISTORICAL REVIEW OF EMPLACEMENT MECHANISMS FURTHER EVALUATION OF EMPLACEMENT MECHANISMS

688

AUREOLE STRUCTURES AND TIMING RELATIONS

707

RECENT METHODS

709

DISCUSSION OF AUREOLE TECTONICS

714

REFERENCES

675 676 677 679 680 681 684 684 685 686 689 692 692 693 695 697 700 701 701 702 703 706 706 707 708 708 708 709 709 710 711 711712 712 712 713 713

Introduction Stoping Cauldron subsidence, ring-dikes and cone sheets Laccoliths Block elevation along faults Diapirs and ballooning plutons Mystery plutons Wall-rock melting Sills/dikes Regional deformation and extension

Foliation patterns Lineation patterns Brittle structures Migmatites in contact aureoles Timing relationships Distinguishing types of structures in granitoids Foliation patterns Additional factors influencing structural patterns and timing Pluton shapes Post-emplacement strain partitioning Poiphyroblast-matrix relations as timing indicators in aureoles PMR s in contact aureoles without regional deformation PMR's in contact auireoles with concurrent regional deformation

Kinematic indicators Strain analyses and mechanical modeling Fractals Rates of processes Crystal size distributions Thermomechanical modeling and field studies

Most ductile processes Brittle/ductile processes Brittle wall rock processes Summary of timing relations Post-cleavage plutons Syn-cleavage plutons Pre-cleavage plutons Concluding remarks

Chapter 14

M.D. Barton, J-M. Staude, E.A. Snow & D.A. Johnson AUREOLE SYSTEMATICS

723 724 725

INTRODUCTION SYMBOLS AND ABBREVIATIONS PHENOMENOLOGY OF CONTACT METAMORPHISM

xv

725 725 728 728

Scale-independent features Isochemical changes Allochemical changes Time-space relationships

730

VARIABLES

732

ENVIRONMENTS OF CONTACT METAMORPHISM

797

SYNOPSIS AND SYNTHESIS

820 821

ACKNOWLEDGMENTS REFERENCES

730 732 732 732 734 735 743 141 756 757 761 770 111 772 115 775 778 779 780 781 790 790 794 795 796 797 797 803 803 808 814 820

Principal variables Depth, composition, and size Other variables

Overview Volcanic and hypabyssal Convergent margin-raltated - terrestrial Convergent margin-related - marine Rift- arid hot-spot related Tholeiitic magmatism Spreading centers Upper crust Intrusions Aureoles Fluidflow and metasomatism Multiple intrusive centers Pegmatites Layered intrusions Mafic layered intrusions Anorthosite massifs Middle crust Lower crust and upper mantle Lower crust Upper mantle Non-igneous types of contact metamorphism Tilted and composite sections

Intrusions and host rocks Distribution with depth Aureoles Isochemical systematics Allochemical systematics Time-space relationships Concluding remarks

xvi

Chapter 1

Derrill M. Kerrick

OVERVIEW OF CONTACT METAMORPHISM INTRODUCTION Since V.M. Goldschmidt's (1911) classic study on the Oslo area of Norway, contact aureoles have provided excellent natural "laboratories" for the elucidation of the processes and conditions of metamorphism. Reverdatto (1973) compiled petrologic studies on contact metamorphism. However, considerable research on the processes and controls of contact metamorphism has been carried out in the last two decades. This volume presents a review of contact metamorphism from a variety of geological subdisciplines (igneous and metamorphic petrology, geochemistry, thermal modeling, and structural geology). This chapter presents a brief overview of the controls and processes of contact metamorphism, and is primarily intended for readers who are relatively unfamiliar with this topic. GEOLOGIC SETTING Contact and regional metamorphism have traditionally been separated according to scale and to the spatial relationship to intrusive heat sources. Accordingly, contact metamorphism occurs in aureoles surrounding intrusives. In contrast, regional metamorphism is of regional extent with no apparent relation to intrusives as heat sources. The association with intrusives provides a global-scale distribution between contact aureoles and intrusive belts and batholiths (Fig. 1). Magmatic arcs at zones of continental collision (Fig. 2) are a locus

Southern California

Figure 1. Exposed granitic rocks in North America. Note the intrusives along the western margin of the continent that form part of the circum-Pacific magmatic arc. The intrusives in the shield area of Canada are generalized (metamoiphic rocks are included). (Modified from Best, 1982).

2 TRENCH \ ™

T

. 2

VOLCANIC FRONT

, 5

1. zeolite 2. prehnite-pumpellyite 3. blueschist 4. eclogite 5. greenschist 6. epidote amphibolite 7. amphibolite 8. granulite

ASTHENOSPHERE

Figure 2. Schematic illustration of a convergent plate boundary showing a paired metamorphic belt. Low-P, high-T metamorphism occurs in the magmatic arc whereas high-P, low-T metamorphism occurs in the subduction zone. (After Ernst, 1976).

for contact metamorphism. Indeed, many published studies of contact metamorphism involve aureoles within these magmatic zones. As shown in Figure 2, intrusives are abundant within the low-P/high-T portion of paired metamorphic belts. Thus, contact aureoles are typically hosted in rocks of the greenschist and amphibolite facies of regional metamorphism. Because intrusives are rare in high-P/low-T rocks closer to the trench (Fig. 2), aureoles are rarely hosted in rocks of prehnite-pumpelleyite and blueschist facies. INTRUSIVES The pressures of contact metamorphism are a function of depth of emplacement of intrusives at the present levels of exposure. As depicted in Figure 2, intrusives occur over a wide range of crustal levels. Shallow intrusives typically have well-developed contact aureoles hosted by country rocks that are unmetamorphosed or enjoyed earlier low-grade regional metamorphism (greenschist facies). Consequently, as shown in Figure 3, aureoles adjacent to shallow intrusives are typically well defined. In contrast, deep plutons typically intrude high grade country rocks and the aureoles are not as obvious as those developed adjacent to shallow intrusives. Accordingly, there is literature bias toward investigations of aureoles adjacent to shallow intrusives. As predicted from a conductive cooling model, there is a general correlation between size (width) of aureoles and size of intrusives (Fig. 3). This correlation is explicable in terms of the relative amounts of heat evolved from large versus small intrusions. Aureoles adjacent to felsic intrusions (e.g., granodiorites) are generally larger and have higher grade metamorphic rocks than aureoles next to basic intrusions (e.g., gabbros). This correlation is compatible with geothermometry which suggests that basic magmas (e.g., gabbros) are hotter than felsic magmas (e.g., granodiorites). The systematic correlation between the nature and size of aureoles and the composition of intrusives is reviewed in Chapter 14 of this volume. The shape of an aureole depends upon the form of the intrusive. Based on a conduction cooling model for the intrusive, isograds in an aureole should generally conform to the intrusive contact. Although this conformance is well displayed in some aureoles (Fig. 3),

3 QUATERNARY ALLUVIUM

SAWTOOTH

MIOCENE OR PLIOCENE ANDESITE FLOWS LATE CRETACEOUS, EARLY TERTIARY TRONDHJEMITE ADAMELLITE GRANODIORITE gP

I

LATE TRIASSIC METAMORPHOSED SEDIMENTS I INNER OUTER

AUREOLE AUREOLE

Figure 3. Contact aureoles in a portion of the Santa Rosa Range, Nevada. As described by Compton (1960), the pelitic hornfelses of the aureoles are readily distinguished from the phyllite country rocks that were subjected to pre-intrusive greenschist facies regional metamorphism. As predicted from conduction thermal modeling, the aureole around the Santa Rosa stock is larger than that of the smaller Sawtooth stock. (From Nagy and Parmentier, 1982, Fig. 1).

Figure 4. Orthographic projection of the Rattlesnake Mountain pluton, San Bemadino Mountains, southern California. (From MacColl, 1964, Plate 7).

4 there are notable exceptions. The shape of the intrusion at depth may provide an explanation for aureoles where the shapes and forms of metamorphic isograds do not conform to the intrusive contact. For example, due to geometric irregularity of the pluton depicted in Figure 4, the geometry of the isograds would be correspondingly complex. An apparent example of this effect is shown in Figure 5. However, an alternative explanation is provided by heat transport due to fluid flow. As discussed in Chapter 10, the presence of a magmatic heat source will cause convective circulation of fluids in the aureole and adjacent country rocks. As depicted in Figure 6, marked convective fluid flow through permeable rocks can significandy affect the thermal regime in aureoles and, thus, markedly affect the spatial locations of isograds. The thermal effects of focused fluid flow through lithologies of relatively high permeability (Fig. 7) could affect the spatial locations of isograds. COMPARISON WITH REGIONAL METAMORPHISM There are notable advantages of studying contact metamoiphism compared to regional metamorphism. First, significant changes in metamorphic grade occur over small distances. In several aureoles, single lithologic units can be followed from outside the aureole to the intrusive contact. As illustrated in Figure 8, the scale of contact metamorphism facilitates the opportunity to track the prograde metamorphism of specified horizons within formational and lithologic units. Second, computations of conductive and convective cooling of intrusives provide important insight into the thermal histories of contact aureoles. Because temperature during contact metamorphism increases with decreasing distance from an intrusive, sampling aureoles in traverses perpendicular to intrusive contact provides a constraint on the relative maximum temperatures of samples collected along each traverse. This is a significant aid for geothermometry. Thus, for example, the relative temperatures of samples along traverses B, C, D and F in Figure 8 are known. Third, intrusives provide a source of volatiles. In light of the considerable research on the physicochemical properties of magmatic volatiles, we are provided with an independent way to further our understanding of fluid transport and metasomatism in aureoles. Fourth, contact aureoles are advantageous in that metamorphism was essentially isobaric. This conclusion is verified by considering the lithostatic pressure variation using the equation: P = p gh (p = rock density, g = acceleration due to gravity, and h = depth). Even in alpine terranes, maximum relief within a given aureole is typically « 1 km. Taking an average metamorphic rock density of 2.8 g/cm 3 (Daly et al., 1966), the maximum pressure gradient would be 280 bars for 1 km of relief. As reviewed in Chapters 4 and 5 of this volume, pressure differences < 280 bars will have a negligible effect on the temperatures most metamorphic equilibria within the pressure range of most contact aureoles. Exclusion of the pressure variable considerably simplifies geothermometry of contact metamorphism compared to regional metamorphism. Fifth, belts of regional metamorphism typically have evidence of complex tectonothermal histories with several periods of metamorphism and deformation. Current models suggest that tectonic imbrication along thrust faults yields considerable differences in the P-T-t histories of rocks within adjacent fault blocks (an excellent example is provided by Spear et al., 1990). In contrast, within the time period of contact metamorphism, rocks in contact aureoles typically have far simpler thermal and tectonic histories. With single magmatic pulses there is a corresponding single contact metamorphic event. Multiple intrusives yield multiple thermal pulses. However, by radiometric age dating of the various intrusive phases, the thermal history of the associated aureole is tractable. Because of the relative size of contact aureoles compared to belts of regional metamorphism, the spatial relationship of aureoles with intrusives, differences in mineral assemblages because of pressure differences (especially pelitic lithologies), and differences in metasomatism (e.g., the restriction of skarns to contact aureoles), contact metamoiphism has traditionally been considered to be distinct from regional metamorphism. However, this distinction has long been questioned. Belts of regional metamorphism typically contain abundant intrusives. Could intrusives collectively increase the regional thermal gradient and thus be a primary cause of regional metamorphism? Within the last few years there has

5 Chlorite Biotite + c h l o r i t e Gornet + b i o t i t e + chlorite S t o u r o l i t e + biotite (chlorite) Andalusite + stourolite + b i o t i t e Andalusite + biotite Sillimanite + biotite C o r d i e r i t e + biotite + c h l o r i t e Andalusite + cordierite + biotite Sillimanite + K - f e l d s p a r + c o r d i e r i t e Gornet ( i n 4 - 7 , B ) All + m u s c o v i t e + q u a r t z + ilmenite Cordierite isogrod (M2n) Stourolite isograd (M2) NORTH

Staurolite-out isograd (M2)

LOBE

Staurolite-out isograd (M2s) Sillimanite isograd (M2s)

2

BINGHAM

CENTRAL

LITTLE

BIGELOW

LOBE

MOUNTAIN

KINGFIELD

Q.

SOUTH

BINGHAM ANSON

Q

Q. Q,

LOBE

METASEDIMENTARY UNITS Seboomook Fm.

Madrid, Fall B r o o k Fms. C a l c a r e o u s Phyllite Smalls Falls, P e r r y Mtn. Fm: Dead River Fm.

SURFACf

Figure 5 (above). Aureole of the Lexington Batholith, Maine, The shapes of the staurolite-in and staurolite-out isograds east of the central and south lobes (in the vicinity of Bingham) are attributed to the presence of a tongue-shaped subjacent intrusive. The conformance of the isograd adjacent to the northern lobe is compatible with a vertical cylinder shape of this lobe. (From Dickerson and Holdaway, 1989, Fig. 2). Figure 6 (left). Schematic illustration of convective hydrothermal fluid flow (dashed lines) produced by an intrusion. (From Fyfe and Henley, 1973, Fig. 8B).

6 LARGE

SCALE

Figure 7. Schematic illustration of heterogeneous fluid flow in rocks undergoing metamoiphism. Top: largescale fluid flow through a shear zone. Bottom', small-scale heterogeniety of fluid flow. In (1) the arrows are vectors illustrating variation in the fluid flux due to varying permeabilities of different lithologies. (From Etheridge et al., 1983, Fig. 7).

Bs

Blown Sand Maas Semipelitic Schists

1

h 1M

Portnoo Limestone C l e e n q o r t ( C l o o n e y ) Pelitic S c h i s t s !

E 2 3

Tonalitic Margin of Ardara Pluton

Milli

Granite-Diorite Complex of Naran Hill

THE A R E A S O U T H OF G W E E B A R R A BAV Co.

DONEGAL

S h o w i n g l o c a t i o n of analysed s a m p l e s in r e l a t i o n t o the geology

y / 'F

Dip of S t r a t a Vertical Strata Faults SCALE

I . l i I 1—1 0 •/,

Geology by

I I Mile

M/cAkaad. A.PGmdy, R.S.Mithol, W.SMtchtr and H.H.Read.

Figure 8 Northern aureole of the Ardara pluton, Donegal, Ireland. The filled rectangles represent samples analyzed by Pitcher and Sinha (1958). (From Pitcher and Sinha, 1958, Fig. 2).

7

|—] Low p r e s s u r e facies series M e d i u m - p r e s s u r e facies series

fy] j

Upper limit ot easily recognizable

metamorphism

intrusives

Figure 9. Schematic illustration of a model for the development of regional metamorphism across the Mesozoic magmatic arc in the western United States. Note the development of regional low-pressure metamorphism above the batholith and the intrusive cluster of the western Great Basin. The relatively isolated intrusives of the eastcentral Great Basin produced localized aureoles. (From Barton et al„ 1988, Fig. 5-9).

been renewed interest in the role of intrusives for low pressure regional metamorphism ("LPM"), which is characterized by metamorphic minerals (especially andalusite and cordierite in metapelites) that are restricted to relatively low pressures (P < 4 kbar). As illustrated in Figure 9, intrusive heat sources have been proposed for LPM because of the abnormally high geothermal gradients ( d P / c f T > 30°C/km). Some workers have regarded LPM as regional-scale contact metamorphism. The issue of heat sources for LPM is the subject of current controversy (Chapter 10). Nevertheless, LPM epitomizes the potential indistinction between contact and regional metamorphism and may argue for a continuum between these two types of metamorphism. Contact aureoles provide a way to track the tectonothermal evolution of metamorphic belts. As seen in Figure 10b, radiometric age dating, coupled with aureole thermobarometry, yields information on the depth at stage (4) of the orogen depicted in Figure 10a.

magmas are generated by anatexis (hachured area). Stage (4) corresponds to magmatic intrusion and development of an adjacent contact aureole, (b) Schematic illustration of analysis of the barometric development of selected stages of the orogen. In stage (4), pressure is determined by geobarometry of the aureole and the time is determined by radiometric age dating of the intrusive. (From Jamieson and Beaumont, 1989, Figs. 5 and 6a; copyright Geological Society).

METAMORPHIC PROCESSES Metamorphism occurs by a combination of the following processes: coarsening, neocrystallization, metasomatism, anatexis, and deformation. The first two processes are ubiquitous in contact aureoles, whereas the last three processes are important to varying degrees in different aureoles. Complete understanding of the overall process of contact metamorphism demands an understanding of the mechanisms and relative roles of all five processes.

8

Distant from contact (m)

Figure 11 (above). Average grain size as a function of distance from intrusive contacts. The curve for the Dashkesan aureole represents calcite in marbles whereas the curve for the Morang aureole is quartz in hornfels. (From Spry, 1969, Fig. 39a; copyrighted by Pergamon Piess). Figure 12 (right). P-T grid showing some equilibria relevant to contact metamorphism of pelitic rocks. The heavy arrow represents the hypothesized prograde path in the aureole of the Anvil batholith in the northern Canadian Cordillera. A => E correspond to prograde metamorphic zones; the filled circles correspond to isograds. (From Smith and Erdmer, 1990, Fig. 15).

o —//—l 450

500

550

600

650

Coarsening Coarsening refers to mineral growth due to processes other than heterogeneous reactions. Perhaps the best known petrologic example of coarsening is the transformation of limestones and dolomites into marble. Indeed, marble ornamental stones owe their beauty to the process of coarsening. In contact aureoles, coarsening is revealed by an increase in average grain diameter with decreasing distance from an intrusive (Fig. 11). This correlation points to the thermal activation of the coarsening process. Considerable research on coarsening is summarized in the metallurgical and ceramic literature. Chapter 11 reviews theory and experiments relevant to coarsening in contact aureoles. Because of the time dependence of the coarsening process, particular emphasis is placed on the kinetics of coarsening. Neocrvstallization Much research in metamorphic petrology has focused on reactions responsible for the formation of new minerals that were not present in the protoliths. Afeocrystallization results from heterogeneous reactions driven by the Gibbs free energy changes of reactions. Most isograds are correlative with heterogeneous reactions. The P-T conditions at which metamorphic reactions proceed, and at which metamorphic mineral assemblages are stable, have been the focus of considerable petrologic attention since Bowen (1940) introduced the concept of the petrogenetic grid. As shown in Figure 12, the petrogenetic grid offers a way to determine the P-T conditions of contact metamorphism. However, this analysis is complicated by multivariancy introduced by variation in mineral and fluid composition. Denoting the mineral and/or fluid compositional variable with X,- (following the usual thermodynamic definition of mole fraction of a component i), much of the quantitative analysis of metamorphic petrology has focused on elucidating the P-T-X, conditions of metamorphism with the assumption that mineral assemblages in contact aureoles represent equilibrium. The P-T-X,- relations of metamorphic mineral assemblages and isograd reactions in major "reactive" lithologies in contact metamorphism are reviewed in Chapters 4 and 5. The assumption of equilibrium implies that there are no kinetic barriers for metamorphic reactions. If, however, sluggish reaction kinetics is significant, much of the quantitative foundation of metamorphic petrology (i.e., the petrogenetic grid) is invalid. The theoretical, experimental and field aspects of heterogeneous metamorphic reaction kinetics are reviewed in Chapter 12. Heterogeneous metamorphic reactions can be

9 forsterite

isograd.

I sum of carbonate, minerals

distance

from

contact

Figure 13. Maximum modal content of noncarbonate minerals in marbles of the Adamello aureole (northern Italy) as a function of distance from the intrusion. Bucher-Nurminen (1982) considered the marked increase in non-carbonate minerals at the forsterite isograd to arise from the influx of Si- and Al-rich magmatic fluids. (From Bucher-Nurminen, 1982, Fig. 2).

subdivided into two sequential processes: nucleation and growth of product minerals. Accordingly, the kinetics of nucleation and growth are separately considered in Chapter 12. Metasomatism Metasomatism refers to changes in bulk-rock composition of one or more chemical elements during metamorphism. Such allochemical metamorphism contrasts with isochemical metamorphism in which there are no changes in bulk-rock chemistry. Metasomatism is particularly evidenced in carbonate host rocks. In some aureoles, there is compelling evidence for significant major-element metasomatism of carbonate rocks (Fig. 13). The presence of skarns developed near contacts between intrusives and carbonate lithologies are products of wholesale metasomatic transformation. Because skarns are host rocks for ore deposits (tungsten is of particular note), the elucidation of the metasomatic controls and processes in the genesis of skams is of interest to both economic geologists and metamorphic petrologists. Anatexis Anatexis is synonymous with partial melting. Migmatites, which are restricted to the high-grade portions of numerous contact aureoles (Fig. 14), are compatible with the

Figure 14. Migmatitic gneiss in the aureole of the Standard pluton, west central Sierra Nevada, California (described by Kemck, 1970).

10

^

I U M P D ZONE 7AUF INNER OF MELTING

++

V

OUTER MELTING ZONE (MIGMAT1TE ZONE)

Figure 15. Schematic illustration of a model for the emplacement of the Cooma Granodiorite, New South Wales, Australia. The migmatitic envelope is dragged upward to higher crustal levels by the intruding diapir. (From Flood and Vernon, 1978, Fig. 2).

hypothesis of anatexis produced by elevated temperatures. Accordingly, the quartzofeldspathic (leucosome) portions of anatectic migmatites represent crystallized partial melt. Experimental studies of anatexis, coupled with studies of anatectic migmatites in terranes of regional metamorphism, are of considerable aid in elucidating anatexis in contact metamorphism. Anatexis can have a considerable influence on the physical and chemical evolution of contact aureoles. For example, anatectic melts are transient sinks for aqueous volatiles. The consequent reduction in a ^ O in the solid residuum (restite) is of significance in analysis of phase equilibria in migmatites. Anatexis in pelites subjected to contact metamorphism is reviewed in Chapter 4. Anatexis yields an significant change in rock rheology. The presence of envelopes of high-grade anatectic migmatites may have an important effect on intrusive mechanisms (Fig. 15). The structural role of migmatites in contact metamorphism is reviewed in Chapter 13. Deformation Because of the "isotropic" fabric, the abundant hornfelses in aureoles have traditionally been interpreted as products of metamorphism under static conditions. However, strain analysis and the presence of rotated porphyroblasts in aureoles counter this traditional assumption. This alternative view is epitomized by the model of "ballooning" intrusives (Fig. 16) whereby the aureole is considered to be attenuated due to forcible distention of an intrusive because of continued supply of magma from depth (akin to the stretching and attenuation of a balloon during continued inflation). Synmetamorphic strain could have important affects on the development of the contact aureole. For example, permeability and porosity (and, thus, fluid flow) would be affected. Significant attenuation of the aureole would affect the spatial reference frame of the aureole. In particular, the interpretation of distances of samples from intrusive contacts would be complicated by significant thinning of the aureole. The hypothesis of ballooning intrusions, and other questions regarding the structural evolution and controls of contact metamorphism, are reviewed and critiqued in Chapter 13. FLUIDS Fluids play a significant role in contact metamorphism and thus deserve particular attention. Petrologic studies of contact aureoles reveal a significant integrated flux of fluids. Fluid-rock interaction has a significant effect on the chemical and mineralogical evolution of contact aureoles. Accordingly, fluids are the focus of several chapters in this volume. Magmatic volatiles are considered in Chapter 2. Chapter 3 reviews the chemical, physical and thermodynamic properties of fluids in contact metamorphism. Phase equilibria of metamorphic reactions involving "mixed volatiles" are reviewed in Chapters 3, 4 and 5. Porosity and permeability in contact aureoles is reviewed in Chapter 6. Understanding of the mechanisms of bulk-rock fluid transport demands an understanding of the microscopic-scale "plumbing system" reviewed in Chapter 6. The transport of

11

\ w

\ \ \ V \ \

c

\ r

\

w

\ \

\s-m ^

€u

/

Marble Canyon Pluton

\ \ \

A \

\

/

1) \ x \ /pi - , * s / (I ^ * ' — / i * \\\ ^ + jy» / ' 7. " ^ n f x w

i^v^hf

/ III

'i'iif N

J

Santa | N. Rita 1 Flat \ / PlutonV,/ 0

r

/

\

>f

\\ \ \ \ \ \ Vc \ \ \ X\ \ \ \ \ x

\

/

\

\\

^ \

C

/

/

Figure 16. Generalized geologic map (top) and cross section (bottom) of the Papoose Flat pluton (California) illustrating attenuation of stratigraphic units by forcible intrusion. (From Sylvester et al., 1978, Fig. 9).

chemical species as solutes in flowing fluids is an important mechanism for metasomatism in contact metamorphism. This topic is reviewed in Chapter 7. The influx of fluids into a metamorphic system will drive heterogeneous metamorphic reactions. As reviewed in Chapter 8, analysis of the progress (extent) of metamorphic reactions provides a way to quantify the extent of fluid-rock interaction during metamorphism. Fluid-rock interaction will also affect the isotopic compositions of rocks and minerals. This topic is reviewed in Chapter 9. Fluid flow will also affect the thermal history of contact metamorphism. Of particular importance is the thermal effect resulting from the influx of relatively cool fluids involved in hydrothermal convection within pemieable rocks adjacent to intrusives (Fig. 6). Were there marked differences in the metamorphic thermal regimes of aureoles developed solely by conductive heat transport versus those with significant heat transport by convective fluid circulation? This and other questions regarding heat transport by fluids are reviewed in Chapter 10.

12

REFERENCES

Barton, M.D., Battles, D.A., Bebout, G.E., Capo, R.C., Christensen, J.N., Davis, S.R., Hanson, R.B., Michelsen, C.J. and Trim, H.E. (1988) Mesozoic contact metamorphism in the western United States. In: Ernst, W.G. (ed.), Metamorphism and Crustal Evolution of the Western United States. Rubey Vol. 7, Prentice Hall, New Jersey, 110-178. Bateman, P.C. and Chappell, B.W. (1979) Crystallization, fractionation, and solidification of the Tuolumne Intrusive Series, Yosemite National Park, California. Geol. Soc. Amer. Bull, Part 1, 90,465-482. Best, M.G. (1982) Igneous and Metamorphic Petrology. W.H. Freeman and Co., San Francisco, 630 p. Bowman, J.R. and Essene, E.J. (1982) P-T-X(C02) conditions of contact metamorphism in the Black Butte aureole, Elkhom, Montana. Amer. J. Sci. 282, 311-340. Bucher-Nurminen, K. (1982) On the mechanism of contact aureole formation in dolomitic country rock by the Adamello intrusion (northern Italy). Amer. Mineral. 67,1101-1117. Compton, R.R. (1960) Contact metamorphism in the Santa Rosa Range, Nevada. Geol. Soc. Amer. Bull. 71, 1383-1416. Dickerson, R.P. and Holdaway, M.J. (1989) Acadian metamorphism associated with the Lexington Batholith, Bingham, Maine. Amer. J. Sci., 289, 945-947. Einaudi, M.T., Meinen, L.D. and Newberry, R J . (1981) Skarn deposits. Econ. Geol. 75, 317-391. Emst, W.G. (1976) Petrologie Phase Equilibria. W.H. Freeman and Co., San Francisco, 333 p. Etheridge, M.A., Wall, V.J. and Vernon, R.H. (1983) The role of the fluid phase during regional metamorphism and deformation. J. Metam. Geol. 1,205-226. Flood, R.H. and Vemon, R.H. (1978) The Cooma Granodiorite, Australia: An example of in situ crustal anatexis? Geology 6, 81-84. Fyfe, W.S. and Henley, R.W. (1973) Some thoughts on chemical transport processes, with particular reference to gold. Mineral. Sci. Engng. 5,295-303. Goldschmidt, V.M. (1911) Die kontact Metamorphose im Kristianigebiet. Kristiania Vidensk. Skr., I, Math-Naturv. Kl. 11. Jamieson, R.A. and Beaumont, C. (1989) Deformation and metamorphism in convergent orogens: a model for uplift and exhumation of metamorphic terrains. In: Daly, J.S., Cliff, R.A. and Yardley, B.W.D. (eds.), Evolution of Metamorphic Belts, Geol. Soc. London Spec. Pub. 43, 117-129. Kerrick, D.M. (1970) Contact metamorphism in some areas of the Sierra Nevada, California. Geol. Soc. Amer. Bull. 81,2913-2938. MacColl, R.S. (1964) Geochemical and structural studies in batholithic rocks of southern California: Part 1, Structural geology of Rattlesnake Mountain Pluton. Geol. Soc. Amer. Bull. 75, 805-822. Nagy, K.L., and Parmentier, E.M. (1982) Oxygen isotopie exchange at an igneous intrusive contact. Earth Planet. Sci. Lett. 59, 1-10. Pitcher, W.S., and Sinha, R.S. (1958) The petrochemistry of the Ardara aureole. Geol. Soc. London Quart. J. 113, 393-408. Reverdatto, V.V. (1973) The Facies of Contact Metamorphism. Australian National University Pub. No. 233, 263 p. Smith, J.M. and Erdmer, P. (1990) The Anvil aureole, an atypical mid-Cretaceous culmination in the northern Canadian Cordillera. Can. J. Earth Sci. 27, 344-356. Spear, F.S., Hickmott, D.D. and Selverstone, J. (1990) Metamorphic consequences of thrust emplacement, Fall Mountain, New Hampshire. Geol. Soc. Amer. Bull. 102, 1344-1360. Spry, A. (1969) Metamorphic Textures. Pergamon Press, 350 p. Sylvester, A.G., Oertel, G„ Nelson, C.A. and Christie, J.M. (1978) Papoose Flat pluton: A granitic blister in the Inyo Mountains, California. Geol. Soc. Amer. Bull. 89,1205-1219.

Chapter 2

George W. Bergantz PHYSICAL AND CHEMICAL CHARACTERIZATION OF PLUTONS INTRODUCTION

The purpose of this chapter is to inventory and describe some of the generic features of intrusive systems which pertain to the understanding of contact metamorphism. Magmas are the sources of heat, mass and mechanical energy that yield contact metamorphism and associated deformation, and an appreciation for the manner in which the intensive and extensive variables vary during plutonism may aid in understanding the temporal and spatial details of contact metamorphism. A wide variety of thermal histories and corresponding styles of self organization are possible given the variety of initial and boundary conditions attendant with plutonism. This review concentrates on the physical and chemical processes in plutons. The scope of this review is limited largely to granitoids; literature on mafic systems or mid-ocean ridges is not included. The purpose of this review is to provide quick access to the current work and paradigms relating to plutonic systems. The first priority is to direct the student of contact metamorphism to the literature useful in constraining magmatic processes. Given the extremely broad nature of the subject and the space available, it was decided to sacrifice detail for scope; a functional understanding of magmatic processes will require further study of the works cited herein. Petrogenetic schemes or compositional-tectonic associations are not discussed. The importance of pedogenesis in the broad characterization of plutons is recognized, however, a detailed treatment of this topic is outside the scope of this review. Pedogenesis and regional tectonics are treated in a number of timely summaries and the interested reader is encouraged to look there. First among them are the general reviews of plutonism by Pitcher (1978, 1979, 1987), Whitney (1988), Zen (1988), and Chappell and Stephens (1988). A number of excellent compilations of magmatism in a regional context are available: reviews of Andean magmatism (Atherton and Tarney, 1979; Harmon and Barreiro, 1984; Pitcher et al., 1985) also see Hildreth's (1987b) review of Pitcher et al. (1985), studies of magmatism in North America (Anderson, 1990; Ernst, 1988), and compilations of papers discussing magmatism around the Pacific margin (Kay and Rapela, 1990; Roddick, 1983). Anthologies of related interest are those edited by Vielzeuf and Vidal (1990) which addresses granulites and crustal evolution, and Mereu et al. (1989) on the physical properties and processes of the lower crust. Contact metamorphism as a conjugate system The transfer of heat and mass from the magma to the country rock comprises what is known as a conjugate, or coupled, system (Bejan, 1984; Bergantz and Lowell, 1987). The important characteristic of conjugate systems is that the quantitative modeling of the transfer of heat and mass from the magma to the country rock requires explicit consideration of the heat transfer systematics on both sides of the contact zone. A direct physical analogy of this is the window of a house on a cold day. Heat is being brought up to the window by whatever processes may be operating in the room, such as a forced air furnace or perhaps natural convection from a wood stove. The heat is being transferred through the window to the cold thermal reservoir outside, and conditions outside the house will determine how efficient the environment is at removing the heat from the window. Thus, the temperature of the window reflects the balance of the heat transfer processes on each side of this coupled system. If the window feels warm to the touch, the heat transfer in the room is able to keep up with the losses to the environment and so the rate-limiting step is the thermal resistance associated with transfer in the outside environment. If the window feels cold to the touch, which is more often the case, that indicates that the rate-limiting processes are associated with heat transfer in the room bringing heat up to the window. The point is that the temperature of the geological window, which is recorded in the contact

14 metamorphism and is time dependent for any case of geological interest, provides a constraint for the possible processes that can occur in the coupled magma-country rock system. The implication of this is that the temperature and spatial extent of contact metamorphism can possibly provide some insight into the coupled nature of the processes operating on both sides of the contact. This suggests that one might be able to "invert" the geologically determined conditions of metamorphism to choose among the possible processes of heat and mass exchange. It is not difficult to imagine a variety of interactions between the pluton and the country rock: (a) conduction on both sides of the contact with or without phase change being explicitly included (Bowers et al., 1990), (b) conduction in the magma with hydrothermal convection in country rock (Cheng, 1978, 1981; Norton and Taylor, 1979; Parmentier, 1979, 1981); and a study by Johnson and Norton (1985) that deserves more attention, (c) convection on both sides of the contact (Bergantz and Lowell, 1987), (d) cracking of the solidifying magmatic rind which permits hydrothermal fluids to cross the contact (Canigan, 1986; Lister, 1974), or (e) conduction in the country rock with convection in the magma, with the possibility of simultaneous crystallization and melting (Bergantz, 1991). It is likely that more than one of these processes may operate during an episode of contact metamorphism, particularily if the magma is subject to open system behavior, e.g., additional magma enters the system. A knowledge of the contact metamorphism, such as the spatial distribution of maximum temperatures, provides a constraint with which the magmatic history must be consistent. This has some appeal for those working in magma dynamics, as it is very difficult to constrain processes in the magma from the temperatures recorded in the pluton. Undoubtedly, this difficulty is largely due to the fact that plutons represent the end result of what is a complex chemical and mechanical history. This complexity is manifested a number of ways: in the ubiquitous disequilibrium mineral textures formed during both growth and subsolidus conditions, the isotopic evidence for assimilation and other open system behavior, mineral fabrics formed during magmatic flow with superposed near or sub-solidus deformation features, and in the diverse sequence of chemistries and eruptive styles exhibited by extrusive systems. The characterization of plutons must borrow substantially from the study of volcanic systems where magmatic conditions can be reconstructed without the veil of subsolidus transformations. Even some of the simplest magmatic systems appear to have complex histories. For example, scientific drilling at the Inyo Dome, California, thought to be a simple shallow silicic system revealed a complex relationship between high and low Si magmas, mixing and transport (Vogel et al., 1989). Complex patterns of re-intrusion, magma mingling and mixing, and segregation can all occur in what is thought to be a single magma chamber as suggested in a number of studies; examples include Vesuvius (Civetta et al., 1991) and the chamber that existed below Mount Mazama (Bacon and Druitt, 1988). A fascinating degree of complexity in space and time of magma types and eruptive styles has been documented at Katmai, Alaska, where a plexus of small, compositionally distinct magma chambers are postulated (Hildreth, 1987a). A study that attempts to explicitly address the compositional, spatial and temporal relationships between plutonism and volcanism is the study of Lipman (1988). Even though volcanic rocks provide the best circumstances from which magmatic intensive variables can be estimated, it can be difficult to demonstrate equilibrium (Frost and Lindsley, 1991). In addition, the depth at which the pluton forms can radically impact the cooling and crystallization history by controlling the timing and extent of volatile exsolution (Swanson et al., 1989; Westrich et al., 1988). The implication of this is that the magmatic conditions that existed at the time of contact metamorphism may be different from those preserved in the pluton, and establishing the conditions which existed when the pluton was a viable "chamber" is subject to uncertainty. Another complication arises in inteipreting magmatic history in light of the contact metamorphism: more than one magmatic history can yield the same metamorphic history. In fact it appears that thermal contact metamorphism is consistent

15 with the simplest scenario imaginable: instantaneous intrusion followed by conductive cooling. This is discussed in more detail below. The most important aspect to consider when addressing the thermal evolution of plutons is the process of solidification. The enthalpy flux that ultimately yields contact metamorphism is invariably accompanied by an increase in crystallinity in the pluton. This can occur at the margins and the magma chamber may form a rind of crystals that propagates inward as cooling continues, perhaps conductivly in a manner analogous to the Hawaiian lava lakes. Alternatively, the crystallinity may increase in the center or in a distributed fashion in the melt, and the pluton will solidify uniformly, perhaps undergoing sustained and vigorous convection. These two end-members can yield very different calculated thermal histories, depending on the assumptions involved in parameterizing the heat transfer (Bergantz, 1991). The solidification process will also have a dramatic effect on the physical properties of the magma namely, the viscosity and the density of the magma will change as the crystallinity and volatile content increase. This in turn influences the processes driving the heat transfer and a feedback is created that is difficult to generalize. Laboratory and numerical experiments of solidification, as shown in Figure 1, reveal a rich diversity of solidification and cooling histories depending on the composition of the liquid,

L L

L

hw»tTTt

L —«..-»»tti,

iLwi«—îrfnttffli 11 ¡¡¡¡u

Figure 1. Output from the numerical simulation of the solidification of a binary, taken from Bennon and Incropera (1987). The left wall of the figure is kept at a constant temperature that is below the solidus, the right wall is insulated. The first panel shows the progression from all solid to a solid-liquid mixture into a pure liquid. The arrows indicate direction and magnitude of flow, note the upward flow near the right margin of the mush due to compositional effects on buoyancy. Also note the complex velocity field. The second panel gives the streamlines, the third is the temperature field and the fourth is the isocomposition contours. Although not directly applicable to magmatic systems, numerical modeling of this kind is currently being adapted to geologically relevant conditions.

16 and the physical properties of the system (Beckermann and Viskanta, 1988; Bennon and Incropera, 1987; McBirney et al., 1985; Oldenburg and Spera, 1991). These studies provide a look at the current state-of-the-art in the formulation and modeling of solidification processes. Although none of the laboratory and few of the numerical experiments use materials that are directly analogous to magmas, they reveal some of the generic processes that occur in plutons. One of the most important of these is the partitioning of the body into a mushy zone near the contact where crystals and melt form a self supporting framework and an adjacent slurry where crystals reside in an expanse of melt. This partitioning has two important implications for contact metamorphism: (a) the contact between a Theologically viable magma and it's solid container propagates inward with time away from the original intrusive contact, thus the solidifying margins of the magma also undergo contact metamorphism in the sense that cracking and interaction of fluids is possible, (b) the conjugate nature of the heat transfer requires explicit consideration of the phase changes in the magma (Bergantz, 1991), and the rate of propagation of the solidification front and any convection in the magma can only occur at rates consistent with the rate of heat loss through the country rock. Thus, the heat transfer can be conceptualized as a heat transfer "circuit" with thermal resistances in a series arrangement (a common analogy used in engineering textbooks) and the magma cannot pump out heat any faster than the country rock can carry it away. In fact the optimal cooling time for a perfectly mixed (convecting) hot fluid body surrounded by a conducting medium is only twice the cooling time if cooling by conduction alone. This partitioning of the magma chamber into a solidifying margin and an adjacent slurry in the interior, which may be subject to re-intrusion, precludes the use of simple dimensionless heat transfer parameters when describing the progress of solidification or the rate of heat loss from the body; the phase change process must be explicitly considered. This review discusses those elements of plutons that influence the physical processes that yield contact metamorphism. The focus is largely on heat transfer and less on mass transfer. Magmatic fluids (volatiles) clearly play an important role in contact metamorphism, as evidenced by ore deposits (Burnham, 1979b). However, quantitative models of the "second boiling" process coupled with the mechanical processes of country rock fracture, with attendant changes in permeability, have yet to be developed. The emphasis on heat transfer is consistent with the classical treatment of contact metamorphism, reflecting in part the quantitative accessibility of this approach. These and other generic aspects of the physical evolution of magmas are discussed in the reviews of Marsh (1989a) and Morse (1988). Establishing the physical history of a pluton requires a careful consideration of the magmatic intensive and extensive variables as the rate at which the pluton loses heat will depend on these in combination. We consider these below. INTENSIVE VARIABLES Sequence of crystallization Establishing the crystallinity at the time of intrusion and the subsequent sequence of crystallization of the magma is difficult but is often one of the only ways to practically bracket magmatic intensive variables. A knowledge of the sequence of crystallization is used to infer the temperature, pressure, water content and rheological properties of the magma. In practice, the order of crystallization is estimated petrographically from crystal morphologies and other textural criteria and then compared to laboratory and computer experiments of solidification. The uncertainties associated with this approach originate in the difficulty of interpreting plutonic textures and in comparing these with experimental systems whose components only partially match those of the pluton being studied. Since the seminal work of Tuttle and Bowen (1958) a number of experiments have

17 A

B

T{°

PI+Af+aO+L+V PI+Af+aQ+V

600

0

2

4

6

8

10

Wt. % H 2 0

600

A' 12

14

A' 0

2

4

6

8

10 12 14

Wt. % H 2 0

Figure 2. Temperature-XH20 diagrams for synthetic granite compositions. From Whitney (1988). These diagrams permit an estimate of the onset of saturation and the paragenetic sequence. The contours in the right hand panel give approximate percent of melt present. Information of this kind is crucial to developing transport and thermal models of magmas.

been done to establish phase relations in the granitic system (Luth, 1976; Wyllie, 1988). Both solidification and melting experiments for common granitoid rock compositions have been done under a variety of pressures and water contents (Huang and Wyllie, 1986; Naney, 1983; Stern and Wyllie, 198 la,b; Whitney, 1975). Of particular interest are the experiments of Naney (1983) and Huang and Wyllie (1986), whose experiments include the ferromagnesian silicates. A revealing and useful format for presenting the experimental data is found in plots of temperature vs. weight percent water for a given composition at a given pressure. Curves are drawn to illustrate saturation of a given phase and thus one can follow the sequence of crystallization for a given water content and even estimate when the system becomes saturated. Whitney (1988) provides crucial additional information: curves of volume percent liquid. An example of this is shown in Figure 2. These are critical to transport modeling of the solidification process (Bergantz, 1990). Melting experiments of broadly granitic materials have been done to simulate the process of granite generation by crustal fusion (Presnall and Bateman, 1973; Wolf and Wyllie, 1989; Wyllie, 1977) which seems to require very high temperatures to generate melt fractions in sufficient quantity to form extractable magmas. This has motivated melting experiments of metaigneous (Beard and Lofgren, 1991; Rushmer, 1991) and metapelitic protoliths (Patino-Douce and Johnston, 1991; Vielzeuf and Holloway, 1988) which generate large amounts of melt at temperatures like those expected in lower crust subject to intrusion by basaltic magma (Bergantz, 1989). Thus, for a given composition and water content, the phase relations for both the crystallization and melting of granitic systems are broadly known, although much work needs to be done with other volatile species and to further refine the thermodynamic database so that the solidification progress can be modeled numerically. Once sufficient experimental data exist to establish the thermodynamic properties of the crystal-melt system, computer algorithms can be developed which allow one to simulate crystallization in some detail (Ghiorso, 1985; Ghiorso and Carmichael, 1985, 1988a; Nekvasil, 1988b; Nielsen, 1990). Each of these authors uses a different approach to the numerical treatment of crystallization, and each algorithm gives good agreement between predicted and naturally occuring assemblages for certain compositional ranges; the formulation of Nekvasil is specifically designed for silicic systems. One element missing from these numerical crystallizers is the ability to model the presence of hydrous phases.

18 This is simply due to the absence of the appropriate thermodynamic data and is not an inherent difficulty in the numerical approach. The numerical approach to crystallization is extremely powerful in that it permits one to couple transport models of time dependent heat and mass transfer to geologically relevant conditions (Bergantz, 1990). One example of this is the work of Bowers et al. (1990) which uses the algorithm of Nekvasil to estimate the temperatures at which the phases appear and their contributions to the time dependent enthalpy changes in the magma. Computer models of phase changes coupled with conjugate heat and mass transport provide an important and exciting direction for the elucidation of the generic features of contact metamorphism. Although laboratory and numerical experiments allow predictions for the appearance of the phases as the pluton cools, it is more difficult to estimate the crystallinity during actual ascent and at the time of intrusion. Extrusive rocks contain up to ~50% crystals (Marsh, 1981). It can be difficult, however, to confidently know when in the magmatic history the crystals grew (compare the phase diagrams in Whitney (1975) for 2 and 8 kbar). Unlike lavas or ash flows which often reveal chilled margins and an estimate of the crystallinity at the time of eruption can be obtained, textures in plutonic rocks invariably represent re-equilibration during prolonged periods near solidus temperatures. Apart from pegmatitic and obviously volatile-rich aplitic apophyses, plutonic rocks most often show a uniformity in grain size, often right up to the contact. There can be little question that nucleation is heterogeneous and there is little evidence for the magma having been in a superheated condition. Crystal growth rates appear constant at 1 0 1 0 - 1 0 1 1 cm/s for a wide range of melt and crystal compositions and cooling conditions as discussed in the comprehensive review by Cashman (1990). These growth rates are consistent with very small amounts of undercooling which indicates that undercoolings are a negligible part of the thermal budget. The heat transfer systematics, whether conductive or convective, will not be influenced by them. Volatiles Volatiles have a substantial influence on magmatic systems: they can substantially lower solidus temperatures and vary the sequence of crystallization, depolymerize the melt structure and hence reduce viscosity and density, influence the ascent history by the onset of saturation, and induce complex patterns of repeated fracturing and partial quenching of the solidifying pluton. Despite the importance of volatiles in understanding magmatic history, it has been very difficult to rigorously quantify the thermodynamic state (speciation) of volatiles in a silicate melt and also very difficult to model the multiphase behavior of a crystal-melt-volatile system from a continuum mechanical approach. These problems are exacerbated when considering actual plutons as one can often only crudely guess the original volatile content, however fluid inclusions may permit estimates of magmatic volatile composition. There are four ways to estimate volatile content (Clemens, 1984): (1) by direct measurement, (2) geological inference, (3) thermodynamic calculation, and (4) experimentally. Of necessity, much of our understanding of magmatic volatiles comes from the study of volcanic systems and extrusive rocks, although it is difficult to confidently extrapolate from eruptive conditions in volatile stratified magma chambers to plutonic conditions. Summary discussions of the role of volatiles in granitoid magmatism can be found in Burnham (1979a,b) and Whitney (1988). Melt inclusions in phenocrysts and the dissolved water in volcanic glass provide a means of determining the water, carbon dioxide, fluorine, sulfur and chlorine contents of magmas at the time of eruption. Water is the most abundant volatile component: measuring 5 wt % in the Fish Canyon Tuff (Johnson and Rutherford, 1989b), 4.3 wt % in the Taupo volcanic center (Dunbar et al., 1989; Hervig et al., 1989), 4-6 wt % in the Bishop Tuff (Anderson et al., 1989), 2-4 wt % in the Bandelier Tuff (Sommer and Schramm, 1983), and 4.1 wt % from Obsidian Dome (Hervig et al., 1989). Many of these volcanic systems appear to have a gradient in water content suggesting that a vertical gradient in volatile content existed in the magma chamber. Understanding the manner in which water is

19 speciated in the melt has been more difficult. The influential works of Burnham (1979a,b) proposed using a solution model based on the simple system albite-water. This model has appeal in that it provides a useful and straightforward means of quantifying the thermodynamic state of water, although there are complexities and questions regarding speciation that remained to be addressed. In their summary article, McMillan and Holloway (1987) demonstrate that molar water solubility increases with decreasing silica in binary and pseudobinary silicates. Silver et al. (1990) note that hydroxyl groups are the dominant hydrous species at low water contents, and that increasing silica content and K over Na leads to an increase in the molecular water relative to hydroxyl. The Henry's law behavior of water provides a means of putting the thermodynamic modeling of water in silicate melts on a firmer basis. Given the chemically complex nature of granitic melts, estimating water content in practice is often done by comparing the petrogenetic sequence as determined petrographically with phase equilibria as determined by laboratory experiments or computer models (Maaloe and Wyllie, 1975); see discussion on the sequence of crystallization given above. Another way to bracket water content is to evaluate the role of water in melt-forming reactions as done by Wyllie et al. (1976), and Clemens (1984). None of these methods yield a precise estimate of volatile content and the presence of hydrous phases can only give minimum values. Carbon dioxide and fluorine are also important volatiles in magmas. The solubility of CO2 in silicate melts appears to be low (Holloway, 1976), hence many magmas may be saturated in CO2 for much of their history which will influence the activity of water. This in turn could have a profound effect on the crystallization and ascent history. Experiments by Peterson and Newton (1990) demonstrates that CO2 in silicate liquids can delay the crystallization of biotite, mafic material is retained in the melt. Stolper et al. (1987) discuss the mechanisms of CO2 dissolution. Anderson et al. (1989) note that there is an inverse correlation between water and CO2 concentration in the Bishop Tuff; the CO2 concentration varying from 0.005 to 0.035 wt %. Assessing the role of CO2 in the thermomechanical history of any particular granitoid pluton is difficult, few experiments with CO2 saturated and water bearing multi-component silicate melts have been done. Fluorine contents can range from tens of ppm to several percent (Bailey, 1977) and is usually held in biotite and hornblende; the F/OH ratio in biotite was used by Ague and Brimhall (1988b) to infer contamination of granitoids in the Sierra Nevada. Dunbar et al. (1989) report fluorine contents of about 450 ppm in rocks from the Taupo volcanic center. Manning (1981) considered the liquidus phase relationships in the water saturated Qz-Ab-Or system and found that the minimum liquidus temperature fell 100°C from that of the fluorine free system. Manning posits that there may be a continuum between magmatic and hydrothermal conditions in mineralized granitoids. Pichavant and Manning (1984) report occurances of up to 3.2 wt % fluorine in tourmaline granites and topaz granites. TTiey provide evidence that these granites formed from highly differentiated residual melts. Pichavant (1987) considered the influence of boron on the phase relations. The boron concentration does not significantly influence the phase relations in the haplogranite system; the boron content of magmas does not exceed about 1 wt %. Based on thermodynamic calculations (as opposed to thermomechanical), it is possible to evaluate the balance between the composition of the magmatic source material and initial depth, the ascent distance, and initial volatile content. Sykes and Holloway (1987), Hyndman (1981) and Marsh (1984) examine the energetics during ascent and demonstrate that water must be considered to yield realistic estimates of magmatic conditions. This limits the range of initial conditions and source compositions of those magmas which can ultimately reach the surface and become erupted: relatively hydrous melts are prohibited from travelling far from their source as the solidus is encountered at depth (Fig. 3). As crystallization proceeds in a magma, volatile elements can be preferentially partitioned into the melt phase until saturation occurs. The pressure increase accompanying boiling can induce fracturing of the crystallizing margins of the pluton,

20 8% H•2', 0

10

4%

1%

8 6 P, kb

4 2 0

800

900

1000 1100 1200 T °C

Figure 3. Ascent trajectories calculated by Sykes and Holloway (1987) for model system albite-H 2 0. Dashed lines represent a constant crystalAnelt ratio and dotted lines represent a melting rate of 3%/kbar.

providing a means for the volatiles to escape. Further solidification permits this process to continue yielding a second and subsequent boilings (Burnham, 1979b). There is no doubt that such processes are important in the thermal history of shallow plutons; a complete quantitative description of this process awaits development. Estimating temperature and pressure There are a variety of methods to estimate pluton temperature and pressure; Zen (1989) discusses the generic features of the two approaches that are commonly used to estimate conditions in plutons. One approach is extrinsic, where pluton temperatures and pressures are determined from the character of the country rock. There can be large uncertainties in this method, as discussed by Hodges and McKenna (1987). The second is intrinsic, which relies on the specific features of the pluton itself. With both methods it is difficult to confidently determine when in the history of the magmatic system the temperature and pressure were set. The temperatures and pressures recorded by magmatic mineral assemblages are usually those of consolidation of the magma and may not represent the same point in time at which contact metamorphism was achieved (Zen, 1989); see discussion on open systems below. In addition, confidently estimating magma temperature and pressure in plutons is difficult due to the re-equilibration of the relevant phases, and the inevitable loss of volatiles. Temperatures thus usually represent minimums and may actually be related to secondary, subsolidus processes. There is a growing appreciation that estimating conditions in volcanic rocks, which ostensibly represent the best sample of a thermodynamically contiguous system, requires a careful assessment of the equilibrium assumption (Frost and Lindsley, 1991). Nonetheless, a number of intrinsic geothermobarometers have been applied to plutonic rocks and we will discuss the systematics of a few of them below. For a general review of the principals behind their application see Bohlen and Lindsley (1987). There are two methods generally used to determine temperature in plutons. One approach is to compare the paragenetic sequence of precipitation of minerals with the phase experiments discussed above, e.g., Naney (1983), and as done by Hill (1988). This is obviously a rather crude way to estimate temperature, however it has the advantages of being inexpensive and usually straightforward. The other approach is to evaluate the

21 compositions of coexisting phases that are thought to represent equilibrium conditions such as coexisting plagioclase and potassium feldspars (Whitney and Stormer, 1977), amphibole and plagioclase (Blundy and Holland, 1990) and/or coexisting oxides (Frost and Lindsley, 1991; Whitney and Stormer, 1976). The experiments of Elkins and Grove (1990) provide the latest attempts at calibration of a two feldspar thermometer following the approach of Fuhrman and Lindsley (1988) and others (Ghiorso, 1984; Green and Udansky, 1986). The experiments of Elkins and Grove (1990) were done at 700°-900°C and 1-3 kbar under water saturated conditions. In most cases their measurements agreed with their thermodynamic model to within 20°C and agree well with Fe-Ti oxide temperatures obtained from volcanic rocks where the temperature could be independently constrained. The oxides are particularily difficult to work with as they commonly have exsolved and the assumptions involved in recreating the equilibrium assemblage often are untenable. Magmatic hornblende has been proposed as a geobarometer. Hammarstrom and Zen (1986) noted that the total A1 content of hornblende from calc-alkalic plutons increases linearly with increasing pressure of crystallization. This led to the development of an empirical barometer based on the assemblage plagioclase + quartz + potassium-feldspar + biotite + amphibole + titanite + Fe-Ti oxides (magnetite or ilmenite). There are two groups of calibrations: those of Hammarstrom and Zen (1986) and Hollister et al. (1987) which are empirical and rely on corroborating pressures as determined by the country rocks or the presence of magmatic epidote, and those of Johnson and Rutherford (1989a) which were derived from reversed experiments with / 02 buffered and vapor present. At present it would appear that the Johnson and Rutherford (1989) calibration would be preferred as the laboratory conditions provide a more confident estimate of total pressure. TTie experiments of Rutter et al. (1989) on partially melted tonalite demonstrate that the total Al content is very sensitive to the specific mineral assemblage and hence caution must be used when garnet or other phases are present. Blundy and Holland (1990) propose a different substitution and argue that the Al substitution is temperature dependent and that amphibole equilibria are not appropriate for geobarometry. They propose an amphibole-plagioclase geothermometer instead. In a regional study of granitoids, Vyhnal et al. (1991) explore both the temperature and pressure dependence of Al in hornblende and conclude that it may be difficult to discriminate both temperature and pressure effects because they are both linked to the solidus. They propose that this might occur due to the occurrence of more than one substitution reaction. It thus seems that the hornblende geobarometer should be used with caution, and that much work remains to be done before the hornblende geobarometer has a more complete thermodynamic basis. The presence of magmatic epidote has been used as a geobarometer by Zen and Hammarstrom (1984) who argue that magmatic epidote is a high pressure, near solidus phase that appears as a reaction product involving hornblende and the melt. TTie laboratory experiments of Naney (1983) produced epidote at a pressure of 8 kbar and epidote has been observed in the chilled margins of dike rocks and rhyolitic lava flows (Dawes and Evans, 1991; Evans and Vance, 1987) establishing that it can occur as a near-liquidus mineral. The difficulty in using the presence of epidote as a geobarometer is in establishing whether the plutonic epidote is indeed magmatic and under what water and total pressures the epidote grew and last equilibrated. Oxygen fugacity also exerts a strong control on epidote stability. Dawes and Evans (1991) note that there are three types of magmatic epidote in the dacitic dikes of the Front Range, which were emplaced at 2 kbar (or less) and provide evidence that epidote formed at pressures of about 8 kbar. As with magmatic hornblende, a complete thermodynamic characterization of epidote awaits development and hence its use as a geobarometer is subject to some uncertainty. Physical properties The ability of magmas to transmit heat to the solidification front will depend on the thermophysical and transport properties of the magma. The two properties that vary the most, and also influence the heat and mass transfer the most, are the viscosity and the

22

Figure 4. Viscosities of some common igneous rocks. From McBimey and Murase (1984).

500

1000

1500

Temperature, °C density. Both will vary strongly as crystallization proceeds and they generally vary in an opposing manner: as the density of the liquid goes down, the viscosity of the melt goes up, although the presence of volatiles complicates this simple picture. Volatiles can depolymerize the melt structure and induce non-monotonic changes in melt density as cooling and crystallization proceeds. Before we begin a discussion on physical properties, the reader should be mindful of what the viscosity and density in the cited works refers to: the property of the melt, or the ensemble melt plus crystals. General reviews of the viscosity of magmas can be found in McBirney and Murase (1984), Ryan and Blevins (1987), and Ryerson et al. (1988). The work of Lange and Carmichael (1990) gives the most recent formulations for computing melt densities. One useful summary of the thermophysical properties, conductivity, specific heat, etc., is the compilation in Touloukian et al. (1981). Viscosity in silicate melts is a strong function of crystallinity, temperature and composition (Fig. 4). For example, the viscosity for basaltic compositions may vary three orders of magnitude over three hundred degrees, for rhyolites (dry) it may be nine orders of magnitude over six hundred degrees. It has been observed that the strong dependence of viscosity on temperature, for a fixed composition, can be given by an Arrhenius type of relation:

where r\ is the shear viscosity (the Greek letter ¡j. is also commonly used), T]0 is a constant, E the activation energy, R the gas constant and T the absolute temperature. In practice, the values of i)0 and E are not known for many compositions and alternative computational methods have been developed. The viscosity of a crystal-free melt (and hence compositionally invariant) liquid can be calculated from the algorithm of Bottinga and Weill (1972) at any temperature. This algorithm is based on summing the contributions to viscosity from the proportions of the partial molar oxides in the melt. We note in passing that the dynamic, or shear, viscosity is often reported in units of poise (gm cm-1 s-1) or Pascal seconds (kg m-1 s-1); the conversion is 10 poise to every 1 Pascal second (Pa s). Current

23 custom discourages the use of cgs units and the reader is encouraged to use Pa s as the proper unit of dynamic viscosity. As cooling of a magma proceeds crystals grow and their presence will influences viscosity. To evaluate the viscosity of a crystal-melt mix, a suspension, requires that the aspect ratio, volumetric concentration and relative particle size distribution of the crystals be known. Metzner (1985) reviews a number of aspects related to the viscosity of suspensions and concludes that for melts with a liquid viscosity of greater than 102 Pa s, the presence of particles does not induce an order of magnitude change in viscosity (at a fixed composition and temperature) until the volume fraction exceeds about 0.5. TTiese results are for a suspension of uniform spheres. Experiments on picritic compositions by Ryerson et al. (1988) reveal that the viscosity of the suspension was independent of the crystallinity up to about 0.25 volume fraction. Marsh (1981) argues that a volume solid fraction of about 0.5 represents a Theological locking up point or critical melt fraction, an idea we develop in more detail below. Metzner (1985) reviews a number of functional forms for suspension viscosity and proposes the following expression: (2) where rjp is the relative viscosity which is the ratio of the crystal-free viscosity at a given temperature and composition, to the actual suspension viscosity,

2 at values greater than ~0.05. The solubility also decreases with increasing temperature, a relation called retrograde solubility, although there is some evidence in the study of Walther and Long (1986) that at 300 MPa, that the solubility increases with temperature. The increase in solubility with small increases in xco 2 have been observed by Ellis (1959), Holland and Borcsik (1965), and Sharp and Kennedy (1965), and Sharp and Kennedy (1965) recognized the maximum in the solubility with increasing xcoj. The initial increase appears to be the result of the decrease in the pH with increasing pco^. but the decrease in solubility with greater pco2 seems to result from the decrease in die activity of H2O. Fein and Walther (1987) indicate that the solubility of calcite in supercritical fluid is consistent with the principal ions Ca+2 and HCO3". The maximum solubility of calcite at 200 MPa, 500°C in CO2-H2O is approximately 25 ppm. The major competing influences on calcite solubility are pressure and temperature. In contact metamorphic aureoles, where the pressure is nearly constant, the decrease in solubility with increasing temperature dominates. In fact, in carbonate aureoles in which aqueous fluids were flowing up-temperature, calcite veins ought to be common where the heating fluid became supersaturated with calcite. Graphite-H2Q Organic matter is a common constituent of shales and limestones. The poorly crystallized organic matter recrystallizes to graphite during metamorphism. Although

71 Solubility of Calcite 1

1

1

„0

1

3CO M P s -

°c

— / / > -

% o ®

u iO O J

200 /j

°c //

J*

—4.0

0

t

Figure 15. The solubility of calcite in H2O-CO2 supercritical fluids, compiled from experimental data of Walther and Long (1986) and Fein and Walther (1987). All experimental data indicate increasing solubility with increasing pressure, decreasing solubility with increasing temperature, and an isothermal-isobaric maximum in solubility at xco2 ~ 0.03.

ICO M P »

i -4.4

MPa

- § 1

1

O.OO

I

0.05

0.10

0.15

Mole Fraction CO 2

Table IX. log I^qf of Commonly Ocurring Fluid-Phase Components, 0.1 M P a

CH4

T (kelvins)

CO

H20(real)

C02

H20(ideal)

H2S

600

1.993

14.32

34.404

18.632

18.631

3.691

700

0.943

12.948

29.505

15.582

15.582

3.298

0.138

2.984

11.916

25.829

13.287

13.287

-0.5

11.109

22.969

11.497

11.496

2.662

1000

-1.018

10.461

20.679

10.06

10.06

2.141

1100

-1.447

9.928

18.805

800 900

8.882

8.881

Data from < COH Equilibrium Constants; Sources of Data Cited in Text. KH-P = - ^ r fe fo2 K c o

= fo2

2

l

log KH2O = ^ f 2 - - 0.979log (7>0.483 1

l o g

Kco =

KCO2 =

log Kco

+0.0421 +0.0276 f ^ i \ T

T

=^

j l KcHi = ^ log Kch, = ^ fä T f. _ T in kelvins, p in bars

+ 4.5771 + 0.0276 - l.ßlllog

(T) + 2.4461 + 0.0276

\ T

Oxygen Fugacity for Some Common Oxygen Buffers Data from Myers and Eugster (1983 ) and Robie et al. (1967) ^

Fe -Fex_xO

to O2

Fei-xO

log fo2 =

-Fe304

FezOi - Fe2Os Si02 • Fe - Fe2SiOi

fogfa

=

=

log fo2 =

z26|34:7 36

+ 6

.471

+

0

.0517(P_l)

' ^ 5 1 ' 2 + 16.092 + 0.0878 i ^ Ü

-23.847.6

+ 1 3 4 8 0 +

Q 0 1 8 5

— ' 7 "il7 z, " ° + 6.396 + 0.0497 ( P - l ) T

_ 94 441 Q

Si02 - Fe3Oi - Fe2SiOi log fo2 = T in kelvins, p in bars

(p-l)

Ç,

+ 8.290 + 0.0937

(P-l)

1.713

72 crystalline graphite is typically found in high-grade metamorphic rocks, whereas "graphite" in low- to medium-grade metamorphic rocks retains much of its organic heritage (Dunn and Valley, 1987; Nabelek et al., 1991), thermodynamic calculations involving C-O-H fluids are made with graphite-H-0 equilibria. The first considerations of the consequences of reactions between graphite and H2O were made by French (1966) and Eugster and Skippen (1967) and were developed from the need to control the oxygen fugacity in hydrothermal experiments on the stabilities of iron-bearing minerals, for example by Eugster and Wones (1962). Several others have since calculated the compositions of fluids in equilibrium with graphite under a variety of circumstances, particularly as the relations between activities and partial pressures of the fluid components became better understood. The more recent works include those of Holloway (1977), Holloway (1981), Ferry and Baumgartner (1987), and Ulmer and Luth (1991). The composition of a fluid in equilibrium with graphite is governed by the equilibrium constants for the various homogeneous equilibria in the fluid. The most abundant species in the equilibrium fluid under metamorphic conditions are invariably H2O, CO2, and CH4, with a lesser amounts of H2 and CO, an extremely small amount of O2, and negligible amounts of higher hydrocarbons. At any temperature and pressure, the composition of the fluid is given by the equilibrium constants for the following reactions: H 2 + V2 0 2 = H 2 0 C + o 2 = C02 C + 1/2 0 2 = CO C + 2 H 2 = CH4. Except for C, all the components occur in the fluid phase. The equilibrium constants of formation for H2O, CO2, CO, and CH4 at temperatures ranging from 600 to 1100 K, taken from Chase et al. (1985), are given in Table 11. Expressions for the dependances of the equilibrium constants on temperature and pressure, tabulated by Ohmoto and Kerrick (1977) and taken from Huebner (1971) and Holland (1965) are also given in Table 11. The oxidation states of rocks and fluids are commonly referenced against a set of redox equilibria in the system Fe-Si-O. At a specific temperature and pressure, equilibrium among quartz + magnetite + fayalite, for example, fixes the value of fo2. The equilibrium constants for solid oxygen buffers have been determined over the range of experimental conditions by several workers. The values of log fo2 determined for the iron-wustite, wiistite-magnetite, magnetite-hematite, quartz-iron-fayalite, and quartz-magnetite-fayalite buffers from Myers and Eugster (1983) and O'Neill (1987), recently determined at temperatures greater than those for most contact-metamorphic aureoles, are given in Table 11. At the temperatures of contact metamorphism, < 750°C, the equilibrium between iron and magnetite is more stable than wiistite. The log / o , expression for iron-magnetite can be determined by the linear combination 0.25 log fo2 (wM) + 0.75 log / o , (IW) = log fo0 (IM). The system of COH equations, above, has two degrees of freedom; that is, there are four equations and six unknowns — the fugacities of H2, O2, H2O, CO2, CO, and CH4. Two additional constraints must be imposed to solve for the composition of the fluid. The first is to assume that the total fluid pressure is equal to the solid pressure. Then the sum of the partial pressures is constrained to equal the total pressure. This is a very reasonable assumption, except in cases where metamorphism is fluid-absent or hydrostatic. The totalpressure constraint is expressed as

73 _ fezO XH2O

+

fcch , feo XCOI Xco

+

f

H2 | foi , fcff4 XO2 XcHi

XH2

in which the c's are the fugacity coefficients. If the fluid components mix non-ideally, then the c's depend on composition, which complicates an already complicated set of non-linear equations. In the previous section, the mixing properties of supercritical fluids were shown to have small deviations from ideality. Thus, ideal mixing will be assumed as a close approximation to the real fluid. This assumption is certainly invalid for those conditions under which there are two immiscible fluids, but in the NaCl-free system, contact metamorphism occurs in the supercritical fluid field. One additional constraint is needed to solve the system of equations. The oxygen partial pressure can be specified as fixed by the environment. For example, the value of fo2 might be fixed by another oxygen buffer, such as FMQ. An example of this assumption is shown in Figure 16a, in which the fluid composition is shown for fo2 fixed at FMQ and 2 log units below fo2 at a pressure of 200 MPa. The system of equations given above is quadratic in fn2 at fixed fo2, but one of the two solutions for / h 2 is negative. The characteristics of the realistic solution for the fluid composition are CH^rich fluid at low temperature, H20-rich fluid at intermediate temperature, and CC>2-rich fluid at high temperature. H2 and CO occur in significant abundance only at temperatures greater than ~800°C. The upper stability of graphite at an oxygen fugacity of FMQ at 200 MPa is approximately 475°C. The stability field is extended to greater temperatures at lower oxygen fugacities, exceeding 600°C at FMQ - 2. Few rocks, except perhaps some meta-iron-formations (Vaniman et al., 1980), buffer the oxygen fugacity along one of the Fe-Si-0 buffers. A different constraint that can be imposed on the system of equations is that of a fixed ratio of H/O. For example, equilibrium between C and H2O in a closed system fixes the ratio of H2 to O2 at 2. This constraint mimics the limiting behavior of fluid/rock interaction of the closed, rockdominated system. The equations reduce to two equations in two unknowns, quadratic in / h 2 and in fo2, with mixed terms. These two equations can only be solved numerically. Because the equilibrium constants strongly favor CO2 and CH4 over CO and H2, the closed system ensures that the abundance of CO2 and CH4 are approximately equal. Figure 16b shows the oxygen fugacity and the abundances of the major constituents of the fluid at 300 MPa for the closed system. H2O dominates the fluid and increases in relative abundance with increasing pressure, but decreases with increasing temperature. The oxygen fugacity is lower than FMQ, except at temperatures less than 375°C, and the buffer curve nearly parallels that for the upper stability limit of graphite. The entire graphite saturation surface in the system COH can be calculated by varying the H2/(H2+02) ratio from 0 to 1. The surface, calculated at 500°C and 200 MPa, is shown in the inset of Figure 16a. Under these conditions, the saturation surface nearly coincides with the binaries CO2-H2O and H2O-CH4. The calculated results shown in Figure 16 were obtained by assuming ideal mixing in the fluid. Ferry and Baumgartner (1987) have compared the ideal-mixing results with those assuming fluid mixing rules associated with the MRK equations of state of Bowers and Helgeson (1983b). The effect of non-ideal mixing is depression of the saturation surface closer to the binaries. In the range 400° to 500°C, 200 to 300 MPa, the saturation surface coincides, for all intents and purposes, with the binaries. The implication is that, if fluid constituents were exclusively COH, graphite-bearing contact-metamorphic rocks were in equilibrium with either a CO2H2O or an H2O-CH4 fluid. Graphitic pelitic hornfelses The graphite equilibria presented in the previous section require an additional constraint

74 c

Temperature (°C)

350

400

450

500

550

600

Figure 16. COH fluid compositions in equilibrium with graphite. Top. Abundances of H2O, CO2, and CH4 at two oxygen fugacities, QMF and two log units below QMF. CH4 is abundant at low temperature, H2O is abundant at intermediate temperature, and CO2 is abundant at high temperature. Top inset shows the graphite saturation surface at 500°C, 200 MPa, calculated assuming ideal fluid mixing. Except along the graphite-H20 join, the saturation surface essentially coincides with the H2O-CO2 and H2O-CH4 binaries. Bottom. GraphiteH2O equilibrium in a closed system in log / o j - T space. Bottom inset. Abundances of the major fluid species at 150 and 300 MPa in the closed graphite-H20 system. CO2 and CH4 are in almost equal abundance.

75 to calculate the fluid composition. Dehydration equilibria in pelitic rocks in contact metamorphic aureoles provide that constraint because the equilibria fix /H2O at constant temperature, pressure, and compositions of the reactant and product minerals. For example, a rock containing the assemblage muscovite + quartz + andalusite + K-feldspar + graphite fixes the /H2O by the equilibrium muscovite + quartz = andalusite + K-feldspar + H2O. The abundances of the other COH species in the fluid are determined by the graphite equilibria in the same sense that an externally imposed oxygen buffer does. Ohmoto and Kemck (1977) calculated the effects of dehydration and decarbonation equilibria on COHS fluid compositions. The examples I present here exclude sulfur-bearing components. Ohmoto and Kerrick (1977), Holloway (1981), and Ferry and Baumgartner (1987), and Poulson and Ohmoto (1989) describe the equilibria involving H2S, SO2, S2, COS, and Ss species and iron sulfides and silicates. The calculations described here are easily extended to include sulfidation equilibria. The system of COH equilibria under the constraint of constant /H2O reduces to a quartic equation in / h 2 with two real solutions. Because one solution is more oxidized than the other, independent information on the oxidation state of the rock is required to determine which is the correct solution. The systematics of dehydration equilibria are illustrated by the following multivariate equilibria: 6 muscovite + 2 phlogopite + 1 5 quartz = 3 cordierite + 8 K-feldspar + 8 H2O, which occurred in the low-pressure aureole of the Duluth complex (Labotka et al., 1981), clinochlore + muscovite + 2 quartz = cordierite + phlogopite + 4 H2O, which occurred in the 300-MPa Ballachulish aureole (Pattison, 1989) and is common in many contact-metamorphosed and low-pressure regionally metamorphosed pelitic schists, and the muscovite + quartz reaction previously mentioned. All three reactions are continuous reactions, even in C-free systems, because Fe and Mg are partitioned among the mafic minerals and because Na and K are partitioned between muscovite and K-feldspar. The effects of partitioning will be examined with the first reaction, after the end-member equilibria with graphite are considered. The three reactions that are used as an example of metamorphosed graphitic, quartzbearing pelitic schist are members of a set of 13 possible reactions in the system K2O AI2O3 - MgO - Si02 - H2O. The three reactions under consideration, labeled by the absent phases, (and, chl), (and, ksp), and (crd, chl, phi), are trivariant and will intersect at invariant points under constant pressure conditions on diagrams depicting temperature and fluid composition. Figure 17A shows the oxygen fugacity over the temperature range 350° to 600°C at 150 MPa for this system. This is the kind of diagram used by Ohmoto and Kerrick (1977). Each equilibrium occurs along a hyperbolic-like curve that is asymptotic to the upper stability of graphite. At any temperature, the equilibrium occurs at two fo2 values. The higher value is that nearly coincident with the value for graphite-C02. The other value corresponds to values for H2O-CH4 fluids. The hyperbolic curves outline regions that are more reducing than FMQ for most of the temperature range under consideration. Inside each curve, the more hydrous reactants are stable; outside, the dehydrated products are stable. This relation reflects the increased abundance of CO2 in the fluid phase with increasing oxygen fugacity and the increased dissociation of H2O with decreasing oxygen fugacity, both of which cause a decrease in the activity of H2O. Two invariant points of the system can be seen in Figure 17A. There are two other stable invariant points, but these occur at low / o 2 and low temperature and are off the figure. Each invariant point also has a mirror image on the high fo2 sides of the curves, but these are compressed along the graphite stability curve and can't be seen. The compositions of the fluids, not shown in the figures, are essentially binary mixtures of H2O with either CO2 or CH4. In the presence of graphite, the fluid composition is never pure H2O. The maximum fraction of H2O, occurring at the highest temperature, ranges from about 0.8 at 500°C to 0.69 at 600°C. The values are the same as those on the graphite-H20 join, Figure 16.

76

A

p = 150 M P a

-20

-

-30

-

muscovite X - - A + biotite / \ A

9» «5 bs

o

(crd.chl)^" (mu.chl)

-40 300

K-feldspar

400

phlogopite

500

600

Temperature (°C)

0.00

1.00

HaO Figure 17. Pelitic equilibria in graphite-bearing systems. A . log f o ~ - T section at 150 MPa showing the reactions in the system K M A S H C with excess Si02. Reaction curves are labeled by the absent phases in parentheses, and invariant points are labeled by the absent phase in brackets. Also shown are the graphite-C02 equilibrium and the standard oxygen buffers IM, F M Q , and M H . B. Projection of the same pelitic equilibria as in A on a T - (1 xh 2 o) plane. The equilibria terminate at the maximum stability of graphite, indicated as 'graphite + H20.'

77 Because the fluids are essentially binary mixtures of H2O and either CO2 or CH4, temperature-fluid composition relations are better displayed in Figure 17B, a T - ( 1 - X H 2 O ) diagram. An alternative representation of phase equilibria in graphite-bearing systems is given by Frost (1979a, b) in which the stabilities of mineral assemblages are depicted on isothermal-isobaric /o 2 -x c diagrams. For those rocks that are relatively oxidized, Figure 17B is essentially a T-xco^ diagram for graphite-bearing systems. The reactions also intersect in a plexus of invariant points on this diagram, all at temperatures less than 400°C and X H , O less than 0 . 1 . The equilibria considered here span nearly the entire range in X H 2 O in fluia composition, except values of X H 2 O greater than the stability limit of graphite. Even in the absence of graphite, metamorphic fluids will be binary H2O-CO2 or H2O-CH4 because the graphite saturation surface is nearly coincident with the binaries. Therefore, assemblages like cordierite + K-feldspar + muscovite + biotite + quartz and quartz + muscovite + cordierite + chlorite + biotite, and quartz + muscovite + andalusite + Kfeldspar, all of which are common assemblages in pelitic hornfelses, ought to be excellent monitors of fluid composition during metamorphism. The natural pelitic system contains iron and sodium, in addition to K20-Mg0-Al203Si02-H20. The additional components increase the variances of the equilibria, and the univariant curves in Figure 17 are replaced by divariant fields. The effects of the additional components are to reduce, but not eliminate, the capacity of the mineral assemblage to buffer the fluid composition and to increase the range in rock compositions that can possess the assemblages. These effects are considered in a following section. The effect of the additional components on the fluid composition is illustrated in Figure 18 for the quartz + muscovite + biotite + cordierite + K-feldspar equilibrium. Calculation of the equilibria involves no more than a consideration of the reduced activity of the Mg end-members in the equilibrium constant. The equilibrium constant for the reaction (and, chl) is given by

jr

_

\

mgi

^ JV ;

- r.

If the relations between activity and composition, temperature, and pressure are known, then the equilibrium constant can be evaluated at constant temperature and pressure. The remainder of the solution of the fluid composition follows as before. This process is greatly simplified if biotite and cordierite have constant distribution coefficients. In this case, the crystalline solutions are ideal, and, for example,

WHO2 •

The Fe-Mg distribution coefficients among chlorite, biotite, and cordierite are near 1.0, and even though KQ is a function of temperature, the value of K D changes very little for these minerals over the temperature range under consideration (Labotka et al., 1981; Labotka, 1981; Pattison, 1987). Natural biotites contain a tschermak component, which also reduces the activity of Mg in the octahedral sites. Individual applications can take the extra component into account by setting xM„ = Mg/3. This, though is not considered in the example here. The partitioning of Na ana K between K-feldspar and muscovite is a strong function of temperature and composition because these minerals are non-ideal solutions (e.g., Eugster et al., 1972; Waldbaum and Thompson, 1969; Green and Usdansky, 1986). These minerals are K-rich, rarely containing less than 0.80 K/(K + Na) (Guidotti, 1984), and for sake of illustration, they will be assumed to have end-member compositions. With these assumptions, the equilibrium constant becomes

78

-10 300

400

500

600

Temperature (°C) 600

B 500

O u a a u 0)

400 ï o H p = 150 M P a _L 0.40

_L 0.20

300 0.00

1 . 0 - X

I 0.60 H 2

J

L.

0.80

1.00

O

Figure 18. The effects of variable Fe-Mg compositions of cordierite and biotite on the log fçu - T and T - (1 XH2O) projections of the reaction 6 muscovite + 2 phlogopite + 15 quartz = 3 cordierite + 8 K-feldspar + 8 H2O. A constant value of K d (Mg/Fe, Crd/Bt) = 1.5 was assumed.

lycrdf

f8 - ( Ko

= Mgi

(

Kd is the ratio

O

(L -

(! which is assumed to be 1.5 in Figure 18.

+ ( 1 •" K o )

f%

2 0

79 The addition of iron to cordierite and biotite increases the stability of the reactants biotite + muscovite + quartz. The Fe-Mg solution makes almost no difference in the fo2 conditions of the reaction, as seen in Figure 18A, and increases the temperature of the reaction only by about 20°C at the upper stability limit of graphite (Fig. 18B). The small effect is the result of the small distribution coefficient between biotite and cordierite. The effect of Fe-Mg partitioning on the temperature of the other mafic reaction (and, ksp) is even less because the K D for chlorite-cordierite is closer to one, -1.3 (Guidotti et al., 1975; Labotka, 1981), than that for biotite-cordierite and because the equilibrium constant is

W+MSK)3 The maximum K ^ i s -1.6 x / ¿ 2 0

at xfij^ = 0 ,

a very small difference from the value at x ^

= 1.

Figures 15 and 16 show that pelitic assemblages containing graphite formed under conditions of reduced an2o- The reduction in activity resulted in the formation dehydrated assemblages at considerably lower temperatures than in graphite-free rocks. Pattison (1989) showed that andalusite-bearing assemblages were developed at the expense of cordierite + K-feldspar in graphitic rocks in the Ballachulish aureole. Andalusite + biotite assemblages in graphitic rocks in the garnet zone of the low-pressure metamorphic terrain in the Black Hills, South Dakota, were also attributed by Helms and Labotka (1991) to a reduced activity of H2O. In the aureole of the Duluth complex, the assemblage andalusite + K-feldspar + muscovite + cordierite + phlogopite occurs in Mg-rich graphitic hornfelses (Labotka et al., 1981). A few calculated values of X H 2 O in graphitic pelitic hornfelses include a value of 0.55-0.75 at 544°C for the Duluth complex aureole, and Pattison (1989) estimated a value of 0.79 in graphitic rocks of the Ballachulish aureole. Calculated values of X H 2 O from regionally metamorphosed graphitic rocks include 0.90 in kyanite-zone rocks from British Columbia (Ghent, 1975), 0.85 in sillimanite-zone rocks from Maine (Guidotti, 1970), and 0.44 to 0.51 in staurolite + andalusite-zone rocks from Maine (Ferry, 1981). The values of 0.79, 0.85, and 0.9 are those expected from closed system graphiteH2O; whereas, the lower values require an external source of CO2 or CH4. Graphitic limestones The occurrence of graphitic pelitic hornfelses that record X H 2 O values other than those that can be produced in the system C-H2O require an external source of carbonic fluids. A source for such fluids is metamorphosed sedimentary carbonate rock. Reactions of calcite and dolomite with silicate minerals provide a ready supply of CO2, which, in turn, mixes with the pore fluid and potentially reduces X H 2 O - Graphite-bearing marbles can fix the composition of the fluid phase in the same manner that the pelitic rocks can. The combination of a decarbonation equilibrium and the graphite-COH fluid reactions at constant temperature and pressure determines the fluid composition. Ohmoto and Kerrick (1977) included decarbonation reactions in their analysis of graphitic systems. Here I will illustrate the method of determining the fluid composition, which is really no different from that already described, with the assemblage calcite + quartz + wollastonite + graphite. The reaction calcite + quartz = wollastonite + CO2 fixes fco2 a t the valued specified by the equilibrium constant for this reaction, which was deteimined from the database of Holland

80

Temperature (°C) Figure 19. Abundances of the major constituents of a COH fluid in equilibrium with calcite + quartz + wollastonite + graphite at a pressure of 150 MPa.

x

co 2

Figure 20. Schematic T - x C o 2 diagram, taken from Greenwood (1967) and Kerrick (1974), showing the slopes of the various types of mixed-volatile equilibria.

81 and Powell (1990). Figure 19 shows the abundances of the major constituents of the fluid in equilibrium with the assemblage at 150 MPa over the temperature range 350° to 600°C. Because the decarbonation reaction, under the conditions pco 2 = Psolid. occurs at a high temperature, ~680°C at 150 MPa, the fluid in the temperature range of Figure 19 has low values of xco2- The composition is dominated by H2O, which shows a maximum value of about 0.8 at 475°C. Below ~510°C, the carbonic fraction of the fluid is mostly CH4; above 510°C, the carbon occurs dominantly as CO2. Above 550°C, CO2 increases significantly as both CH4 and H2O decrease. H2 generally makes up less than 1% of the fluid, and the oxygen fugacity is about 1.5 log units below FMQ. Mixed volatile equilibria in CCk-HUO systems Many, if not most, carbonate and pelitic rocks in metamorphic aureoles lack graphite, particularly at high grades. Graphite could occur at high temperatures, but it commonly has been exhausted by reaction with the pore fluid. This happened in the Notch Peak aureole in carbonate rocks and in the aureole of the Duluth complex in argillaceous rocks. Because the graphite saturation surface nearly coincides with the CO2-H2O binary in relatively oxidized environments, equilibria in the binary CO2-H2O fluid system will be considered. The thermodynamics and phase equilibria in mixed-volatile systems have been investigated by numerous workers since Greenwood (1961) and Greenwood (1962). Kerrick (1974) reviewed the systematics of mixed-volatile equilibria pertinent to metamorphism of carbonate rocks, including some of the theoretical work of Greenwood (1967) and experimental work of Slaughter et al. (1975), Johannes (1969), and many others to that time. There have been numerous experimental investigations of reactions in the siliceous dolomite system Ca0-Mg0-C02-H20, including Eggert and Kerrick (1981), Gordon and Greenwood (1970), Jacobs and Kerrick (1981b), Käse and Metz (1980), Metz (1967), Metz (1970), Metz (1976), Metz and Puhan (1970), Skippen (1971), and Skippen (1974). There have also been several experimental investigations of reactions pertinent to the metamorphism of micaceous limestones and calcareous argillites, such as Allen and Fawcett (1982), Gordon and Greenwood (1971), Hewitt (1973b), Hewitt (1975), Hochella et al. (1982), Hoschek (1973), Liou (1973), Puhan (1978), and Storre and Nitsch (1972). Theoretical calculations of mixed-volatile equilibria have been presented by Greenwood (1967), Kerrick and Slaughter (1976), Skippen and Carmichael (1977), Flowers and Helgeson (1983), Trommsdorff and Connolly (1990), and others. The types of mineral-fluid equilibria that occur in CO2-H2O fluid systems are outlined in Figure 20, taken from Greenwood (1967) and Kerrick (1974), which is a constant pressure T-xccb diagram. The derivations of the shapes and slopes of the curves on the diagram can be found in their papers. There are six general cases. I. Decarbonation reaction curves, represented by calcite + quartz = wollastonite + CO2, have positive slopes and negative curvature. II. Dehydration reaction curves, exemplified by 4 zoisite + quartz = 5 anorthite + grossular + 2 H2O, have negative slopes and curvature. ni.Dehydration-decarbonation reaction curves, illustrated by tremolite + 3 calcite + 2 quartz = 5 diopside + 3 CO2 + H2O, have a maximum at the stoichiometric fluid composition. IV Carbonation-dehydration reaction curves, such as 2 zoisite + CO2 = 3 anorthite + calcite + H2O, have negative slopes and inflection points. V. Hydration-decarbonation reaction curves, for example, 3 dolomite + 4 quartz + H2O = talc + 3 calcite + 3 CO2, have positive slopes and inflection points.

82 VI Solid-solid reactions, such as grossular + 2 quartz = 2 wollastonite + anorthite, are isothermal lines. The numerous experimental studies of reactions that occur in metamoiphosed carbonate systems have provided a relatively well determined petrogenetic grid and well characterized thermodynamic properties for many calc-silicate minerals (Skippen, 1974). There remain, though, some uncertainties about the relative stability of magnesite-bearing assemblages in metamorphosed calcareous ultramafic rocks (Trommsdorff and Connolly, 1990). Figure 21 shows some of the reactions that can occur in contact metamorphosed siliceous dolomites. The numerous reactions that can occur among phases in the system Ca0-Mg0-Si02-CC>2-H20 can be calculated realtively painlessly by computer programs such as GeO Calc by Berman et al. (1987). Most reactions on Figure 21 have low positive slopes over much of the range in xco?- Only at values of x c q j less than 0.1 or greater than 0.9 is there a great temperature dependence. The lowest-temperature reactions that occur in siliceous dolomites are those that form talc and tremolite-bearing assemblages. The reactions intersect at an invariant point, and talc-forming reactions occur on the low-xco 2 side of the invariant point. Talc forms in rocks that are in equilibrium with water-rich fluids only. This restriction is even greater at high pressure because the position of the invariant point shifts from about xco 2 = 0.65 at 200 MPa to xco 2 = 0.20 at 800 MPa (Eggert and Kerrick, 1981). The value in Figure 21 calculated by Souza (1991) is -0.55 at 250 MPa. Some reactions occur only in the presence of CC>2-rich fluids. Two shown on Figure 21 are dolomite + 2 quartz = diopside + 2 CO2 and 3 dolomite + diopside = 4 calcite + 2 forsterite + 2 CO2. Assemblages related by these reactions, particularly dolomite + diopside, are found in some high-grade contact-metamorphosed dolomites, e.g., Trommsdorff (1966). Skippen and Trommsdorff (1975) argue that occurrences of this assemblage in metamorphosed ultramafic rocks implies that a second invariant point involving dolomite, diopside, calcite, tremolite, and forsterite also occurs in the H20-rich region of the diagram. This invariant point, though, must be at least metastable with respect to antigorite-bearing assemblages (Evans et al., 1976). T

_j

0.1

1

0.2

1

0.3

1

0.4

1

x0.5

co 2

1

0.6

1

0.7

1

0.8

1—

0.9

Figure 21. The 250 MPa T-xco 2 section of common equilibria in contact-metamorphosed siliceous dolomites, taken from Souza (1991). The few equilibria involving phlogopite and K-feldspar are shown as dashed lines. These equilibria can also be seen in Figure 22. The unlabeled reaction on the C02-rich side of invariant point II is dolomite + quartz = diopside + CO2.

83 Reactions involving magnesite can affect the sequence of assemblages observed in metamorphosed carbonate rocks, even though the reactions themselves can only be recorded in ultramafic rocks. Some recent calculations by Chernosky et al. (1988) and Evans and Guggenheim (1988) indicate a large field of stability for the assemblage tremolite + magnesite at the expense of the assemblage talc + forsterite + dolomite. Trommsdorff and Evans (1972), Trommsdorff and Evans (1977), and Trommsdorff and Connolly (1990) provide field evidence and phase-equilibrium calculations that indicate an extremely limited stability of tremolite + magnesite in CC>2-rich fluids. Enstatite-bearing assemblages also have stabilities restricted to CC>2-rich fluids, but these are generally found only in metamorphosed ultramafic rocks. In contrast to the few types of assemblages restricted to CC>2-rich fluids, there are numerous assemblages in metamorphosed carbonate rocks that are stable only in H20-rich fluids. As illustrated in Figure 21, these include talc + calcite and wollastonite. Wollastonite is stable over the entire range in xco2> but only at temperatures exceeding 600°C. In many contact-metamorphic aureoles, the temperature rarely exceeded 600°C, and wollastonite is restricted to rocks in equilibrium with fluids having xco2< ~0.2. Many other assemblages stable in water-rich fluids occur in aluminous carbonate rocks. Some of these are shown in Figure 22. The reactions shown in Figure 22 involve the minerals anorthite, zoisite, vesuvianite, grossular, phlogopite, and K-feldspar, in addition

Figure 22. H20-rich portion of the 200 MPa T-xcoj section, showing equilibria that occur in calcareous argillites metamorphosed in the presence of water-rich fluids, taken from Labotka et al. (1988a).

84 to those shown in Figure 21. The reactions involving vesuvianite are schematic, drawn by Labotka et al. (1988a) to be consistent with experimental work of Hochella et al. (1982). The topology of the reactions is somewhat different if the thermodynamic database of Holland and Powell (1990) is used to calculate vesuvianite reactions, containing the stable reaction calcite + anorthite + diopside + quartz + H2O = vesuvianite + CO2 and making the analogous grossular-forming reaction metastable in diopside-bearing rocks. Although the details of vesuvianite-forming reactions are yet to be determined, all work indicates that vesuvianite-bearing assemblages are restricted to H20-rich environments (Rice, 1983; Valley et al., 1985; Labotka et al., 1988). There are several other hallmarks of an H20-rich fluid that can be noted on Figure 22. Zoisite-bearing and grossular-bearing assemblages are indicative of water-rich fluid. As noted in the previous paragraph, wollastonite in contact-metamorphic rocks generally indicates a water-rich environment. The transition from phlogopite + calcite + quartz to diopside + K-feldspar also occurs in water-rich fluids at low temperature. The stabilities of many epidote-bearing mineral assemblages in micaceous marbles and calcareous argillites seem to be restricted to H20-rich fluids. Epidote + quartz + muscovite + calcite + chlorite and epidote + quartz + muscovite + calcite + phlogopite are common in low-pressure regional metamorphic terrains in the Panamint Mountains, California, (Labotka, 1981), in southeastern Maine (Ferry, 1976), and south-central Connecticut, although epidote is absent from most rocks (Hewitt, 1973a). Reactions among these minerals and tremolite and K-feldspar are crowded into the H20-rich portion of a T-x diagram. The topology of the reactions and their related invariant points is given by the Hcc>2 - M-H2O diagram, Figure 23. The diagram is qualitative because chemical potentials are always measured relative to a standard state, although the diagram is essentially the same as a quantitative In fco2 - fiho diagram. Despite this disadvantage, the |ico 2 - M-H2O diagram has the advantages of expanding the low-xco 2 region of the T-x diagram, and of having reaction slopes that are easy to calculate. The slopes can be derived from the GibbsDuhem equations for the minerals participating in the reactions. For each reactant and product phase the Gibbs-Duhem equation at constant T and p is

X m 4l; =0 The change in the number of moles of the components during reaction can be obtained by taking the difference between the sums of the products and the sums of the reactants. Because the only components lost or gained are the volatile components CO2 and H2O, all Awj-'s are 0 except for the volatile components. Ancch = - v c q ,

and

h2o =

~vh2o

The sum of the Gibbs-Duhem equations reduces to -VCO2 dV-CCh -VH20

d\lH20 = 0

from which the slope of the reaction is d^CQz _ _ v H 2 Q 9(J-H2O vcoz The |ico 2 - M-H2O diagram is a useful tool for exploring the possible reactions that relate assemblages in multicomponent systems for which thermodynamic data are limited or have large associated uncertainties. Natural occurrences of assemblages can be used to

85

03 O u A

* h

2

O

Figure 23. Schematic isothermal-isobaric Hcc>2 - HH2O section showing the reactions that occur in H20-rich fluids in quartz + epidote-bearing calc-argillites.

600

Plag + Qtz +•

Cal » Di » H20 4) ti s «4

= Vea * C02

500

•h

I»

H

400

0.10

0.00 Zo + Qtz 4- Cal + Oi • H20 » Vss •

coa

CO,

Figure 24. The 200 MPa T - x c o j section, showing the effects of Na and NaCl on scapolite stability in calcareous rocks. The pure carbonate scapolite is stable only at extremely high temperatures, but small amounts of marialite component expand the stability limit to contact-metamorphic temperatures. Reactions are shown for the scapolite composition meioniteo.775. Tlie compositions in the box with the arrow pointing to a reaction curve intersecting the wollastonite reaction are those appropriate for the scapolite in the Notch Peak aureole.

86 determine which of the possible invariant points are stable. The |icc>2 " M-HiO diagram in Figure 23 shows that muscovite + tremolite assemblages are stable only in relatively H2Orich fluids. Assemblages in NaCl-CCk-FkO fluids Additional components, such as Fe in the magnesian minerals, Mg in calcite, and Na in plagioclase, affect the phase equilibria presented in Figures 19 and 20 by reducing the activities of the end-member components. The effects of the additional components are to expand the stability fields of those minerals that preferentially contain the components. For example, plagioclase + calcite stability is expanded to low xcoj values relative to zoisite with increasing albite content. The resulting increase in variance makes the assemblage more common in a wide range in rock compositions. The effects on the equilibrium constant can be computed in the manner already described in conjunction with Figure 18. Rice (1977a) has shown that the effects of substitution of Fe for Mg has a minuscule effect on the T-xcc>2 conditions for siliceous dolomite equilibria. Na can have a greater effect on plagioclase stability, and NaCl stabilizes scapolite. The effects of plagioclase solid solution, calculated with the activity model of Newton et al. (1980), are the restriction of zoisite to xcc>2 values less than 0.01 at temperatures greater than 500°C at a plagioclase composition or An 40 . Zoisite stability is restricted to lower XCO2 values at higher temperatures and at lower an-contents. Scapolite is a common mineral in some contact and medium temperature regional metamorphic terrains. It is abundant in some rocks at Notch Peak, Utah, (Hover-Granath et al. (1983), at Lone Mountain, Nevada, (Richards and Labotka, 1991), at the northwest Idaho batholith (Hietanen, 1967; Mora and Valley, 1989), in the Sierra Nevada (Kerrick et al., 1973; Ferry, 1989), in southeastern Maine (Ferry, 1976), and in the Panamint Mountains, California (Labotka, 1981). The formation of the pure calcium carbonate scapolite from calcite + anorthite occurs only at temperatures exceeding 800°C (eg., Moecher and Essene, 1990); the common occurrence of scapolite must be the result of substitution of the sodium component, NaSiCa_iAl_i, or the substitution of NaCl for CaCC>3. Figure 24 shows the H20-rich portion of the 200-MPa T-xco2 diagram for the formation of scapolite at Notch Peak (Nabelek et al., 1991; Labotka, in prep.). Small dilution of the meionite component expands the scapolite stability field to temperatures less than 500°C. The equilibria for the composition x m e = 0.77 are shown. A side effect of the breakdown of the calcic scapolite component is the expansion of the stability field of vesuvianite, which may account for some of the ambiguities about the relative stabilities of vesuvianite and grossular previously noted. The potential value of natural scapolite compositions is as a monitor of the NaCl content of the coexisting fluid through the equilibrium 3 albite + NaCl (aq) = marialite. Once the relation between the activity and composition of marialite are determined, as reported by Moecher and Essene (1990) and Vanko and Bishop (1980), quantitative assessment of the NaCl contents of metamorphic fluids can be made, independent of fluid inclusion studies. The effects of NaCl on phase equilibria, particularly in the mixed volatile siliceous dolomite system, were addressed by Bowers and Helgeson (1983a). Under the conditions for the single-phase supercritical fluid, the principal effect of NaCl is the reduction of the activities of CO2 and H2O. Because NaCl is sparingly soluble in CO2, the region of immiscibility is greatly expanded by adding NaCl to a CO2-H2O fluid. Fluids with a salinity of 7.15% (wt % NaCl/(NaCl + H2O)) increases the solvus to temperatures in excess of 500°C. Figure 25 shows this solvus on the 200 MPa T-xco 2 diagram. At this pressure and salinity, the crest of the solvus lies at ~530°C. All reactions involving talc + tremolite + calcite + dolomite + quartz are intersected by the solvus. If Figure 25 A were a true two-component phase diagram, each reaction that intersects the solvus has a corresponding invariant point on the other side of the solvus connected by an isothermal tie-line. Bulk fluid compositions within the solvus at the temperature of the tie-line consists

87

0.0

0.2

0.4

0.6

0.8

1.0

NaCl0.o7i5 H2O0.9285

X C

° 2

n c o

2

Figure 25. The 200 MPa T-xcoi section, showing the effects of NaCl on immiscibility in CO2-H2O fluids after Bowers and Helgeson (1983a). The diagram A is a section from CO2 to NaClo.o7i5 H20o.9285- The solvus intersects several equilibria in the system CMASHC, including the commonly occurring reactions about the invariant point I from Figure 21. B. shows the intersections of these equilibria with the solvus at 200 MPa. Isotherms on this diagram represent the isothermal projection of the solvus on the CO2-H2O side of the NaClCO2-H2O triangle. Ncojis the ratio CC>2/(C02 + H2O).

88 of a liquid and a gas phase in equilibrium with the reactants and products. The intersections on Figure 25A cannot be connected by isotherms, though, because the compositions of the liquid and gas lie outside the plane of the diagram. That is, the section is not parallel to the tie-lines in the three-component fluid system. The C02-rich phase is less dense than the FkO-NaCl phase, and H20-rich fluid compositions could be produced from immiscibility and separation of the gas phase. The intersections of the reactions and the solvus at 200 MPa are shown in Figure 25B. All reactions that intersect the solvus have temperature minima where the reaction curves become tangent to the crest of the solvus. These are dehydration-decarbonation, hydrationdecarbonation, and decarbonation reactions — all the types shown on Figure 25B. Dehydration and carbonation-dehydration reactions will show temperature maxima where the reaction curve crosses the critical curve. Small changes in the salinity of bulk fluid compositions in the two-phase region can have a major effect on the mineral assemblage. For example, at a temperature of ~460°C, just below that of the invariant point I in Figure 25B, mineral assemblages stable along the solvus are tremolite + calcite near Nco 2 (= CO2/CO2+H2O) = 0.25, talc + calcite + tremolite at N C o 2 between -0.30 and 0.35 and between -0.20 and 0.18, talc + calcite + quartz at Nco 2 between -0.35 and 0.38 and between -0.18 and 0.15, and dolomite + quartz at Nco 2 > 0.38 and Nco ? < 0.15. Small changes in the amount of NaCl at constant Nco 2 . or small changes in Nco 2 at constant x NaCi can cause a change in the mineral assemblage. A 200 MPa isothermal section of the NaCl-COo-F^O system is shown in Figure 26 for the temperatures of 400 and 500°C, modified fromLabotka et al. (1988a). Three reactions are shown, illustrating the relations between the solvus and the reaction surfaces. Two reactions, calcite + quartz = wollastonite + CO2 and tremolite + 3 calcite + 2 quartz = 5 diopside + 3 CO2 + H2O, cross the critical point on the solvus, but the temperature is lower than that of the intersection. The third reaction, 5 dolomite + 8 quartz + H2O = tremolite + 3 calcite + 7 CO2 intersects the 500°C solvus. All three reactions are decarbonation reactions, and their surfaces become tangent to the NaCl-H20 side of the triangle with decreasing temperature. The inset of the diagram shows schematically the three reactions terminating at halite saturation. At the temperatures of the diagram, similar to contactmetamorphic temperatures, NaCl has little effect on the wollastonite and diopside reactions, except to depress the curves to lower values of xco 2 - The tremolite-forming reaction, though, is affected in two significant ways. First, talc-forming reactions are completely suppressed (compare Figure 25B), and second, the tremolite-forming reaction occurs in I0W-XCO2 fluids, contrary to the relations in Figure 21 for the NaCl-free system. A serious error could be made in tacitly assuming that the formation of tremolite from dolomite + quartz implies a C02-rich fluid composition. Clearly, an assessment of the composition of the metamorphic fluid in equilibrium with siliceous dolomites requires information about the salinity, information that can be obtained, in the best of all worlds, from fluid inclusions. The compositions of metamorphic fluids Throughout this section, I have presented some of the chemical systematics relating the composition of the fluid phase to the equilibrium mineral assemblage. While it is comparatively easy to determine what the fluid composition was during metamorphism, it is much more difficult to determine how it got that way. In the broadest sense, the fluid compositions were established by a combination of the volatile components released to the fluid by mineral reactions and of components from external fluid sources. The nature of the interaction between fluid and rock has been described as either openor closed-system, depending on whether the rock has been infiltrated by externally derived fluids. This division is artificial and can be misleading because rocks rarely appear to have behaved as truly closed systems and because the same resulting mineral assemblage can be

89 produced in either system. Instead, I will use the terms rock-dominated and fluiddominated to describe the relation between the mineral assemblage and the chemical potentials of the fluid components. The term rock-dominated behavior is used to describe cases in which the fluid-phase components have abundances buffered by the mineral assemblage. If the amount of infiltrating fluid is less than the buffer capacity of the rock through which it is flowing, then the equilibrium fluid composition evolves toward the stoichiometric composition of fluids released by the rock with increasing reaction progress. The fluid composition evolution in this case was described by Greenwood (1975). The fluid evolved by mineral reaction makes up a large proportion of the pore fluid. Fluid-dominated behavior describes the case in which the chemical potentials of the volatile components are fixed outside the rock system. In this case, the amount of externally derived fluid that interacts with a rock exceeds the buffer capacity of the rock. The rock has a high-variance assemblage that does not buffer the fluid composition, and the equilibrium fluid composition may bear little similarity to the composition of the reactionevolved fluids.

Figure 26. Isothermal-isobaric sections of the system NaCl-C02-H20 at 200 MPa, showing the stability regions for mineral assemblages in siliceous dolomites at 500°C (solid lines) and 400°C (dashed lines), adapted from Labotka et al. (1988a). At 400 and 500°C, the two equilibria calcite + quartz = wollastonite + CO2 and tremolite + 3 calcite + 2 quartz = 5 diopside + 3 CO2 + H2O do not intersect the solvus until halite saturation (inset). The reaction 5 dolomite + 8 quartz + H2O = tremolite + 3 calcite + 7 CO2 intersects the solvus at 500°C and is connected on either side of the solvus by the tie-line connecting the compositions of the coexisting saline liquid and C02-rich gas. Talc- bearing assemblages are precluded at 500°C. The tie-lines are schematic, the solvus and fugacity coefficients for the calculated positions of the curves are from Bowers and Helgeson (1983b).

90 The magnitude of the influence of externally derived fluids on the mineral assemblage is obscured if the original pore fluid has the same composition as that of the evolved fluid or if the infiltrating fluid has the same composition as the reaction-evolved fluid. This makes the study of fluid/rock interaction difficult in pelitic terrains where evolved, pore, and infiltrating fluids are all H20-rich. Many pelitic rocks contain graphite or are open to carbonic fluids, giving the pore fluids compositions very different from those evolved by reaction. An example from the aureole of the Duluth complex is shown in Figure 27. Pelitic hornfelses in this aureole were in equilibrium with an initially 1 O W - X H 2 O pore fluid, probably resulting from equilibrium with graphite or from a CC>2-rich fluid derived from an underlying carbonate iron-formation. The assemblage in the hornfelses, muscovite + biotite + cordierite + K-feldspar + quartz, buffered the fluid composition to values of X H O = 1.0 along the isobarically trivariant reaction curves muscovite + biotite + quartz = cordierite + K-feldspar + H2O. In this aureole, the fluid/rock interaction was rockdominated, even though the mineral assemblage was high-variance. 2

Most cases of rock-dominated behavior have been documented for progressive metamorphism of siliceous dolomites. Rice (1977a,b) documented in detail the change in fluid composition in two aureoles in the Boulder batholith, Montana, from an initial H2Orich fluid composition to compositions of isobaric invariant points or thermal maxima for reactions in T-xco? space. He, in fact, corroborated the suggestion of Greenwood (1975) that petrographicaily significant changes in mineralogy occur at the invariant points. In these aureoles, the fluid composition was buffered to values of xcoj as great as 0.9, with diopside and diopside + phlogopite isograds marking fluid compositions of xcc>2 = 0.75, the stoichiometric value. There are numerous other examples of contact-metamorphosed dolomitic rocks that have buffered fluid compositions during much of the thermal history of metamorphism. These include the massive dolomites in the Alta aureole (Moore and Kerrick, 1976), marbles in the Twin Lakes and Hope Valley roof pendants, Sierra Nevada (Kerrick et al.,

7

Temperature (°C) Figure 27. The 150 MPa In / H 2 0 - T section for the equilibrium 6 muscovite + 2 phlogopite + 15 quartz = 3 cordierite + 8 K-feldspar + 8 H2O, from Labotka et al. (1984), showing the buffering of the fluid composition to H20-rich compositions by the reaction. The contours are lines of constant XMg in cordierite and biotite, as in Figure 18. The points with error bars are values calculated from mineral compositions in contact-metamorphosed Proterozoic Rove Formation.

91 1973; Ferry, 1989; Davis and Ferry, 1990), dolomites in the aureole of the Adamello pluton, Italy (Bucher-Nurminen, 1982), and dolomitic marbles in the Notch Peak aureole, Utah (Hover-Granath et al., 1983). All these record strong control of the fluid composition by the mineral assemblage. Fluid-dominated behavior, as especially related to open-system fluid infiltration, has been described primarily in aureoles containing argillaceous limestones or calcareous argillites. The reason this occurs in this lithology is that the reactions in these rocks are primarily decarbonation reactions and occurred in the presence of an aqueous fluid. No doubt, the open system behavior of fluids is the rule in most rock types, but the documentation of this behavior is more difficult for those rocks that evolve aqueous rather than carbonic fluids. Fluid flow through pelitic rocks could, perhaps, be more closely documented if the stratigraphy contained thin interbeds of calcareous rock that could act as passive markers of infiltration. Examples in which large volumes of H20-rich fluid interacted with calcareous rocks are typically marked by the occurrence of wollastonite, grossular, epidote, and vesuvianite in the mineral assemblage. Reactions that produce these minerals evolve CO2, yet the assemblages are stable only in H20-rich fluids under contact-metamorphic conditions. Calcareous argillites in the Notch Peak aureole contain these assemblages, as do some of the aluminous limestones in the Sierra Nevada roof pendants. The occurrence of periclase in dolomitic rocks also indicates an H20-rich fluid because the equilibrium dolomite = periclase + calcite + CO2 occurs only at temperatures in excess of 700°C, unless the fluid composition is less than xco 2 = 0.2 at 100 MPa (Kerrick, 1970). The Alta aureole, Utah, (Moore and Kerrick, 1976) and the Black Butte aureole, Montana, (Bowman and Essene, 1982) contain periclase marbles indicating the presence H20-rich fluids during metamorphism. There are some aureoles that contain the more exotic minerals from Bowen's (1940) song. Limestones in the Christmas Mountains aureole (Joesten, 1974; Joesten, 1976) contain melilite, merwinite, rankinite, perovskite, larnite, and tilleyite-bearing assemblages. These are stable in CC>2-rich fluids only at temperatures near or exceeding 1000°C. The blue-calcite marbles at Crestmore, California, also contain exotic minerals like monticellite, melilite, merwinite, and spunite (Burnham, 1959). The occurrence of these minerals is probably facilitated by very low pressure (Joesten estimates the pressure in the Christmas Mountains at ~35 MPa), but the stabilities are also extended to low temperature by an H20-rich fluid composition. The source of H20-rich fluids in metacarbonate rocks must be external to the system, which has been invoked by many, including me, as evidence for infiltration or diffusion metasomatism of carbonate rocks. The tacit acceptance of an external source for H^O-rich fluids must be tempered by the effects of NaCl on the miscibility of CO2 and H2O, as discussed in the previous section. Halite-bearing fluid inclusions in metamorphic rocks were described by Rich (1979), who believed they were derived during dissolution of an evaporite horizon. Numerous fluid-inclusion studies of metacarbonate rocks, for example, by Crawford et al. (1979) and Hollister and Burruss (1976) in rocks from British Columbia, and by Sisson et al. (1981) in rocks from the Grenville Province, have indicated that fluid-phase separation occurred during metamorphism. Trommsdorff and Skippen (1986) have suggested that selective loss or effervescence of the less dense CC>2-rich gas leaves behind an aqueous brine, which becomes ever more saline until halite saturation with increased loss of the gaseous phase. Consideration of Figures 23 and 24 show that immiscibility under contact-metamorphic conditions significandy affects reactions about the tremolite + talc + calcite + dolomite + quartz invariant point, but that wollastonite stability requires fluids less C02-rich than those on the solvus. Immiscibility might be more common than previously recognized, particularly at low-grades of metamorphism.

92 TRANSPORT PROPERTIES OF AQUEOUS FLUIDS Viscosity All transport processes are governed by a law that relates flow to a potential gradient and by a conservation law. Aqueous fluids flowing through rocks transport momentum, heat flowing through rocks transports energy, and diffusion through fluids and rocks transports mass. All three processes are observed to behave similarly — the flux is proportional to the negative of a gradient. Momentum is transported by fluid flow according to Newton's law of viscosity, which in one form can be written as

TyX is the shear stress in the plane perpendicular to y in the x direction, r\ is the viscosity, and the derivative is the change in the ^-component of the fluid velocity in the y direction. The analogous law for fluid flow through porous media is Darcy's law Q=v=-Kq(Vp-pg) Q is the discharge or flux with units of m 3 /s m 2 , which is equivalent to the velocity v, K„ is the hydraulic conductivity, which is a proportionality constant that depends on both fluid and rock properties, p is pressure, and g is the acceleration of gravity. Kq is related to the viscosity of the fluid through the permeability of the rock medium: K0 = ^ = i k is the permeability and v is the kinematic viscosity. Fluid dynamics are described in greater detail in Chapter 10, and many texts discuss the details of fluid flow through porous media, including Bear (1972), Bird et al. (1960), Strack (1989), and Domenico and Schwartz (1990). A summary of the geologic consequences of fluid flow in the Earth's crust is presented in Bredehoeft and Norton (1990). A recent description of the prediction of transport properties, as well as thermodynamic potential functions, of H2O near its critical point is given by Johnson and Norton (1991). The dynamic viscosity, r\, of H2O at pressures up to 100 MPa are given in Table 12, taken from Haar et al. (1984). At surface conditions, the viscosity is ~1 cp, or 10 -3 Pa s. At the temperatures of contact metamorphism, the viscosity is only -5% of the 25°C value, approximately 4 to 5 times the viscosity of air at 1 atm (Bird et al., 1960). There is a minimum in the value at ~600°C, and increased pressure raises the viscosity by ~0.33%/MPa. Thermal conductivity A significant amount of the energy transport in contact-metamorphic aureoles occurs by conduction. Heat conduction is governed by Fourier's law, in which the heat flux is proportional to the negative thermal gradient. The constant of proportionality is the thermal conductivity, K . Values of the thermal conductivity of H2O are given in Table 12, and range from 550 to 600 mW/m K at 350°C to a minimum of -150 to 200 mW/m K at 650 to 800°C. There is an increase in the conductivity with pressure of ~0.3 to 0.5 %/MPa. Rocks typically have conductivities that are an order of magnitude greater than those of H2O. Values listed by Clark (1966) range from 1.7 W/m K for limestones and shales to 6.7 W/m K for quartzite. Granitic rocks and gneisses have conductivities of approximately 3 W/m K.

93

IN CO oi-H CO in 00 i-H CO 5, Grt + Crn, S p l + Q t z *

Invariant points: 0 * * Univariant curves (invariant points in KFASH or KMASH): 5 * * Name Spl,Crn,Hy]** Spl,Crn,Al 2 Si0 5 ] Qtz,Hy,Grtj Qtz,Crn,Hy] Qtz,Cm,Al2Si05]

Reaction B t + A b S i C k + Q t z = Crd+Grt + Kfs + L B t + ( G r t ) + Q t z = (Crd) + H y + K f s + L * * B t + A l 2 S i 0 5 = S p l + C r d + C r n + K f s + V or L Bt+Grt+Al 2 SiC>5 = C r d + S p l + K f s + L B t + C r d + G r t = Hy+Spl + K f s + L

Reaction No. (6) (7) (23) (24) (25)

Divariant curves (univariant curves in KFASH or KMASH): 15 Spl,Crn,Hy,Grt] Spl,Crn,Hy,Crdl Spl,Crn,Hy,Al 2 Si05] Spl,Crn,Al 2 Si0 5 ,Grt] Spl,Crn,Al 2 Si0 5 ,Crd] Qtz,Hy,Grt,Spl) Qtz,Hy,Grt,AJ 2 Si05] Qtz,Hy,Grt,Crd] Qtz,Hy,Grt,Crn] Qtz,Hy,Spl,Crn] Qtz,Crn,Hy,CrdJ Qtz,Crn,HyAl2Si05] Qtz,Crn,A] 2 Si0 5 ,Spl] Qtz,Crn,AI 2 Si05,Grt] Qtz,Crn,Al 2 Si0 5 ,Crd]

Bt+Al 2 SiC>5 + Q.tz = Crd + Kfs + V or L B t + A l 2 S i 0 5 + Qtz = G r t + K f s + V or L Bt+Crd+Qtz = Grt+Kfs + L B t + Q t z = Hy(+Crd) + K f s + L " Bt+Qtz = Hy(+Grt)+Kfs + L * * Bt+Al 2 SiC>5 = C r d + C r n + K f s + V or L B t + C r d + C r n = Spl + K f s + V or L B t + A l 2 S i 0 5 + Crn = S p l + K f s + V or L Bt+Al 2 SiC>5 = C r d + S p l + K f s + V or L Bt+Crd = G rt+Al2Si05 + K f s + L B t + A l 2 S i C k = G r t + S p l + K f s + V or L Bt+Crd = G r t + S p l + K f s + V or L B t + C r d + Hy = G r t + K f s + V or L Bt+Crd = Spl + H y + K f s + V or L B t + G r t = Spl + H y + K f s + V or L

* Spl + Q t z has been reported as a stable association in the Laramie Aureole, Wyoming (Grant & Frost, 1990), but in other aureoles S p l + Q t z is rare and so it has been excluded as a general stable association even though at higher pressures it is relatively common (e.g. Hensen & Harley, 1990). * * See footnotes at the bottom of Table 4.

Figure 13. Sketch m a p of the C a m Chois diorite and surrounding aureole (Comrie aureole). (Used by permission of Allen and Unwin, from Mason, 1978, Fig. 5.6, p. 68, b a s e d on data in Tilley, 1924.)

120 CONTACT AUREOLES IN PELITES Metapelitic sequences have been examined from 76 contact aureoles in 14 countries (Appendices 1, 3 and 4), supplemented by data from 42 low pressure regional metamorphic sequences (Appendices 2 and 5). In many of these settings, the distinction between contact and regional metamorphism is obscure, so that in all cases the classification of the author(s) was accepted. In some low-pressure regional terranes, contact effects around plutons have been interpreted as local metamorphic highs, or overprints on an already elevated regional thermal regime, e.g., Huntly-Portsoy intrusions in the western Buchan zones, Scotland (Ashworth, 1975) and the Lexington batholith, northwestern Maine (Dickerson and Holdaway, 1989). In these cases, it may be that development of the lower-grade regional metamorphic zones preceded development of the higher-grade zones immediately adjacent to the intrusions. For a number of reasons, difficulties were encountered in the extraction of assemblage information from the literature. Many articles give a general account of the range of minerals that were found in a given metamorphic zone, instead of listing the mineral assemblages of individual rocks. Failure in some articles to distinguish between stable and metastable minerals has resulted in apparent assemblages that contain too many phases to be in equilibrium. Many accounts do not explicitly list the presence or absence of chlorite in low grade assemblages, and some lump K-feldspar and plagioclase together as feldspar. Consequently, some interpretation was required in the compilation of Appendices 3-5. The aureoles are grouped together on the basis of repeated prograde sequences of mineral assemblages, i.e., facies series. The diagnostic sequences of assemblages from each facies series type are summarised in Table 2. Not all aureoles show a complete prograde sequence. Aureoles surrounding relatively cool granites tend to show only the low- to intermediate-grade parts of the sequence (e.g., Shap, Kisokoma), compared to aureoles around intermediate to mafic intrusions (e.g., Ballachulish, Ronda, Insch) which are commonly more complete. Plutons emplaced in high-grade metapelitic protoliths may show only high-grade assemblages (e.g., Belhelvie, Lochnagar, Laramie, Cortlandt), and there may be gaps in either exposure or sample control (e.g., Sierra Nevada, Haddo House, Etive). Consequently, some of the placements are more tentative than others. Some aureoles contain assemblages that can be placed in more than one facies series type. In these cases, the most commonly reported assemblages are used to assign the aureoles to a given facies series. In Appendices 4 and 5 arrows indicate the possible adjacent facies series to which aureoles or regional settings with non-diagnostic mineral assemblages may be assigned. We emphasize that these facies series types are no more than approximate guidelines, and require back-up chemical analysis to be more rigorous. Nevertheless, we feel there is an impressive consistency that shows the value of such firstorder paragenetic analysis. CONTACT METAMORPHIC FACIES SERIES The different sequences of assemblages listed in Appendix 3 can be divided into four facies series types. Diagnostic features of each facies series are described starting from the lowest pressure. Short descriptions of representative aureoles from each facies series are given, followed by discussion of other aureoles and regional sequences from the same facies series. Model reaction sequences are inferred from the documented mineral assemblages, and are used to build up a schematic petrogenetic grid. The reaction numbering is based on Tables 3 and 4. Reactions are balanced using the mineral compositions in Table 1. Because a low-grade biotite zone is not present in all aureoles, the descriptions below concentrate on assemblages upgrade of the first appearance of biotite. Guidotti's (1984) review of biotite-forming reactions in low-grade metapelites applies to both contact and regional metamorphic settings.

121

122

tf ¡r

VSf' SCOTLAND c ^ T ^ Bollochulish,

MAP AREA

i

Edinburgh

Kilometers SEDIMENTARY UNITS \

Ouartzite

Granite

Phyllite

Bt - Am p h - Q t z Diorite

Ballachulish Slate I ^ I Leven Schist

i s

ill

PROGRADE R E A C T I O N S

IGNEOUS UNITS

C p x - A m p h - Q t z Diorite

Opx-Cpx-Monzodiorite

I Mu Chi Qtz

&

11 ^Chl out Crd Bt Ms Bt Qtz III Ms Crd 1 'Qtz Bt And ¡ A Crd Kfs » 11 T r—\ I V b Ms Qtz I V a ^ •pty i i And Kfs * i~Wf

Vb

»I

Kfs Qtz|±Bt,Crd,AI-sil1PI) I

tiiiiniiiiiDiNWiiiiink^'i i±crd, 4|-si| llllllllllllllllllllllHIINillllllllllllllllllBII Figure 15. Simplified map of the Ballachulish Igneous Complex and aureole, Scotland, showing mineral assemblage zones and distribution of andalusite and sillimanite. Open diamonds - andalusite. Filled diamonds sillimanite. Half-filled diamonds - andalusite + sillimanite. Model reactions separating mineral assemblage zones are discussed in the text (for numbering of reactions see Table 3). (Modified from Pattison, in press)

123 Zones I S l a

Zones II a n i

60 40 20

0 -20

10

20

30

40

50

60

-40

70

reaction 8

30

40

50

M/FM

60

70

80

M/FM

Zone 121(b) "250 -627v 280|l4b 59 4/9? 286 7g\ I \!24 38. \ /26 / 33 ,

60 50

40 20

s

0

«1


2. Aluminous hypersthene required still higher aSiC>2 for formation, and this phase is limited to the close proximity of the veins. The final silicate in the sequence is cordierite, which required the highest aSiC>2 and is spatially restricted to vein margins. Evidence cited by Tracy and McLellan (1985) suggests that SiC>2 had limited diffusive mobility in this dry

151 environment, and was probably less mobile than AI2O3, FeO and MgO. The reactivity of aluminous spinels with quartz is dependent on Mg/(Mg+Fe). At the pressure of emery formation (7.5 kbar) magnesian spinel is not stable with SiC>2 (see Caporuscio and Morse, 1978), whereas more hercynite-rich spinel occurs adjacent to quartz crystals, apparently stably. Tracy and McLellan (1985) reported an upper limit of spinel Mg/(Mg+Fe) of 0.350.40 in stable association with quartz, but it undoubtedly varies with pressure. In summary, emeries are rare and perhaps unique rocks, but they are potentially quite useful as indicators both of the stable mineral equilibria that prevail in the realm of ultrametamorphism and of the characteristics of disequilibrium in rapidly heated and cooled rocks. They are also especially helpful in tracking the more advanced stages of anatexis and residuum formation. PROPOSED PETROGENETIC GRID In this section, documented mineral assemblages from the different facies series types are rationalised in a schematic petrogenetic grid in part of KFMASH. In contrast to rigorous grids based on thermodynamic data for the end member KFASH and KMASH subsystems (e.g., Spear and Cheney, 1989; Powell and Holland, 1990), the aim of this grid is to account for the relative relationships between the majority of natural assemblages in pelites of normal composition. The grid does not, nor is it intended to, account for metapelitic assemblages in all compositions. The approach is based on Schreinemakers' analysis, augmented with simple composition, entropy and volume considerations. We believe that establishing the relative relationships amongst repeated occurrences of natural assemblages is a necessary first step toward the establishment of more complex grids and is the ultimate test of the usefulness of thermodynamically predicted grids. A general criticism of many existing grids is that they provide poor models for mineral assemblages in the low-pressure part of P-T space, corresponding to facies series Types la-2a. Table 3 is a summary of the phases, components and assemblage restrictions used in the construction of the schematic grids of Figures 32 and 33. The phases involved are those discussed in the 'Chemography' section, which represent the vast majority of minerals found in the observed metapelitic sequences. Two schematic grids have been constructed, one for either of the Fe- or Mg-end member systems (KFASH or KMASH) (Fig. 32), and one for the full KFMASH system (Fig. 33). Visualisation of relationships in the KFMASH grid is aided by consideration of the simpler end member grid. In the KFMASH grid, univariant curves correspond to invariant points in KFASH or KMASH; divariant fields in KFMASH correspond to univariant curves in KFASH or KMASH, etc. Restrictions All assemblages below the muscovite + quartz breakdown reaction are assumed to contain ms-qtz-bt, in addition to a hydrous supercritical fluid (V) during times of reaction. Thus, low-grade biotite-forming reactions are not considered in the grid. All assemblages above the muscovite + quartz breakdown reaction are assumed to contain kfs-bt-qtz, in addition to a silicate melt phase (L) or hydrous fluid phase during times of reaction. All quartz-absent reactions, as well as biotite-absent AFM reactions (e.g., Crd = Grt + Sil + Qtz) therefore cannot be represented in this grid. A grid for high-grade quartz-bearing and quartz-absent assemblages is provided separately below. Based on the observed mineral assemblages, several parageneses are considered to be unstable in the pure KFMASH system, each of which is discussed below. The low- to intermediate-grade assemblages include grt-crd-ms, grt-chl-and-ms, st-crd-ms, chl-kfs and st-kfs. At high grade, the following parageneses are considered unstable: hy-chl, hy-ms, hy+Al 2 Si05 and hy-st.

152

Figure 32. Schematic KMASH (or KFASH) petrogenetic grid for reactions listed in Table 3. Numbering of reactions is from Table 3. At P-T conditions above the dotted fluid-consuming anatectic reaction, reactions may produce melt (L) rather than fluid (V) (see text for discussion).

153

Figure 33. Schematic petrogenetic grid for a portion of KFMASH involving reactions listed in Table 3. Model univariant reactions are shown in heavy lines; bounding Fe-richer and Mg-richer end member reactions are shown in medium and light lines, respectively. Dashed lines are for the Ky = Sil, Ky = And and And = Sil equilibria, the latter shown occuning over an interval (see discussion in text). The net is oriented approximately in P-T space based upon experimental and thermodynamic data discussed in the text. Univariant curves have been drawn parallel for simplicity, even though they may intersect at higher pressures (see grids of Spear and Cheney, 1989; Powell and Holland, 1990). See text for details of the construction of the diagram.

154 A v e r a g e pelite

P = 5 kbars

staur

All with quartz albite muscovite biotite H 2 0 fluid

garnet chlorite

^Mn

I garnet chlorite

chlorite

| staurolite + chlorite)-^^

• I j

staur

+

I •

! Sill C i garnet |

8arnet

staurolite

r 500

600

Temperature ("C) Figure 34. Mineral assemblages predicted for 'average pelite' (Shaw, 1956) at 5 kbar from 400°-650°C. Wholerock MnO contents vary between 0.0 and 0.07 wt %. All assemblages contain quartz, muscovite, plagioclase, biotite and fluid. The shaded region is the range of normal pelite MnO contents. (From Symmes and Ferry, in press)

Garnet stability—the effect of Ca and Mn. Qualitatively, Mn and Ca have long been recognized to stabilize garnet (e.g., Hollister, 1966; Atherton, 1968; Harte and Hudson, 1979; Spear and Cheney, 1989). Spear and Cheney (1989) and Symmes and Ferry (in press) calculated that Fe-Mg garnet should not be stable in low grade metapelites except for very Fe-rich bulk compositions. Symmes and Ferry showed that the small amount of MnO present in most pelites is sufficient to stabilize garnet in many assemblages (Fig. 34). Consequently, garnet is regarded as an 'extra' phase stabilized by MnO in lowtemperature and low-pressure assemblages of crd-and facies series (Types la, 2a). Typically, garnet in such assemblages is Mn-rich and we therefore feel it is reasonable to regard grt-crd-ms and grt-A^SiOs-chl-ms as unstable parageneses in KFMASH. The sporadic occurrence of garnet in many low-pressure (Types la, 2a) contact metapelitic sequences, commonly showing no consistent relationship with grade, supports this reasoning. In contrast, garnet is typically a product of the breakdown of staurolite in reaction (4) in muscovite-bearing metapelites in higher-pressure sequences (facies series Types 2b4). Garnet in such rocks typically contains Mn/(Fe+Mg+Ca+Mn) < 0 . 1 (e.g., Waldron, 1986). Thus, it has been decided to include Grt+Ms+Bt as a stable paragenesis. Stability fields for grt-ms-chl-bt, ms-chl-grt-st-bt and ms-grt-st-bt are also included because these parageneses are ubiquitous in higher-pressure (Types 2b-4) sequences. Staurolite-cordierite-muscovite. The exclusion of st-crd-ms as a stable paragenesis is arguably the most controversial exclusion of all, because it eliminates an invariant point that appears in almost every metapelitic petrogenetic grid dating from Albee's (1965) grid (e.g., Hess, 1969; Thompson, 1976; Harte and Hudson, 1979; Carmichael, 1981, unpublished; Powell and Holland, 1990), with the one exception of Spear and Cheney (1989). As argued above, st-crd-ms is either absent or very rare in contact aureoles. However, it has been reported from a few well-known regional localities such as the Pyrenees (Zwart, 1958, 1962; Guitard, 1965) and the Waterville area, Maine (Osberg, 1968), possibly explaining why it has come to be an accepted paragenesis. If st-crd-ms-bt is stable, there must be a stable invariant point involving staurolite, cordierite, chlorite, A^SiOs, muscovite, biotite, quartz and fluid. In the above grids, this invariant point has been located in the range of 3-5 kbar.

155

Figure 35. P - T diagrams showing model uni-, di- and trivariant assemblages near the K F M A S H invariant point involving Ms, Chi, Bt, Qtz, St, Crd, And and V. The three diagrams are for fixed compositions in A F M projection with A below the Grt-Chl tie line and the following Mg/(Mg+Fe) values: (a) more Mg-rich than, (b) within, and (c) more Fe-rich than, the composition field of the invariant point. Open dot—invariant point metastable; solid dot—invariant point stable. Petrogenetic grids in Figures 32, 33 and 36 correspond to the topology in (c); however, in contrast to Fig. 36, Hudson (1980) considered that in A F M projection, chl plots to the Mg-richer side of the bt-crd tie-line, accounting for the topology of divariant fields involving chl-bt-crdAI2S1O5, as compared to Fig. 36. (Used by permission of the editors of the Contributions to Mineralogy and Petrology, from Hudson, 1980, Fig. 9, p. 48.)

Figure 35 (Hudson, 1980) shows the invariant point and associated univariant and divariant curves for different M g / ( M g + F e ) values. A major reason to reject the presence of the above invariant point, and therefore the stability of ms-crd-st, for most (not necessarily all) pelitic compositions is that the model univariant reactions M s + Chl + St + Qtz = A l 2 S i 0 5 + H 2 0 (2) and M s + Chl + Qtz = Crd + A l 2 S i 0 5 + Bt + H 2 0 (1), along with the linking model divariant reaction M s + Chl = A n d + Bt + Qtz + H 2 0 (9), can only be stable above the invariant point. Progressive prograde metapelitic assemblages with increasing pressure in contact aureoles show the following sequences: crd-chl-bt

crd-bt —> crd-kfs-bt

crd-chl-bt —> crd-bt —» crd-and-bt

(Types la, lb; lowest pressure); (Type l c ) ;

crd-chl-bt —> crd-and-bt-chl —»crd-and-bt —>crd-and-sil-bt and-chl-bt —> crd-and-bt-chl—>crd-and-bt—»crd-and-sil-bt st-chl-bt —> st-and-chl-bt —> st-and-bt—> st-and-sil-bt

(Types l c , 2a); (Type 2a);

( T y p e 2b; highest pressure).

Referring to Figures 35 and 36, this sequence with increasing pressure is consistent only with P - T paths above the invariant point. A P - T path below the invariant point would mean that the reaction M s + Chl = A n d + Bt + Qtz + H 2 0 (9) would not be possible, and therefore the common assemblage ms-and-bt±chl would not be seen. Holdaway (1991, pers. c o m m . ) pointed out that the widespread occurrence o f s t - A l 2 S i 0 5 - b t and crdAl 2 SiC>5-bt assemblages precludes stable st-crd-bt (refer to A F M diagram in Fig. 12a). It is possible that the invariant point exists at l o w pressures and for very Fe-rich compositions. Considering the f e w documented occurrences of st-crd-ms, it may be that the these assemblages were stabilized by minor components. Zn and L i fractionate strongly into staurolite, thereby expanding its stability field (Holdaway et al., 1986). The result would be to displace reaction ( 2 ) (Ms + Chl + St + Chl = Al 2 SiC>5 + Bt + H 2 0 ) up in temperature toward reaction (1) (Ms + Chl + Qtz = Crd + Al 2 SiC>5 + Bt + H 2 0 ) . Because reaction (2) has a steeper slope than reaction (1) (Harte and Hudson, 1979; Powell and Holland, 1990), this would cause the invariant point formed by the intersection o f these two reactions to be displaced along reaction (1) to higher pressures. K-feldspar + chlorite. In most K-feldspar-bearing pelites at low grade, K-feldspar is typically consumed in the reaction K f s + Chl = M s + Bt + Qtz + H 2 0 , producing the

156 common low-grade assemblage ms-chl-bt-qtz. If Chi was consumed, most of the higher grade reactions would not be possible. K-feldspar + staurolite. The terminal reaction of staurolite (see above) results in the loss of staurolite in muscovite- and quartz-bearing metapelites. With two reported exceptions, the Ronda aureole (Loomis, 1972a) and the Huntly-Portsoy area, Scotland (Ashworth, 1975a), staurolite does not persist beyond muscovite + quartz breakdown in the contact aureoles or regional localities examined. Zn might have stabilized staurolite to anomalously high grade, as suggested by Ashworth (1975) and Grant (1973). Hvpersthene. The exclusions hy-chl, hy-ms and hy-st reflect the restriction of hypersthene-bearing assemblages to very high grade. The exclusion hy-A^SiOs, although valid for the aureoles observed, does not remain applicable at higher pressure (Grant, 1985; Tracy and McLellan, 1985; Hensen and Harley, 1990). Univariant and divariant reactions It is emphasized that the above restrictions are based on common natural assemblages only, and should not be taken to represent all possibilities in KFMASH. With these restrictions, the range of possible invariant points, univariant reactions and divariant reactions in KFMASH is drastically reduced (Table 3). There are no stable non-degenerate KFMASH invariant points, 7 univariant curves and 16 divariant reactions, resulting in the relatively simple petrogenetic grids of Figures 32 (KFASH-KMASH) and 33 (KFMASH). It should be noted, however, that superimposition of the A^SiC^ phase diagram on the grid introduces numerous degenerate invariant points, including the and-ky-sil triple point and invariant points at each intersection of an A^SiC^-bearing equilibrium with one of the three ALjSiC^ equilibria. Orientation of the grid The slopes of reactions in the grid in Figures 32 and 33 are broadly consistent with the thermodynamically calculated reactions of Spear and Cheney (1989), Powell and Holland, 1990) and Berman (1988), all of which are based on experimental data sources listed in each paper. The main exception is reaction (10) (Ms + Crd = A^SiC^ + Bt + Qtz + H2O), which is shown to have a negative slope. Field evidence for a negative slope includes the development of and-crd-bt-ms-qtz assemblages upgrade along strike from mscrd-qtz-bt assemblages in the Ballachulish aureole (Pattison, 1989), the Cooma region (Vernon, 1988) and the Rosebud Syncline (Reinhardt and Rubenach, 1989), as well as increase in andalusite abundance with increasing grade in the assemblage ms-qtz-crd-and-bt in the Slave-Bear region, Canada (Frith, 1978), the northern Bavarian Forest, Germany (Bliimel and Schreyer, 1976) and the Bugaboo aureole (Pattison and Jones, unpubl.). The experimental studies of Hirschberg and Winkler (1968), Seifert (1970) and Bird and Fawcett (1973) also indicate a negative slope. In Figure 33, the And = Sil reaction has been drawn schematically as a band rather than a unique line. This recognizes the possibility that factors such as minor element composition, grain size and defect density may result in a family of stable And = Sil curves (Kerrick, 1990). Reactions at the onset of anatexis It is assumed that when fluid-consuming anatexis commences, all free fluid is consumed, resulting in fluid-absent, melt-bearing assemblages at higher grade (see 'Chemography' section). This situation is considered to be generally applicable because of the vanishingly small porosity of high grade rocks. Although the variance of assemblages across the anatectic boundary is therefore maintained, the general condition of uniform, and

157 generally high, aH20 becomes untenable because melting reactions will tend to buffer aH20 to lower values as temperature increases (Powell, 1983). From Schreinemakers' constraints, reactions which produce free hydrous fluid (dehydration reactions such as Ms + Qtz = Sil + Kfs + H2O) must have flatter slopes in PT space than those that produce melt (dehydration-melting reactions: Thompson, 1982; e.g., Ms + Qtz = Sil + Kfs + L). This is qualitatively shown by the steepening of reactions (5) and (18) as they cross the anatectic boundary in Figure 33. In some situations, sufficient externally-derived hydrous fluid may have been supplied to the site of melting, allowing greater degrees of fluid-consuming melting (e.g., Wickham, 1987; Pattison and Harte, 1988), perhaps even allowing melt to be H2Osaturated and free fluid to exist in the surroundings. In such cases, reactions in the vicinity of the anatectic boundary may be considerably more complex than illustrated in Figure 33 (e.g., Abbott and Clark, 1979). Facies series and bathozones It is useful to show the variation in mineral assemblages in rocks of fixed composition as a function of P and T (e.g., Hensen, 1971). Figure 36 shows a schematic P-T diagram (derived from Fig. 33) for uni-, di- and trivariant assemblages assuming a fixed bulk composition below the garnet-chlorite tie line at intermediate Mg/(Mg+Fe) in AFM projection. Documented mineral assemblages from different prograde contact metamorphic sequences are plotted on this diagram. Notwithstanding the fact that metapelites are not exactly the same in composition, Figure 36 indicates that pressure differences alone can explain most of the different facies series types documented above. Because this diagram assumes a fixed bulk composition, however, not all documented assemblages can be represented. Figure 36 also shows that although divariant reactions are less useful as isograds than univariant reactions, they are still important to understanding the evolution of mineral assemblages with grade. To show more clearly the effects of pressure and Fe-Mg variation, three schematic isobaric T-Xpe_Mg diagrams are shown in Figures 37-39. These represent isobaric sections appropriate for facies series Type la (Fig. 37), Type lb/c (Fig. 38) and Type 2a/b (Fig. 39). The most common range of bulk rock Mg/(Mg+Fe) for metapelites is shown as a broad arrow in each diagram. Examining Figure 39, it could be argued that the apparently lowest-pressure facies series characterized by cordierite (e.g., Types la, lb) could simply reflect Mg-richer bulk compositions than apparently higher-pressure facies series characterized by andalusite (e.g., Types lc and 2a) or staurolite (e.g., Type 2b). However, analyzed mineral compositions from aureoles of different facies series typically show bulk compositional ranges that overlap substantially (compare Figs. 14, 16, 19, 24-26), demonstrating that pressure is of primary importance in the development of different facies series. Comparison with the bathozone scheme of Carmichael (19781. Our proposed facies series 1 corresponds to Carmichael's bathozone 1. Within facies series 1, the boundary between Type lc (crd-and at intermediate grade—higher pressure) and Types la and lb (crd-kfs at intermediate grade—lower pressure) is in part dependent on composition as well as pressure (andalusite is favored in more Fe-rich compositions, cordierite is favored in Mg-rich compositions). Facies series 2 roughly corresponds to Carmichael's bathozone 2 and the lower part of bathozone 3. The distinction between facies series Types 2 and 3 is defined by a bathograd arising from the intersection of And = Sil with the reaction Ms + Chi + St + Qtz = Al 2 Si0 5 + Bt + H 2 0 (2). This is slightly different from Carmichael (1978), who defined

158

159 his bathograd 3/2 on the intersection of And = Sil with the reaction Ms + St + Qtz = Grt + Al 2 Si0 5 + Bt + H 2 0 (4). The occurrence of grt-and-bt assemblages, definitive of Carmichael's bathozone 2, is complicated by stabilization of garnet by Ca or Mn, and in many cases may therefore be unrelated to the breakdown of ms-st-qtz (see detailed discussion above in "Further subdivision of facies series Type 2b"). Moreover, with seemingly few exceptions (e.g., Lexington South—Dickerson and Holdaway, 1989), most Type 2b sequences in which the breakdown of ms-st-qtz is carefully documented show that this subassemblage gives way upgrade to grt-bt-sil assemblages, even when grt-bt-and assemblages are found at lower grades in the same sequence. In contrast, reaction (2) is not as severely affected as reaction (4) by non-KFMASH components (with the principal exception of Zn in staurolite). The intersection of reaction (2) with And = Sil defines a bathograd that appears to successfully distinguish between facies series 2b (characterized by and-bt upgrade of st-chl; see examples in Appendices 3 and 4) and facies series 3 (characterized by sil-bt upgrade of stchl; see examples in Appendix 5). Within facies series 2, the boundary between Type 2a (characterized by crd-and at low to intermediate grade) and Type 2b (characterized by st-and at low to intermediate grade) is partly dependent on composition (staurolite is favored in Fe-rich compositions; cordierite is favored in Mg-rich compositions; see Fig. 39). Consequently, in a few aureoles and regional settings that contain metapelites with a wide range of bulk compositions (e.g., Santa Rosa and Sawtooth aureoles: Compton, 1960), staurolite-, andalusite- and cordierite-bearing ms-chl-qtz-bt assemblages may occur at the same grade, allowing classification into Type 2a or 2b. In general, however, there is surprisingly little overlap, suggesting that these facies series can be generally effective in discriminating between higher pressure (staurolite-bearing) and lower pressure (cordierite-bearing) sequences. The distinction between facies series 3 and 4 is based on the intersection of reaction (2) with the Ky = Sil equilibrium. This bathograd is probably less useful at higher pressures than at lower pressure because of the greater modal abundance of low-grade garnet in higher pressure sequences, typically resulting in the consumption of chlorite at reaction (1). Consequently, Carmichael's 4/5 bathograd involving the intersection of reaction (4) and Ky = Sil may be of more practical use. Nevertheless, the bathograd proposed here gives a minimum pressure for facies series 4. High-grade assemblages in the anatectic zone. In high-grade quartz-bearing assemblages with K-feldspar and biotite, there is also good correlation between facies series types and pressure (Fig. 36). In facies series Type la, high-grade assemblages are gametabsent. In Types lb and lc, garnet is present but typically not with sillimanite. In facies series 2, crd-grt-sil is a common high-grade subassemblage. In facies series 3 and 4, grtsil is common, with and without cordierite. Even though the boundaries between these assemblages are strongly dependent on bulk composition (garnet is favored in Fe-rich compositions, cordierite is favored in Mg-rich compositions, e.g., Hensen, 1971), the observed distribution of assemblages can be explained by variations in pressure alone, e.g., Duluth (Labotka et al., 1981)-—grt-absent: 1.5 kbar; Laramie (Grant and Frost, 1990) and Ballachulish (Pattison, 1989)—grt-present and grt+sil-absent: both 3.0 kbar; Belhelvie (Droop and Charnley, 1985)—grt-sil present: 4.0 to 5.0 kbar. Table 4 (above) summarizes the construction of a schematic petrogenetic grid for both quartz-bearing and quartz-absent high grade assemblages in contact aureoles (see more detailed treatment in Pattison and Harte, 1991). Biotite, K-feldspar and either a fluid or melt phase are considered to be present in all assemblages. Associations that are not considered to be stable include qtz-crn, hy-A^SiC^, grt-crn and spl-qtz. At higher pressures, probably the only exclusion that still holds is qtz-crn (Hensen and Harley, 1990). Spl-qtz is not considered to be a stable paragenesis based on its overall rarity in

160 contact aureoles, with the exceptions of Laramie (Grant and Frost, 1990) and Cortlandt (Waldron, 1986; Tracy and McLellan, 1985). With these exclusions, there are no stable non-degenerate KFMASH invariant points and five stable univariant curves. Resultant petrogenetic grids for the end member KFASH or KMASH systems and the full KFMASH system are plotted in Figures 40 and 41, respectively. The stoichiometrics of the reactions Bt + Grt + Qtz = Crd + Hy + Kfs + L (7), Bt + Al 2 Si0 5 = Crd + Crn + Spl + Kfs + V or L (23), and Bt + Grt + Al 2 Si0 5 = Crd + Spl + Kfs + L (24), and the topology of Figure 41, have been corrected from Pattison and Harte (1985). Plotting documented assemblages from the Comrie, Ballachulish, Lochnagar, Laramie, Nain and Belhelvie aureoles (Appendix 3; see also Grant and Frost, 1990), results in the same overall correlation with pressure as with the grids in Figures 33 and 36. Stability of kvanite + biotite in andalusite-sillimanite sequences. In intermediatepressure facies series (e.g., facies series 3 and 4), sil-bt and ky-bt first appear by either reaction (2) (Ms + St + Chi + Qtz = Sil/Ky + Bt + H 2 0 ) or (4) (Ms + St + Qtz = Grt + Sil,Ky + Bt + H 2 0). Both of these reactions must occur at higher temperatures than the Al2SiC>5 triple point because of the occurrence of regional prograde sequences in which sillimanite is produced by reaction (2) (e.g., Rangeley quadrangle, Maine—Guidotti, 1974). Thermodynamic calculations give the same result (e.g., Spear and Cheney, 1989; Powell and Holland, 1990). In general, it should therefore not be possible to have a biotite-bearing prograde sequence of kyanite to andalusite to sillimanite, which means that interpretations of kyanite-bearing aureoles as representing normal isobaric paths through the stability fields of all three polymorphs (e.g., Pitcher and Berger, 1972; Atherton et al., 1975) must be reassessed. There are several possible alternative explanations. Kyanite might have been produced from the breakdown of pyrophyllite at low grade, although evidence for pyrophyllite is typically lacking. Reaction (2) may have proceeded at reduced aH 2 0, resulting in its displacement to lower temperatures into the kyanite and andalusite fields. However, this requires an efficient fluid buffer to maintain low aH20 throughout progress of an H 2 0-producing reaction. Loomis (1972b) noted that kyanite seems to be restricted to the inner zones of aureoles showing strong fabrics, typically parallel to the margins of the intrusions. He therefore concluded that kyanite-bearing rocks were dragged up from deeper levels along the margins of rising intrusions to be ultimately emplaced at depths where andalusite and sillimanite were stable (for a rebuttal of this suggestion, see Atherton et al., 1975). Finally, it is possible that the kyanite is a relict regional metamoiphic mineral overprinted by contact metamorphic andalusite and sillimanite. ANATEXIS In the highest grade parts of many aureoles, a variety of distinctive textures and structures suggest that anatexis has occurred. Macroscopic features include: "swirling textures" (Akaad, 1956); disruption of layering; development of segregations (leucosomes) rich in quartz and K-feldspar, with or without mafic-enriched selvages; anastomosing leucocratic veins and patches; and disoriented rigid fragments (schollen) suspended in a relatively homogeneous semipelitic matrix. Two examples from the Ballachulish aureole, Scotland are shown in Figures 42 and 43 (from Pattison and Harte, 1988). Very similar features are found in the Laramie aureole (Grant and Frost, 1990), Anmatjira-Reynolds range (Vernon and Collins, 1988, 1990), and Foyers aureole (Mould, 1946). Although these textures are characteristic of anatexis, studies by Sawyer and Barnes (1988) and McLellan (1989) show that similar features may be developed by subsolidus processes, showing the need for caution in their interpretation. Microscopic textures indicative of anatexis include: the development of skeletal (or

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5 + Qtz = Crd + Kfs + H 2 0 are of key importance to contact aureole studies, and deserve further careful experimental study. Resetting of geothermobarometers Low-pressure contact aureoles heat up and cool down relatively quickly compared to regional terrains, and so they are ideal settings in which to investigate problems of resetting of high grade exchange thermometers. In theory, if resetting generally occurs during slow cooling (Frost and Chacko, 1989; Spear, 1991), then the effects should be less severe in contact aureoles than in more slowly cooled regional terrains. Ultimately, diffusion rates might be constrained from the analysis of compositional profiles in coexisting minerals in aureoles whose thermal history is well known. Testing of multi-equilibrium techniques The most recent versions of GeO-Calc (Berman and Brown, 1988) and Theimocalc (Powell and Holland, 1990) allow for solid solution effects in calculating P-T locations of equilibria, and thus can and should be tested against quantitative thermobarometry in contact aureoles. The isobaric nature of most contact aureole settings should at least permit checking for similar pressures in various zones of a single aureole, while also checking for the correct sense of temperature change across an aureole. ACKNOWLEDGMENTS We are greatly indebted to Drs. M.J. Holdaway, D.M. Kerrick and J.A. Speer for providing critical reviews of this manuscript on short notice. Their comments have materially improved the final product. We also thank Drs. R.H. Vernon (Macquarie University) and T. Oba (Joetsu University) for sending difficult-to-obtain references from Australia and Japan, respectively, and Dr. D.M. Carmichael for a stimulating discussion on the topic of bathograds. Financial support was provided by NSERC Operating Grant 0037233 (DRMP) and NSF Grant EAR-8816382 (RJT). BIBLIOGRAPHY AND REFERENCES Abbott, R.N., Jr. and Clarke, D.B. (1979) Hypothetical liquidus relationships in the subsystems AI2O3FeO-MgO projected from quartz, alkali feldspar and plagioclase for a(H 2 0) less than or equal to 1. Can. Mineral. 17, 549-560. Akaad, M.K. (1956) The northern aureole of the Ardara Pluton of County Donegal. Geol. Mag. 93, 377392. Albee, A.L. (1965) A petrogenetic grid for the Fe-Mg silicates of pelitic schists. Amer. J. Sci. 263, 512536. Albee, A.L. (1972) Metamorphism of pelitic schists: reaction relations of chloritoid and staurolite. Geol. Soc. Amer. Bull. 83, 3249-3268. Allaart, J.H. (1965) The geology and petrology of the Trois Seignieurs massif, France. Leidse Geol. Med. 11,97-214. Allan, W.C. (1970) The Morven-Cabrach basic intrusion. Scottish J. Geol. 6, 53-72. Archibald, D.A., et al. (1983) Geochronology and tectonic implications of magmatism and metamorphism. Southern Kootenay Arc and neighbouring regions, southeastern British Columbia. Canadian J. Earth Sci. 20, 1891-1913. Arzi, A.A. (1978) Critical phenomena in the rheology of partially melted rocks. Tectonophys. 44, 173-184. Ashworth, J.R. (1975a) Staurolite at anomously high grade. Contrib. Mineral, and Petrol. 53,281-291. Ashworth, J.R. (1975b) The sillimanite zones of the Huntley-Portsoy area in northeast Dalradian, Scotland. Geol. Mag. 112(2), 113-224. Ashworth, J.R. (1976) Petrogenesis of migmatites in the Huntley-Portsoy area, northeast Scotland. Mineral. Mag. 40,661-682.

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Appendix 1. Contact aureoles containing metapelitic assemblages Name

Location

References

Ardara

N. Ireland

Pitcher & Sinha (1958), Pitcher & R e a d (1963), Naggar & Atherton (1970), Pitcher & Berger (1972), K e r r i c k ( 1987)

Arthursleigh

New S. Wales, Australia

Chenhall et al. (1988)

Ballachulish

Scotland

Pattison (1987; 1989; 1991; 1992), Pattison & H a r t e (1985; 1988; 1991), H a r t e et al. (1991), Voll et al. (1991).

Barnesmore

N. Ireland

Smart (1962), Pitcher & R e a d (1963), Naggar & A t h e r t o n (1970), Pitcher & Berger (1972)

Barr-Andlau (Steiger Schiefer)

Vosges, France

Rosenbuch (1877), Bosma (1964)

Belhelvie

Scotland

Stewart (1946), D r o o p & Charnley (1985)

Bodmin Moor

S.W. England

Ussher et al. (1909), Flett in Reid et al. (1910), Barrow in Ussher (1912)

Bruinbun

New S. Wales, Australia

B a t e m a n (1982)

Bugaboo

British Columbia, C a n a d a

Pattison & Jones (unpubl.)

Burke

Vermont, U.S.A.

Woodland (1963)

Bushveld

South Africa

Hall (1908), Hall (1915), Hall (1932)

Caplongue

France

Delor et al. (1984)

Cashel-Lough Wheelan

N. Ireland

L e a k e & Skirrow (1960), Treloar (1981)

Cligga H e a d

S.W. England

Reid & Scrivener (1906)

Comrie

Scotland

Tilley (1924)

Cortlandt Complex

New York, U.S.A.

Waldron (1986)

Cupsuptic

Maine, U.S.A.

Harwood (1966; 1973), Harwood & Larson (1969), Tcwhey & Hess (1974), Ryerson & Hess (1978), Bowers et al. (1990)

Dartmoor

S.W. England

Reid et al. (1912)

Donegal (Main G r a n i t e )

N.Ireland

Pitcher et al. (1959), Pitcher & R e a d (1960), Pitcher & Read (1963), Naggan & Atherton (1970), Pitcher & Berger (1972)

Duluth

Minnesota, U S A

G r o u t (1933), Labotka et. al. (1981; 1984), Labotka (1983)

Easky

Ireland

Yardley & Long (1981)

Errol

New Hampshire, U.S.A.

G r e e n (1963)

Etive (Cruachan)

Scotland

Kynaston & Hill (1908), D r o o p & Treloar (1981)

Fanad

N. Ireland

Pitcher & Read (1963), Naggar & Atherton (1970), E d m u n d s & Atherton (1971), Pitcher & Berger (1972), Kerrick (1987)

Foyers

Scotland

Mould (1946), Tyler & Ashworth (1983)

Gunnislake

S.W. England

Reid et al. (1911)

Haddo House

Scotland

Gribble (1966; 1968), D r o o p & Charnley (1985)

Hidaka

Japan

Shiba (1988)

Hingston Down

S.W. England

Reid et al. (1911)

198

Name

Location

References

H u n tly/Kn ock/Po rtsoy

Scotland

R e a d (1923; 1952), Ashworth (1975; 1976), Ashworth & Chinner (1978), D r o o p & Charnley (1985)

Insch/Boganclogh

Scotland

Read (1923; 1952), Fettes (1970), D r o o p & Charnley (1985)

Kiglapaii

Labrador, Canada

Berg (1977a,b), Speer (1982), Berg & Docka (1983), Kerrick & S p e e r (1988) Oki (1957; 1961), Katada (1964)

Kisokoma

Japan

Kwoiek

British Columbia, C a n a d a

Hollister (1969a; b)

Land's End

S.W. England

Reid & Flett (1907)

Laramie

Wyoming, U.S.A.

Grant & Frost (1990)

Leinster

Ireland

Brindley (1957)

Lexington

Maine, U.S.A.

Dickerson & Holdaway (1989)

Liberty Hill

S. Carolina, U.S.A.

Speer(1981)

Lilesville

N. Carolina, U.S.A.

Evans & S p e e r (1984)

Lochnagar

Scotland

Chinner (1962), Ashworth & Chinner (1978)

McGerrigle

Quebec, C a n a d a

Van Bosse & Williams-Jones (1988)

Mine/Wall/Shaw Creek

British Columbia, C a n a d a

Reesor (1973), Glover (1978), Archibald et al. (1983)

Morven-Cabrach

Scotland

Allan (1970), D r o o p & Charnley (1985)

Moy

Scotland

Zaleski (1985)

Mt. Raleigh

British Columbia, C a n a d a

Woodsworth (1979)

Nain aureoles

Labrador, C a n a d a

Berg (1977a,b)

Omey

Ireland

Ferguson & Harvey (1979)

Onawa

Maine, U.S.A.

Philbrick (1936), M o o r e (1960)

Orijarvi

Finland

Eskola (1914)

Osgood

Nevada, U.S.A.

H o t z & Willden (1964)

Oslo

Norway

Goldschmidt (1911)

Ronda

Spain

Loomis (1972a; b)

Ross of Mull

Scotland

Bosworth (1910), Barnicoat & Prior (1991)

Santa Rosa (several stocks)

Nevada

C o m p t o n (1960)

Shap

N. England

H a r k e r & M a r r (1891)

Shelburne

Nova Scotia, C a n a d a

H w a n g et al. (in press)

Sierra Nevada (southern)

California

Durrell (1940), Best & Weiss (1964)

Skiddaw

N. England

Rastall (1910)

St. Agnes

S.W. England

Reid et al. (1906)

St. Austell

S.W. England

Ussher et al. (1909)

Steinach

Bavaria, G e r m a n y

Okrusch (1969; 1971), Okrusch & Evans (1970)

Stillwater

M o n t a n a , U.S.A.

Page & Zientek (1985), Labotka (1985), Kath & Labotka (1986)

Strontian

Scotland

Tyler & Ashworth (1982), Ashworth & Tyler (1983)

199

Name

Location

Summit

British Columbia, C a n a d a

R e e s o r (1958), Archibald et al. (1983)

Tanohata

Japan

Shimazu (1962)

Thorr

N. Ireland

Iyengar et al. (1954), Pitcher & R e a d (1963), Naggar & Atherton (1970), Pitcher & B e r g e r (1972)

References

Tinaroo

Queensland, Australia

R u b e n a c h & Bell (1988)

Tono

Japan

Seki (1957), Okuyama-Kusunose (1985), O k u y a m a - k u s u n o s e & Itaya (1987)

Vogtland

Saxony, G e r m a n y

Hussak (1887), Bosma (1964)

Willi Willi

New S. Wales, Australia

KJnny e t a l . (1985)

Appendix 2. Low pressure (andalusite-sillimanite type) regional metapelitic sequences

Name

Location

References

Abukuinu Plateau

Japan

M i y a s h i r o (1958; 1 % 1 )

Anmatjira-Reynolds Range

Northern Territory, Australia

V e r n o n & Collins (1988), V e r n o n et al. (1990), C l a r k e et al. (1990)

Northwest Territories, Canada

Nielsen (1978)

B e a r Province - Arseno Lake

Northwest Territories, Canada

S t - O n g e (1984)

B u c h a n ( Y t h a n Valley)

Scotland

R e a d (1923; 1952), H u d s o n (1980)

Chugach Complex

Alaska, U.S.A.

H u d s o n & P l a f k e r (1982), Sisson et al. (1989)

Cooma Region

New South Wales, Australia

J o p l i n (1942), H o p w o o d (1976), V e r n o n (1978; 1979; 1982; 1988), V e r n o n & H o b b s (1984) M u n k s g a a r d (1988)

English River Belt

Ontario, Canada

T h u r s t o n & B r e a k s (1978)

H i d a k a Belt

Japan

S h i b a (1988) S c h a u (1978)

- Wopmay Orogen

Keewatin Region

Northwest Territories, Canada

Klin Flon Belt-Niblock L a k e

Manitoba, Canada

Briggs (1990)

Massif C e n t r a l

France

T h o m p s o n & B a r d (1982)

M e g u i n a Belt

N o v a Scotia, C a n a d a

R a e s i d e et aJ. (1988)

Michigan

U.S.A.

J a m e s (1955)

200

Name

Location

References

Mt. Raleigh Pendant

British C o l u m b i a , C a n a d a

W o o d s w o r t h (1979), Kerrick & W o o d s w o r t h (1989)

New England - Augusta

M a i n e , U.S.A.

- O r r s Island

M a i n e , U.S.A.

- Waterville

M a i n e , U.S.A.

- Errol

N e w H a m p s h i r e , U.S.A.

N o v a k & H o l d a w a y (1981), H o l d a w a y e t al. (1982; 1988) D u n n A L a n g (1988), L a n g A D u n n (1990) O s b e r g (1968; 1971), F e r r y ( 1 9 8 0 ) G r e e n (1963)

N o r t h e r n Bavarian F o r e s t

Germany

O m e o Region

N e w S. Wales, A u s t r a l i a

Panamint Mountain

California

Liibotka (1981)

Peninsula Ranges

California, U.S.A.

S c h w a r c z (1969)

Placitas-Juan Taho R a n g e

N e w Mexico, U . S . A .

V e r n o n (1987)

Pyrenees-Canigou

France

G u i t a r d (1965)

- East Northern

France

G o l b e r g A L e y l e r o u p (1990)

- Trois Seigneurs

France

A l l a a r t (1959), W i c k h a m (1987)

- Valle d e A r a n

France/Spain

Z w a r t (1958; 1962)

Q u e t i c o Belt

Ontario, Canada

Pirie A M a c k a s e y (1978)

R o b e r t s o n River Syncline

Queensland, Australia

R e i n h a r d t A R u b e n a c h (1989)

R o s e b u d Syncline

Queensland, Australia

R e i n h a r d t A R u b e n a c h (1989)

-Migo

Japan

Tsuji (1965)

- Kiso

Japan

O k i (1957; 1961), K a t a d a (1964)

- Kii

Japan

Suwa (1961), M i y a s h i r o (1961)

Slave Province

Northwest Territories, Canada

R o s s (1967), H e y w o o d A D a v i d s o n (1969), Kretz (1968; 1973), K a m i n e n i (1975), R a m s a y (1974), R a m s a y A K a m i n e n i (1977), T h o m p s o n (1978). Frith (1978)

S o u t h e r n Province

Ontario, Canada

C a r d (1978)

Truchas Peaks

N e w Mexico, C a n a d a

Grambling (1981)

Wantabndgery-Tumbarumba

N e w S. W a l e s , Australia

Vallance (1953; 1967)

Wongwibinda

N e w S. W a l e s , A u s t r a l i a

Vernon (1982)

B l u m e l & S c h r e y e r (1976; 1977) M o r a n d (1990)

R y o k e Belt

201

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