Stable Isotopes in High Temperature Geological Processes 9781501508936, 9780939950201

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REVIEWS in MINERALOGY Volume 16

STABLE ISOTOPES in HIGH TEMPERATURE GEOLOGICAL PROCESSES Editors:

John W. Valley, Hugh P. Taylor, Jr., James R. O'Neil Authors: ROBERT N. CLAYTON Department of Geophysical Sciences University of Chicago Chicago, Illinois 60637

HIROSHIOHMOTO Department of Geosciences Pennsylvania State University University Park, Pennsylvania 16802

DAVID R. COLE Geosciences Group, Chemistry Division Oak Ridge National Laboratory Oak Ridge, Tennessee 37831

JAMES R. O'NEIL United States Geological Survey 345 Middlefxeld Road Menlo Park, California 94025

ROBERT E. CRISS United States Geological Survey National Center, Mail Stop 981 Reston, Virginia 22092

SIMON M.F. SHEPPARD C.R.P.G., BP 20, 54501 15 rue Notre Dame de Pauvres Vandoeuvre-les-Nancy Cedex, France

ROBERT T. GREGORY Department of Earth Sciences Monash University Clayton, Victoria 3168 Australia

BRUCE E. TAYLOR Geological Survey of Canada 601 Booth Street Ottawa, Ontario, Canada K1A0E8

T. KURTIS KYSER Department of Geology University of Saskatchewan Saskatoon, Saskatchewan, Canada S7N0W0

HUGH P. TAYLOR, J r . Division of Geological and Planetary Sciences California Institute of Technology Pasadena, California 91125

KARLIS MUEHLENBACHS Department of Geology University of Alberta Edmonton, Alberta, Canada T6G 2E3

JOHN W.VALLEY Department of Geology and Geophysics University of Wisconsin Madison, Wisconsin 53706

Series Editor: PAUL H. RIBBE Department of Geological Sciences Virginia Polytechnic Institute & State University Blacksburg, Virginia 24061

COPYRIGHT

1986

MINERALOGICAL SOCIETY of AMERICA

REVIEWS in MINERALOGY (Formerly: SHORT COURSE NOTES ) ISSN 0275-0279 Volume 16: STABLE

ISOTOPES

in High Temperature Geologic Processes ISBN 0-939950-20-0 ADDITIONAL COPIES of this volume as well as those listed below may be obtained from the MINERALOGICAL SOCIETY of AMERICA, 1625 I Street, N.W., Suite 414, Washington, D.C. 20006 U.S.A. Reviews in Mineralogy Volume 1: Sulfide Mineralogy, 1974; P. H. Ribbe, Ed. 284 pp. Six chapters on the structures of sulfides and sulfosalts: the crystal chemistry and chemical bonding of sulfides, synthesis, phase equilibria, and petrology. I S B N # 0-939950-01-4. Volume 2: Feldspar Mineralogy, 2nd Edition, 1983; P. H. Ribbe, Ed. 362 pp. Thirteen chapters on feldspar chemistry, structure and nomenclature; At,Si order/disorder in relation to domain textures, diffraction patterns, lattice parameters and optical properties: determinative methods; subsolidus phase relations, microstructures, kinetics and mechanisms of exsolution, and diffusion; color and interference colors; chemical properties; deformation. I S B N # 0-939950-14-6. Volume 3: Oxide Minerals, 1976; 0. Rumble III, Ed. 502 pp. . Eight chapters on experimental studies, crystal chemistry, and structures of oxide minerals; oxide minerals in metamorphic and igneous terrestrial rocks, lunar rocks, and meteorites. I S B N * 0-939950-03-0. Volume 4: Mineralogy and Geology of Natural Zeolites, 1977; F. A. Mumpton, Ed. 232 pp. Ten chapters on the crystal chemistry and structure of natural zeolites, their occurrence in sedimentary and low-grade metamorphic rocks and closed hydrologic systems, their commercial properties and utilization. I S B N # 0-939950-04-9. Volume 5: Orthosilicates, 2nd Edition, 1982; P. H. Ribbe, Ed. 450 pp. Liebau's "Classification of Silicates" plus 12 chapters on silicate garnets, olivines, spinels and humites; zircon and the actinide orthosilicates; titanite (sphene), chloritoid, staurolite, the aluminum silicates, topaz, and scores of miscellaneous orthosilicates. Indexed. I S B N # 0-939950-13-8. Volume 6: Marine Minerals, 1979; R. Q. Burns, Ed. 380 pp. Ten chapters on manganese and iron oxides, the silica polymorphs, zeolites, clay minerals, marine phosphorites, barites and placer minerals; evaporite mineralogy and chemistry. I S B N * 0-939950-06-5. Volume 7: Pyroxenes, 1980; C. T. Prewitt, Ed. 525 pp. Nine chapters on pyroxene crystal chemistry, spectroscopy, phase equilibria, subsolidus phenomena and thermodynamics; composition and mineralogy of terrestrial, lunar, and meteoritic pyroxenes. I S B N # 0-939950-07-3. Volume 8: Kinetics of Geochemical Processes, 1981; A. C. Lasaga and R. J. Kirkpatrick, Eds. 398 pp. Eight chapters on transition state theory and the rate laws of chemical reactions; kinetics of weathering, diagenesis. igneous crystallization and geochemical cycles; diffusion in electrolytes: irreversible thermodynamics. I S B N * 0-939950-08-1.

Volume 9A: Amphiboles and Other Hydrous Pyriboles—Mineralogy, 1984; D. R. Veblen, Ed. 372 pp. Seven chapters on biopynbote mineralogy and polysomatism; the crystal chemistry, structures and spectroscopy of amphiboles; subsolidus relations; amphibole and serpentine asbestos—mineralogy, occurrences. and health hazards. I S B N # 0-939950-10-3. Volume 9B: Amphiboles: Petrology and Experimental Phase Relations, 1982; 0. R. Veblen and P. H. Ribbe, Eds. 390 pp. Three chapters on phase relations of metamorphic amphiboles (occurrences and theory); igneous amphiboles; experimental studies. I S B N # 0-939950-11-1. Volume 10: Characterization of Metamorphism through Mineral Equilibria, 1982; J. M. Ferry, Ed. 397 pp. Nine chapters on an algebraic approach to composition and reaction spaoes and their manipulation; the Gibbs' formulation of phase equilibria; geologic thermobarometry; buffering, infiltration, isotope fractionation, compositional zoning and inclusions; characterization of metamorphic fluids. I S B N * 0-939950-12-X. Volume 11: Carbonates: Mineralogy and Chemistry, 1983; R. J. Reeder, Ed. 394 pp. Nine chapters on crystal chemistry, polymorphism, microstructures and phase relations of the rhombohedral and orthorhombic carbonates; the kinetics of C a C O a dissolution and precipitation; trace elements and isotopes in sedimentary carbonates; the occurrence, solubility and solid solution behavior of Mg-calcites; geologic thermobarometry using metamorphic carbonates. ! S B N # 0-939950-15-4. Volume 12: Fluid Inclusions, 1984; by E. Roedder. 644 pp. Nineteen chapters providing an introduction to studies of ail types of fluid inclusions, gas, liquid or melt, trapped in materials from the earth and space, and their application to the understanding of geological processes. I S B N # 0-939950-16-2. Volume 13: Micas, 1984; S. W. Bailey, Ed. 584 pp. Thirteen chapters on structures, crystal chemistry, spectroscopic and optical properties, occurrences, paragenesis, geochemistry and petrology of micas. I S B N # 0-939950-17-0. Volume 14: Microscopic to Macroscopic: Atomic Environments to Mineral Thermodynamics, 1985; S. W. Kieffer and A. Navrotsky, Eds. 428 pp. Eleven chapters attempt to answer the question. "What minerals exist under given constraints of pressure, temperature, and composition, and w h y ? " Includes worked examples at the end of some chapters. I S B N # 0-939950-18-9. Volume 15: Mathematical Crystallography, 1985; by M. B. Boisen, Jr. and G. V. Gibbs. Approx. 450 pp. A matrix and group theoretic treatment of the point groups. Bravais lattices. and space groups presented with numerous examples and problem sets, including solutions to common crystallographic problems involving the geometry and symmetry of crystal structures. I S B N * 0-939950-19-7.

FOREWORD The editors and authors of this volume presented the thirteenth in a series of short courses on behalf of the Mineralogical Society of America in November 1986, just prior to the annual meetings of MSA and the Geological Society of America in San Antonio, Texas. "STABLE ISOTOPES in HIGH TEMPERATURE GEOLOGICAL PROCESSES" was prepared for the course, and it is the seventeenth volume published by MSA in its now well established series, R E V I E W S in M I N E R A L O G Y [see detailed list and ordering information on the page opposite]. The text of this book was assembled from authorprepared, camera-ready copy — thus the wide variety of font types. The Mineralogical Association of Canada will sponsor a short course in May 1987 on the use of stable isotopes in the study of low temperature geological processes, and readers interested in the volume that will result from this undertaking (involving several of the authors of this volume) should write to: MAC, Department of Mineralogy and Geology, Royal Ontario Museum, 100 Queen's Park, Toronto, Ontario, Canada M5S 2C6. Paul H. Ribbe Series Editor September 16, 1986

DEDICATION This volume on the stable isotope geochemistry of high temperature geologic processes is respectfully dedicated to Samuel Epstein, Professor of Geochemistry at the California Institute of Technology for the past 30 years. Although the topics of this volume encompass only a small range of the applications of stable isotope geochemistry to which Sam has been a major contributor, this sub-field was to a large degree originated by Sam and his first Ph.D. student, Bob Clayton, in the middle 1950's when they made their pioneering studies of oxygen isotope geochemistry and geothermometry of coexisting minerals in igneous and metamorphic rocks and ore deposits. Sam is now near the age of mandatory retirement, but is still extremely active in a wide range of disciplines, as evidenced by his recent (1985/1986) papers on D/H and 13 12 C/ C studies of meteorites, isotopic studies of plants and animals, D/H studies of water in silicate melts and igneous rocks, and his work on a phosphate-chert-H 2 0 paleotemperature scale, to cite just a few. The field of stable isotope geochemistry was established by H. C. Urey, and A.O.C. Nier provided us with the basic mass spectrometer to do the work. But it is Sam Epstein who was mainly responsible for the flowering and maturing of this field into the enormous range of sub-fields and specialties that we know today (ocean paleotemperatures, high-temperature geothermometry, origins of natural waters, paleoclimatology, glacier research, biologic and geobiologic processes including plant and animal physiology, meteorology, ore deposits, oceanography, weathering and soil formation, and studies of the origin of igneous, metamorphic, and sedimentary rocks, meteorites, and tektites, etc.). Of course, starting with Harry Thode, who might be termed the father of sulfur isotope geochemistry, there are other eminent scientists who contributed immensely to the early development of this field, most of whom were also disciples of Urey: Harmon Craig (carbon isotopes), Irving Friedman (hydrogen isotopes), Cesare Emiliani (paleotemperatures), and John McCrea, Charles McKinney, Sol Silverman, Peter Baertschi, among others. The importance of the various men listed above to the field of geochemistry as a whole is demonstrated by the list of prizes they have been awarded: one Nobel Prize, 4 Day Medals, and 6 Goldschmidt Medals (out of the total of 14 awarded so far). In this dedication, we wish to recognize all of these pioneers. The editors and authors of this volume, in particular, owe a sizable debt to Samuel Epstein. The linkage with the past is clearly seen when one traces the scientific lineage of the contributors to this volume. Of the three editors, one is a scientific 'son' of Sam's, iii

one is a 'grandson', and the other is a scientific 'great-grandson'. Of the 12 authors of chapters, 10 are direct descendants (2 'sons', 2 'grandsons', and 6 'great-grandsons'). Several of S a m ' s scientific 'grandsons' have in fact worked directly with him as postdoctoral fellows, and 6 of the 12 contributors to this volume received their early training in stable isotopes either as Ph.D. students or as post-doctoral fellows at Caltech. T h e list does not at all do justice to the extended and widely ranging members of Sam's scientific family who permeate the other sub-fields throughout the field of stable isotope geochemistry and who are not among the authors of this book. Sam takes pride in all of the s t u d e n t s and post-doctoral fellows who have worked with him over the years, and it is fair to say t h a t he has left his mark on all of them in some way, particularly through his intuitive 'feel' for certain seemingly intractable but very important problems t h a t in fact can be solved with a little clever laboratory work, a refusal to get bogged down in extraneous and u n i m p o r t a n t details, and an understanding of the intrinsic accuracy required for a given measurement to be decisive. There is very little work in stable isotope geochemistry t h a t doesn't utilize or make reference to some aspect of Sam's collected scientific works.

BIOGRAPHICAL SKETCH of SAMUEL EPSTEIN Samuel Epstein was born in Kobryn, Poland (now U.S.S.R.) in 1919. His family moved to Winnipeg, Manitoba, C a n a d a in September 1927, thereby escaping the fate of all of his relatives when the holocaust descended on Poland 14 years later. Sam became a Canadian citizen and grew up in Winnipeg, attending the University of M a n i t o b a where he received a B.Sc. and M.Sc. in chemistry and geology in 1941 and 1942. At this point he switched totally to chemistry and began work on the kinetic reactions of the high explosive RDX at McGill University in Montreal, where he received his Ph.D. in September 1944. Sam then joined the Canadian Atomic Energy Project in Montreal, where he worked on rare-gas fission products. Here he met Harry T h o d e and also his f u t u r e wife Diane, whom he married in September 1946, while he was a visiting scientist in T h o d e ' s group at McMaster University in Hamilton, Ontario. In 1947, Harold Urey was looking for someone to carry out the paleotemperature project he had started at Chicago. In October 1947, at T h o d e ' s recommendation, Sam and Diane moved to Chicago where they established residence in an a p a r t m e n t above Harold Urey's garage (formerly the coachman's quarters). Recognizing the golden opportunity he had in this fascinating research problem, Sam went to work with immense drive and dedication. Some of the most difficult mass spectrometric and analytical problems were solved t h a t first year, b u t because of visa problems, the Epsteins went back to C a n a d a in October 1948. Sam had become so valuable to the project t h a t Urey personally visited the Immigration Bureau and was finally able to resolve the visa difficulty so t h a t Sam could return to the project in the Spring of 1949. By this time McCrea and McKinney had left the project and Urey's interests had largely turned to the origin of the solar system. So, basically with only the help of Toshiko Mayeda, Sam had to solve most of the remaining problems associated with the development of the carbonate paleotemperature scale. During this period he also started several other projects, including the first survey of the oxygen isotope compositions of natural waters. In June 1952, when Harrison Brown moved from Chicago to Caltech, he invited Sam to come along. T h i s transfer from the University of Chicago, which also included Claire P a t t e r s o n and Charles McKinney, was the beginning of the geochemistry operation at Caltech. Lee Silver, a graduate student at the time, immediately became integrated into the operation, and Bob Clayton came over from the Chemistry Division to become S a m ' s first g r a d u a t e student. T h e University of Chicago exodus to Caltech was concluded when Heinz Lowenstam and Gerry Wasserburg emigrated in 1953 and 1955, respectively. Sam and Diane became citizens of the U.S.A. in 1953, raised two sons, and are now proud grandparents of three. During his career Sam Epstein has initiated several new sub-fields of isotope geochemistry and has written more than 100 research papers. In recent years he has been widely recognized for his monumental scientific achievements by receiving the Goldschmidt Medal (1977), the Day Medal (1978), and being elected to the National Academy of Sciences (1977) and the American Academy of A r t s and Sciences (1977). iv

PREFACE T h e development of modern isotope geochemistry is without doubt attributed to the efforts, begun in the 1930's and 1940's, of Harold Urey (Columbia University and the University of Chicago) and Alfred O.C. Nier (University of Minnesota). Urey provided the ideas, theoretical foundation, the drive, and the enthusiasm, b u t none of this would have made a m a j o r impact on E a r t h Sciences without the marvelous instrument developed by Nier and later modified and improved upon by Urey, Epstein, McKinney, and McCrea at the University of Chicago. Harold Urey's interest in isotope chemistry goes back to the late 1920's when he and I.I. Rabi returned from Europe and established themselves at Columbia to introduce the then brand-new concepts of quantum mechanics to s t u d e n t s in the United States. Urey, of course, rapidly made an impact with his discovery of deuterium in 1932, the 'magical' year in which the neutron and positron were also discovered. Urey followed up his initial i m p o r t a n t discovery with many other experimental and theoretical contributions to isotope chemistry. During this period, A1 Nier developed the most sophisticated mass spectrometer then available anywhere in the world, and made a series of surveys of the isotopic ratios of as many elements as he could. Through these studies, which were carried out mainly to obtain accurate atomic weights of the various elements, Nier and his co-workers clearly demonstrated t h a t there were some fairly large variations in the isotopic ratios of the lighter elements. However, the first inkling of a true application to the E a r t h Sciences d i d n ' t come until 1946 when Urey presented his Royal Society of London lecture on ' T h e Thermodynamic Properties of Isotopic Substances' (now a classic paper referenced in most of the published papers on stable isotope geochemistry). With the information discovered by Nier and his co-workers t h a t limestones were a b o u t 3 percent richer in l s O than ocean water, and with his calculations of the temperature coefficient for the isotope exchange reaction between C a C 0 3 and H 2 0 , Urey realized t h a t it might be possible to apply these concepts to determining the paleotemperatures of the oceans. Urey was never one to overlook important scientific problems, regardless of the field of scientific inquiry involved. In fact, he always admonished his s t u d e n t s to 'work only on truly important problems!' Urey, then a Professor at the University of Chicago, decided to take a hard look into the experimental problems of developing an oxygen isotope paleotemperature scale. Although the necessary accuracy had not yet been attained, the design of the Nier instrument seemed to offer a good possibility, with suitable modifications, of making the kinds of precise measurements necessary for a sufficiently accurate determination of the I8 0 / i e O ratios of both C a C 0 3 (limestone) and ocean water. Enormous efforts would be required to do this, because even if all the mass spectrometric problems could be solved, every analytical and experimental procedure would have to be invented from scratch, including the experimental calibration of the temperature coefficient of the equilibrium fractionation factor between calcite and water at low temperatures. T o carry out this formidable study, Urey gathered around himself a remarkable group of students, postdoctoral fellows, and technicians, as well as his paleontologist colleague Heinz Lowenstam. With Sam Epstein at the center of the effort and acting as the principal driving force, the rest, as they say, 'is history.' T h e marvelous nature of the Nier-Urey mass spectrometer is attested to by the fact t h a t the basic design is still being used, and t h a t there are now h u n d r e d s of laboratories throughout the world where this kind of work is being done. For example, the original instrument built by Sam Epstein and Chuck McKinney at Caltech in 1953 is still in use and has to date produced more than 90,000 analyses. University, government, and industrial laboratories have found these instruments to be an indispensable tool. Enormous and widely varying application of the original concepts have been made throughout the whole panoply of Earth, Atmospheric, and Planetary Sciences. In the

v

present volume we concentrate on an important sub-field of this effort. T h a t particular sub-field was inaugurated in Urey's laboratories at Chicago by P e t e r Baertschi and Sol Silverman, who developed the fluorination technique for extracting oxygen from silicate rocks and minerals. This technique was later refined and improved in the late 1950's by Sam Epstein, Hugh Taylor, Bob Clayton, and Toshiko Mayeda, and has become the prime analytical method for studying the oxygen isotope composition of rocks and minerals. T h e original concepts and potentialities of high-temperature oxygen isotope geochemistry were developed by Samuel Epstein and his first s t u d e n t , Bob Clayton. Also, Bob Clayton, A.E.J. Engel, and Sam Epstein carried o u t the first application of these techniques to the study of ore deposits. T h e first useful experimental calibrations of the high-temperature oxygen isotope geothermometers quartz-calcite-magnetite-H„0 were carried out initially by Bob Clayton, and later with his first s t u d e n t Jim O'Neil. In the meantime, Sam Epstein and his second student, Hugh Taylor, had begun a systematic s t u d y of 1 8 O/ 1 0 O variations in igneous and metamorphic rocks, and were the first to point out the regular order of l s O / l e O fractionations among coexisting minerals, as well as their potential use as geochemical tracers of petrologic processes. During this period, a parallel development of sulfur isotope geochemistry was being carried out by Harry T h o d e and his group at McMaster University in C a n a d a . They developed all the mass spectrometric and extraction techniques for this element, and also provided the theoretical and experimental foundation for understanding the equilibrium and kinetic isotope chemistry of sulfur. S t a r t i n g from these beginnings, most of which took place either at the University of Chicago, Caltech, or McMaster University (but also with i m p o r t a n t input f r o m Irving Friedman's laboratory at the U.S. Geological Survey, from Athol R a f t e r ' s laboratory in New Zealand, and from Columbia, Penn State, and the Vernadsky Institute in Moscow), there followed during the decades of the late 60's, 70's, and early 80's the development and m a t u r i n g of the sub-field of high-temperature stable isotope geochemistry. This discipline is now recognized as an indispensable a d j u n c t to all studies of igneous and metamorphic rocks and meteorites, particularly in cases where fluid-rock interactions are a m a j o r focus of the study. T h e twin sciences of ore deposits and the study of hydrothermal systems, both largely concerned with such fluid-rock interactions, have been profoundly and completely transformed. Virtually no issue of Economic Geology now appears w i t h o u t 3 or 4 papers dealing with stable isotope variations. No one writes papers on the development of the hydrosphere, hydrothermal alteration, ore deposits, melt-fluid-solid interactions, etc. without taking into account the ideas and concepts of stable isotope geochemistry. Although the present volume represents only a first effort to fill the need for a general survey of this sub-field for students and for workers in other disciplines, and although it is still obviously not completely comprehensive, it should give t h e interested s t u d e n t an idea of the present 'state-of-the-art' in the field. It should also provide an entry into the pertinent literature, as well as some understanding of the basic concepts and potential applications. Some t h o u g h t went into the arrangement and choice of chapters for this volume. T h e first three chapters focus on the theory and experimental d a t a base for equilibrium, disequilibrium, and kinetics of stable isotope exchange reactions among geologically i m p o r t a n t minerals and fluids. T h e fourth chapter discusses the primordial oxygen isotope variations in the solar system prior to formation of the E a r t h , along with a discussion of isotopic anomalies in meteorites. T h e fifth chapter discusses isotopic variations in the E a r t h ' s mantle and the sixth chapter reviews the variations in the isotopic compositions of n a t u r a l waters on our planet. In C h a p t e r s 7, 8, 9 and 10, these isotopic constraints and concepts are applied to various facets of the origin and evolution of igneous rocks, bringing in much material on radiogenic isotopes as well, because these problems require a multi-dimensional attack for their solution. In C h a p t e r s 11 and 12, the problems of hydrothermal alteration by meteoric waters and ocean water are considered, together with discussions of the physics and chemistry of hydrothermal systems and the l s O / l e O history of ocean water. Finally, in C h a p t e r s 13 and 14, these concepts are applied to problems of metamorphic petrology and ore deposits, particularly

vi

with respect to the origins of the fluids involved in those processes. It seems clear to us (the editors) t h a t this sub-field of stable isotope geochemistry can only grow and become even more pertinent and dominant in the f u t u r e . One of the most fruitful areas to pursue is the development of microanalytical techniques so t h a t isotopic analyses can be accurately determined on ever smaller and smaller samples. Such techniques would open up vast new territories for exploitation in every aspect of stable isotope geochemistry. Exciting new methods have recently been developed whereby a few micromoles of C 0 2 and SO s can be liberated for isotopic analyses from polished sections of carbonates and sulfides by laser impact. There are also new developments in m a s s spectrometry like RIMS (resonance ionization mass spectrometry), Fourier t r a n s f o r m mass spectrometry and the ion microprobe t h a t offer considerable promise for these purposes. Stable isotope analyses of large-sized samples (even those t h a t must be obtained by reactions of silicates with fluorinating reagents) have now become so routine and so rapid t h a t they represent an 'easy' way to gather a lot of d a t a in a hurry. In fact 'mass production' techniques for rapidly processing samples are starting to become prevalent, so much so t h a t one of the biggest worries in the f u t u r e may be t h a t a flood of d a t a will overwhelm us and outstrip our abilities to carefully define and carry out sampling strategies, as well as to think carefully and in depth about the data. An organized system of handling the D / H , 1 3 C/ 1 2 C, 1 6 N / U N , 1 8 0 / l e 0 , and 3 4 S/ 3 2 S data, a n d / o r a computerized d a t a base t h a t could be manipulated and added to would be a useful p a t h to follow in the future, particularly if it were integrated into a larger d a t a base containing radiogenic isotope data, major- and trace-element analyses, electron microprobe data, x-ray crystallographic data, and petrographic d a t a (particularly modal d a t a on mineral abundances in the rocks). T h e editors wish to t h a n k the authors for meeting most of the deadlines associated with t h e preparation of this volume and for preparing the final typed versions of their manuscripts. We also t h a n k all the individuals who graciously gave of their time and energy in reviewing and helping to improve the various manuscripts: R. Becker, P. Brown, R. Criss, B. Giletti, C. Johnson, B. Marsh, T. Mayeda, W. McKenzie, J. Morrison, P. Nabelek, E. Ripley, R. Rye, S. Savin, D. Stakes, L. T o r a n , and H. Wang. A particularly noteworthy contribution to this volume was made by Robert E. Criss whose careful reviews and critical comments substantially improved several chapters. We also t h a n k A. G u n t h e r , M. Hass, L. Marnoch, L. McMonagle, S. Morris, J. Murochick, and M. Wilson for their major editorial and typing efforts. W e are greatly indebted to Paul H. Ribbe, who as series editor of Reviews in Mineralogy, was responsible for preparing camera-ready copy, and to B a r b a r a Minich at M.S.A. Headquarters, who provided much needed assistance on various logistical and administrative problems. We are also grateful to M. Strickler and C. Ribbe for their assistance with additional typing, drafting and editorial work, all carried out at the D e p a r t m e n t of Geological Sciences, Virginia Polytechnic Institute and S t a t e University.

James R. O'Neil Menlo P a r k , California

Hugh P. Taylor, Jr. Pasadena, California September 1, 1986

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John W. Valley Madison, Wisconsin

TABLE of CONTENTS Page ii iii v

COPYRIGHT; ADDITIONAL COPES FOREWORD; DEDICATION PREFACE

Chapter 1 James R. O'Neil THEORETICAL and EXPERIMENTAL ASPECTS of ISOTOPIC FRACTIONATION 1 2 2 3 4 5 7 7 8 9 9 11 13 13 16 18 19 20 20 22 23 23 24 25 28 28 28 30 31 33 36 37

INTRODUCTION KINETIC and EQUILIBRIUM ISOTOPE EFFECTS Kinetic Isotope Effects Equilibrium Isotope Effects The FRACTIONATION FACTOR The PARTITION FUNCTION Translational Partition Function Rotational Partition Function Vibrational Partition Function CALCULATION of EQUILIBRIUM CONSTANTS for ISOTOPE EXCHANGE REACTION Gases Condensed Phases FACTORS INFLUENCING the SIGN and MAGNITUDE of a Temperature Chemical Composition Crystal Structure Press ure LABORATORY DETERMINATIONS of ISOTOPIC FRACTIONATION FACTORS Two-direction Approach Sample calculation ofl&lna Pseudo isotope exchange reactions Problems with the two-directional method Partial Exchange Technique Three-Isotope Method Stable Isotope Fractionation Curves STABLE ISOTOPE THERMOMETRY Scope Tests for Equilibrium Present Status CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

Chapter 2 David R.Cole & Hiroshi Ohmoto KINETICS of ISOTOPIC EXCHANGE at ELEVATED TEMPERATURES and PRESSURES 41 42 42 44 47

INTRODUCTION BASIC CONCEPTS in ISOTOPE EXCHANGE REACTIONS Homogeneous versus Heterogeneous Reactions Rate Law for Isotope Exchange Reactions Determination of Rf from F and Its Relation to kf, the True Rate Constant viii

49 52 55 57 58 58 60 61 63 63 74 14 76 77 78 81 83 83 84 86 87

Determination of D, the Diffusion Coefficient The bulk exchange technique Microbeam analytical techniques MECHANISMS and RATES of ISOTOPE EXCHANGE in HOMOGENEOUS SYSTEMS Kinetics of Isotopic Exchange Reactions in Solutions The sulfate-sulfide system The sulfate-water system Kinetics of Isotopic Exchange Reactions Between Gases MECHANISMS and RATES of ISOTOPE EXCHANGE REACTIONS in HETEROGENEOUS SYSTEMS Isotope Exchange Accompanying Surface Reactions Rate model Rates and activation parameters The relationship between rf and water/solid ratio Isotope Exchange Accompanying Diffusion The effect of temperature on rates of diffusion Effects of pressure on the rates of diffusion Anisotropy in diffusion Water in minerals Comparison of the Surface-Reaction and Diffusion Models SUMMARY ACKNOWLEDGMENTS

Chapter 3 Robert T. Gregory & Robert E. Criss ISOTOPIC EXCHANGE in OPEN and CLOSED SYSTEMS 91 92 92 92 93 93 93 93 97 98 99 99 101 101 101 101 103 103 103 104 106 107 107 108 110 111 111 111 112 113

INTRODUCTION BASIC PRINCIPLES Isotopic Exchange Reactions Delta Space Conservation of Mass CLOSED SYSTEMS General Statement Representation in Delta Space Temperature Effects Transformation to Other Coordinate Systems A-A plots S-Aplots Isotherm plots Summary: Closed Systems OPEN SYSTEMS General Statement Kinetic Effects Elementary rate law "Closed" system exchange model Open system exchange model "Buffered" open system exchange model Interpretation of the Kinetic Models Fluid!rock ratios Exchange trajectories versus isochronous arrays Calculation of effective fluid/rock ratios from disequilibrium arrays Summary: Open Systems APPLICATIONS to NATURAL SYSTEMS General Statement Plagioclase-Pyroxene from Layered Gabbros Primary magmatic compositions ix

113 114 114 114 116 120 120 122 123 124 124 125 126

Temperatures of fluid-rock interactions Fluid!rock ratios Fluid isotopic compositions and pathlines Mineral Pairs from Granitic Rocks Precambrian Siliceous Iron Formation: Quartz-Magnetite Mineral Pair Systematics Applied to Upper Mantle Assemblages Eclogite mineral pairs Peridotite mineral pairs Impact of an open system model on the evolution of the upper mantle Summary: Natural Systems CLOSING STATEMENT ACKNOWLEDGMENTS REFERENCES

Chapter 4 Robert N. Clayton HIGH TEMPERATURE ISOTOPE EFFECTS in the EARLY SOLAR SYSTEM 129 129 129 131 131 132 132 134 137 139

INTRODUCTION PLANETARY PROCESSES Achondrites and the Moon Ordinary Chondrites Carbonaceous Chondrites NEBULAR PROCESSES Evaporation and Condensation Isotopic Anomalies FUN Inclusions REFERENCES

Chapter 5 T. Kurtis Kyser STABLE ISOTOPE VARIATIONS in the MANTLE 141 142 142 144 146 152 152 155 157 158 160 162

INTRODUCTION OXYGEN ISOTOPE COMPOSITIONS Mafic Lavas Mantle Xenoliths Processes Controlling the 1 8 0/ 1 6 0 Ratio of the Mantle Ancient Oxygen Isotope Compositions CARBON NITROGEN SULFUR HYDROGEN CONCLUDING REMARKS REFERENCES

Chapter 6 Simon M.F. Sheppard CHARACTERIZATION and ISOTOPIC VARIATIONS in NATURAL WATERS 165 165 166 167 167

INTRODUCTION Concept of Combined Isotope Approach Definition of Principal Water Types ISOTOPIC CHARACTERISTICS of NATURAL WATERS Ocean Waters X

167 167 168 168 170 172 173 174 175 177 178 178 179 180 181

Present-day Ancient Meteoric Waters Present-day Ancient Connate and Formation Waters Metamorphic Waters Magmatic Waters Organic Waters Hydrothermal Waters Exotic Waters Other Waters CONCLUSIONS ACKNOWLEDGMENTS REFERENCES Chapter 7

185 185 185 185 186

187 187 187 189 190 190 190 190 190 191 195 195 196 196 196 197 198 198 198 201 201 201 204 204 204 207 207 207 207 211 213 213

Bruce E. Taylor M A G M A T I C VOLATILES: ISOTOPIC VARIATION o f C , H, and S

INTRODUCTION SOLUBILITY, SPECIATION, and CHEMICAL COMPOSITION Solubility and Speciation Carbon Hydrogen Sulfur Composition of Magmatic Gases C, H, and S and Mafic Glasses C, H, and S in Felsic Glasses ISOTOPIC FRACTIONATION and MIGRATION of VOLATILES Isotopic Fractionation Carbon Hydrogen HjO and hydrous minerals H20 and hydrous magmas Sulfur Isotopic Variation During Magmatic Processes Contamination-assimilation Contamination-exchange Degassing Volatile Migration in Magmas CARBON ISOTOPES Volcanic Gases Carbon dioxide Methane Carbon Dioxide in Vesicles Carbon Isotopes in Basaltic Glass Carbon in Mafic and Felsic Rocks HYDROGEN ISOTOPES Volcanic Gases HjO in Fluid Inclusions Mafic Magmas Felsic Magmas Hydrogen Isotopes in Basaltic Glasses Hydrogen Isotopes in Felsic Plutons SULFUR ISOTOPES Volcanic Gases xi

214 218 218 219 220

Basalts Andésites CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

Chapter 8

Hugh P.Taylor, Jr. & Simon M.F. Sheppard IGNEOUS ROCKS: I. PROCESSES of ISOTOPIC FRACTIONATION and ISOTOPE SYSTEMATIC S

227 229 230 230 231 233 233 234 237 239 241 241 242 244 246 249 249 252 252 252 252 253 253 253 253 254 254 254 254 256 256 259 260 260 260 263 265 266 268 269

INTRODUCTION GENERAL 1 8 0/ 1 6 0 and D/H VARIATIONS in IGNEOUS ROCKS PRIMORDIAL ISOTOPE RATIOS Oxygen Isotopes Hydrogen Isotopes CLOSED-SYSTEM EFFECTS Rayleigh Fractionation versus Equilibrium Crystallization Equilibrium 1 8 0/ 1 6 0 Fractionations between Crystals and Melt Evidence from Natural Igneous-Rock Suites Mineral-Melt D/H Fractionations and Magmatic Differentiation Effects OPEN-SYSTEM PROCESSES The Effects of Assimilation Oxygen Isotope Effects Radiogenic Isotope Effects 87 Sr- ß O Effects during Assimilation-Fractional Crystallization Adamello Massif, Northern Italy EXCHANGE EFFECTS at the MARGINS of MAGMA BODIES SOURCE-ROCK RESERVOIRS AND MELT GENERATION Isotopic Compositions of Possible Source Rocks Oceanic crust Sedimentary rocks Metasedimentary rocks Archean cratons Granulite-facies lower continental crust Upper mantle Preexisting igneous rocks Low-180 rocks formed by meteoric-hydrothermal alteration Formation waters or marine waters Source Contamination versus Crustal Contamination Summary LOW- l s O MAGMAS LOW- and HIGH-DEUTERIUM MAGMAS The VOLCANIC ROCKS of ITALY General Features Alban Hills Volcanic Center M. Vulsini Volcanic Center Mixing Models Involving Tuscan Basement Rocks Summary ACKNOWLEDGMENTS REFERENCES

Chapter 9 Hugh P. Taylor, Jr. IGNEOUS ROCKS: II. ISOTOPIC CASE STUDIES of CIRCUMPACIFIC MAGMATISM 273

INTRODUCTION xii

273 273 277 278 279 281 284 286 286 288 291 292 294 294 295 296 297 300 300 300 301 303 304 306 306 306 308 312 314 314 315 315 316

CALC-ALKALINE VOLCANIC ROCKS in the ANDEAN CORDILLERA 18 0 / 1 6 0 Ratios in Late Cenozoic Andean Volcanic Rocks The Central Volcanic Zone (CVZ) in Peru The Northern Volcanic Zone (NVZ) in Colombia and Ecuador Comparison of the NVZ and the Peruvian CVZ Comparison of the Southern CVZ with the SVZ and AVZ Summary The PENINSULAR RANGES BATHOLITH (PRB) of SOUTHERN and BAJA CALIFORNIA General Statement 18 0 / 1 6 0 Ratios Other Isotopic and Geochemical Gradients Origin of the Isotopic Variations in the PRB San Jacinto-Santa Rosa Mountains Block OTHER CRETACEOUS GRANITIC B ATHOLITHS in the UNITED STATES Relationship to the Peninsular Ranges Batholith Regional Isotopic Systematics Idaho Batholith Summary OXYGEN and STRONTIUM ISOTOPE STUDIES of WESTERN PACIFIC ISLAND ARCS Geologic Setting and Sources of Data Mariana and Volcano Arcs Japan and Izu Arcs SiO, versus ICO ORIGIN of LOW^ s O RHYOLITE MAGMAS in WESTERN NORTH AMERICA General Statement Catastrophic Isotopic Changes in Magmas During Caldera Collapse, Yellowstone Volcanic Field Origins of the Low- l s O Magmas in the Yellowstone Caldera Complex L o w - 1 8 0 Rhyolite Magmas Elsewhere in the Western U.S.A. The Low- l s O Magma Problem Emplacement into a rift-zone tectonic setting Chemical composition ACKNOWLEDGMENTS REFERENCES

Chapter 10 Simon M.F. Sheppard IGNEOUS ROCKS: III. ISOTOPIC CASE STUDIES of MAGMATISM in AFRICA, EURASIA and OCEANIC ISLANDS 319 319 319 323 324 326 329 332 335 336 337 340 342 342 344 346

INTRODUCTION ANOROGENIC MAGMATISM Oceanic Islands North Atlantic or Brito-Arctic Igneous Province Primary unmodified magmatic isotopic compositions Low- O magmas Gardar Igneous Province Granitic Rocks of East and South China Alkaline Ring - Complexes in Africa and Arabia Arabian complexes Ring-complexes of Cameroun and Nigeria Summary OROGENIC MAGMATISM Plutonic Belts of the Himalaya-Transhimalaya Transhimalaya batholith High, north and "Lesser Himalaya" belts xiii

349 349 353 356 358 359 362 365 367 368

Variscan Magmatism 18 16 0l 0 ratios and type of magmatism Nature of source materials Caledonian Magmatism of Northern Britain Younger basic intrusions 'Older' and 'younger' granites Archaean Granites of Southern Africa Concluding Remarks ACKNOWLEDGMENTS REFERENCES

Chapter 11 Robert E. Criss & Hugh P. Taylor, Jr. METEORIC-HYDROTHERMAL SYSTEMS 373 374 374 375 379 383 387 387 387 391 393 393 395 397 397 400 402 405 407 408 411 411 413 413 415 417 418 420 421 422

INTRODUCTION FLUID DYNAMICS in a PERMEABLE MEDIUM Basic Principles Free Convection Permeability and Porosity Theoretical Scaling Law, Cooling Times, and Water/Rock Ratios MODERN METEORIC-HYDROTHERMAL SYSTEMS Geologic Settings of Geothermal Systems Origin and Composition of Geothermal Fluids Physical State of Geothermal Fluids ISOTOPIC EFFECTS in HYDROTHERMALLY-ALTERED ROCKS Stable Isotopic Systematics of Altered Rocks Effects of Alteration on Geochronologic Systems FOSSIL METEORIC-HYDROTHERMAL SYSTEMS General Occurrence and Character Isotopic Relationships in the Skaergaard Intrusion Numerical Modeling of the Skaergaard Intrusion Water/Rock Ratios in the Skaergaard Intrusion Lake City Caldera Idaho Batholith Yankee Fork District Meteoric-Hydrothermal Ore Deposits PETROGRAPHIC, CHEMICAL, and PHYSICAL ASPECTS of ROCK ALTERATION Volcanic Country Rocks Granitic Plutons Layered Gabbro Bodies Contrasting Effects in Gabbros and Granites SUMMARY ACKNOWLEDGMENTS REFERENCES

Chapter 12 Karlis Muehlenbachs ALTERATION of the OCEANIC CRUST and the l s O HISTORY of SEAWATER 425 426 426 428 428 429

INTRODUCTION CURRENT SEA FLOOR PROCESSES Low Temperature Processes High Temperature Processes Metavolcanic rocks Dikes xiv

429 431 431 435 436 437 439 443

Gabbros Plagiogrcmites ON-LAND EXPOSURES of the SEA FLOOR MID-OCEAN RIDGE HOT SPRINGS MODELLING OF MID-OCEAN RIDGE HYDROTHERMAL SYSTEMS ls O BUDGET at MID-OCEAN RIDGES 8 1 8 0 of Ancient Oceans REFERENCES

Chapter 13 JohnW.Valle: STABLE ISOTOPE GEOCHEMISTRY of METAMORPHIC ROC 445 445 446 447 448 450 451 451 452 454 458 461 461 463 464 465 465 465 467 470 471 473 475 475 476 478 480 481 486

INTRODUCTION ISOTOPIC THERMOMETRY METAMORPHIC VOLATILIZATION Batch Volatilization Rayleigh Volatilization Dehydration Decarbonation Mixed Volatile Reactions Coupled O-C Depletions CONTACT METAMORPHISM Volatilization During Contact Metamorphism Fluid Infiltration During Contact Metamorphism Controls ofpermeability Calculation offluid/rock ratios Skarn The Effects of Variable P-T or Disequilibrium Polythermal exchange Graphitization REGIONAL METAMORPHISM Fluid Convection During Regional Metamorphism Isotopic Exchange by Diffusion and Recry stallization The Scale of Oxygen Isotope Exchange During Regional Metamorphism Granulite Facies Metamorphism The role of fluids in granulite genesis 8lsO and Sf3C in granulites The Adirondacks as a case study Low- and high-180 granulites; low-I80 eclogites ACKNOWLEDGMENTS REFERENCES

Chapter 14 Hiroshi Ohmoto STABLE ISOTOPE GEOCHEMISTRY of ORE DEPOSITS 491 493 493 494 494 494 495 495 496 498

INTRODUCTION APPLICATIONS of STABLE ISOTOPES as GEOTHERMOMETERS Hydrogen Systems Oxygen Systems Carbon Systems Sulfur Systems HYDROGEN and OXYGEN ISOTOPE GEOCHEMISTRY of HYDROTHERMAL SYSTEMS Methods of Estimating the 8D and 5 1 8 0 of Ore Fluids Hydrogen and Oxygen Isotopic Characteristics of Reference Waters Hydrogen and Oxygen Isotopic Characteristics of Recycled Waters XV

502 502 505 505 506 506 506 507 508 508 509 509 512 513 513 513 518 520 523 528 528 530 532 534 537 537 539 540 542 544 545 551 553 555 556

Isotopie Zoning in Wall Rocks Problems in the Quantification of Water/Rock Ratios Quantification of Hydrological and Geochemical Nature of Ore-forming Systems CONTRASTS BETWEEN CARBON-SULFUR and HYDROGEN-OXYGEN SYSTEMATICS Causes of Isotopie Variation Isotopie Equilibrium Isotopie Effects during Mineralization Multiple Sources CARBON ISOTOPE GEOCHEMISTRY of HYDROTHERMAL SYSTEMS Isotopie Variation during Geochemical Cycles of Carbon in Near-surface Environments Carbon Isotopie Composition of the Mantle Isotopie Relationships among Aqueous Carbon Species Methods of Determining the Sources of Carbon in Ore Deposits SULFUR ISOTOPE GEOCHEMISTRY of HYDROTHERMAL SYSTEMS Sulfur Isotopie Characteristics of Reference Reservoirs Sulfur Isotopie Characteristics of Recycled Seawater-sulfur Isotopie Relationships Among Aqueous Sulfur Species Sulfur Isotopie Relationships Between Fluid Species and Minerals and Between Co-existing Minerals Methods of Determining the Sources of Sulfur in Ore Deposits SULFUR ISOTOPES in MAGMATIC SYSTEMS Speciation and Solubility of Sulfur in Silicate Melts 5 S Values of Mantle-derived Igneous Rocks Assimilation of Crustal Sulfur by Mafic Magmas and the Formation of Cu-Ni Sulfide Ores Assimilation of Crustal Sulfur by Felsic Magmas Sulfur Isotopie Characteristics of Magmatic Fluids GENETIC MODELS for ORE DEPOSIT TYPES Porphyry and Skam Type Deposits Base- and Precious-metal Veins, and Replacement Deposits Massive Sulfide Deposits of Submarine Volcanic Association Shale/Carbonate-hosted Zn-Pb Deposits Mississippi Valley-type Deposits Red-bed Associated Cu Deposits CONCLUDING REMARKS ACKNOWLEDGMENTS REFERENCES

Appendix: TERMINOLOGY and STANDARDS 561 561 561 562 563 564 564 565 565 565 565 567 567 569 569

TERMINOLOGY The 5 Value Isotope Exchange Reactions The Fractionation Factor, a 103/«a and the A Value The e Value Fractionation Curves Crossovers Reversals STANDARDS Oxygen Hydrogen Carbon Sulfur REFERENCES xvi

James R. O'Neil

Chapter 1

James R. O'Neil

THEORETICAL and EXPERIMENTAL ASPECTS of ISOTOPIC FRACTIONATION INTRODUCTION Light stable isotope geochemistry is concerned with variations in the isotopic ratios of only six elements: H, C, N, O, Si, and S. Boron is sometimes added to this list, b u t its isotopic ratios are measured by techniques t h a t are different f r o m those used in conventional stable isotope laboratories. These elements have several characteristics in common: (1) T h e y have low atomic mass. Isotopic variations have been looked for b u t not found for heavy elements like Cu, Sn, and Fe. (2) T h e relative mass difference between the rare (heavy) and a b u n d a n t isotope is large. For example, compare the values of 12.5 and 8.3 percent for the pairs 1 8 0 - 1 6 0 and 13 C- 12 C, respectively, with the value of only 1.2 per cent for 87 Sr- 86 Sr. T h e relative mass difference between D and H is almost 100 percent and hydrogen isotope fractionations are, accordingly, about ten times larger t h a n those of the other elements of interest. (3) T h e y form chemical bonds t h a t have a high degree of covalent character. A t t e s t i n g to the importance of bond type to isotopic fractionation, the 4 8 C a / 4 0 C a ratio varies little in terrestrial rocks despite the large relative mass difference (only D-H is larger) between the isotopes (Russell et al., 1978). (4) T h e elements exist in more t h a n one oxidation s t a t e (C, N, and S), form a wide variety of compounds (notably O), and are i m p o r t a n t constituents of naturally occurring solids and fluids. (5) T h e abundance of the rare isotope is sufficiently high (tenths to a few per cent) " to assure the ability to make precise determinations of the isotopic ratio by mass spectrometry. Depending on the instrument used, the analytical error of deuterium analyses is up to ten times larger than those of the other elements because of the low a b u n d a n c e of deuterium (about 160 ppm) in nature. N a t u r a l variations in isotopic ratios of terrestrial materials have been reported for other light elements like Mg and K, b u t such variations usually t u r n o u t to be laboratory artifacts. T h e case of magnesium is fairly straightforward. Aside f r o m the fact t h a t its bonds are dominantly ionic in character, magnesium is almost always surrounded by the same atomic environment (an octahedron of oxygen) in nature. T h u s with little or no possibility of site preference in magnesium compounds, conditions are not favorable for isotopic fractionation of this element in nature. In any event, variations in stable isotope ratios of light elements other than the seven mentioned above are small in terrestrial substances. T h e reasons for this are not completely understood and are only loosely discussed in terms of characteristics such as those noted above. These characteristics are only observed and are not rigorously tied to theoretical principles. Except for the case of stable isotope relations in extraterrestrial materials ( C h a p t e r 4), the discipline of stable isotope geochemistry deals only with those isotopic variations or effects t h a t arise either from isotope exchange reactions or from mass-dependent fractionations t h a t accompany physical and chemical processes occurring in n a t u r e or in the laboratory. While ultimately q u a n t u m mechanical in origin, such isotope effects are governed by kinetic theory and the laws of thermodynamics. N a t u r a l variations in the stable isotope ratios of the heavy elements of geologic interest like Sr, Nd, and P b involve nuclear reactions and are governed by other factors including the ratio of radioactive p a r e n t and daughter, decay constants, and time. Essential to the interpretation of natural variations of light stable isotope ratios is knowledge of the magnitude and temperature dependence of isotopic fractionation factors between the common minerals and fluids. These fractionation factors are obtained in three ways: 1

(1) Semi-empirical calculations using spectroscopic d a t a and the methods of statistical mechanics. (2) Laboratory calibration studies. (3) Measurements of n a t u r a l samples whose formation conditions are well-known or highly constrained. In this chapter methods (1) and (2) are evaluated and a review is given of our present s t a t e of knowledge of the theory of isotopic fractionation and the factors t h a t influence the isotopic properties of minerals. For f u r t h e r information on nuclear properties and isotopes in general, consult the texts by Faure (1977) and Hoefs (1980).

K I N E T I C A N D EQUILIBRIUM I S O T O P E E F F E C T S Kinetic Isotope Effects Kinetic isotope effects are common both in n a t u r e and in the laboratory and their magnitudes are comparable to and sometimes significantly larger t h a n those of equilibrium isotope effects. Kinetic isotope effects are normally associated with fast, incomplete, or unidirectional processes like evaporation, diffusion, and dissociation reactions. T h e examples of diffusion and evaporation are explained by the different translational velocities of isotopic molecules moving through a phase or across a phase boundary. Kinetic theory tells us t h a t the average kinetic energy (K.E.) per molecule is the same for all ideal gases at a given temperature. Consider the isotopic molecules 12 le C O and 1 2 C l s O t h a t have molecular weights of 28 and 30, respectively. Solving the expression equating the kinetic energies (K.E. = 1/2 Mv 2 ) of both isotopic species, the ratio of velocities of the light to heavy isotopic species is (30/28) 1 / 2 , or 1.034. T h a t is, regardless of T , the average velocity of 1 2 C l e O molecules is 3.4 percent greater than the average velocity of 1 2 C l s O molecules in the same system. T h i s and other such velocity differences lead to isotopic fractionations in a variety of ways. Isotopically light molecules can preferentially diffuse out of a system and leave the reservoir enriched in the heavy isotope. In the case of evaporation, the greater average translational velocities of lighter molecules allows t h e m to break through the liquid surface preferentially, resulting in an isotopic fractionation between vapor and liquid. For example, the ft 1/ < v E'

< E

Relative to the dissociated atoms, molecules containing the heavy isotope are more stable, i.e. have greater binding energies, than molecules with the lighter isotope. This is a quantum mechanical effect, depending on the existence of the ZPE. At higher temperatures molecules are excited to higher vibrational levels (n > 0), and the difference in energy between isotopic species diminishes. Note that the potential energy curve for a real molecule like hydrogen becomes increasingly different from the harmonic oscillator curve as the temperature increases (Fig. 1). Inasmuch as the harmonic oscillator approximation plays a key role in the development of the theory of isotopic fractionation, it is well to keep in mind that 'anharmonicity' corrections may be important considerations, particularly at high temperatures. The quantitative treatment of the effect of temperature requires the methods of statistical mechanics and will be discussed in a later section. It is important to point out that the free energy differences associated with isotope effects (either chemical or physical) are about 1000 times smaller than typical bond energies or heats of chemical reactions. Take, for example, the isotope exchange reaction between C 0 2 and H 2 0 (written such that one atom of oxygen is exchanged between the two molecules): i/2C 1 6 0 2 + H 2 18 0 = I/2C 1 8 0 2 + H 2 16 0 The equilibrium constant for this exchange reaction at 25°C is 1.0412. The free energy change, given by -RTlnK, is only -23.9 cal/mole, a typical value for isotope exchange reactions at low temperature. Clearly such low free energy differences could never be a driving force for chemical reactions. On the other hand it seems reasonable that isotopic exchange equilibrium should be established during reactions whose products are in chemical (mineralogical) equilibrium. In fact demonstration that the minerals in a rock are in oxygen isotope equilibrium is strong evidence that the rock is in chemical equilibrium. After all, conditions so profound as to break and reform Al-O and Si-O bonds to allow equilibrium distribution of l s O among the minerals should be sufficient to effect chemical equilibrium in the system.

THE FRACTIONATION FACTOR The fractionation factor between two substances A and B, a A . B , is defined as follows: 4

a A-B

=

-Kjk.b

.

(3)

where R A = (D/H) A , ( I 3 C/ 1 2 C) A , ( 1 8 0 / 1 8 0 ) A , etc. A t equilibrium, a is related, to the very good approximation that the isotopes are randomly distributed among all possible sites in the molecule, to the equilibrium constant K for the isotope exchange reaction between the two substances (Bigeleisen, 1955). In general, a =

K 1 /"

,

(4)

where n is the number of atoms exchanged in the reaction as written. It is a factor analogous to a distribution coefficient and is the most important quantity used in evaluating stable isotope variations observed in nature. Different authors report equilibrium fractionation factors as a, lna, 10 3 lna, K , InK, e, and A (see Appendix 1). Each of these symbols will be used in this chapter in order to familiarize the reader with all usages. A s we saw above, isotopic fractionations can be described in terms of isotope exchange reactions, an example of which is: 12co

+

13

CH4 =

13co

12

CH4

+

.

T h e equilibrium constant K for this reaction is written in the usual way: =

(13CO)(12CH4) (12CO)(13CH4)

•

1

'

Concentrations are used rather than activities or fugacities because ratios of activity coefficients for isotopically substituted molecules are equal to unity. W h e n there is more than one atom of the isotopic element in the chemical formula, as in Si 1 8 0 2 or CD 4 , all atoms in the substance are the isotope indicated. T h i s is by custom and meaningful only for simplifying calculations. Substances with such chemical formulas do not exist in nature. For isotope exchange reactions, the total energy change is just the difference in A Z P E between the t w o molecular species. A s mentioned above, a typical value for the change in energy of an isotope exchange reaction is only about 20 cal/mole. N o t e that like all equilibrium constants, the equilibrium constants for isotope exchange reactions are temperature dependent. T h i s is the basis for their use in thermometry.

THE PARTITION FUNCTION T h e principal thermodynamic functions, including the equilibrium constant, can be expressed in terms of the partition function, a mathematical relation arising from statistical mechanics. Partition functions contain all the energy information about a molecule and they are used to calculate equilibrium constants (fractionation factors) for isotope exchange reactions. Recall that the internal energy of a molecule E i n l can, to a first approximation, be represented as a simple sum of all forms of energy: translation (tr), rotation (rot), vibration (vib), electronic (el) and nuclear spin (sp). Eint =

E t r + Erot + Evib + Eel + Esp

.

(6)

For our purposes the last t w o terms are negligible and will be ignored in the treatment that follows. A t any instant, the value of E int and the distribution of internal energy among the individual terms may vary in different molecules of the same kind. In an equilibrium assemblage at temperature T , the fraction of molecules having energy E iDt can be expressed as ^ = no

gie"E^T

5

,

(7)

where n 0 = nE = g. = k'=

number of molecules with zero-point energy only number of molecules having E i n t statistical weight Boltzmann's constant = 1.381 x 10 1 9 erg/°K.

It is possible t h a t there may be more than one state corresponding to the energy level E.. If so, the level is said to be 'degenerate' and must be assigned a statistical weight (g.) that is equal to the number of superimposed levels. The summation of all terms on the right hand side of equation 7 is by definition the partition function Q, also called the 'sum-over-states,' and the 'statistical sum': Q = £g,e"E'/kT

•

(8)

This function is first met in statistical mechanics as the denominator of the Boltzmann distribution law written in its most general form. The average energy of an assemblage of molecules will be the ordinary internal energy E. If the allowed energies of the whole system are E^ E 2 , ...E P the average energy will be £niEi

EgiEie-

E /kT

'

=

•

Substituting Q for the denominator in equation 9 and recasting in terms of partial derivatives of Q with respect to T , we have 2

E = k T

^ . dT

(10) '

v

and one connection between the thermodynamic functions and partition functions is seen. Statistical mechanics deals with large numbers of individual particles whereas thermodynamics deals with large-scale systems containing very many particles, the usual measure being the mole, or 6.02 x 1023 molecules. T h u s a distinction is made between molecular and molar partition functions. You will often see the thermodynamic functions written in terms of molar partition functions (and R rather than k). For example, S =

+ RlnQ

.

(11)

Equilibrium constants for any reaction can be written in terms of the partition functions of the reactants and products. K is written in the normal way except t h a t Q is used rather than the activity. For the carbon isotope exchange reaction between CO and CH 4 discussed above, the equilibrium constant can be written K

_

Q( 1 3 CO)/Q( 1 2 CO) Q( 1 3 CH 4 )/Q( 1 2 CH 4 )

(12]

For the general case of an isotope exchange reaction between substances A and B, aA x + bB 2 = aA 2 + bBj

,

(13)

an expression can be written for the equilibrium constant, or fractionation factor, as follows: R

=

(QA,) a (Q Bl ) b (Q A i r(QB 2 )

b

=

(Qa/QtK (Q 2 /Qi) B b

'

where the subscripts 2 and 1 refer to molecules totally substituted by the heavy and light isotopes of an element, respectively. In order to calculate equilibrium constants for isotope exchange reactions, our ultimate goal will be to calculate partition function ratios for the different substances of interest. T o begin with we will write expressions for the partition functions (and partition function ratios) of diatomic molecules that are (1) ideal gases, (2) rigid rotors, and (3) harmonic oscillators. We will then generalize to the polyatomic case, and finally, introduce the complications of treating condensed phases. 6

It follows from equation 6, and to the same approximation, that the total partition function for an ideal gas is equal to the product of partition functions for each form of energy: Q total =

QtrQrotQvib •

(15)

Partition functions have been developed for each of these forms of energy and they will be examined in turn. Through the use of very good approximations, we will be able to express (^¡/C^ as a function only of vibrational frequencies and temperature. Translational Partition Function All forms of energy are quantized; that is, they can assume only certain discrete values. Above about 2°K, the quanta of translational energy are so small, and the energy levels are so closely spaced, that the summation in the expression for the translational partition function can be evaluated as an integral Qtr =

EStre~E"/kT =

/ e - E " / k T dn

,

(16)

where dn is the number of energy levels in the energy range dE. From quantum mechanics we know that the energy of translational motion of a particle in a cubical box of side a is, for each degree of freedom,

and E,%r kT where, if T =

(18)

8a2mkT

constant, X=

—= 8a m k T

constant

.

Substituting these values for E t r / k T into equation 16, we have for each degree of freedom Qtr =

Je^dn =

i

y

f

=

.

(19)

For three degrees of freedom, =

(2™kT)3/*v

h13 where V = a 3 , the volume of the cube. Furthermore the result is valid for a volume of any shape. We may replace m by the molecular weight M, and when ratios are made to obtain the translational partition function ratio, everything cancels but the molecular weights. (Qi/QlJtr = ( M ! / m 2 ) 3 / 2

.

(21)

Thus the isotopic fractionation arising from translation is a function only of the ratio of molecular weights raised to the 3 / 2 power and is independent of temperature. Rotational Partition Function The solution to the Schrodinger equation for a diatomic molecule in the rigidrotator model (consider the molecule as a dumbbell rotating about a center of mass) allows the following energies: E

=

j(J + Dh 2 87rV2

=

7

j(j + l ) h 2 S^2!

(22)

where j = rotational q u a n t u m number I = m o m e n t of inertia = fir2 H = reduced m a s s r = interatomic distance T h e r e are two complications t h a t enter into the rotational partition function: the degeneracy of the rotational levels and the s y m m e t r y of the molecular eigenfunctions. Diatomic molecules have t w o axes of rotation around which r o t a t i o n a l energy can be distributed. T h e number of w a y s of distributing j q u a n t a of energy between the two axes equals 2j + 1, because in every case except j = 0 there are two possible alternatives for each a d d e d q u a n t u m . T h e statistical weight of a rotational level j is therefore 2j + 1. T h e rotational partition function now becomes Qrot =

E(2j + l)e*

+

1

»h2/8*2lkT

•

(23)

E x c e p t for the case of hydrogen isotopes, the spacings between energy levels are less than k T . A g a i n because of these small s p a c i n g s the s u m can be a p p r o x i m a t e d by an integral and e v a l u a t e d in closed form to give the classical partition function 8;r 2 IkT —72— h

n

Qrot =

'

(24)

In homonuclear diatomic molecules ( 1 4 N 1 4 N , 1 6 O l e O , etc.) only all o d d or all even j ' s are allowed, depending on the s y m m e t r y properties of the molecular eigenfunctions. If the nuclei are different ( H N 1 6 N , 1 6 0 1 8 0 , etc.) there are no restrictions on the allowed j's. Therefore a s y m m e t r y number a m u s t be introduced. F o r diatomic molecules a is either 1 (heteronuclear) or 2 (homonuclear). Equation 24 also holds for linear polyatomic molecules, with 9 0 % ) reservoir of oxygen in t h e s y s t e m . U n d e r these conditions t h e 6 18 O value of t h e final w a t e r can be d e t e r m i n e d precisely by a m a t e r i a l balance calculation r a t h e r t h a n by direct m e a s u r e m e n t . Alternatively, and requiring considerably more care, t h e noble-metal reaction t u b e s can be p u n c t u r e d in a v a c u u m and t h e liberated microliter a m o u n t s of w a t e r can be q u a n t i t a t i v e l y collected and analyzed by s t a n d a r d techniques. Direct m e a s u r e m e n t of t h e isotopic composition of t h e final w a t e r is necessary if t h e e x p e r i m e n t is done in a piston cylinder a p p a r a t u s . A sample calculation is given here for companion exchange experiments (1 and 2) between calcite and w a t e r at 500°C. T h e weights, initial isotopic compositions of calcite, w a t e r s 1 and 2, and t h e final isotopic compositions of calcite a f t e r experiments 1 and 2 were completed are given in T a b l e 3. T h e molecular weights of calcite and w a t e r are 100 and 18, respectively. T h e m a t e r i a l balance equations for calculating t h e isotopic composition of w a t e r after equilibration are as follows: (0-atom% c a l c i t e )(ACO 2 0)

Grootes et al. (1969)

do

SO 2 -S-H 2 S

Thode et al. (1971)

HC0,--(aq>C0 2 (g)

Malinin et al. (1967)

Zns-HS'-PbS

Kiyosu (1973)

ZnS-PbS

Czamanske and Rye (1974)

Mook et al. (1974) Emrich et al. (1970)

Elcombe and Hulston (1975)

C0,(aq)-C02(g)

Thode et al. (1965)

HSO 4 -S-H 2 S

Robinson (1973)

CO 2 -CH 4

Bottinga (1964)

PbS-S

Puchelt and Kullerud (1970)

C0 2 -Diamond

H^gJ-H^aq)

Sharan et al. (1974)

C0 2 -Calcite

SO^H^l) SO^-(aq>H 2 S(g)

Sakai (1957)

do Baertschi (1957) Vogel (1959)

Thode et al. (1977)

Northrop and Clayton (1966)

Sakai (1968)

Bottinga (1968b)

S^-SO^g)

do

HS'-S 2 "

do

Emrich et al. (1970) Graphite-CH 4

Bottinga (1969a)

so^-s 2

Miyoshi et al. (1984)

C0 2 -tholeiite melt

Javoy et al. (1978)

BI^S-S 0 Series of Sulfides

Bente and Nielson (1982)

Series of Carbonates

Golyshev et al. (1981)

Bachinski (1969)

CO 2 -OCS-HCN-CS 2

so 3 -so 2 -ocs-cs 2 CS-SJ-HJS

CH 4 -CO-CS Richet et al. (1977)

35

Richet (1977)

methods of making empirical calculations of fractionation factors (e.g., Hulston, 1978; Kieffer, 1982). T h e methods await testing with forthcoming modern spectral d a t a including those taken on synthetic minerals containing a significant a m o u n t of the heavy isotopes.

ACKNOWLEDGMENTS T h e a u t h o r is indebted to Bob Clayton, Sam Epstein, Hugh Taylor, Yan Bottinga and Richard Becker for many years of illuminating discussions of t h e subtleties of isotopic fractionation. Helpful comments on earlier versions of this chapter were made by Richard Becker, Bill McKenzie, John Valley, and Tosh Mayeda. Special t h a n k s are given to A n i t a Grunder, P a t Dobson and John Chesley for their long-suffering instructions on the use of the Stanford Vax system used in preparing this manuscript.

36

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T(°C) Figure 8. Arrhenius plot of log k versus 1/T for sulfur isotope exchange between aqueous sulfates and sulfides (from Ohmoto and Lasaga, 1982). k in units of kg/mole/hr. Numbers in parentheses are the calculated in-situ pH values. For the meaning of the symbols and letters« refer to Ohmoto and Lasaga (1982). Arrows indicate minimum_or maximum k values. Scale on right gives the time in years to achieve F = 0.9 if ES = 10 2 m; *pH computed at 350°C.

1

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pH in situ

Figure 9. Relationship between pH and half-life (in hours) of oxygen Isotope exchange between dissolved sulfate and water for ENa - 0.1 and ES = 0.01 m. (Solid lines, from Chiba and Sakai, 198S). Dotted line is the relationship between pH and the half-life of sulfur isotope exchange in aqueous sulfate-sulfide systems for ENa " 0.1 and ES • 0.01 m (Ohmoto and Lasaga, 1982).

59

librium in about 4 years. A decrease in p H to 2 results in an equilibration time of only 3 days, whereas the time increases to 10 s years for a pH = 9. A drop in temperature to 150°C results in equilibration times of about 0.5 yrs, 4000 yrs and 10 9 yrs for pH's of 2, 4-7 and 9, respectively. Also, an order of magnitude decrease in ES produces an order of magnitude increase in the equilibration time. The sulfate-water system. The rate of oxygen isotope exchange between dissolved sulfate and water has been measured recently by Chiba and Sakai (1985) at 100, 200 and 300°C. Their rate equation for the exchange of oxygen isotopes between sulfate and water is given in the general form as TJ _ -Rn(l - F)4XY R r„ = (4X + Y ) t '

(26)

where X and Y are the concentrations of ESOj and E H 2 0 , respectively, t is time (sec), F is the fraction of oxygen isotope exchange (see eqns. 10 and 11), the constant 4 accounts for the number of oxygen atoms in sulfate, and R is the overall reverse (r) rate of oxygen isotope exchange for the reaction

S18O316O + H218O

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^ «O O CQ • ON H -H H « « •OH O Q H H s a ) w © J w M I J w r t 55 û£ h c ^ a) w - o . < í^ a ( i n " v o quartz > phlogopite > amphiboles > magnetite. Hydrogen diffusion rates have also been measured, and some of these data for selected silicate-fluid systems are presented on an Arrhenius plot in Figure 13. Despite comparable activation energies (^15 to 30 kcal/mole for H), the diffusivities of hydrogen and oxygen measured in hydrothermal experiments are quite different. The rates of hydrogen diffusion are faster by as much as 3 orders of magnitude — e.g., compare hornblende rate lines in Figures 12 and 13. If the anhydrous exchange experiments represent predominantly volume diffusion, which seems likely given the high and varied activation energies, the addition of water to the system must act to lower the energy barrier for oxygen and hydrogen isotope exchange and assist in the transport rates. The uniformity of the activation energies for oxygen and hydrogen isotope exchange in these hydrothermal experiments seems to indicate the dominating effect of water, suggesting that similar exchange mechanisms operate in most minerals. On a plot of E versus log A (Figs. 1A and 15), data falling on a common trend may indicate a common diffusion mechanism (Giletti, 1985). Examples of this behavior are observed for olivine-02 and quartz-H20 data in Figure 1A, and the feldspar-solution data in Figure 15. Distribution of data along a particular trend is attributable, in part, to compositional variations, and to the nature of defects in the crystal. Winchell (1969) and Winchell and Norman (1969) demonstrated that various diffusing species in silicate glasses showed a positive correlation on a log A versus E plot. A similar effect was noted by Hofmann (1980) for cation diffusion in basalt. This implies that diffusion rates of different species tend to converge at a particular (universal) temperature; this effect is known as the "compensation effect" or "compensation law." Hart (1981) showed that this effect is also present for diffusion in feldspars and olivines. However, Lasaga (1981b) notes that the use of the compensation law in silicates and other complex phases must be modified to account for the variations in structure and the types of defects which are possible in minerals. 78

TCC)

10 3 /T(K)

Figure 12. A c o m p o s i t e A r r h e n i u s p l o t of s e l e c t e d o x y g e n d i f f u s i o n coefficients for a v a r i e t y of m i n e r a l s r e a c t e d w i t h e i t h e r h y d r o t h e r m a l fluids (solid lines) or 0 2 (or C 0 2 ) (dashed lines). S e e T a b l e 1 in order to m a t c h the source of the d a t a w i t h the appropria t e rate line. c, J-c a n d refer to t r a n s p o r t o r i e n t a t i o n s in q u a r t z — i.e., p a r a l l e l to the c - a x i s , n o r m a l to c, a n d n o r m a l to the rhombohedral face (1011).

TCC) 800

0.8

700

1.0

600

500

1.2

400

1.4

350

1.6

300

1.8

10 3 /T(K)

Figure 13. A c o m p o s i t e A r r h e n i u s p l o t of h y d r o g e n d i f f u s i o n c o e f f i c i e n t s from r e p r e s e n t ative h y d r o u s m i n e r a l - f l u i d exchange experiments. Solid lines refer to t r a n s p o r t p a r a l l e l to the c - a x i s ; d a s h e d lines refer to transport n o r m a l to the c - a x i s . See Table 1 for the source o f d a t a u s e d to c o n s t r u c t these lines.

79

(00

h

r

QUARTZ. FORSTERITE 80

-

60

-

40 -

20

-

J

I

I

-e

L_

-4

_l

I

-2

I

I

l

I

log A 0 ( c m 2 / s e c )

Figure 14. Summary plot of activation parameters (E versus log A ) for oxygen diffusion in quartz and forsterite, and hydrogen diffusion 3 in quartz (fong dashes). Quartz: circle - a-quartz; square s JJ-quartz; open * anhydrous; solid * hydrothermal; Si0 2 glass-0 2 : triangle; Quartz-D 2 0: diamond and inverted triangle; Forsterite-0j: hexagon. Refer to Table 1 for more details.

log AQ (cm2/sec)

Figure 15. Summary plot of activation parameters for oxygen diffusion in feldspar-fluid systems. Alblte: circle; Microcline: triangle; Adularia: square; Anorthite: inverted triangle. See Table 1 for more details. Solid line is a least-squares regression of data.

80

Using techniques outlined by Lasaga (1981b), we have calculated examples of universal temperatures (T*) and associated universal diffusion coefficients (D*) for certain mineral groups. The results are presented in Table 2. While there are significant variations in the a and b coefficients in these examples, the high correlation factors (R 2 ) suggest that, if the law is applied within restricted classes of minerals for a particular diffusing species (e.g., H 2 0), it may be a useful empirical tool (Lasaga, 1981b). The concept of closure temperature (Dodson, 1973), i.e., the temperature (or range of temperature) at which geologically significant diffusion ceases, depends on the magnitudes of E and A . Giletti (1986) has used this concept to examine the effects of diffusion on oxygen isotope temperatures of slowly cooled igneous and metamorphic rocks. Effects of pressure on the rates of diffusion. Increases in the rates of oxygen and hydrogen diffusion with increasing temperature are expected from theory, and are well documented by experimentation. The dependence of diffusion in solids on pressure is commonly expressed as D(P) = D

o

exp(-PAV*/RT)

(38)

where AV* is the activation volume for diffusion, P is pressure, T is absolute temperature, and R is the gas constant (Lazarus and Nachtrieb, 1963). ¿V* is equal to the sum of the activation volume of motion, AV , and the activation volume of formation of imperfections, AV.. ffn increase in pressure results in a systematic decrease in the diffusion coefficient only if AV is positive. Yund and Anderson (1978) measured a 10-fold increase in the oxygen diffusivity in adularia reacted with 2 M KC1 at 650°C as water pressure (fugacity) increased from 125 to A,000 bars (see Fig. 16). This effect was explained by differences in the concentration of "water" in the feldspar. However, Freer and Dennis (1982) observed no significant pressure dependence in oxygen diffusion experiments where albite was reacted with water at 600°C. They claim that the surface 1 8 0 concentration increased with increasing pressure in a non-linear manner, implying that a "surface exchange" mechanism, and not the oxygen diffusion rates, varies as a function of pressure over most of the range examined (0.5 to 8 kbars). Similar inconsistencies are evident in the work of Giletti and Yund (1984) and Dennis (1984) on quartz-water isotope exchange. Dennis (1984) observed no pressure effect in his experiments. In contrast, Giletti and Yund (1984) observed a pronounced increase in D with an increase in total water pressure. They attributed this behavior to an increase in the concentration of some impurity such as OH , H , etc., in the crystal with increasing pressure. There are several reasons for the different results. Ewald (1985) suggests that the bulk of the pressure effect is manifested in a surface reaction process which is fast compared to diffusion, in agreement with Dennis (1984). Freer and Dennis (1982) previously suggested that a surface mechanism could account for the differences between their single-crystal experiments and the experiments of Yund and Anderson

81

TABLE 2.

EMPIRICAL OBSERVATIONS BASED ON THE COMPENSATION LAW FOR SELECTED DATA GIVEN IN TABLE 1

Phase

Isotope

Feldspars

18

0/160

Mica/Amphibole

18

0/160

D/H

E

R2

a =

50,.92 + 3,.42 log

A

o

74,.91 + 4,.91 log

A

o

c

37,.48 + 2..38 log

A

o

c

Forsterite

18

0/160

105..98 + 6..61 log

A

o

Corundum

18

0/160

127..37 + 9..27 log

A

o

T*(°C) (a)

D*(cm 2 /sec) (b)

0.94

474

1.3 X 10~ 1 5

0.98

800

5.5 X 1 0 " 1 6

0.87

247

1.8 x 1 0 " 1 6

0.98

1171

9.2 x 10" 1 7

0.99

1753

1.8 x I D ' 1 4

(a) T*< is the universal temperature of diffusion for a particular phase; see text. (b) D* is the universal diffusion coefficient for a particular phase; see text.

log f H z 0 ( b o r s )

Figure 16. E f f e c t of w a t e r pressure o n d i f f u s i o n in q u a r t z (Giletti a n d Yund, 1984) a n d adularia (Yund a n d A n d e r s o n , 1978). D a t a a r e p l o t t e d a s log D versus log w a t e r f u g a c i t y . Transport is p a r a l l e l to the c - a x i s for b o t h natural q u a r t z (open circles) a n d s y n t h e t i c quartz (solid c i r c l e s ) . T h e s o l i d a n d d a s h e d lines are least-squares r e g r e s s i o n of the data.

82

(1978), w h i c h utilized powdered materials. The discrepancy can also be partly attributed to differences in hydrogen contents (i.e., defects) of the samples in the various experiments. Dennis (1984) used annealed (850°C) quartz in his experiment, whereas the unannealed quartz of Giletti and Yund (1984) would contain significant hydrogen impurities and defects (see Aines and Rossman, 1984). Similarly, the albite sample used by Freer and Dennis (1982) would be expected to contain only a small amount of water that is present as O H , whereas adularia has the highest concentrations of water (both OH and H 2 0 ) present in any feldspar (A.M. Hofmeister, pers. comm., 1986). The difficulty of introducing water into the defect-free (dry) samples even under compression apparently precludes the existence of a pressure effect on diffusion rates (Hofmeister, pers. comm., 1986). Anisotropy in diffusion. Diffusion in anisotropic crystals is controlled by diffusion coefficients that are not simple constants but rather are second-order tensors. In such a case, material transport rates will differ in different crystallographic directions, and the direction of net transport will, in general, not coincide w i t h the concentration gradient. Microbeam techniques make it possible to evalute such effects for specific crystal surfaces using oriented samples. For example, Giletti and Y u n d (1984) report that, for quartz, oxygen diffusion coefficients for transport parallel to the c-axis (D ) are approximately 2-3 orders of magnitude greater than those for transport normal to c (Di ) (see Fig. 12). A significant anisotropy, similar in magnitude to that observed by Giletti and Y u n d (1984), w a s also observed by Dennis (1984) for transport parallel to the c-axis and normal to the rhombohedral face in quartz. W e have seen that w h e n water is present, the Water in minerals. rates of diffusion are considerably faster than diffusion rates m e a s u r e d under anhydrous conditions. In addition, diffusion rates in hydrothermal systems appear to be very pressure dependent. These relationships are related, in large part, to the amount and speciation of "water" (i.e., hydrogen-bonded molecules) accommodated by the minerals during diffusion. Infrared (IR) spectroscopy has contributed m u c h to: (1) determining the actual speciation of "water", (2) determining the crystallographic environment of that species, and (3) determining its analytical concentration (Aines and Rossman, 1984). IR can recognize water in minerals in the form of fluid inclusions, and as isolated molecules which can be distinguished from hydroxide ions. Kats (1962) showed that three types of 0 - H related defects m a y be distinguished in natural quartz: (1) those in which alkali and h y d r o g e n ions are associated, (2) defects involving two or more hydrogens, and (3) associations of aluminum-hydrogen. Recent work by Aines and R o s s m a n (1984) using near-infrared spectroscopy demonstrated that water m o l e cules are the dominant hydrogen-bearing species in synthetic quartz. IR studies by Aines and Rossman (1982) indicate that OH is a common substitute in garnet. Minor amounts of water are commonly found in feldspars (Solomon and Rossman, 1979, 1982; Hofmeister and Rossman, 1985). The concentration, speciation, and position of "water" in feldspars varies with factors related to: (1) availability of water, (2) feldspar composition (bulk as well as impurity), and (3) subsolidus history (Solomon and Rossman, 1982). In natural feldspars, higher water content is associated w i t h a greater extent of Al/Si ordering. Water in adularia is dominantly H 2 0 and is usually h i g h in concentration, w h e r e a s 83

water in albite is as OH and meister, pers. comm., 1986).

is typically

low

in concentration

(Hof-

These IR studies, and numerous others not cited, clearly indicate that water is present in natural minerals, sometimes in significant quantities. If the mineral-fluid system was in chemical equilibrium, then diffusion played a major role in fluid transport (i.e., water uptake) and subsequent isotope exchange. If, however, recrystallization occurred, then water m a y be incorporated into the mineral during growth — e.g., in fluid inclusions or as molecular species. Giletti (1985) notes that there are m a n y natural cases where it is difficult to distinguish between these two mechanisms. Comparison of the Surface-Reaction and Diffusion Models In order to illustrate the differences and similarities between the surface-reaction and diffusion models, w e will examine the interaction of feldspars and micas w i t h fluids. These phases are particularly important because of their ubiquitous occurrence in natural environments. In addition, most of our knowledge concerning mechanisms and rates of isotopic exchange are derived from experiments on feldspars and micas. The results of model calculations are depicted graphically in Figure 17, which presents the fraction of isotopic exchange represented as £n(l - F) plotted against time in years. Oxygen isotopic exchange by means of diffusion in the systems albite-water (#30, Table 1) and phlogopite-water (#51, Table 1) was modeled for: (1) 400°C, (2) grain sizes of 0.01, 0.1 and 1.0 cm, and (3) varying fluid/mineral mass ratios (0.1, 1.0 and 10). Albite and phlogopite were m o d e l e d as a sphere and plate, respectively, using equations provided by Crank (1975). Similarly, surface-controlled isotopic exchange was modeled using equation 37 (or 36) for albite altering to K-feldspar (#37, Table 1) and paragonite altering to muscovite (#57, Table 1) at 400°C, w i t h variable grain sizes and fluid/mineral m a s s ratios (X/Y). It is important to reiterate that the volume ratios u s e d in the diffusion equations must be corrected for concentration by use of the equilibrium isotope partition ratios ( K 1 ) (Cole et al., 1983). Results of the surface exchange calculations plotted in the upper portion of Figure 17 indicate that the time required to attain a certain fraction of exchange is less for lower fluid/mineral mass ratios. For example, a 0.01 cm albite grain altered to K-feldspar at A00°C reaches 90% exchange in 0.17 years for a fluid/mineral mass ratio of 1.0, but the same degree of exchange is achieved in only 0.04 years w i t h a mass ratio of 0.1 (Fig. 17a). A similar decrease in time is observed for paragonite altering to muscovite as the fluid/mineral mass ratio is decreased (Fig. 17b). The faster rate occurs at low water/solid mass ratios because the solid doesn't have to shift as far in composition. A n increase in the grain size produces a similar result — i.e., increasing grain size increases the time required to achieve a given degree of exchange. For example, at a constant fluid/mineral m a s s ratio of 0.1, the time required to reach 50% exchange for albite reacted to K-feldspar at 400°C (Fig. 17a) increases from 0.09 to about 0.9 years as a consequence of increasing the grain size from 0.10 to 1 cm. For

84

0.1

(8

26

34

42

t*102)

4

TIME (yeors)

f.O -0-'-1 6

8 (x 10®)

TIME (years)

Figure 17. Fraction of oxygen isotopic exchange given as £n(l - F) versus time in years for surface reactions (af b) and diffusion (c, d) in feldspar and mica-fluid systems at 400'C. Rate constants for albite altering to K-feldspars and paragonite altering to muscovite are from Cole et al. (1983). (See Table 1, }'s 37 and 57.) Diffusion coefficients from Giletti et al. (1978) and Giletti and Anderson (1975) were used to model albite and phlogopite, respectively. W/S is the fluid to solid mass ratio, a is the grain radius (cm), and v is the plate thickness (cm). The top number next to a line or curve is (W/S) and the bottom number is the a or w.

another example, but this time at a constant fluid/mineral mass ratio of 1.0, a change in grain thickness from 0.001 to 0.01 cm for paragonite altered to muscovite (Fig. 17b) increases the time required to attain 50% exchange from A years to nearly AO years. Diffusion-controlled isotopic exchange shown in the lower portions of Figure 17 (c and d) behaves in a somewhat similar manner. The time required to attain a given fraction of exchange increases with an increase in either grain size or fluid/mineral mass ratio. The times for diffusion exchange, however, are several orders of magnitude slower than surface exchange. A A00oC system with a grain radius of 0.01 cm and a fluid/mineral mass ratio of one reaches 90% exchange in 800 years for oxygen diffusion in albite (Fig. 17c), very long compared to the above-mentioned period of 0.17 years when the albite is reacting to K-feldspar. Another significant difference is in the shape of the curves. Straight lines are produced for the surface-reaction model, whereas curves exhibiting a less rapid increase in fraction of exchange with increaseing time are observed for the diffusion model. These data suggest that the oxygen isotopic exchange reactions between fluids and rocks in natural systems may proceed in two stages: the first through a surface-reaction mechanism during mineralogic alteration, and then through a diffusional mechanism once the system attains chemical equi85

librium.

SUMMARY The brief discussion of items listed below highlights the more generally useful results described in this review:

several

of

1.

Isotope exchange commonly results from the collision of molecules leading to the formation of an intermediate complex, w h i c h then breaks up to give products. For example, exchange m a y occur between species undergoing collision in solution, or between a volatile and a solid as the former is adsorbed onto the surface(s) of the solid.

2.

In view of the convincing evidence developed from field and laboratory studies, it is evident that many exchange pairs are slow to equilibrate isotopically. of isotopic equilibration m a y be u s e d to quantify the history of a particular geologic system.

3.

Slow rates of isotope exchange are not restricted to heterogeneous mineral-volatile systems; they m a y also occur in homogenegous systems — i.e., in solutions or gases.

4.

A generalized isotope exchange rate law is applicable to a wide variety of systems, but m u s t be adjusted to suit the particular application at hand.

5.

Experimental determination of the rate of isotope exchange is typically achieved by measuring the change in isotope composition as a function of time at various temperatures, pressures, and initial composition.

6.

The bulk exchange technique m a y be used to measure the rates of isotope exchange regardless of the mechanism. In comparison, microbeam profiling methods are used predominantly to determine diffusion coefficients.

7.

Rates of isotope exchange between aqueous species are strongly dependent on temperature and pH. The relationship between pH and the dominant aqueous species must be established through chemical speciation calculations in order to evaluate the mechanistic p a t h of reaction — i.e., formation of intermediate species.

8.

Rates of isotope exchange between gas species can be very slow for reactions occurring in the absence of catalytic agents (e.g., a metallic surface, H 2 0 , etc.). For example, this appears to be the case for exchange of carbon isotopes between C 0 2 and CH,,, w h i c h have been reacted under essentially anhydrous conditions at elevated temperatures.

9.

Isotope exchange between minerals and fluids can be attributed to two major processes, diffusion and surface reactions. The former refers to systems where minerals exhibit no signs of alteration, whereas the latter connotes mineral-fluid interac86

numerous types of The lack temporal

tion involving recrystallization or the formation of a new phase. Interaction between a solid and a gas phase (fluid absent) occurs only through a diffusion process, and generally results in the slowest rates of isotope exchange for a heterogeneous process. 10.

Rates of isotope exchange in heterogeneous systems are controlled principally by the temperature, grain size, and fluid (or volatile)/solid ratio regardless of the mechanism. Additionally, water pressure and crystal anisotropy may have a profound effect on the rate of isotope exchange via the diffusion process. Some researchers have observed a faster rate of diffusion-controlled exchange at higher water pressures .

11.

Systems experiencing surface-controlled isotope exchange reach equilibrium well before diffusion-controlled isotope exchange can contribute significantly to the overall exchange. The extent to which surface reactions influence the overall isotope exchange during mineral (or rock)-fluid interaction depends on the abundance of the product phases and their composition. ACKNOWLEDGMENTS

We wish to thank Drs. A.C. Lasaga, D.A. Palmer, and D. Wesolowski for many valuable discussions. The critical reviews by Drs. R.E. Criss, B.J. Giletti, and J.G. Blencoe are gratefully acknowledged. Their suggestions and comments have contributed immeasurably to improving this chapter. The authors would like to single out and thank Dr. Criss for the extra time and care he took to review all versions of this chapter. Comments by A.M. Hofmeister and C. Kaiser are also gratefully acknowledged. This paper couldn't have been completed without the typing talents of Ceci Steele, Joe Ellen Rast, and Betty Benton. This study was supported by grant EAR 85-08379 from the National Science Foundation (Ohmoto) and by the Division of Engineering and Geosciences, Office of Basic Energy Sciences, U.S. Department of Energy, under contract DE-AC05840R21400 with Martin Marietta Energy Systems, Inc. (Cole).

87

88

Si

o M X V 41

3 41 4J a VI O 4J ai m a -4 0

O 1 fi X c . M 4) JS « O «T 4J

4J 4) ae C -H o i-i n «1 £> «N•H ^ z

Si

»a

AU O X o w 00 fi Mfi-4 o c O 41 1 4J 41 o U) ••

fi Qu i o : •H vi m fi in

« 4) e 4i 00 JS -

of ma in t ovotry

i-i a JS

» ° ^41 Xi k 3 A N4J O ri ,

c z o fi •H c • UDÌ 9 0

gen : Omar

M 00 c

O «1 -o "O •-4 Q « a «

nta] in t età

hi i ¡3 V 1 c

;i980) i fract Lm. Cos

C .2 » 3 E b

0 2 S § Q y fi • h ^ o X m IN 41 ^ i r^ O 4J h. >

n fi 3 w (« O -H (a. c w X o a 1*4 41 O fi c« C O •H O *J M o

0 3 f S

H

C

0

o

41 a O «1 A O» u -h o n e : « o e u c « -h n v n 9 « •o 41 « U C < l tf> w O -H « JS %o 3 O •• M « I • Si • >> "O C C « 44-1» -H • 9>*4X fCa, i I — a= « o ^ • • • J "S « « c 5 A >> M «H U >H l V c A 9 U « -4 /-V . | > C M 41 C c o « « Q4 si O • 4) 41 « M • 0 19 N u 4i > • « 9 : Z -4 C «C A 00 00*0 fH •-> fi • 01 *J 41 « 4J 9>£B r-p» u « /-v u s 4) «• a e , u o C O « *j x 3 M « • • r 3 -H O«m c« VO V4 (0 O Q 4J h 0 -H U 9 K (4 • 4> O -9 X -4 t>N *J V 01 -o « O 4J 3 « e c p • (3 «H the exchange trajectories will be straight lines, two of which are the legs of the open system triangle shown in Figure 8a, and the isochronous arrays will be parallel to isotherms. Alternatively, i f k i / k 2 » 1, and i f the more rapidly exchanging mineral is plotted on the y-axis, the exchange trajectories are strongly curved, but the arrays are linear or nearly linear with positive slopes that vary from ki/k2, at f2 = 0, to 1 at f2 = 1 (Fig. 8b; see below). Exchange trajectories and isochronous arrays that l i e outside the closed system triangular areas must have been formed through exchange in an open system. Kinetic Effects Elementary rate law. Kinetic effects of isotopic exchange reactions in fluidmineral systems were originally considered in experimental studies because of the com mon lack of attainment of equilibrium in laboratory experiments designed to measure isotopic fractionation factors (e.g., Northrop and Clayton, 1966). Recently, the determination of isotopic exchange rates themselves has become a major goal of the experiments (e.g., Cole et al., 1983; see Chapter 2). Disequilibrium effects attributed to relative exchange rate differences between minerals and fluids are prominently displayed in a number of suites of hydrothermally exchanged plutonic rocks (e.g., Taylor and Forester, 1979; Gregory and Taylor, 1981; Criss and Taylor, 1983). Numerical simulations of 1 8 0 exchange associated with cooling gabbro plutons (Norton and Taylor, 1979) and the cooling of the oceanic crust (Cathles, 1983) require some knowledge of exchange rate parameters. Thus f a r these studies have only modeled two-phase systems (solid plus fluid). In the future, the slopes and positions of the disequilibrium arrays observed in real multi-mineral systems may provide a major constraint on the numerical solutions derived through simulation studies. In the discussion that follows, we will use the approach of Criss et al. (1986) to model disequilibrium exchange. For any exchange reaction involving fluid and mineral, the rate of change of the isotopic ratio of the ith mineral will be related to the difference between the forward and reverse reaction rates given by dR-j/dt = a i w k i R w - k i R i

.

(16)

Equation 16 states that the rate of change of R^ is proportional to the difference between the actual isotopic ratio of the mineral (R^) and the value (o« w R w ) t h a t i t would have i f i t were in equilibrium with fluid R w . It is highly probable that in the heterogeneous mineral-fluid system, the rate constant k.i is directly proportional to the surface area per mole of the mineral (see Chapter 2), but we make no particular assumptions about reaction mechanisms here. Equation 16 has striking affinities to a Prigogine rate law, useful for the description of reaction ratesin systems close to equilibrium (e.g., see Benson, 1960). "Closed" system exchange model. In a two-phase "closed" system consisting of a single mineral plus fluid in fixed proportions, conservation of mass dictates that X i dR1 = -X w dRw

.

The solution of equations 16 and 17 for such a system has the form 1 - f1 = = exp - a i w X i x + XW J k . d t

103

(17) .

(18)

Because a is close to unity, the first term of the argument of the exponential i n equation 18 becomes 1 / X w . If k^ is taken to be constant, equation 18 can be reduced to 1 - f , = exp - k j t / X w . (19) The value of f is determined by equation 15 and i n i t i a l conditions, and k^ can be determined in experimental runs by utilizing plots of l n ( l - f ) versus t , independent of the reaction mechanism or the form of k^ M c K a y , 1938). For small X w , the equilibration t i m e is shortest because the exponential term goes to zero rapidly (i.e., the isotopic composition of the solid does not change much). Longer equilibrium times are required for X w approaching one, because i n this case all of the isotopic change is taken up by the solid. To extend this treatment to n-phase systems, we assume t h a t the minerals only communicate through the fluid phase, and write as many equations 16 as there are phases, and modify equation 17 t o take into account the other phases (see equations in Criss et al., 1986). This more general n-phase "closed" system becomes a special case of an open system i n f i l t r a t e d by external fluid t h a t travels at a normalized flow rate u (in see"*), as discussed below. Open system exchange model. To model the open system isotopic exchange between fluids and n crystalline phases (Fig. 9), we must know: the mole fractions X.j; the fluid flow rate (u); the fractionation factors (ctj w ); the exchange rate parameters (k^), and the i n i t i a l conditions (R i , . . . , R n , R ). Following Criss et al. (1986), the simultaneous differential equations that describe the system are dR^dt

=

ki

Xw dR w /dt + ö^dR-j/dt

=

( R w 1 n - R w )u

a

R

iw

w "

R

i

, and

(20a) (20b)

Equations 20a and 20b can be written in matrix format and then rearranged in a standard form dR/dt = AR + W , (21) where R and W are column vectors ( R i R n , R ) and ( 0 , ..., 0, u R ™ / X w ) , respectively. The square matrix A is of order n+1 a n a i s given lay: -ki 0 0 -kiXi/Xw

0 . -k 2 0

0 . -k3

0 0 0 . 0 - k n„ "knVxw

k

i > *2W

2

3 k,n°riw > - ( £K.j ®jwX-j /X w + u/X w )

Equation 21 becomes a standard eigenvalue problem (e.g., Boyce and DiPrima, 1969) that has the solution f o r two mineral phases and fluid: rk,aiw k i - Xi R =

k

2«2W k2 - X i

k

!aiw k l - X2

k

k

k

2 a 2W k2 - x 2

iaiw k l - X3

2 a 2W k2 " X3

e

-\it

0

0

e

0

0

104

-\2t

0

Ci

Ri e ( 1

0

C2

R2e(l

C3

D in K w

-\3t

(22)

S18O2 Figure 9. Solutions to the general open-system case (eq. 22) for the following arbitrary choice of parameters for phase 1 and phase 2, respectively: X- = (0.4995, 0.4995); k, = ( 1 0 " i o " 5 ) f o r a relative rate constant of 100; a, w » (1.001, 0.999). Also, X w = 10"3. The lighter curves correspond to a family of open system exchange trajectories for different choices of the fluid flux u, whereas the heavy lines represent isochrons that connect points of equal time on these trajectories.

Figure 10. The 6-values of solid phases 1 and 2 are shown along with the fluid 6 1 8 0 composition as a function of time for the calculations of Figure 9. Values of the fluid flow rate u are indicated on the various curves. In this example, the fluid and mineral 1, the fastest exchanging mineral, are close to isotopic equilibrium (as shown by the similarity of their respective curves), whereas mineral 2 1s not in equilibrium with either phase. For small values of u/ki, large isotopic shifts 1n the fluid are possible. 105

-2.0

-1.0

0

2

3

4

5

6

S18 0 - B

Figure 11. Comparison of the exchange trajectories generated for the "closed" and "buffered" opensystem models for three relative rate constants (kiA2 = 1, 2, and 5). For short time scales or for small values of fluid/rock, the two models are difficult to distinguish. Interruption of the isotoplc exchange process before equilibrium 1s achieved will result 1n large errors In Isotoplc temperature estimates, even for equal exchange rates (I.e., k i A 2 * 1)- Modified from Gregory et al. (1981).

where C is a co1u mn vector with scalar components calculated from the initial conditions. The eigenvalues (\) are real, and are found by solving the matrix equation A - I\ = 0, where I is the identity matrix. Examples of the solution are shown in Figures 9 and 10 for some arbitrary choices of parameters. Because the actual values of the constants are not well known for real geologic examples, i t is useful to consider some simple end member cases to gain further insights into the behavior of the natural systems, and to reconcile the general exchange models with earlier models. "Buffered" open system exchange model. R w is effectively constant (R™) when u is large (u » k^j. For large u, equations 20a have the same solution as equation 19 with X = 1: 1 - f , = exp - k , t

(23)

This expression, hereafter termed the "buffered" open system equation, assumes that the fluid flux is so large that a boundless supply of fresh unreacted fluid (with a constant 6 value) is readily available to exchange with the rock. In a "buffered" open system, the kineticalty-controlled 6-values of two coexisting minerals may be directly compared by rearranging all equations 23 to eliminate t, and exponentiating the results: ki/k 2 (24) 1 - f i = (1 - f 2 ) In Figure 11, graphs of equation 24 are shown for comparison against results for a "closed" system (Gregory et al., 1981; Criss et al., 1986). The similarity of the "closed" system and "buffered" open system solutions for short times indicate that disequilibrium mineral-pair arrays can be generated by either model. The above solutions represent mathematical statements of the obvious fact that if an isotopie exchange process is interrupted, the measured mineral fractionations will have no equilibrium temperature significance. This result 106

suggests that stable isotope geothermometry can be highly suspect, and inherently unreliable, in terranes that have been subjected to infiltration by transient fluids, particularly atlow temperatures where equilibrium is difficult to attain. Interpretation of the Kinetic Models Fluid/rock ratios. Considerably different fluid/rock ratios are implied by the various kinetic models described above. In the "closed" system (eq. 16, 17, 18), u is considered to be zero. The fluid/rock ratio for any such "closed" system is si m ply */r = y ( E *j] . (25) Equation 25 gives the actual fluid/rock ratio for a specially defined system, where the fluid and rock are present in proportions that remain fixed in time. On the other hand, for the general open-system case described by equation 22, the fluid/rock ratio becomes an explicit function of time w/r = (Xw + u t ) / ( r X1-)

.

(26)

Note that this actual fluid/rock ratio is independent of any isotopic quantities. If the k^ are all small, large amounts of fluid therefore could "see" the rock with little or no isotopic effect. For geologic systems, i t is most reasonable to interpret X w in equations 25 and 26 in terms of the porosity. This would suggest that "closed" system w/r ratios are always quite small. However, in geothermal systems i t is possible to get high "closed" system fluid/rock ratios by repeatedly cycling fluid from cold unreactive bodies of rock through zones where the temperature is high enough to drive exchange reactions. In this case, w/r becomes related to NCw/(l - )Cr, where is the porosity, N is the number of times the fluid in the pore volume of the hot rock has been replaced, and Cw and Cr are oxygen concentration factors. In this situation, "closed" system w/r ratios can also be considered to increase with ti m e. It is interesting to compare the above "actual" w/r ratios with the "effective" (i.e. "isotopic") w/r ratios of Taylor (1974a, 19/7). Equation 17 states that for a "closed" system, the oxygen w/r ratio is equal to the ratio of the changes in the isotopic compositions, -(6r - 6°r)/(6w - 6°w), that would be observed in the fluid and rock undergoing exchange. The actual and the effective w/r ratios are accordingly identical in this case. In practice, however, because the final isotopic composition of the fluid cannot generally be measured (particularly in fossil hydrothermal systems), i t is assumed that equilibrium is ultimately attained, so that the "closed" system isotopic w/r ratio becomes (Taylor, 1977): w/r = - ( 6 r - 6 ° r ) / U r - A r w - 6 ° w )

.

(27)

It proves more difficult to compare the actual and the effective w/r ratios in the open system case. However, this can to some degree be achieved i f we consider making a simple variable substitution of dw/^ for ^dt in equation 20a. Here dw represents conceptual incremental packets of incoming fluid with a constant 6 value that completely exchange and equilibrate with an amount r,- of each mineral in time dt, and then are discarded. On integration, the open system isotopic fluid/rock equation of Taylor (1977) may be derived: loge

( 1 - f ) = -w/r

,

(28)

For each mineral an effective water/mineral ratio can be calculated from the 107

value of f (see below), and the cumulative water/rock ratio is the sum of these individual effective water/mineral ratios weighted according to the mole fractions of the minerals (see Fig. 7 of Gregory and Taylor, 1981). Note that the derivations of equations 27 and 28 for isotopic w/r ratios both involve the assumption that equilibrium is attained, whereas no such assumptions were made in derivations of equations 25 and 26 for actual w/r ratios. As a result, the isotopic w/r ratio given by equation 27 will always underestimate the actual "closed" system w/r ratio given by 25, and the isotopic w/r ratio given by equation 28 will always be less than the actual open system w/r ratio given by 26. Isotopic fluid/rock ratios for natural systems lie between the two extremes indicated for the open (eq. 28) and "closed" (eq. 27) systems. Taylor (1977) has shown for small fluid/rock ratios that these equations give the same result. The "closed" system equation (27) gives a larger "isotopic" or effective fluid/rock ratio for a given difference in 6 between final and initial rock than does the "open" system equation 28 (see Fig. 7 in Taylor, 1977). The actual fluid/rock ratios need not bear a close relation to these isotopic estimates of the fluid/rock ratio. In all single mineral plus fluid methods of estimating the fluid/rock ratio, the temperature of exchange and the initial isotopic compositions of the fluid and mineral must be assumed. However, by analyzing coexisting minerals from a suite of samples, the fluid/rock ratio can also be estimated from the slope of the disequilibrium array for cases where u/ki is large, provided that the relative rate constantis qualitatively known (see below). Exchange trajectories vs. isochronous arrays. Given appropriate rate constants and isotopic parameters, and in most cases small f values, the calculated fluid-rock exchange curves generally mimic the 6-6 arrays observed in rocks of hydrothermally-altered terranes (see below). Along a single-trajectory exchange curve, samples that plot farthest from the initial point have suffered the highest fluid/rock ratio. However, in natural systems such as the Skaergaard intrusion (Taylor and Forester, 1979; Norton and Taylor, 1979) and the Sam ail ophiolite (Gregory and Taylor, 1981), some samples with nearly "normal" ó 1 8 0 values have, on other geologic grounds, clearly exchanged with large amounts of circulating fluid. Simple single-trajectory exchange models cannot account for these observations, which appear to require that the isotopic composition of the fluid was not constant throughout the rock mass. Numerical simulations, such as those of Norton and Taylor (1979), have shown that the 1 8 0 composition of the fluid is strongly dependent on the prior history of the fluid packet as it circulates through the temperature field of the system. Consideration of the "buffered" open-system case (eq. 24) for ranges of fluid isotopic compositions provides important clues into the probable origins of the commonly observed disequilibrium arrays. By artifically allowing the fluid 6 1 8 0 composition to vary in different parts of a hydrothermal system, the different exchange histories of distinguishable parts of the rock mass may be qualitatively simulated. Calculations for the plagioclase-pyroxene system (Fig. 12) indicate that by connecting points of equal f along a family of such curves, each of which represents an individual subsystem that has exchanged with a distinct fluid, a positive-sloped, isochronous linear array results (Gregory, 1986; Gregory and laylor, 198bb; Criss et al., 198b). The slopes of these lines of constant f vary in the same way as in Figure 8, and these lines are related through equations 23 and 28 to the elapsed time and to the effective fluid/rock ratio in the "buffered" open system. Furthermore, the position of a point along the array is related to the isotopic composition of the fluid that the rock actually "sees". For this "buffered" open system (X v + 1 or u + »), where different rocks are assumed to have exchanged with fluids naving a range of 6 values, the slope of a 180

s 1 8 0 PYROXENE Figure 12. A family of exchange trajectories (curved lines) is shown for a "buffered" open system (eq. 24). The positive-sloped straight lines have constant f p y _ , and two of these lines are isotherms. The diagonally-ruled area represents the part of 6-6 spadr that would be accessible to a closed system having the assumed initial isotopic compositions (5.5, 6.0), and the stippled area shows the disequilibrium field of layered gabbro samples from the Samail ophiolite, Oman. Modified after Gregory e t a l . (1986).

constant f array in 8-8 space can be calculated from equations 15 and 24: Array slope = 1 - (1- f 2 ) f2

ki/k2

= fi/f2

.

(29)

Equation 29 proves that lines of constant f are always positive, because 0 < f 580° C, Hagstrum and Johnson, 1986), the near vertical (slightly positive?) slope of the magnetite-quartz data (solid triangles, Fig. 15) may also be indicative of open system conditions. However, i f magnetite is present in very low abundance in this pluton, a closed-system quartz-magnetite exchange vector could have an extremely steep (almost vertical) negative slope. In principle, application of equation 29, the slope of a constant f line, and equation 24 (relating the f values of two minerals) to the natural data should allow 115

crude estimates of the average relative rates of reaction for minerals of the Rio Hondo pluton. In order to accomplish this task, one of the relative rates has to be assumed and the value f of one of the minerals (e.g., quartz) has to be estimated from the slope of the disequilibrium array. For k^/k t z > 10, virtually all of the disequilibrium arrays involving quartz will have thè same slope, l / f f l t z An additional constraint provided by the system is that the final choice or relative rate constants should not result in isotopie fluid/rock ratios that are excessively large or so small that impossibly light values of external meteoric fluid are necessary to explain the observed isotopie changes in the rock. Because of the difficulties associated with assigning slopes to the arrays observed in such natural samples, there is considerable uncertainty in this type of estimate of the relative exchange rate constants. However, i f the relative rate constant for feldspar-quartz is assumed to be 10, and i f f q t z is estimated to be between .06 and 0.1, then calculated values of relative rate constants are: 2.32.0, mgt-qtz; 23-19, bt-qtz; 11-9.5, bt-mgt; 2.6-1.9, bt-fsp; 4.4-5.0, fsp-mgt. For this choice of parameters, the effective fluid/rock ratio for feldspar is close to 1. Note that the observed quartz-magnetite pairs lie along a steep near-vertical array (Fig. 15), which for small f t z suggests that the array slopes should be equal to k m q t / k The estimated value k m q t / k 2 is at odds with the observed near-vertical slope (Fig. 15), suggesting t n a t t n e original choice of the k^s /k t z may have been too low. On the basis of their hydrogen isotopie studies, Hagstrflm and Johnson (1986) estimate that the meteoric fluid that interacted with the Rio Hondo pluton had an initial ó 1 8 0 value of about-12. If such 1 8 0 - depleted meteoric water could reach the pluton without undergoing an 1 8 0 - s h i f t along the way, effective fluid/rock ratios of 0.85 would be required to produce the most 18 0 - depleted feldspar in the intrusion ( 6 1 8 0 = -1.2). Alternatively, i f the relative exchange rate between feldspar and quartz is assumed to be 100 and i f 0.01 < f t < 0.06, then the relative rates are: 6-20, mgtqtz; 99-101, bt-qtz; 18-5, bt-mgr, 1-4, bt-fsp; 18-5, fsp-mgt. For this choice of parameters, the effective fluid/rock ratio for feldspar is 1-5 (in oxygen units). The calculations thus strongly suggest that very large (kf s p / k qtz > ^ relative rate constants are not reasonable for any array with a non-vertical slope, because the isotopie fluid/rock ratios get unreasonably high. The spread in the values of the relative rate constants is qualitatively consistent with the expected spread in possible exchange rate parameters estimated through experimental studies (e.g., Cole et al., 1983; Wood and Walther, 1983; Chapter 2). These estimated rangesin relative rate constants suggest that only a few orders of magnitude separate all of the different silicate-fluid exchange rate constants at any given temperature. This analysis suggests some potential for this approach i f the method is practiced with extreme caution! Precambrian Siliceous Iron Formation: 18

Quartz-Magnetite

16

Because of the large 0 / 0 fractionation between quartz and magnetite, siliceous iron formations are ideal targets for oxygen isotope geothermometry (James and Clayton, 1962). The sedimentary origin of these rocks also made them attractive as indicators of ancient surface conditions (e.g., Cloud, 1968). Perry and coworkers, in a series of papers (e.g., Perry et al., 1978; Perry and Ahmad, 1983) made extensive use of 6 - a diagrams to infer the bulk 1 8 0 distribution of Archean cherts and hence infer the 1 8 0 history of the oceans. Perry and coworkers were impressed by the remarkable linear trends displayed by quartzmagnetite data on 6 - A diagrams (Fig. 16a,b). As discussed above, in such a two-phase system, all available isotopie information is represented by a single point on a 6- A plot. Furthermore, any 116

Figure 17. 6 - 6 plots of the same quartz-magnetite data shown In Figure 16. Also shown are the Hamersley data of Becker and Clayton (1976) for assemblages with < 70 wt. % magnetite. The steep slopes of the data arrays are inconsistent with a closed-system interpretation of the Hamersley data. For a "buffered" open system, the position of the array would be independent of the modal mineralogy. In the Hamersley samples, there is a slight dependence of the 6-values on modal composition, with the more easily exchanged magnetites of quartz-rich layers showing the greatest enrichment in 6 1 8 0 . The iron formation data from Minnesota show strikingly different relationships on this 6-6 plot than were displayed when the same data were plotted on the 6 - A plot. (cf. Fig. 16a). Modified from Gregory et al. (1986).

linear array on a 6-A plot transforms to a linear array in 6 - 6 space. In Figures 17ab, 6-6 representations are shown for the same Minnesota iron formation data (Perry and Ahmad, 1983) and Hamersely basin data (Becker and Clayton, 1976) shown in Figures 16a,b. On the 6-6 plot, the Minnesota iron formations can be grouped into three types of samples: (1) contact metamorphosed samples with an array slope close to 45° (also see Fig. 19); (2) samples from low-grade iron formation that exhibit a positive-sloped disequilibrium array; and (3) samples transitional to groups 1 and 2. These differences are not well displayed in the 6A graph of the same data. In Figure 16a, the intercept of these data on the 6axis contains no information about the bulk 1 8 0 composition, and certainly does not correspond to the original primary sedimentary bulk composition of this iron formation. This point is emphasized by the Hamersley data in Figure 16b, which show that the closed-system exchange tregectory appropriate for the approximately fixed modal composition of the samples (~ 85% quartz) differs significantly from the actual data. The Hamersley system was clearly open to external 1 8 0 reservoirs. Examination of the Hamersley data in more detail (Fig. 17b) indicates that there is a weak dependence of array position on modal mineralogy. In the 118

S 1 8 0 QUARTZ Figure 18. Exchange trajectories using the "buffered" open-system equation (24) for a variety of fluid 6 f f l 0 values (-9, -6, -3 f 0, 3, 6, 9, 12, and 15). Isochronous Tines of constant of f have positive slopes, even 1n a system where most of the exchange trajectories Initially have negative slopes. In the calculation, the constant-f lines intersect the exchange Isotherm very close to the point giving the Initial 6 l a 0 value of the quartz. The near vertical or positive-sloped data point arrays are predictable consequences of open-system disequilibrium exchange models. Modified from Gregory (1986).

"buffered" open system, the exchange trajectory is independent of the modal mineralogy. However, for a "closed" system, the exchange trajectory is strongly dependent upon the mode, and in the general open system case (eq. 22) there is a slight dependence on the modes because the eigenvalues and eigenvectors of matrix A depend on the X^'s. If magnetite is the more easily exchanged mineral, then the more magnetite-poor assemblages should plot higher up on Figure 17b. If the temperature of exchange is known, one can estimate the initial composition of the quartz from the intersection of the array with the exchange isotherm ("buffered" open system case; Fig. 18). As in Figure 18, isochrons in the general case would also have positive slopes even though the exchange paths of individual samples in these prograde metamorphic systems are likely to have had negative rather than positive slopes. When the existing data for quartz-magnetite pairs from Precambrian iron formations are considered together (Fig. 19), the data can be divided into two broad classes: deformed or higher-grade assemblages that lie close to, or along 45° arrays, and very low-grade assemblages that lie along steep disequilibrium arrays. Re-equilibration of a disequilibrium assemblage to a new equilibrium array is possible, i f not probable, under the higher temperature conditions inferred for some of these rocks. All of these samples may have undergone an earlier, open system burial metamorphic event. Given the slopes of the disequilibrium 119

8 1 8 0 QUARTZ

Figure 19. Summary 6-6 plot of Precambrian quartz-m agnetite data from the North American, African, Australian, and Asian continents. NA(D), contact aureole metacherts associated with the Duluth gabbro; WG, 3800 Ma metamorphosed cherts from West Greenland; KR, metamorphic cherts from Krlvoy Rog, Russia; K, low-grade cherts from the Kuroman formation, South Africa; H and WA, low-grade Hamersley and Weld Range cherts from Western Australia; and B, low-grade Biwabik cherts from the Great Lakes region of North America. Most of these rock suites exhibit either the steep, positive-sloped disequilibrium arrays that are indicative of open system behavior, or the 45° arrays in0.5 for the burial-metamorphic open systems are not unreasonable. Also important are the highly distinct positions of the disequilibrium arrays, indicating that the 6 1 8 0 values of the quartz in these terranes were considerably different before the burial metamorphic event. The important and unresolved question is whether this isotopic heterogeneity represents a true secular variation in the surface conditions of the Earth or whether i t is an artifact of the earlier diagenetic history (remembering that the quartz and magnetite may have started out as amorphous silica and some form of hydratediron oxide, respectively)? Mineral Pair Systematics Applied to Upper Mantle Assemblages Eclogite mineral pairs. 6 1 8 0 data from mantle eclogites provide an important constraint on the possible variability.of mantle assemblages (see review by Gregory and Taylor, 1986b). Figures 20ab illustrate the mineral pair system atics of mantle eclogite as well as of massif eclogite and blueschist Given our plate tectonic understanding of the 6 1 8 0 heterogeneity of oceanic crust, the spread in the data is hardly surprising. We now know from a number of studies of intact sections of oceanic lithosphere (Spooner et al., 1974; Heaton and Sheppard, 1977; Gregory and Taylor, 1981) that there is sufficient 6 1 8 0 variation (2 < 6 1 8 0 < 20) in such rocks to account for the variation seen in eclogites, which are presumably derived from them. It is interesting to note that at this point, mantle 120

8

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Figure 20a,b. Coexisting mineral 6 1 8 0 data f o r eclogite from massifs and mantle nodules, along with data from blueschlsts and garnet peridotite nodules. Note t h a t few of the eclogites l i e within the f i e l d of closed-system space f o r normal 6 1 S 0 = +6 mantle. Eclogite nodules have much closer a f f i n i t i e s with massif eclogites and with blueschist terranes related to convergent plate boundaries than they have with "normal" mantle-derived magmas. From Gregory and T a y l o r (1986b).

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4000 m) at low latitudes, as found in the high Andes of Peru (e.g., Landis and Rye, 1974) or mountains of New Guinea (Chivas et al., 1984). Major mountains near the sea cannot be excluded (e.g., New Guinea) nor a rapid increase in altitude. In a study of the very young (1.2 Ma) Ok Tedi porphyry copper deposit in New Guinea, Chivas et al. (1984) showed that 8D meteoric water changed from >-15 at the time of mineralization to -94 to -73 today, a major change in a very short period of time. For Panasqueira, interpreting the very large range in 8D values as a meteoric water-time effect implies that initially it was near sea level, but during the development of main stage mineralization, altitudes greater than about 4000 m were attained at or between Panasqueira and the ocean vapor source. An alternative interpretation invokes the role of organic waters. Fluid inclusion studies on Hercynian tungsten deposits have shown that not only are the fluids chemically complex, but that they are also quite variable during the evolution of the deposit (Turpin et al., 1981; Cheilletz, 1984; Giuliani, 1984; Ramboz et al., 1985). Fluid constituents include N 2 , C0 2 , CH 4 , etc., strongly suggesting the role of organic matter and sedimentary formations. "Graphitic" or organic-rich shales are closely associated with all of these deposits. If hydrogen in the graphitoids or organic matter (Sheppard and Charef, 1986) or ammonium micas (Dubessy and Ramboz, 1986) dominates the total hydrogen in some of the fluids through dehydration, oxidation or exchange reactions, this H 2 0-bearing fluid, called organic water by Sheppard and Charef (1986), is likely to be strongly depleted in deuterium (Fig. 6; see below). Further data are needed to test this hypothesis. The Panasqueira example has been treated in some detail because it illustrates a number of important points. Although ancient meteoric waters probably approximately followed the MWL, the problem is to know where the locality plotted on the MWL at a given time. The paleolatitude and paleoaltitude implications of such information are evident. To arrive at a coherent interpretation, a number of different approaches may have to be considered, including studying the surrounding environment, even on a regional scale. A choice among several alternative interpretations may not always be evident Connate and Formation Waters Retaining the definition that connate waters are fossil seawaters trapped in the sediments at the time of formation, their initial isotopic compositions would be identical to those of ancient seawater (Fig. 1). The H- and O-isotope compositions of formation waters from different sedimentary basins in North America (Clayton et al., 1966; Hitchon and Friedman, 1969; Kharaka et al., 1973), Europe (Dowgiallo and Tongiorgi, 1972; Sheppard and Langley, 1984) and southern Israel (Fleischer et al., 1977) are summarized on Figure 4. The majority of these formation waters come from depths between 500 and 3700 m. Most of the reservoir rocks are of marine origin and range back to at least the Ordovician. Most of these waters are in fact brines with 5 to 30 wt % total dissolved solids (e.g., Hitchon and Friedman, 1969). Formation waters from a given basin plot within a fairly well-defined isotopic field. The most lsO-depleted waters are usually the least saline and plot on or near the MWL (Fig. 4). Because of the geographic dependence of the 5D value at the low- ls O end of the individual fields (compare Fig. 3a with Fig. 4), Clayton et al. (1966) argued that formation waters are dominantly of meteoric water origin and not of seawater origin. This is an important conclusion with far-reaching hydrodynamic implications. Despite this and subsequent supporting studies, interstitial sedimentary brines are still sometimes considered to be dominantly of connate or evolved seawater origin. The 5 18 0-values of formation waters tend to increase with increase in temperature (Fig. 5) and increase in salinity. For many basins, 8D also tends to increase with increase in temperature or 5 l s O (Fig. 4). The degree of scatter, however, emphasizes that the interpretation of the isotopic and chemical properties of each sample is complex, assuming 172

that the samples are representative of the interstitial waters (sampling conditions are rarely stated). Nearby samples either from the same general sedimentary formation or from adjacent formations can be distinctly different (Sheppard and Langley, 1984). These observations imply that within a relatively restricted volume, brines may be of varying ages, chemistries, and histories. Although meteoric waters must often be a major component of formation waters, a diagenetically modified seawater (Hitchon and Friedman, 1969; Flesicher et al., 1977) and metamorphic water contribution (Kharaka et al., 1973) may also be present. The 5D data for the Alberta basin indicate that even deep and hot formation waters (2000-3500 m; 70-120°C) contain 30% or more of meteoric waters (Hitchon and Friedman, 1969). The characteristic ls O-shift to the right of the MWL is predominantly due to isotopic exchange with 1 8 0-rich sedimentary minerals and particularly carbonates (Fig. 1). The 8D variations are less well understood; mixing of different waters, exchange with H 2 S (depleted in D relative to water (Richet et al., 1977)), fractionation during membrane filtration, and exchange with hydrous minerals are some of the possibilities. The hydrothermal phase of carbonate-hosted Pb-Zn-(F-Ba) deposits such as the Mississippi Valley type or those associated with saliferous diapirs often involve fluids of formation-water origin (Hall and Friedman, 1963; Charef and Sheppard, 1985; Sheppard, 1984; Charef and Sheppard, 1986a,b; Chapter 14). Formation waters may also migrate and become a source of hydrothermal fluids (1) during magma emplacement in sedimentary formations and (2) when rock units are thrust over sediments containing formation waters (see Exotic Waters below). Metamorphic Waters The isotopic compositions of metamorphic waters are usually calculated from the isotopic compositions of the rocks at the temperature of interest (Taylor, 1974; Sheppard, 1977b, 1981). Metasediments typically have 5D ~ -40 to -100 and 5 l s O ~ +8 to +26 (e.g., Shieh and Taylor, 1969; Taylor, 1974; O'Neil and Ghent, 1975; Magaritz and Taylor, 1976; Hoernes and Friedrichsen, 1978; Chapter 13). Metabasalts and other associated rocks from the ocean crust and ophiolite complexes typically have 8D ~ -35 to -70 and 8 l s O ~ +3 to + 1 4 (e.g., Muehlenbachs and Clayton, 1971, 1972; Wenner and Taylor, 1973; Heaton and Sheppard, 1977; Sheppard, 1980; Gregory and Taylor, 1981; Chapter 12). Applying the H- and O-isotope fractionation factors (Chapter 1), the calculated isotopic field for metamorphic waters is 8D ~ 0 to -70 and S O ~ +3 to +20 (Figs. 4 and 6). The 8D range of this field has been extended up to 0 per mil to include values of metamorphic waters that could conceivably be derived from the dehydration of oceanic crust (Sheppard, 1981). The relatively large 1 8 0 / 1 6 0 variations reflect the wide range of compositions of the sedimentary (authigenic and detritic minerals) and meta-igneous parent rocks and the range of temperatures where isotopic fractionation factors tend to be quite variable. The partial overlap of the metamorphic water field with the higher temperature and more ls O-rich formation waters (Fig. 4) is to be expected, as there is a continuum between diagenesis and low-grade metamorphism. The boundaries of the metamorphic water field are clearly not absolute. H-isotope analyses of fluid inclusions of metamorphic origin in quartz lenses and veins in metamorphic terrains have compositions that are usually within the metamorphic water field (Rye et al., 1976). O-isotope values of the fluids are invariably calculated from an analysis of the host mineral and knowledge of the temperature (often the pressurecorrected homogenization temperature), because of O-isotope exchange reactions between the fluid inclusion and host mineral during cooling (Rye and O'Neil, 1968). It should be noted that metamorphic fluid inclusions studies are usually carried out on pod or vein minerals. These fluid samples may only represent a part of the fluid system.

173

Figure 6. Isotopic compositions and fields for seawaters, meteoric waters, primary magmatic waters, Comubian magmatic waters, metamorphic waters, and organic waters. The kaolinite weathering line is given for reference (Savin and Epstein, 1970). The 1 8 0-shift trends due to water-rock interaction, and exchange in hydrothermal systems are shown for seawater and meteoric waters of initial compositions A and B. Fields for pore waters in older oceanic crust and brines from the Canadian Shield are also given. Wilbur Spring is a possible modem metamorphic water (White et al., 1963).

Isotopic data for metamorphic springs are extremely limited because of problems of recognition and mixing with local ground waters. In the Coast Range, California, White et al. (1973) have interpreted the warm (~53°C) and overpressured Wilbur Springs as being of metamorphic origin because their D/H ratios are distinctly different from the local meteoric waters with 5D ~ -60 (Fig. 6). Magmatic Waters The H- and O-isotope compositions of magmatic waters are calculated from the measured isotopic compositions of fresh unaltered igneous rocks (or minerals) by applying the mineral-H 2 0 isotopic fractionation factors at temperatures of about 700°C to 1200°C. This indirect method is used because no water of undisputed magmatic origin has yet been directly analyzed isotopically. The great majority of nonaltered plutonic rocks of I-type (Chappell and White, 1974), cafemic association (Debon and LeFort, 1983) or magnetite series (Takahashi et al., 1980) have well-defined H- and O-isotope compositions: 8D ~ -50 to -90 and 5 l s O ~ +5.5 to +10 (see Figs. 1 and 2 in Chapter 8). Such values have often been referred to as "normal" igneous values. Waters in equilibrium with such magmas are estimated to have compositions in the range 5D = -40 to -80 and 8 l s O ~ +5.5 to +9.5 and were defined as primary magmatic water by Sheppard et al. (1969). Compared with the initial definition, the range of S l s O values has been extended from 7.0 to 5.5 to include the complete range of "normal" igneous values. The word "primary" implies that the water was transported with the magma from its deep-seated source. This primary magmatic water field is often used as a reference even for discussing Precambrian igneous rocks where the data base is still rather limited. 174

The H-isotope composition of mantle rocks that were apparently not modified during their emplacement in the crust have a similar range of compositions, 8D = -45 to -90 (see Fig. 2 in Chapter 8; see Chapter 5). Mantle hydrogen from oceanic and rift zone environments, however, generally has 5D values of -70 to -90 (Sheppard et al., 1977; Kyser and O'Neil, 1984). Whether this narrow range also represents that of subcontinental mantle, with all 8D values > -70 being values that were modified in the crust, has not yet been resolved. Some primary magmatic amphiboles in oceanic periodites, basalts, or their accompanying nodules are more D-rich (-23 > 5D > -45), and the involvement of seawater has been proposed for some of these samples (Sheppard and Epstein, 1970; Kuroda et al., 1977; Boettcher and O'Neil, 1980; Graham et al., 1982). Mantle-derived water is therefore indistinguishable from the range of values presented above for I-type magmatism. Many of the isotopic compositions of magmatic waters calculated to be in equilibrium with the high- ls O peraluminous granites do not plot within the primary magmatic water field. However, a magmatic water field can usually be defined for a given igneous complex. For example, Figure 6 shows the calculated field for Cornubian magmatic waters (Sheppard, 1977a) which is also generally applicable to aluminous associations in the Hercynian. In a number of special situations, magmatic waters with other isotopic compositions could be produced. These include, for example, low- 18Q magmas (see Chapter 8) and magmas that have undergone major devolatilization (Nabelek et al., 1983; Taylor et al., 1983; Chapter 7). Although the definition of magmatic water carries no implication concerning its ultimate origin, the isotopic ratios may contain some relevant information. The isotopic compositions of unaltered igneous rocks probably reflect in large part the isotopic ratios of the source region of the magma plus any modifications brought about by the assimilation and/or magma-wall-rock interaction processes and later unmixing processes. Primary magmatic water values are probably largely related to both mantle values and seawater-hydrothermally altered oceanic crust as it becomes involved in subduction processes, returning part of its water to the continents during magmatism and carrying some down into the mantle. The ls O-rich magmatic waters reflect the importance of low-temperature sedimentary minerals in the source region. The more variable D/H ratios of these rocks also indicate that minerals that have exchanged or formed in the presence of surface waters are involved. A magmatic-hydrothermal fluid will retain its magmatic H- and O-isotope signature only so long as its composition is controlled or buffered by exchange with the igneous magma or silicates at magmatic temperatures. Exchange during post-magmatic processes can modify both the 8D and 8 l s O values of magmatic water. Organic Waters Organic water has recently been defined to be a water whose D/H ratio is derived from the direct or indirect transformation of organic matter, bitumen, coal, kerogen, petroleum, organic gases, etc. by processes such as dehydration, dehydrogenation, oxidation and/or exchange (Sheppard and Charef, 1986). Dubessy and Ramboz (1986) have suggested that the destabilization of NH4-bearing micas could also contribute "organic" hydrogen to the system. The H-isotope composition of organic water is related to that of its organic source material by the appropriate isotopic fractionation factors. Sheppard and Charef (1986) have proposed H-isotope compositions for organic water in the range -90 > 8D > -250. The O-isotope compositions of organic waters are probably controlled by the 8 1 8 0 values of the reservoir rocks and the temperature of exchange; they will therefore be similar to formation or metamorphic water values (Fig. 6). The 8D-values of organic waters could be higher than -90, but then they become indistinguishable from formation and metamorphic waters. The upper limit of -90 is therefore a practical cut-off. 175

The production of organic water can be considered to occur during at least four stages during the evolution of an organic-matter-bearing sediment to a graphite-bearing metamorphic or igneous rock. (1) During very early diagenesis near the sediment-water interface, a fraction of the freshly deposited organic matter can undergo reactions that liberate water. This is not of immediate interest here. (2) The diagenetic and catagenetic transformation of kerogen (usually 80-90% of sedimentary organic carbon (Tissot and Welte, 1984)) and related hydrocarbons (bitumen, oil, and gas) can generate water by dehydration reactions represented here by a simplified overall reaction of the type [C20H20O6]

(kerogen 1) -> [C20H10O]

(kerogen 2) + 5H 2 0 ,

(2)

where the material in [ ] is a schematic representation of the substance, in this case kerogen. (3) Oxidation of kerogen, methane, petroleum, or hydrogen-bearing graphitoids during metagenesis and metamorphism releases water following possible simplified overall reactions of the type [C^O]

(kerogen) + 2[Oz] -> 20C (graphite) + 5H z O ,

(3)

with the oxygen being supplied, for example, during inorganic sulphate reduction. (4) At higher metamorphic temperatures, H-isotope exchange reactions between water or mineral hydrogen and organic hydrogen (kerogen and graphitoids) or H j or CH4 of organic origin can become effective. The occurrence of pure organic water in geologic systems is unlikely; in the sedimentary and diagenetic environments, formation waters are invariably present, and in metamorphic systems, water can be generated during prograde mineral reactions. Therefore, the term organic water only implies that there is a recognizable organic water contribution. Similarly, many formation waters and some metamorphic waters probably contain an organic water contribution that is not readily identifiable. Part of the variability of 8D in formation waters from a given basin may be related to variations in the organic water contribution. The concept of organic waters is perhaps more novel in the hydrothermal and metamorphic environment than during diagenesis. The evolution of sedimentary organic matter is often discussed with the aid of a van Krevelen diagram (see Fig. VI-3 in van Krevelen, 1961, or Fig. II.8.7 in Tissot and Welte, 1984). The van Krevelen diagram portrays some possible evolutionary trajectories for the production of H2O (our organic water), CH4, and CO2 from organic matter including coals and kerogen. There is extremely little discussion of what happens to organic matter under low-grade metamorphic conditions after the production of petroleum and dry gas and before arriving at pure graphite. It is noted, for example, that during metagenesis a reaction of the type (2) could generate about 1 gm HoO for every 3 gm of kerogen with H/C = 0.5 if a suitable source of oxygen is available. An organic-rich sediment, after dewatering and the dehydration of water-rich minerals like smectites, could yield a low-grade metamorphic rock where an important fraction of the total hydrogen is of possible organic origin. Subsequent prograde reactions could generate organic waters. Organic-hydrothermal waters have been proposed for the massive Pb mineralization that formed in the contact zone to a Triassic saliferous diapir at Fedi-el-Adoum, Tunisia, at 70°C to 90°C from saline waters with 8D ~ -110 and 5 l s O ~ +10 (Charef and Sheppard, 1986b; Sheppard and Charef, 1986)). Organic-hydrothermal fluids were proposed above as an alternative interpretation for the D-depleted, 180-rich fluids at Panasqueira and some other tungsten deposits. Further studies are needed to test this proposition and to determine the role of organic waters in systems at temperatures above 75°C. 176

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Hydrothermal Waters The isotopic trends of a typical seawater-hydrothermal and two different meteorichydrothermal fluids for two different localities A and B are shown on Figure 6. In these examples neither mixing phenomena nor unmixing or evaporation processes are considered. The constant 8D values of the waters indicate that H-isotope composition of the water overwhelmingly dominates the total hydrogen in the system (i.e., despite H- isotope exchange, the D/H ratio of the water buffers the system). The variable l S 0-shifts are a result of differences in the water-to-rock ratio at comparable temperatures of exchange. Magmatic-, metamorphic- or organic-hydrothermal waters generally have isotopic compositions that plot within their respective "boxes" (Fig. 6) unless modified by exchange or mixing processes. A water derived from the mixing of two different waters has a composition that plots on the line joining the two end members unless modified by isotopic exchange processes. The H- and O-isotope trends for a wide variety of active near-neutral hydrothermal systems are summarized on Figure 7. This figure illustrates a number of important points that can aid the interpretation of fossil geothermal systems (see Chapters 11 and 12). Excluding the Red Sea Brines, Reykjanes, Iceland, and Shimogamo, Japan, the 8D values in a given system are essentially identical to those of the local meteoric waters. These results indicate the overwhelming dominance (>95%) of meteoric waters in all of these thermal waters. The size of the "O-shift is highly variable — negligible at Wairakei but very large at Salton Sea and Lanzarote (Canary Islands). The chemical compositions and salinities are also variable with salinities tending to increase with increase in l s O-shift at comparable temperatures. The Salton Sea system illustrates that meteoric-hydrothermal waters can acquire isotopic characteristics that are very similar to primary magmatic waters. 177

The Reykjanes and Shimogamo systems are located on peninsulas surrounded by the sea. These thermal waters are of mixed meteroic-seawater origin (Arnasson, 1976; Sakai and Matsubaya, 1974). The thermal waters do not lie on simple mixing lines, indicating that water-rock interactions have modified their O-isotope compositions. The Red Sea hot brine pools do not have isotopic compositions that are similar to immediately surrounding Red Sea waters (Craig, 1966). Mixing processes are probably also important here, possibly with deep circulating coastal meteoric waters of Saudia Arabia (White, 1974). In addition to the above near-neutral thermal waters, acid hot springs are also dominantly of meteoric origin (Craig, 1963). During the evaporation processes, both the H- and O-isotope compositions are modified from the meteoric water value. Such evaporated waters typically plot about a line of slope three (Fig. 1). Exotic Waters Meteoric- and seawater-hydrothermal systems can be considered to be special cases of exotic-water hydrothermal systems because the source of the water can be identified as of meteoric or seawater origin. During the study of some systems, however, it can be demonstrated from the O- and/or H-isotope analyses that the origin of the fluid is external to the immediate system, but the specific origin cannot be identified with certainty. This is principally because waters of different origins can have overlapping isotopic compositions. For example, a fluid with 8D = -50 and 5YsO = +6 could be a modified meteoric water, a formation water, a metamorphic water, or a magmatic water (Figs. 4 and 6). In such a situation, additional isotopic and other data may not be available or sufficient to arrive at a specific identification of the origin of the water. In this case, it is convenient to use the term "exotic water" because it emphasizes that a fluid of unknown origin entered the system. In a study of late-Hercynian uranium-bearing veins within a high-grade metamorphic complex in the western French Alps, Negga et al. (1986) showed that two externally-derived fluids were involved during vein development. A D-rich (~-35 per mil), low f0 2 , saline fluid at 350°C to 400°C mixed with a less D-rich (-50 to -65 per mil), more oxidizing and less saline fluid that is consistent with a meteoric origin. The former fluid was interpreted to be an exotic fluid that was released from underlying formations into the hotter overlying metamorphic slab. The underlying formations, which are not locally exposed, could be sedimentary rocks containing formation waters or low-grade metamorphic rocks undergoing dehydration on being heated up by the overthrusted hot high-grade metamorphic slab. These exotic waters are therefore of probable formation water or metamorphic origin. Similarly, in a study of fluids in deep fault zones in Canada and Brazil, Keirich et al. (1984) proposed that two different fluids, both of external origin, were involved during the evolution of the fault system. Based on O-isotope data alone, they interpreted the higher temperature high- 18 0 exotic fluid to be of metamorphic origin followed by the entry of meteoric waters at lower temperature and at a shallower level. The thrust zone environment is a particularly favorable situation for generating exotic fluids by provoking dewatering or prograde metamorphic reactions in the underlying units as they become compacted and/or heated (reverse metamorphism) during burial (P^-^ > Pload). The efflux of these exotic fluids may be focused along fault and shear zones aiding movement or be released through the overlying formations via infiltration or hydrofracturing processes. If the overlying units are hot enough, the introduction of this exotic fluid may promote melting. These conditions can be met during subduction or during continent-continent collision events as exemplified by the generation of the High Himalaya leucogranites (LeFort, 1981; LeFort et al., 1986). Other Waters Waters can have isotopic compositions that plot to the left of the meteoric water line. 178

There are two well documented cases: (1) deep saline brines of the Canadian Shield, and (2) pore waters in sediments of the older oceanic crust. Saline Ca-Na-Cl brines with 10,000 to 330,000 mg l 1 total dissolved solids (TDS) occur in fault and shear zones at depths greater than 600 m throughout the Canadian Shield (Frape et al., 1984; Sheppard, Guha and Leroy, in preparation). In a given mine, 5D values increase with increase in salinity or increase in displacement from the meteoric water line, indicating mixing with local meteoric waters. The most saline brines are depleted in 8 1 8 0 by up to 8 per mil relative to meteoric waters with the same 8D value (Fig. 6). Brines with salinities greater than 200,000 mg l" 1 TDS typically have SD > -55 whether they come from Chibougamau, Sudbury, Thompson, or Yellowknife. Based on fluid inclusion and isotopic studies of zoned calcites directly associated with the brines at Chibougamau, Sheppard et al. (in preparation) have shown that the brines could be late-stage hydrothermal fluids, initially with T > 120°C, which reequilibrated with the wall-rock minerals during cooling to their present temperatures of about 25°C. The O-isotope compositions of the brines plot to the left of the meteoric water line because their compositions are controlled by the mineral oxygen and the low temperature of exchange (Fig. 1). The most D-rich brines could be very ancient but probably not as old as the main-stage mineralization whose 8D values are about -20 at Chibougamau. Pore waters in oceanic sediments at depths of 100 m or more (Fig. 6) are often depleted in l s O and D relative to actual seawaters (e.g., Lawrence and Gieskes, 1981). Decrease in 5 l s O and 8D with increase in the Ca 2 + ion concentration with depth are typically observed. Muehlenbachs and Clayton (1972), Lawrence and Gieskes (1981), and others have shown that the changes in the isotopic compositions of the pore waters are related to the alteration of basalts or of volcanic ash in the sediments. The isotopic compositions of the pore fluids are therefore controlled by those of the once high-temperature minerals and the low temperatures of exchange. The nature and extent of these low-temperature alteration processes in Layers 1 and 2 have important implications concerning the isotopic composition of subducted oceanic crust. CONCLUSIONS It is evident that coupled H- and O-isotope studies of natural waters and hydrous phases can identify, or at least place tight constraints on, the source or sources of waters. In turn, this has clarified our understanding of many geologic processes both at high and low temperatures because of the general importance of water-mineral reactions throughout the crust. Certain waters, such as seawater or meteoric water, have well-characterized isotopic compositions that clearly distinguish them. Other waters such as some high-temperature formation waters or magmatic or metamorphic waters, however, do not necessarily have distinctive compositions; such waters are usually represented by a field or "box" that may partly overlap other fields. Also, the "boundaries" of these fields are not absolute. Because of this situation, additional geologic data and arguments may be necessary to arrive at a sound interpretation of a given system. It is also evident that the areas of uncertainty are numerous and increase exponentially as we go back in time. This applies in particular to the Precambrian where the number of detailed multi-isotope studies is extremely limited. Inevitably, the interpretation of ancient systems involving aqueous fluids of seawater or meteoric origin is strongly influenced by our more detailed knowledge of the actual ocean water-meteoric system. Data on today's situation, however, represent but a single frame at the end of a long movie. When trying to reconstruct earlier frames, we need to employ multiple working hypotheses and examine the available data with an open but critical mind.

179

ACKNOWLEDGMENTS It is a pleasure to acknowledge the benefit of discussion with S. Epstein, to whom this volume is dedicated, and with H.P. Taylor, Jr. I am grateful to R. E. Criss and L. Toran for critical reviews of the manuscript and to J. Gorau, A. Legros, and Ch. Lehmann for technical assistance. Contribution No. 653, Centre de Recherches Petrographiques et Geochimiques.

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Chapter 7

Bruce E. Taylor MAGMATIC VOLATILES: ISOTOPIC VARIATION ofC, H, and S

This chapter discusses the application of carbon, hydrogen, and sulfur isotope measurements to the identification and tracing of volatiles in magmas (primarily H2O, CO2, SO2, and H 2 S). Magmatic volatiles are defined as those which are in chemical equilibrium with, or which exsolve from silicate liquids. Similar to the magmas themselves, magmatic volatiles may have complex histories reflecting source material, contamination and fractionation. Juvenile volatiles are those which have not previously participated in geological or biological processes in the earth's crust or hydrosphere (e.g., Chapter 6). Many magmas are emplaced in close proximity to the biosphere. This enhances the liklihood of contamination of the (e.g., Chapter 11) and, in some cases, of the magmas themselves. recent investigations of magmatic volatiles have focussed on volcanic glasses.

hydrosphere and magmatic rocks For this reason, the analysis of

One of the main points to be made in this chapter is that in most cases, isotopic variations are associated with gain or loss of volatile constituents. Therefore, assessment of volatile-related magmatic processes requires calculation of isotopic material balance. As a result of degassing, magmatic rocks may contain only a fraction of their original C, H, and S, and the retained volatiles will differ isotopically from the original volatiles. Consequently, the isotopic compositions of the exsolved, or separated volatiles must either be measured by tracing their e f f e c t s in the rocks, or be calculated using experimental or theoretical fractionation equations. SOLUBILITY, SPECIATION, AND CHEMICAL COMPOSITION Compositions and speciation of magmatic volatiles are determined by analysis of volcanic gases, glasses, minerals and fluid inclusions, or estimated by calculations of equilibria. This section reviews briefly solubility data (25 per mil. Taylor and Westrich (1985) determined a fractionation of +23.6 per mil at 950°C and 50 MPa between H 2 0 and rhyolite melt using both synthesis and reversed isotope exchange experiments. Thus, as H 2 0 separates from a vapor-saturated granitic magma, the 6D of the melt is lowered by material balance. If hydrogen speciation in hydrous glasses (Persikov, 1972; Stolper, 1982a, b) reflects speciation in vapor-saturated melts, then «Dmelt = 6 D g l a s s = X H 2 q * « D h 2 q * + U - X ) 0 H * « D q H * 191

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Table 1.

Isotope Fractionation Factors for Hydrogen, Sulfur, and Carbon Determined at Temperatures >400*C.

Phases A-B

Fractionation* 1000 In =

T Range CC)

Ref.

HYDROGEN H2O-H2 H 2 O-H 2 S H 2 O-HCI H 2 0-hydrous min. -muscovite -biotite -hornblende -tremolite -hornblende H 2 0-rhyolite melt

= -11.4KX 1 ) + 91.769(X 2 ) + 215.785(X)-87.44 = 3.03(X ») - 29.477(X ' ) + 316.50(X) - 98.97 = -1.15(X') + 11.636(X 2 ) + 227.143(X)-63.77 = 22.4(10'/T 2 )-28.2 - (2XAl-4XMg-68XFe) = = = = = = =

22.1 ( 1 0 ' / T 2 ) - 19.1 21.3 ( 1 0 ' / T 2 ) + 2.8 23.9 ( L O ' / T 2 ) - 7.9 21.1 ±2 31.0( l O ' / T 2 ) - 14.9 23.1 ± 2.5 23.6

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Figure 15. Comparison of 6D and { 1 ' O of volcanic gases with local meteoric waters, magmatic water from felsic continental magmas, and water in fresh MORB. Modified from Allard C1983>; additional data from Barnes (1984); Baskov et al. (1973). Abbreviations: Mo, Momotombo; W.Is, White Island; K, Krakatoa; M, Merapi; EA, Erta Ale; St.V, St. Vincent; St.H, Mount Saint Helens; S-I, SatsumaIwojima; El. M., El Misti; U, Usu; Sh, Showashinzan; E, Mount Etna; S, Surtsey; T, Tolbachik; A, Avacha. Table 6. Estimates of the Hydrogen Isotope Compositions of Magmatic and Deep-Seated (Mantle) Water Range of SD

Basis

Ref.

Deep-Seated Water -48 -70 -85 -80

± 20 to -40 ± 10 to -60

phlogopite ( S D phlogopite ( 6 D phlogopite ( 6 D phlogopite ( 6 D kimberlites

= -40 = -40 =-70 = -81

to - 7 0 ) in kimberlites to -70) in kimberlites t o - 2 8 ) in UM* and carbonatite t o -47) in UM xenoliths and

(1) (2) (3) (5)

M agmatic Water - M af ic M agmas -70 to -60 -80 ± 5 -60

biotite ( 6 D = - 7 0 to -80) hornblende ( 6 D = -81 to t o -70) in gabbro (plus calculation using (4)) MORB, OIB glass ( 6 D = - 8 3 to -33) MORB vesicle fluid

(6) (7) (8)

M agmatic Water - Felsic M agmas - 8 5 to - 5 0 -62 t o - 5 3 -37 to -29 -45 t o - 4 0 -18 -50 ± 4 -59 -32

biotite and hornblende in fresh granites - granodiorites (9) biotite ( 6 D = -115 to -84) Ibaragi granitic complex, (10) Japan (calculated using (4) plus subsolidus " O temperatures) hornblende ( 6 D = -102 t o -52) biotite ( 6 D = -86 to (II) -59) Kitakami Mtns. granites, Japan (calculated using (4)) obsidian tephra clasts ( 5 D = -51 to -90); western USA; (12) estimated 6 D h , q upon reaching saturation (see text) biotite ( 6 D = -103 t o - 6 6 ) muscovite ( i D = -61 t o (13) -41) Papoose Flat Pluton; estimated 6 D h 2 o u P ° n reaching saturation (see text) fluid inclusion H 2 0 in pocket quartz in layered aplite (14) pegmatite dikes range of 6D = - 3 6 t o - 7 9 average fluid inclusion H 2O in primary inclusions in (15) various granitic pegmatites; range of 6D = -32 to -94 extrapolation of high-temperature fumarole gas data; dacitic (16) Showashizan Volcano, Japan

* UM = ultramafic. References: (l)Sheppard and Epstein, 1970; (2) Sheppard and Dawson, 1975; (3) Kuroda et al., 1977a; (4) Suzuoki and Epstein, 1976; (5) Sheppard et al., 1977 (6) Sheppard et a l , 1971; (7) Kyser and O'Neil, 1984; (8) Moore, 1970; (9) Taylor, 1974; (10) Kuroda et al., 1977; ( l l ) K u r o d a et al., 1974; (12) Taylor and Westrich, 1985; see also Taylor et a l , 1983; (13)Brigham and O'Neil, 1985; (14) Taylor et a l , 1979; (15) Taylor and Friedrichsen, 1983; (16) Mizutani, 1978. 206

host rocks (Table 5; Allard, 1983). The dominance of meteoric H 2 0 and evolved brines in geothermal fluids has been well-known since the early work of Craig (1963); also see Truesdell and Hulston (1980) for a summary. Waters of several origins (e.g., evolved meteoric and ocean waters, magmatic water) are present in arc volcanic complexes, and their contributions to fumaroles vary with volcanic activity (e.g., Stewart and Rafter, 1975; Mizutani, 1978; Matsuo et al., 1985). For example, fumaroles in Tolbachik volcano, Kamchatka (T = 1000°C, 6D = -82 to -55) may contain perhaps 60-80% meteoric water (6D = -110 ± 5). Barnes (1984) deduced a mixture of 51% volcanic vapor (6D = -33) and 49% atmospheric water vapor (estimated 6D = -167) in a plume above a large fumarole at Mount St. Helens (6D = -97.9). From heat balance calculations, Gerlach and Casadevall (1986) suggested increasing admixture of meteoric water: from 20 to 30% after one year; only below 600 to 700°C is meteoric H 2 0 thought to make up to 70% of the total H 2 0 in the fumarole fluid. H 2 Q in Fluid Inclusions Mafic magma Kyser and O'Neil (1984) extracted small amounts of H 2 0 (generally with an average value of 37.2 per mil. Felsic magma Fluid inclusions in minerals from miarolitic cavities in layered granitic aplite-pegmatite dikes (San Diego Co., California) contain H 2 0 with 6D values from ca. -49 to -77; values of about -50 are common to many dikes (Taylor et al., 1979; Table 5). Hydrogen isotope fractionations between "coexisting" inclusion H 2 0 and muscovite (ca. 15 per mil) in some dikes indicate temperatures (ca. 525°C; Fig. 6) which are compatible with magmatichydrothermal conditions for tourmaline-precipitating melts (Taylor et al., 1979; London, 1986). Isotopic studies of other granitic pegmatites (Taylor and Friedrichsen, 1983) document an average 6D of -59 for inclusion H 2 0 , and rnagmatic temperatures based on hydrogen isotope fractionations between H 2 0 and micas. Hydrogen Isotopes in Basaltic Glasses Hydrogen isotope and water content data for MORB, OIB, and BAB glasses are compared in Figure 16. OIB glasses were both air-(pumice) and water-quenched (KER flows). The OIB and BAB glasses analyzed thus far exhibit generally higher and positively correlated H 2 0 and 6D, than found for most MORB. However, pumice samples are characterized by a very narrow range of water content but nearly a 30 per mil variation in 6D. Three processes could potentially contribute to these effects: (1) degassing, (2) contamination/assimilation or contamination/exchange with non-magmatic water, and (3) variations in source compositions and volatile solubilities of the magmas. The 6D of sequentially-erupted pumice from Kilauea was shown by Friedman (1967) to increase with decrease in wt % H 2 0 . Craig and Lupton (1976) interpreted this trend to be the result of open-system degassing, with an estimated H 2 0 - m e l t 207

Wt. % H 2 0

Figure 16 (left). Variation of SD and water content for MORB ( • ) , OIB ( • * ) , and BAB ( • ) ; PB = primary basalt. Data f r o m Poreda et al. (1986); Poreda (1985); Kyser and O'Neil (198*); Craig and Lupton (1976). Figure 17 (right). Plot of 6D versus 1 / H 2 0 for submarine and subaerial basalts showing contamination at different temperatures and with different water sources. Alternative explanation for OIBS and some MORBs is discussed in the t e x t . Abbreviations: PB, primary basalt; O + B, OIB and BAB; Soc. Is., Society Islands; H, KER flows, Hawaii ( O ) ; MT, Mariana Trough ( • ) and Lau Basin ( • M, Molokai ( • ) ; metabasalts ( A ) . Data sources as in Figure 16, plus Satake and Matsuda (1984).

fractionation of -30 per mil, i.e., H2O was depleted by 30 per mil relative to the magma. Their estimate of the initial 6D of the magma was == -79, based on the first-erupted pumice (with 0.099 wt % H 2 0). Diffusive loss of H 2 0 (e.g., to C 0 2 rich bubbles) during degassing might also explain enrichment of the melt in D, however, since the kinetic fractionation factor (proportional to (20/18)^) is 0.9487, and the HzO-melt fractionation is -52.7. 6 D - H 2 0 relations for the Kapoho pumice (flank eruption; 6D = -91 to -57) are unclear, perhaps as a result of a more complex degassing history (see Moore, 1965; Gerlach and Graeber, 1985). The magnitude and sign of the apparent fractionation require experimental confirmation to clarify these hydrogen isotope variations. The 6D-H z O relations are generally similar for glass from the KER flows, the Lau basin, and the Mariana Trough BAB (Poreda, 1985), and for fresh and altered basalt (Satake and Matsuda, 1979) (Figs. 16 and 17). The 6D versus 1 / H 2 0 trends for basalts and BAB glasses (Fig. 17) project to 6D = -27 to -25 at infinite H 2 0 (Satake and Matsuda, 1979; Poreda, 1985), whereas KER data project to slightly higher (less negative) 6 D values. These projections are consistent with contamination of magma and/or glass by seawater, where 1000 In a n 2 o - m e l t / g l a s s = 0 to 25 per mil, but also with magmatic degassing (if the H 2 0 - m e l t fractionation is positive). Kyser and O'Neil (1984) suggested absorption of seawater by magmas during emplacement of the KER basalts. Poreda (1985) interpret similar variations in BAB glasses (Lau Basin and Mariana Trough) as various mixtures of MORB and H 2 0-rich, high 5D "slab" components. Rare gas systematics may support contamination of basalt (He, Poreda, 1985; Ar, Jambon et al., 1985; Kyser and O'Neil, 1984), but the timing of contamination is unresolved. Limited H 2 0 diffusivity (Fig. 9), plus 6D versus 1/H 2 0 relationships for pillow rims and cores (not shown), suggest most H 2 0 enrichment occurs prior to emplacement and quenching of pillow margins. 208

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A degassing process would require an H 2 0 - and D-enriched mantle source. Helium isotope evidence (high values of 'He/^He in support of such a source on the Rekjanes Ridge were considered recently by Poreda et al. (1986). Accordingly, 6D values for MORB of ca. -80 might more nearly represent degassed than undegassed magma. Hydrogen Isotopes in Felsic Magmas The results of hydrogen isotope studies of water in quenched felsic magmas (obsidian) from young, subaerial volcanoes are best interpreted (aided) if the sampling is done in conjunction with careful geologic studies. The 6 D - H 2 O relations described below are of particular genetic significance because of the known stratigraphic relations of the obsidian clasts. A temporal variation in isotopic composition and water content can be established, which signifies a degassing process. Eichelberger and Westrich (1982) documented a decrease in the water content of obsidian clasts from successively-erupted tephra layers (wt % H2O > OA), followed by dome-forming, obsidian flows (

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C r-i E 1 O™ •* . I s « E Ü * 4-1 (M IQ s í s ¿ .H C i « o> i - " ?S 5 Q ûb ^U !.uj O i- 5 rt Il i o S £ .5 J | U E 'o S 2 < O . U 00 Ë VO ¿ . U 4» ¡j Q¡ «x s « s ì > 00 oE c a. 8. . o >noo ¿r > 8 -s o. 4 x ra o 4> .2 S V a i X> îa 7 u » u ! o >o S -o ç " ¡ 1 8 . O M 1 1 \ . g < 3 1 ~ ».a C-" »I-) g S-" !. "ra ! u 2. «s S| s " Í5 ^ r?5 ¡ s » - E ° .S00 E ö y c ai oo "gjäS in E n u > 8. .a ! £ 2 t .£ 'g o 'S £ u>- s l i a i e ° s - 5P "5. •o "3H o J= — JZ s* 2 C 4) b uj Jt oiUi = U » Q.. co -C &u = -g.S-5 i o ó I J¿ o w > ° \ ¿2 uj _ i s -c J X g. E S 2 | < So 8 h ^ TI i •Q,S ^Í fl O E S_ O-OII M- C I S 3 d » : ° Ä c u .3 00 O O e^ s " _! O U 3o „¿ £I s ' aO „-E" « e CN ^ w £ 1 , o -o •o X g ^ Í4) isfc4) Coo = •s Í Éftj tú ra . .C ! S l ' S -u ; «.SE O ç - t - lTP£ Ss l K , W3 4» MoO ^ M rî ^ E* 3 „ " -Je c t 0.7100) is thus definitely primary, and cannot be the result of assimilation. There is no way that such uniform isotope ratios could be produced within the continental crust by varying amounts of interaction with such isotopically heterogeneous country rocks (see Fig. 7). Some process such as metasomatism must have enriched and homogenized a large part of the upper mantle beneath Central Italy in radiogenic strontium (and presumably in other LIL elements, as well). M. Vulsini Volcanic Center Ferrara et al. (1986) divided the M. Vulsini samples into 3 main groups, arbitrarily termed A, B, and C (Fig. 18). Group A would correspond to the High-K series, and portions of Groups B and C would correspond to the Low-K series, as defined by Appleton (1972). Appleton (1972) attributed great im portance to the Ca-rich (sub-group A*) lavas, because these are among the best candidates for a primitive magma in the Roman Province. These form a compact grouping in the lower left-hand corner of Figure 18. Fractional crystallization of olivine and clinopyroxene will drive such magmas upward and to the right on Figure 18, toward higher SiO2 and K2O (Appleton, 1972). This process can explain much of the major element variation in the Vulsini samples, butit cannot account for most of the trace-element and isotopic variations. On Figure 19, the above-described fractional crystallization process drives the residual magmas nearly horizontally off to the l e f t As discussed above, there will only be at most a slight enrichment of S 1 8 0 during this process. Thus, a diagram like Figure 19 provides a sensitive test for closed-system fractional crystallization. The primitive, Ca-rich (A*) samples have relatively low 6 1 8 0 values, and i t is clear that virtually none of the other sam pies from Vulsini could have formed from these "primitive" Vulsini magmas by a simple, closed-system process. The most highly differentiated samples (Aa, A-C, and C) display the highest 6 1 8 0 values, suggesting either an AFC process (Taylor, 1980) and/or a process of mixing of high-180 anatectic magmas with the undersaturated Roman magmas (Taylor and Turi, 1976). Mixing curves are essentially straight lines on diagrams like Figure 19, because the oxygen contents of most rocks and magmas are similar. Note that the simple mixing line between average Tuscan metasedimentary basement and the primitive Vulsini A* magma-type passes very close to two lava samples from the Torre Alfina locality, which is one of the earliest Vulsini lavas, and is a lava locality well known for its high content of sedimentary xenoliths.

263

C a O Cwt. %) Figure 19. Plot of 6 l e 0 vs. CaO for the M. Vulsini samples analyzed by Ferrara etal. (1986) and by Holm and Munksgaard (1982). The symbols are the same as used in Figure 18. The stippled area represents a plausible (< 1 per mil) enrichment in 1 8 0 during strong, closed-system fractional crystallization (see text).

0.710

0.711 87

0.712

Sr/86Sr

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264

The Vulsini 6 1 8 0 and 8 7 Sr/ 8 6 Sr data are compared with the data from the Alban Hills (Ferrara et al., 1985) on Figure 20. Although all of the fields overlap, note the steep positive correlation shown by the A-Group samples, similar to the trend displayed by the Alban Hills data set The B and C samples show an even more clear-cut positive correlation trend with a shallower slope. The primitive, Ca-rich, A* samples overlap the Alban Hills data-points, particularly the Group II Alban Hills samples (Ferrara et al., 1985), which are the ones chemically most similar to the Vulsini samples. Mixing Models Involving Tuscan Basement Rocks Two characteristic, hiqh- 18 0, high- 87 Sr Tuscan basement schists are shown on Figure 21, one with 8 7 Sr/ 8 6 Sr = 0.725 and S 1 8 0 = +15.5, containing 180 ppm Sr, and the other with 8 7 Sr/ 8 6 Sr = 0.733 and 6 1 8 0 = +15.0, containing 200 ppm Sr. Two low- 1 8 0, low- 8 7 Sr end-members are also shown: (1) an H KS end member with a 6 1 8 0 = +5.8 and 8 7 Sr/ 8 6 Sr = 0.71024 (2500 ppm Sr) which plots at the low - 1 8 0 end of the Alban Hills field; and (2) a purely hypothetical LKS end member with 6 1 8 0 = +5.8 and 8 7 Sr/ 8 6 Sr = 0.7085 (1250 ppm Sr). Mixing curves are constructed for all 4 of the above end-member compositions (Fig. 21). The convex-upward mixing curve shown in Figure 21 is in striking contrast to the convex-downward "mantle mixing" or "source-contamination" curve proposed by Holm and Munksgaard (1982) for these same rocks, based on the very high 6 1 8 0 values of the lavas at Cupaello and San Venanzo (Fig. 15). Turi et al. (1986) demonstrate clearly that the convex-downward curve has no validity. I t i s obvious that the Ca-rich A* magmas at Vulsini and the abundant primitive lavas and tuffs at the Alban Hills represent much more plausible High-K Series parent magmas in the Roman Province than the tiny lava flows at San Venanzo and Cupaello; the latter magmas were apparently strongly enriched in 1 8 0 by exchange with the continental crust as they penetrated upward along very narrow, freshly-generated, volcanic conduits. The most highly contaminated lavas at M. Vulsini in the present study are also from tiny eruptions, namely the Torre Alfina lavas, which are among the oldest volcanic products at M. Vulsini. The Torre Alfina activity followed directly after an earlier period of anatexis that produced Tuscan rhyolitic magmas in this general area, including the voluminous hybrid magmas at the nearby M. Cimini complex. This relationship supports the idea (Ferrara et al., 1985; Taylor et al., 1984) that, other factors being equal, the most highly contaminated magmas are typically: (1) the earliest ones that come up through the conduits; (2) the smallest volume eruptions; and (3) those that encountered the hottest country rocks. Figure 22 is analogous to Figures 7 and 17, and it compares strontium data on some nearby Tuscan volcanic centers with the data fnim M. Vulsini and the Alban Hills. It is obvious that the high- 1 8 0, high- ° ' S r Vulsini lavas are geochemically transitional to the even higher- 18 0, higher- 87 Sr Tuscan volcanoes at M. Cimini, M. Amiata, and Radicofani (Taylor and Turi, 1976; Poli et al., 1984). The M. Amiata data envelope forms a linear array (AFC or mixing line?) that points toward that part of Figure 22 that contains both the bulk of the AGroup samples from Vulsini and the Alban Hills lavas. Similarly, the M. Cimini lavas also form an approximately linear array that points toward the same set of High-K Series data-points, and these M. Cimini lavas lie very close to the mixing line that links this HKS end member with typical Tuscan metasedimentary basement rocks. All of the data-points on Figure 22 lie within the triangular region bounded by the 0.71024 "Alban Hills line" and the "Torre Alfina mixing line", which intersect at the M ML or "Mantle Mixing Line" from Figure 17. In other words, virtual265

Figure 21. A graph similar to Figure 20, modified after Ferrara et al. (1986), and expanded to include some data on Tuscan metasedimentary basement rocks (schists and slates, SV-lla, 12b, 12c). The various fields delineated in Figure 20 are also shown, together with 4 hypothetical mixing curves (see text).

ly all of the High-K Series lavas in the Roman Province to the north of the Alban Hills can be explained by a combination of fractional crystallization and mixing of Tuscan magmas and/or Tuscan basement rocks with a unique High-K magma end member whose 8 7 Sr/ 8 6 Sr ratio and Sr content lie at the intersection of all these mixing lines on Figure 22. Also shown on Figure 22 are all the available data from Tuscan basement rocks and from the Tuscan rhyolites at Roccastrada and San Vincenzo (Fig. 15). The diagram is largely self-explanatory: all of the Tuscan rocks plot between the "Torre Alfina line" and the "Alban Hills line", within the triangle that has its lower-left vertex on the "Mantle Mixing Line" (Fig. 23). It is evident that either a variety of AFC processes are involved (e.g. Fig. 7), or that the Tuscan "end member1' involved in these mixing processes is extremely heterogeneous in 87 Sr/ 8 6 Sr and Sr content (and 6 1 8 0 , ~ +11 to + 20, see Taylor and Turi, 1976 and Taylor et al., 1984). It is likely that both effects are present, illustrating one of the main problems of applying the simple AFC models; namely, because of its heterogeneity it is very difficult to accurately constrain the isotopic and other geochemical characteristics of the assimilated country-rock end member. Summary The HKS parent magmas at M. Vulsini had an extremely uniform 8 7 Sr/ 8 6 Sr = 0.7102 to 0.7103, identical to that of the primitive High-K Series magmas at the Alban Hills volcanic center, 120 km to the south (Ferrara et al., 1985; 1986). These 8 7 Sr/ 8 6 Sr values are enormously more radiogenic than either mid-ocean ridge basalts (MORB) or the bulk Earth, indicating a major upper mantle enrichment event. Thus, during the last few million years, the upper mantle beneath central Italy apparently underwent a large-scale (metasomatic) mixing process that introduced radiogenic strontium into the source regions of the leucite-bearing volcanic rocks, leading to the production of prodigious volumes of High-K Series magmas. Another mixing process is evident in the upper mantle to the south of the Alban Hills, between this HKS magma or parent material and a characteristic LKS material (Fig. 23). 266

1/Sr(ppm"1)x10 4 Figure 22. Plot of 8 7 S r / 8 6 S r vs. 1/Sr (an expanded version of Fig. 17), modified after Ferrara et al. (1986), showing the fields of M. Vulsini data points and Alban Hills data points, and including the fields of the Tuscan volcanoes of M. Cimini, Radicofani, and M. Amiata (Hawkesworth and Vollmer, 1979; Vollmer, 1977; and Poli et al., 1984), as well as additional data from Tuscan rhyolites, Tuscan metasedimentary basement, and country rocks encountered in a drill hole in the Larderello geothermal area in Tuscany (G. Ferrara, unpublished data). M M l = Mantle Mixing Line of Ferrara et al. (1985); see Figure 23. T.A. = Torre Alfina. The positions of some data-points plotting outside the figure are indicated by the arrows.

Although the 6 1 8 0 values of these primitive High-K Series magmas are not known precisely, they clearly range from values as low as + 5.5 to values as high as +7.5. Can it just be a coincidence that this is the identical range observed in the mantle peridotite xenoliths analyzed by Kyser et al. (1981, 1982) and the mantle phologopites and amphiboles analyzed by Sheppard and Epstein (1970) and Boettcher and O'Neil (1981)? We think not, and in that respectit may be possible to integrate the conclusions of Ferrara et al. (1985; 1986) with those of Gregory and Taylor (1986a; 1986b), all of whom conclude that metasomatic " O enrichment events of one to two per mil must occur in the upper mantle. Most of the Roman leucite-bearing magmas had

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