Water in Nominally Anhydrous Minerals 093995074X

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REVIEWS in MINERALOGY and Geochemistry Volume 62

2006

Water in Nominally Anhydrous Minerals EDITORS

Hans Keppler

Universität Bayreuth Bayreuth, Germany

Joseph R. Smyth University Colorado Boulder, Colorado

Cover Photograph : Thin section of a garnet lherzolite mantle xenolith

from Pali-Aike, Patagonia. The almost colorless grains are olivine, orthopyroxene is brownish-green, clinopyroxene bright green and garnet is red. Grain size is about 1 mm. Photograph courtesy of Sylvie Demouchy.

Series Editor: Jodi J. Rosso GEOCHEMICAL SOCIETY MINERALOGICAL SOCIETY of AMERICA

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“Journey to the Center of the Earth”: In 1864, the French writer Jules Verne published his novel “Voyage au Centre de la Terre” (Journey to the Center of the Earth). In this novel, Otto Lidenbrock, a German professor of mineralogy (!) discovers a secret handwriting inside an old manuscript. The handwriting appears to show a way to get to the center of the Earth by climbing down the conduit of an extinct volcano on Iceland. Together with his nephew and a guide from Iceland, Lidenbrock follows this trail. On his way to the center of the Earth, he discovers, among many other things, a gigantic ocean in the Earth´s interior. This scene is depicted in the woodcut illustration overleaf, which is taken from the first complete edition of the works of Jules Verne in French. The idea of a major water reservoir in the Earth´s interior has therefore been anticipated already more than one century ago. Research in the last decades has entirely confirmed this idea, aside from some rather minor details, which are described in this book.

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Short Course Series Dedication Dr. William C. Luth has had a long and distinguished career in research, education and in the government. He was a leader in experimental petrology and in training graduate students at Stanford University. His efforts at Sandia National Laboratory and at the Department of Energy’s headquarters resulted in the initiation and long-term support of many of the cutting edge research projects whose results form the foundations of these short courses. Bill’s broad interest in understanding fundamental geochemical processes and their applications to national problems is a continuous thread through both his university and government career. He retired in 1996, but his efforts to foster excellent basic research, and to promote the development of advanced analytical capabilities gave a unique focus to the basic research portfolio in Geosciences at the Department of Energy. He has been, and continues to be, a friend and mentor to many of us. It is appropriate to celebrate his career in education and government service with this series of courses.

Reviews in Mineralogy and Geochemistry, Volume 62 Water in Nominally Anhydrous Minerals ISSN 1529-6466 ISBN 0-939950-74-X

Copyright 2006

The MINERALOGICAL SOCIETY of AMERICA 3635 Concorde Parkway, Suite 500 Chantilly, Virginia, 20151-1125, U.S.A. www.minsocam.org

The appearance of the code at the bottom of the first page of each chapter in this volume indicates the copyright owner’s consent that copies of the article can be made for personal use or internal use or for the personal use or internal use of specific clients, provided the original publication is cited. The consent is given on the condition, however, that the copier pay the stated per-copy fee through the Copyright Clearance Center, Inc. for copying beyond that permitted by Sections 107 or 108 of the U.S. Copyright Law. This consent does not extend to other types of copying for general distribution, for advertising or promotional purposes, for creating new collective works, or for resale. For permission to reprint entire articles in these cases and the like, consult the Administrator of the Mineralogical Society of America as to the royalty due to the Society.

WATER in NOMINALLY ANHYDROUS MINERALS 62

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From the Series Editor The review chapters in this volume were the basis for a four day short course on Water in Nominally Anhydrous Minerals held in Verbania, Lago Maggiore, Italy (October 1-4, 2006). The editors Hans Keppler and Joe Smyth have done an excellent job organizing this volume and the associated short course. Meeting deadlines (often ahead of schedule!) and keeping track of so many authors can be a thankless job at times but I truly appreciate all their hard work. Hans’ “friendly reminder” e-mails certainly kept us all on task and his eye for detail (small and large) made my job much more enjoyable! I extend my sincere thanks to him for his efforts! Any supplemental material and errata (if any) can be found at the MSA website www.minsocam.org. Jodi J. Rosso, Series Editor West Richland, Washington August 2006

PREFACE Earth is a water planet. Oceans of liquid water dominate the surface processes of the planet. On the surface, water controls weathering as well as transport and deposition of sediments. Liquid water is necessary for life. In the interior, water fluxes melting and controls the solidstate viscosity of the convecting mantle and so controls volcanism and tectonics. Oceans cover more than 70% of the surface but make up only about 0.025% of the planet’s mass. Hydrogen is the most abundant element in the cosmos, but in the bulk Earth, it is one of the most poorly constrained chemical compositional variables. Almost all of the nominally anhydrous minerals that compose the Earth’s crust and mantle can incorporate measurable amounts of hydrogen. Because these are minerals that contain oxygen as the principal anion, the major incorporation mechanism is as hydroxyl, OH−, and the chemical component is equivalent to water, H2O. Although the hydrogen proton can be considered a monovalent cation, it does not occupy same structural position as a typical cation in a mineral structure, but rather forms a hydrogen bond with the oxygens on the edge of the coordination polyhedron. The amount incorporated is thus quite sensitive to pressure and the amount of H that can be incorporated in these phases generally increases with pressure and sometimes with temperature. Hydrogen solubility in nominally anhydrous minerals is thus much more sensitive to temperature and pressure than that of other elements. Because the mass of rock in the mantle is so large relative to ocean mass, the amount that is incorporated the nominally anhydrous phases of the interior may constitute the largest reservoir of water in the planet. Understanding the behavior and chemistry of hydrogen in minerals at the atomic scale is thus central to understanding the geology of the planet. There have been significant recent advances in the detection, measurement, and location of H in the nominally anhydrous silicate and oxide minerals that compose the planet. There have also been advances in experimental methods for measurement of H diffusion and the effects of H on the phase 1529-6466/06/0062-0000$05.00

DOI: 10.2138/rmg.2006.62.0

Water in Nominally Anhydrous Minerals ‒ Preface boundaries and physical properties whereby the presence of H in the interior may be inferred from seismic or other geophysical studies. It is the objective of this volume to consolidate these advances with reviews of recent research in the geochemistry and mineral physics of hydrogen in the principal mineral phases of the Earth’s crust and mantle.

The chapters We begin with a review of analytical methods for measuring and calibrating water contents in nominally anhydrous minerals by George Rossman. While infrared spectroscopy is still the most sensitive and most convenient method for detecting water in minerals, it is not intrinsically quantitative but requires calibration by some other, independent analytical method, such as nuclear reaction analysis, hydrogen manometry, or SIMS. A particular advantage of infrared spectroscopy, however, is the fact that it does not only probe the concentration, but also the structure of hydrous species in a mineral and in many cases the precise location of a proton in a mineral structure can be worked out based on infrared spectra alone. The methods and principles behind this are reviewed by Eugen Libowitzky and Anton Beran, with many illustrative examples. Compared to infrared spectroscopy, NMR is much less used in studying hydrogen in minerals, mostly due to its lower sensitivity, the requirement of samples free of paramagnetic ions such as Fe2+ and because of the more complicated instrumentation required for NMR measurements. However, NMR could be very useful under some circumstances. It could detect any hydrogen species in a sample, including such species as H2 that would be invisible with infrared. Potential applications of NMR to the study of hydrogen in minerals are reviewed by Simon Kohn. While structural models of “water” in minerals have already been deduced from infrared spectra several decades ago, in recent years atomistic modeling has become a powerful tool for predicting potential sites for hydrogen in minerals. The review by Kate Wright gives an overview over both quantum mechanical methods and classical methods based on interatomic potentials. Joseph Smyth then summarizes the crystal chemistry of hydrogen in high-pressure silicate and oxide minerals. As a general rule, the incorporation of hydrogen is not controlled by the size of potential sites in the crystal lattice; rather, the protons will preferentially attach to oxygen atoms that are electrostatically underbonded, such as the non-silicate oxygen atoms in some high-pressure phases. Moreover, heterovalent substitutions, e.g., the substitution of Al3+ for Si4+, can have a major effect on the incorporation of hydrogen. Data on water in natural minerals from crust and mantle are compiled and discussed in three reviews by Elisabeth Johnson, Henrik Skogby and by Anton Beran and Eugen Libowitzky. Among the major mantle minerals, clinopyroxenes usually retain the highest water contents, followed by orthopyroxenes and olivine, while the water contents in garnets are generally low. Most of these water contents need to be considered as minimum values, as many of the mantle xenoliths may have lost water during ascent. However, there are some cases where the correlation between the water contents and other geochemical parameters suggest that the measured water concentrations reflect the true original water content in the mantle. The basic thermodynamics as well as experimental data on water solubility and partitioning are reviewed by Hans Keppler and Nathalie Bolfan Casanova. Water solubility in minerals depends in a complicated way on pressure, temperature, water fugacity and bulk composition. For example, water solubility in the same mineral can increase or decrease with temperature, depending on the pressure of the experiments. Nevertheless, the pressure and temperature dependence of water solubility can be described by a rather simple thermodynamic formalism and for most minerals of the upper mantle, the relevant thermodynamic parameters are known. The highest water solubilities are reached in the minerals wadsleyite and ringwoodite stable in the transition zone, while the minerals of the lower mantle are probably mostly dry. The rather limited experimental data on water partitioning between silicate melts and minerals are reviewed by Simon Kohn and Kevin Grant. One important observation here is that comparing vi

Water in Nominally Anhydrous Minerals ‒ Preface the compatibility of hydrogen with that of some rare earth element is misleading, as such correlations are always limited to a small range of pressure and temperature for a given mineral. The stabilities of hydrous phases in the peridotite mantle and in subducted slabs are reviewed by Daniel Frost and by Tatsuhiko Kawamoto. While most of the water in the mantle is certainly stored in the nominally anhydrous minerals, hydrous phases can be important storage sites of water in certain environments. Amphibole and phlogopite require a significant metasomatic enrichment of Na and K in order to be stabilized in the upper mantle, but serpentine may be an important carrier of water in cold subducted slabs. The diffusion of hydrogen in minerals is reviewed by Jannick Ingrin and Marc Blanchard. An important general observation here is that natural minerals usually do not loose hydrogen as water, but as H2 generated by redox reaction of OH with Fe2+. Moreover, diffusion coefficients of different mantle minerals can vary by orders of magnitude, often with significant anisotropy. While some minerals in a mantle xenolith may therefore have lost virtually all of their water during ascent, other minerals may still preserve the original water content and in general, the apparent partition coefficients of water between the minerals of the same xenolith can be totally out of equilibrium. Accordingly, it would be highly desirable to directly deduce the water content in the mantle from geophysical data. One strategy, based on seismic velocities and therefore ultimately on the effect of water on the equation of state of minerals, is outlined by Steve Jacobsen. The dissolution of water in minerals usually increases the number of cation vacancies, yielding reduced bulk and shear moduli and seismic velocities. Particularly, the effect on shear velocities is strong and probably larger than the effect expected from local temperature variations. Accordingly, the vS / vP ratio could be a sensitive indicator of mantle hydration. A more general approach towards remote sensing of hydrogen in the Earth’s mantle, including effects of seismic anisotropy due to lattice preferred orientation and the use of electrical conductivity data is presented by Shun-ichiro Karato. Probably the most important effect of water on geodynamics is related to the fact that even traces of water dramatically reduce the mechanical strength of rocks during deformation. The physics behind this effect is discussed by David Kohlstedt. Interestingly, it appears that the main mechanism behind “hydrolytic weakening” is related to the effect of water on the concentration and mobility of Si vacancies, rather than to the protons themselves. Water may have major effects on the location of mantle discontinuities, as reviewed by Eiji Ohtani and Konstantin Litasov. Most of these effects can be rationalized as being due to the expansion of the stability fields of those phases (e.g., wadsleyite) that preferentially incorporate water. Together with other geophysical data, the changes in the depths of discontinuities are a promising tool for the remote sensing of water contents in the mantle. The global effects of water on the evolution of our planet are reviewed in the last two chapters by Bernard Marty, Reika Yokochi and Klaus Regenauer-Lieb. By combining hydrogen und nitrogen isotope data, Marty and Yokochi demonstrate convincingly that most of the Earth´s water very likely originated from a chondritic source. Water may have had a profound effect on the early evolution of our planet, since a water-rich dense atmosphere could have favored melting by a thermal blanketing effect. However, Marty and Yokochi also show very clearly that it is pretty much impossible to derive reliable estimates of the Earth´s present-day water content from cosmochemical arguments, since many factors affecting the loss of water during and after accretion are poorly constrained or not constrained at all. In the last chapter, Klaus Regenauer-Lieb investigates the effect of water on the style of global tectonics. He demonstrates that plate tectonics as we know it is only possible if the water content of the mantle is above a threshold value. The different tectonic style observed on Mars and Venus may therefore be directly related to differences in mantle water content. Earth is the water planet – not just because of its oceans, but also because of its tectonic evolution. vii

Water in Nominally Anhydrous Minerals ‒ Preface Acknowledgments This volume and the accompanying short course in Verbania were made possible by generous support from the Mineralogical Society of America, the Geochemical Society, the United State Department of Energy, the German Mineralogical Society and Bayerisches Geoinstitut. The Verbania short course is the first MSA/GS short course ever held in Italy. We are very grateful for the generosity and the international spirit of the supporting institutions, which made this project possible. The preparation of the short course benefited enormously from the permanent advice by Alex Speer. Finally, we would like to thank Jodi Rosso for the efficient and professional handling of the manuscript and for her patience with authors and editors who ignore deadlines. August 2006 Hans Keppler Bayreuth, Germany

Joseph R. Smyth Boulder, Colorado, USA

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TABLE of CONTENTS

1

Analytical Methods for Measuring Water in Nominally Anhydrous Minerals George R. Rossman

INTRODUCTION.....................................................................................................................1 ANALYTICAL METHODS......................................................................................................2 Early infrared studies......................................................................................................2 Quantitative IR methods.................................................................................................3 Mineral specific calibrations..........................................................................................8 Thermogravimetric methods..........................................................................................9 P2O5 cell coulometry ......................................................................................................9 Hydrogen extraction with uranium reduction methods................................................10 Nuclear methods for hydrogen determination..............................................................11 Nuclear magnetic resonance.........................................................................................16 Secondary ion mass spectrometry (SIMS)...................................................................18 PREVIOUS REVIEWS OF METHODS.................................................................................20 SURFACE WATER..................................................................................................................21 CURRENT STATUS OF CALIBRATIONS............................................................................23 GLASSES................................................................................................................................23 ACKNOWLEDGMENTS........................................................................................................24 REFERENCES........................................................................................................................24

2 The Structure of Hydrous Species in

Nominally Anhydrous Minerals: Information from Polarized IR Spectroscopy Eugen Libowitzky and Anton Beran

INTRODUCTION...................................................................................................................29 The importance of hydrous species in NAMs..............................................................29 Why use IR spectroscopy?...........................................................................................30 History..........................................................................................................................30 CONCEPTS OF INFRARED SPECTROSCOPY...................................................................31 Introduction to IR spectroscopy...................................................................................31 Sample requirements....................................................................................................32 ix

Water in Nominally Anhydrous Minerals ‒ Table of Contents Experimental equipment ..............................................................................................32 QUANTITATIVE DATA FROM INFRARED SPECTROSCOPY.........................................33 The distance - frequency correlation of hydrogen bonds.............................................33 The spatial orientation of hydrous species...................................................................34 Total absorbance: a first step towards quantitative water analysis...............................35 CONCEPTS OF STRUCTURAL MODELS FROM INFRARED DATA..............................36 Charge balance and substitution...................................................................................36 Electrostatic considerations on defect geometry..........................................................37 Space requirements: ideal and distorted models..........................................................38 Influence on band energies from cation substitution....................................................38 Discrimination among hydrous defects........................................................................39 Deuteration...................................................................................................................40 EXAMPLES............................................................................................................................40 Vesuvianite: orientation and hydrogen bonding of hydroxyl groups...........................40 Hydrogarnet substitution - the (OH)44− cluster.............................................................41 Water molecules in structural cavities: beryl and cordierite.........................................43 OH substitution in topaz...............................................................................................44 OH incorporation in diopside.......................................................................................44 OH defects in perovskite..............................................................................................47 OH traces in corundum.................................................................................................48 ACKNOWLEDGMENTS........................................................................................................49 REFERENCES........................................................................................................................49

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Structural Studies of OH in NominallyAnhydrous Minerals Using NMR Simon C. Kohn

INTRODUCTION...................................................................................................................53 PRINCIPLES OF SOLID STATE NMR.................................................................................54 Positions of 1H MAS NMR resonances........................................................................55 Widths of 1H MAS NMR resonances...........................................................................55 Intensity of 1H MAS NMR resonances........................................................................56 APPLICATION OF 1H MAS NMR TO NOMINALLY ANHYDROUS MINERALS...........58 Attractive features of 1H MAS NMR for studies of NAMs.........................................58 Problems and difficulties in applying 1H MAS NMR to NAMs..................................58 1 H MAS NMR studies of orthopyroxene......................................................................59 1 H MAS NMR studies of clinopyroxene......................................................................60 1 H MAS NMR studies of olivine..................................................................................60 1 H MAS NMR studies of garnet...................................................................................61 1 H MAS NMR studies of SiO2 polymorphs.................................................................61 1 H MAS NMR studies of feldspars and other aluminosilicate framework minerals....61 1 H MAS NMR studies of wadsleyite............................................................................62 NON-SPINNING 1H NMR EXPERIMENTS.........................................................................62 STUDIES OF OTHER NUCLEI IN NAMS...........................................................................63 PROSPECTS FOR FUTURE DEVELOPMENT OF NMR FOR STUDIES OF NAMS.......65 ACKNOWLEDGMENTS........................................................................................................65 REFERENCES........................................................................................................................65 x

Water in Nominally Anhydrous Minerals ‒ Table of Contents

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Atomistic Models of OH Defects in Nominally Anhydrous Minerals Kate Wright

INTRODUCTION...................................................................................................................67 POINT DEFECTS IN MINERALS.........................................................................................67 THEORETICAL BACKGROUND.........................................................................................69 Quantum mechanical methods.....................................................................................69 Classical methods.........................................................................................................70 Treatment of defects.....................................................................................................71 OH DEFECTS IN MANTLE SILICATES..............................................................................72 The Mg2SiO4 polymorphs.............................................................................................73 Pyroxene.......................................................................................................................79 General remarks and future directions........................................................80 ACKNOWLEDGMENTS........................................................................................................81 REFERENCES........................................................................................................................81

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Hydrogen in High Pressure Silicate and Oxide Mineral Structures Joseph R. Smyth

INTRODUCTION...................................................................................................................85 Geochemistry of H........................................................................................................85 Crystal Chemistry of H..............................................................................................86 Nominally Hydrous High-Pressure Silicate Phases...................................87 Brucite..........................................................................................................................90 Serpentine.....................................................................................................................90 Talc...............................................................................................................................90 True micas....................................................................................................................91 Chlorite.........................................................................................................................92 Amphiboles..................................................................................................................92 Lawsonite.....................................................................................................................92 Epidote..........................................................................................................................93 Humite..........................................................................................................................94 Clinohumite..................................................................................................................94 Chondrite......................................................................................................................94 Phase A.........................................................................................................................95 Phase B.........................................................................................................................96 Superhydrous Phase B..................................................................................................96 Phase D.........................................................................................................................96 Phase E ........................................................................................................................97 Phase Pi .......................................................................................................................97 Topaz-OH ....................................................................................................................98 Phase Egg ....................................................................................................................98 K-cymrite......................................................................................................................98 xi

Water in Nominally Anhydrous Minerals ‒ Table of Contents Nominally Anhydrous High-Pressure Silicate and Oxide Phases ......99 Periclase-wüstite ..........................................................................................................99 Corundum ....................................................................................................................99 Coesite .......................................................................................................................101 Stishovite and rutile ...................................................................................................101 Pyroxenes ..................................................................................................................102 Akimotoite .................................................................................................................103 Garnet.........................................................................................................................104 Olivine .......................................................................................................................104 Wadsleyite .................................................................................................................105 Wadsleyite II ..............................................................................................................106 Ringwoodite ..............................................................................................................106 Anhydrous phase B ...................................................................................................107 Kyanite.......................................................................................................................107 Perovskite...................................................................................................................108 Post-perovskite ..........................................................................................................108 Zircon ........................................................................................................................109 Titanite .......................................................................................................................110 Conclusions....................................................................................................................110 ACKNOWLEDGMENT........................................................................................................110 REFERENCES......................................................................................................................110

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Water in Nominally Anhydrous Crustal Minerals: Speciation, Concentration, and Geologic Significance Elizabeth A. Johnson

INTRODUCTION.................................................................................................................117 Importance of nominally anhydrous minerals in the crust.........................................117 Scope and goals of this chapter..................................................................................117 HYDROUS SPECIES AND CONCENTRATIONS IN CRUSTAL MINERALS................118 Quartz and coesite......................................................................................................118 Feldspars and nepheline.............................................................................................119 Pyroxenes...................................................................................................................122 Garnets........................................................................................................................123 Al2SiO5 polymorphs ..................................................................................................124 Rutile and cassiterite..................................................................................................126 Zircon and titanite......................................................................................................128 Cordierite and beryl....................................................................................................129 UNDERSTANDING GEOLOGIC SYSTEMS.....................................................................130 Thermodynamic properties.........................................................................................130 Physical properties.....................................................................................................134 The water budget of the Earth....................................................................................136 SUMMARY AND FUTURE POSSIBILITIES.....................................................................136 ACKNOWLEDGMENTS......................................................................................................137 REFERENCES......................................................................................................................137 APPENDIX............................................................................................................................142 Studies of hydrogen in quartz. ...................................................................................143 xii

Water in Nominally Anhydrous Minerals ‒ Table of Contents Geological studies of H2O and CO2 in cordierite.......................................................143 Hydrous species in feldspars......................................................................................144 Structural hydroxyl concentrations in crustal and mantle pyroxenes.........................146 Structural hydroxyl concentrations in crustal garnets................................................148 Structural hydroxyl concentrations in kyanite............................................................152 Structural hydroxyl concentrations in sillimanite......................................................152 Structural hydroxyl concentration in andalusite.........................................................153 Structural hydroxyl concentrations in rutile...............................................................153 Structural hydroxyl concentrations in cassiterite.......................................................154

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Water in Natural Mantle Minerals I: Pyroxenes Henrik Skogby

INTRODUCTION.................................................................................................................155 OH ABSORPTION BANDS IN IR SPECTRA.....................................................................156 Diopside......................................................................................................................156 Augite.........................................................................................................................156 Omphacite..................................................................................................................157 Orthopyroxene............................................................................................................157 Absorption from inclusions .......................................................................................158 Correlations of OH and sample chemistry...................................................159 Water concentration in mantle pyroxenes.................................................162 Implications for water in the upper mantle................................................164 Acknowledgments......................................................................................................165 References......................................................................................................................166

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Water in Natural Mantle Minerals II: Olivine, Garnet and Accessory Minerals Anton Beran and Eugen Libowitzky INTRODUCTION.................................................................................................................169 OLIVINE...............................................................................................................................170 Basic structure and possible sites of hydrogen incorporation....................................170 Defect types in mantle-related olivines from different localities...............................171 Calibration approaches and summary of hydrogen contents......................................176 GARNET...............................................................................................................................176 Structural and spectral features..................................................................................176 Calibration and hydrogen content..............................................................................180 ACCESSORY MINERALS...................................................................................................182 Kyanite.......................................................................................................................182 Rutile..........................................................................................................................184 Coesite........................................................................................................................185 Spinel..........................................................................................................................186 Zircon.........................................................................................................................186 ACKNOWLEDGMENTS......................................................................................................188 REFERENCES......................................................................................................................188 xiii

Water in Nominally Anhydrous Minerals ‒ Table of Contents

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Thermodynamics of Water Solubility and Partitioning Hans Keppler and Nathalie Bolfan-Casanova

INTRODUCTION.................................................................................................................193 BASIC THERMODYNAMICS OF WATER SOLUBILITY AND PARTITIONING .........194 The meaning of the term “water solubility”...............................................................194 Thermodynamics of water solubility..........................................................................195 Relationship between water solubility and partitioning.............................................198 EXPERIMENTAL STRATEGIES FOR MEASURING WATER SOLUBILITY AND WATER PARTITION COEFFICIENTS..........................................................................199 Annealing experiments...............................................................................................199 Crystallization experiments........................................................................................200 WATER IN UPPER MANTLE MINERALS........................................................................201 Water solubility in and the Al content of orthopyroxenes as “geohygrometer”.........201 Water solubility in olivine..........................................................................................205 Water solubility in garnet...........................................................................................211 Water solubility in clinopyroxene...............................................................................211 Water partitioning among upper mantle minerals......................................................212 Water storage capacity of the upper mantle and the origin of the Earth´s asthenosphere.......................................................................................................214 Water recycling by subducted slabs............................................................................215 WATER IN TRANSITION ZONE MINERALS...................................................................216 Water solubility in wadsleyite and water partitioning between wadsleyite and olivine..........................................................................................216 Partitioning of water between wadsleyite and ringwoodite........................................219 Partition coefficients of water between other high-pressure phases...........................220 WATER IN MINERALS OF THE LOWER MANTLE........................................................222 Water in ferropericlase...............................................................................................222 Water in magnesium silicate perovskite.....................................................................223 The distribution of water at the 660 km discontinuity...............................................225 THE EQUILIBRIUM DISTRIBUTION OF WATER IN THE EARTH’s INTERIOR........226 ACKNOWLEDGMENTS......................................................................................................227 REFERENCES......................................................................................................................227

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The Partitioning of Water Between Nominally Anhydrous Minerals and Silicate Melts Simon C. Kohn and Kevin J. Grant

INTRODUCTION.................................................................................................................231 Partitioning of water between NAMs and melts; methodology and approach............................................................................233 Experimental studies of water partitioning between NAMs and melts......................233 Summary, Implications and future research..............................................238 Acknowledgments......................................................................................................239 REFERENCES......................................................................................................................239 xiv

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The Stability of Hydrous Mantle Phases Daniel J. Frost

INTRODUCTION.................................................................................................................243 Mantle Metasomatism..............................................................................................244 Evidence from mantle xenoliths......................................................................246 Peridotite massifs and xenoliths from alkaline basalts...............................................247 Xenoliths from kimberlites.........................................................................................247 Mantle amphibole mineralogy ...................................................................................248 Mantle mica mineralogy.............................................................................................250 Experimental studies on the stability of known mantle hydrous minerals...............................................................................251 Pargasitic amphiboles.................................................................................................251 Apatite........................................................................................................................255 Phlogopite...................................................................................................................256 K-richterite.................................................................................................................259 Experimental studies on the stability of potential high pressure hydrous mantle minerals................................................260 Phase X.......................................................................................................................261 Humite and dense hydrous magnesium silicate phases..............................................262 The stability of hydrous phases in ultramafic lithosphere and the convecting mantle................................................262 Acknowledgments......................................................................................................265 References......................................................................................................................266

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Hydrous Phases and Water Transport in the Subducting Slab Tatsuhiko Kawamoto

Introduction.................................................................................................................273 LOW-PRESSURE HYDROUS MINERALS AND HIGH-PRESSURE HYDROUS PHASES.......................................................................................................273 STABILITY OF HYDROUS PHASES IN DOWNGOING PERIDOTITE..........................277 Stability of hydrous phases in downgoing basalt and sediment............................................................................................279 PRESSURE - TEMPERATURE CONDITIONS AND DEHYDRATION REACTIONS IN THE SUBDUCTING SLAB................................................................280 COMPOSITION AND DIHEDRAL ANGLES OF AQUEOUS FLUIDS IN MANTLE PERIDOTITE.............................................................................281 SECOND CRITICAL ENDPOINT BETWEEN MAGMAS AND AQUEOUS FLUID: IMPLICATIONS FOR SLAB-DERIVED COMPONENT..282 CONCLUDING REMARKS.................................................................................................285 ACKNOWLEDGMENT........................................................................................................286 REFERENCES......................................................................................................................286 xv

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Diffusion of Hydrogen in Minerals Jannick Ingrin and Marc Blanchard

INTRODUCTION.................................................................................................................291 BASIC CONCEPTS OF DIFFUSION IN MINERALS........................................................291 EXPERIMENTAL METHODS.............................................................................................292 MEASUREMENT TECHNIQUES.......................................................................................294 Infrared spectroscopy ................................................................................................295 Mass spectrometry......................................................................................................296 Thermogravimetry......................................................................................................297 Nuclear reaction analysis............................................................................................297 Liquid scintillation counting......................................................................................298 Proton-proton scattering.............................................................................................298 Theoretical techniques................................................................................................298 DETECTION OF H DIFFUSION THROUGH ISOTOPE EXCHANGE.............................299 Anhydrous minerals...................................................................................................299 Hydrous minerals.......................................................................................................304 EXTRACTION/INCORPORATION REACTIONS IN ANHYDROUS MINERALS.........307 Olivine........................................................................................................................307 Diopside......................................................................................................................310 Enstatite......................................................................................................................313 Garnets........................................................................................................................314 Quartz.........................................................................................................................316 Feldspars.....................................................................................................................316 CONCLUSION AND FUTURE DIRECTIONS...................................................................317 ACKNOWLEDGMENTS......................................................................................................318 REFERENCES......................................................................................................................318

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Effect of Water on the Equation of State of Nominally Anhydrous Minerals Steven D. Jacobsen

INTRODUCTION.................................................................................................................321 ELASTIC PROPERTIES OF NOMINALLY ANHYDROUS MINERALS IN THE UPPER MANTLE..............................................................................................322 Olivine........................................................................................................................322 Humite-group minerals along the forsterite-brucite join............................................323 Garnet.........................................................................................................................323 Grossular-hydrogrossular...........................................................................................326 Pyroxene.....................................................................................................................326 ELASTIC PROPERTIES OF NOMINALLY ANHYDROUS MINERALS IN THE TRANSITION ZONE........................................................................................328 Wadsleyite..................................................................................................................328 Wadsleyite-II..............................................................................................................330 Ringwoodite...............................................................................................................330 DENSE HYDROUS MAGNESIUM SILICATES................................................................332 xvi

Water in Nominally Anhydrous Minerals ‒ Table of Contents Phase A.......................................................................................................................332 Phase-B group minerals..............................................................................................332 Phase D.......................................................................................................................332 Phase E.......................................................................................................................334 WATER-ELASTICITY SYSTEMATICS..............................................................................334 CALCULATED HYDROUS VELOCITIES IN THE UPPER MANTLE AND TRANSITION ZONE......................................................................................................335 CONCLUSIONS AND FUTURE RESEARCH OPPORTUNITIES....................................338 ACKNOWLEDGMENTS......................................................................................................338 REFERENCES......................................................................................................................338

15

Remote Sensing of Hydrogen in Earth’s Mantle Shun-ichiro Karato

INTRODUCTION.................................................................................................................343 GEOPHYSICAL OBSERVATIONS......................................................................................344 Electrical conductivity................................................................................................344 Seismic wave velocities..............................................................................................346 Seismic wave attenuation...........................................................................................346 Seismic anisotropy.....................................................................................................347 Topography of discontinuity.......................................................................................348 Sharpness of discontinuities.......................................................................................349 PHYSICAL BASIS FOR INFERRING HYDROGEN CONTENT FROM GEOPHYSICAL OBSERVATIONS...................................................................349 Electrical conductivity................................................................................................349 Seismic properties......................................................................................................351 Partial melting?...........................................................................................................362 SOME EXAMPLES..............................................................................................................363 Water content in the transition zone...........................................................................363 Distribution of hydrogen in the upper mantle............................................................366 Hydrogen in the lower mantle....................................................................................369 SUMMARY AND OUTLOOK.............................................................................................370 ACKNOWLEDGMENTS......................................................................................................370 REFERENCES......................................................................................................................371

16

The Role of Water in High-Temperature Rock Deformation David L. Kohlstedt

INTRODUCTION.................................................................................................................377 BACKGROUND....................................................................................................................379 MODELS OF CLIMB-CONTROLLED CREEP..................................................................379 THE CASE FOR OLIVINE...................................................................................................381 DISLOCATION CLIMB.......................................................................................................383 DIFFUSION...........................................................................................................................384 DEPENDENCE OF CREEP RATE ON WATER FUGACITY.............................................386 xvii

Water in Nominally Anhydrous Minerals ‒ Table of Contents CONCLUDING REMARKS.................................................................................................388 ACKNOWLEDGMENTS......................................................................................................390 REFERENCES......................................................................................................................390 APPENDIX............................................................................................................................394 Charge neutrality........................................................................................................394 Flux equations for a semi-conducting silicate............................................................395

17

The Effect of Water on Mantle Phase Transitions Eiji Ohtani and K. D. Litasov

INTRODUCTION.................................................................................................................397 RECENT PROGRESS ON PRESSURE SCALES FOR THE DETERMINATION OF PHASE BOUNDARIES IN MANTLE MINERALS...............398 EFFECT OF WATER ON PHASE TRANSFORMATION...................................................400 Dry and wet phase boundaries in the olivine-wadsleyite transformation ..................400 Wadsleyite-ringwoodite transformation ....................................................................401 Post-spinel transformation .........................................................................................401 Post-garnet transformation in basalt (MORB) ..........................................................404 EFFECT OF WATER ON PHASE TRANSFORMATION KINETICS................................406 Olivine-wadsleyite phase transformation kinetics .....................................................406 Post-spinel and post garnet phase transformation kinetics ........................................408 IMPLICATION FOR SEISMIC DISCONTINUITIES AND PHASE TRANSFORMATION BOUNDARIES UNDER DRY AND WET CONDITIONS......409 410 km seismic discontinuity and olivine-wadsleyite phase boundary......................409 The 660 km seismic discontinuity and the post-spinel transformation: average depth and topography of the 660 km seismic discontinuity....................410 The density relation of basalt and peridotite near the 660 km discontinuity ............412 Seismic reflectors: the possible existence of fluid in the lower mantle .....................413 CONCLUDING REMARKS.................................................................................................415 Acknowledgments......................................................................................................415 REFERENCES .....................................................................................................................415

18

Water in the Early Earth Bernard Marty and Reika Yokochi

INTRODUCTION.................................................................................................................421 ISOTOPIC CONSTRAINTS ON THE ORIGIN OF TERRESTRIAL WATER...................422 Hydrogen isotopic ratios............................................................................................422 Nitrogen and carbon isotopic ratios............................................................................424 Noble gas isotopic ratios............................................................................................424 Other isotopic constraints...........................................................................................427 POTENTIAL WATER CONTRIBUTORS............................................................................428 Contribution of water-rich planetary embryos...........................................................429 Asteroid contribution .................................................................................................429 Constraints on water delivery by asteroidal material from the terrestrial highly xviii

Water in Nominally Anhydrous Minerals ‒ Table of Contents siderophile element budget.........................................................................................430 The case of interplanetary dust as a source of terrestrial water..................................432 PROCESSES OF WATER INCORPORATION IN EARTH.................................................435 Solar nebula................................................................................................................435 Impact degassing .......................................................................................................437 Impact erosion............................................................................................................438 Post-accretional role of a proto-atmosphere in the Early Earth’s evolution ..............438 A summary of volatile delivery processes and of their inherent uncertainties ..........439 Cooling of the primordial Earth ................................................................................439 THE WATER CYCLE IN THE HADEAN............................................................................440 WATER CONTENT OF THE ARCHEAN MANTLE FROM THE COMPOSITION OF KOMATIITES...............................................................................443 CONCLUSIONS....................................................................................................................444 ACKNOWLEDGMENTS......................................................................................................444 REFERENCES......................................................................................................................445

19

Water and Geodynamics Klaus Regenauer-Lieb

INTRODUCTION.................................................................................................................451 WATER IN THE LITHOSPHERE........................................................................................452 Water and the rigidity of (oceanic) plates...................................................................452 Water and the nucleation of (new) plate boundaries..................................................454 Water and the evolution of plate boundaries..............................................................455 Water and the stored energy potential Ψ....................................................................455 Water and the dissipated energy potential Φ..............................................................457 Solid versus fluid dynamic modeling setups..............................................................461 Application to subduction initiation...........................................................................462 WATER IN THE CONVECTING MANTLE........................................................................465 DISCUSSION AND CONCLUSIONS .................................................................................465 REFERENCES......................................................................................................................467 APPENDIX: THERMOMECHANICAL APPROACH........................................................471

Additional Volume Content: Color Plate 1..................................................................................................................475 Color Plate 2..................................................................................................................476 Color Plate 3..................................................................................................................477 Color Plate 4..................................................................................................................478

xix

1

Reviews in Mineralogy & Geochemistry Vol. 62, pp. 1-28, 2006 Copyright © Mineralogical Society of America

Analytical Methods for Measuring Water in Nominally Anhydrous Minerals George R. Rossman Division of Geological and Planetary Sciences California Institute of Technology Pasadena, California, 91125-2500, U.S.A. e-mail: [email protected]

INTRODUCTION Decades of work have shown that trace- to minor-amounts of hydrous components commonly occur in minerals whose chemical formula would be normally written without any hydrogen, namely, the nominally anhydrous minerals (NAMs). When the concentrations of the hydrous components are several tenths of a percent by weight or higher, a variety of analytical methods such as weight loss on heating, X-ray cell parameters, X-ray structure refinement, Karl-Fischer titrations, or even careful electron microprobe analyses can be used to establish their concentrations (e.g., Aines and Rossman 1991). However, for most NAMs, accurate determinations with these common analytical methods prove difficult if not impossible. For this reason, infrared (IR) spectroscopy has become, and remains, the most widely used method to detect and analyze hydrous components (OH or H2O) in minerals and glasses because it is both highly sensitive and can be done rapidly with a commonly available, modestly priced instrument and at dimensions of just a few tens of micrometers. A change in the electric dipole occurs when the OH bond in either water and hydroxyl ions vibrate. This motion has a resonance coupling with electromagnetic radiation generally in the 3500 cm−1 region of the infrared spectrum. In addition, bending motions of the water molecule, and overtones and combination of these motions produce absorption in the infrared. Under favorable conditions, namely a sharp band in a single orientation, just a few nanometers equivalent thickness of a hydroxyl species such as an amphibole can be detected in an otherwise anhydrous mineral such a pyroxene (Skogby et al. 1990). Routinely, detection limits of a few to tens of ppm wt of H2O in a mineral can be detected and often quantitatively determined. The overtone and combination modes of OH and H2O behave in predictable fashion in minerals (Rossman 1975) so that the two species can usually be separated from each other. Infrared spectra, however easily obtained, are not rigorously self-calibrating, so independent methods of analysis have been necessary to calibrate the spectroscopic work. A couple general correlations of IR band intensity with the absorption energy have proven useful, if approximate. Various absolute hydrogen extraction methods have proven highly useful for purpose of rigorous calibration. More recently, nuclear methods that rely upon specific resonant reactions with the hydrogen nucleus or nuclear scattering specific to hydrogen have gained importance and have provided critical absolute calibrations of the infrared spectra. Secondary Ion Mass Spectroscopy (SIMS) for hydrogen is still in the early stages of development but once calibrated, and with established protocols, should play an ever-expanding role in the future. NanoSIMS promises to bring hydrogen analyses to ever finer spatial dimensions but will require significant effort before it can be regarded as an accurate analytical technique for small concentrations of hydrogen. The purpose of this chapter is to review the various methods that have been used to analyze hydrous components in the NAMs. 1529-6466/06/0062-0001$05.00

DOI: 10.2138/rmg.2006.62.1

2

Rossman ANALYTICAL METHODS

Early infrared studies Much of the early interest in OH in minerals came from the study of synthetic minerals used in the electronics industry. Quartz, in particular, was an important phase used for frequency control in telecommunications and radio circuits. Consequently, much effort was directed towards the understanding of factors that influenced the efficiency and cost of these devices. Water in quartz was one of the most important factors. The OH bond is dipolar with a partial negative charge on the oxide ion and a partial positive charge on the hydrogen ion. Thus, the vibrations of the OH bond coupled well to infrared radiation and infrared spectroscopy quickly became the tool of choice to study OH in both natural and synthetic minerals. An important early study was conducted by Kats and Haven (1960) who used deuteration to demonstrate which bands in the complex quartz spectrum in the 3000 to 4000 cm−1 region originated from 1H as opposed to overtone or combination bands of the quartz vibrational spectrum that appeared in the same region. Once the OH vibrations were positively identified, Kats (1962) performed a comprehensive study of OH in quartz and identified which of the sharp band absorptions in the 3000-3600 cm−1 region are due to O-H stretching vibrations. Kats further showed that most of the absorptions are primarily due to the presence of Al3+ substitution for Si4+ with charge compensating cations (such as H+, Li+, Na+) in defects in the crystal. Other studies were taking place at Bell Labs in the United States where elastic properties and dielectric loss in synthetic quartz was related to H defects (King et al. 1960; Dodd and Fraser 1965, 1967). In these studies, the relationship between infrared absorption, and hydroxyl and water defects in quartz was also being established. During these times, Brunner et al. (1961) concluded that H enters defects in clear, natural quartz in the form of OH ions and estimated the amount of H as 1018 per cc (corresponding to about 15 ppm H2O wt). These early estimates showed that small amounts of hydrous components could have a large impact on the physical properties of the host phase. As work on synthetic quartz progressed, studies of quartz also used natural samples and ultimately, the results were reported in the mineralogical literature through the work of Dodd and Fraser (1965). Simultaneously, interest in the low concentrations of water in ring silicate minerals was generated by infrared (Schreyer and Yoder 1964; Wood and Nassau 1967; Farrell and Newnham 1967) and NMR (proton nuclear magnetic resonance) (Pare and Ducros 1964; Sugitani et al. 1966) studies of beryl that demonstrated that water molecules occur in the c-axis channels. The NMR work showed that the water molecules were in motion and the IR studies showed that the water molecule existed in two independent crystallographic orientations in the crystal. In Austria, in the late 1960’s and early 1970’s, Beran and Zemann obtained the IR spectra of a number of minerals such as titanite, kyanite, axinite, titanium oxides, cassiterite (Beran 1970a,b,c,d; Beran and Zemann 1969a,b, 1971) and demonstrated that they had structurally bound, crystallographically oriented OH groups. These studied demonstrated that polarized infrared radiation could establish the orientation of the OH groups in minerals and demonstrated that trace amounts of hydroxyl occur broadly in a number of nominally anhydrous minerals. A couple of significant motivations to develop quantitative understanding of the H-content of nominally anhydrous minerals appeared in the early 1970’s. Martin and Donnay (1972) suggested that hydrogen may be stored as OH groups in minerals in the deep earth, and Wilkins and Sabine (1973) initiated a broad effort to determine the amount of hydrous components in a variety of minerals by combining infrared absorption with independent water analysis (P2O5 electrolytic coulometry). Although we now recognize that many of the analyses of Wilkins and Sabine included alteration products and water in micro-inclusions, they did set the quantitative stage for further detailed studies.

Analytical Methods

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Another major impetus to the study of water in the nominally anhydrous minerals came from the studies of the rheological properties of, first, quartz (Griggs and Blacic 1965; Kirby and McCormick 1979), and then olivine (Mackwell et al. 1985). To study how water weakens minerals, it was necessary to know both the chemical species of the hydrous components that enter nominally anhydrous minerals, and to know their absolute concentrations.

Quantitative IR methods The determination of the concentration of OH or H2O in an “anhydrous” mineral depends upon accurate measurement of the infrared spectrum and ultimately on an independent calibration. Infrared spectra are intrinsically not self-calibrating. A number attempts have been made to develop generic calibrations. These often may be good as an initial estimate of the water concentration, but, for many systems, have been shown to be inadequate for precise work. Thus, mineral-specific calibrations have been developed. Once such calibrations are established and properly published, they can be used by other labs worldwide, even if an inhouse standard is not available. The well-established Beer-Lambert law is used to determine the concentration of hydrous species in a mineral from the infrared spectra: Absorbance = ε × c × t (1)



This relates Absorbance (A), the band height in the region of interest (corrected for baseline), c, the concentration of hydrous species expressed in moles of H2O per liter of mineral, and t, the thickness of the path (in cm) through which the measurement is made where ε is a mineral-specific calibration factor. In the classical chemical applications, the sample is in solution, so only one measurement is made. In the case of anisotropic solids, it is necessary to make the measurement in multiple directions (Libowitzky and Rossman 1996). Typically, linearly polarized light would be used and measurements would be made along the three principal extinction directions, X, Y, and Z. In this case, the intensities would be summed so A becomes AX + AY + AZ (where AX is the absorbance obtained with light polarized in the X direction, etc.). This approach tends to work best with phases that have one or a small number of narrow bands in the OH region. It also requires knowledge of the density of the mineral to convert from moles per liter to weight percent (or ppm) water. For most minerals, it is usually more useful to use a modified version of the Beer-Lambert law that uses integrated band areas rather than band heights. Band heights can vary depending on both the quality of the polarizer in the instrument and on the spectroscopic resolution of the instrument whereas band areas are less dependent on these parameters. The band height measured by the Absorbance is replaced by the total integrated area of bands in the region of interest Absorbancetotal (also written as Abstotal or Atotal). The concentration, c, remains expressed as moles of H2O per liter of mineral. In this case, the absorption coefficient, ε, is replaced by the integral molar absorption coefficient, I, in units of L/(mol·cm2). When c is expressed as ppm H2O by weight, the absorption coefficient becomes the integral specific absorption coefficient (I’, ppm−1·cm−2). The absorption coefficient for each species of hydrogen is found by determining the concentration, c, by an independent, absolute method and measuring Abstotal from polarized IR spectra in the three principal optical directions (X, Y, and Z) for the mineral of interest. For an orthorhombic mineral such as olivine: Abstotal =

1 ta

ν2

1

v2

1

v2

∫ Absa dν + tb ∫ Absb dv + tc ∫ Absc dv

v1

v1

(2)

v1

Here, the equation specifies measuring the integrated area of an orthorhombic crystal with light polarized in the E||a, E||b, and E||c directions between the appropriate wavenumber limits of the OH bands, ν1 and ν2. For lower symmetry crystals (monoclinic, and triclinic)

4

Rossman

Abstotal = ∫AbsX + ∫AbsY + ∫AbsZ, and for a uniaxial crystal (hexagonal or tetragonal) Abstotal = 2∫Abs⊥c + ∫Absc (e.g., Libowitzky and Rossman 1996). To be comparable to measurements on lower symmetry crystals, an isotropic crystal would need to have Atotal = 3∫Absa. Paterson’s method. If the absorption frequency and intensity of a unit concentration of OH were a constant, then a single calibration of the OH spectrum would be all that is needed to conduct quantitative analysis with IR spectroscopy. Unfortunately, that is not the case. First of all, while the fundamental stretching vibration of a free (gaseous) hydroxide ion occurs at 3555.59 cm−1 (Lutz 1995), the OH stretching frequency in a mineral commonly can occur over a range of several hundred wavenumbers and can vary by nearly 2000 cm−1. A variety of studies (Nakamoto et al. 1955; Bellamy and Owen 1969; Novak 1974) showed that for a variety of chemical elements, the stretching frequency of an X-H bond in an X-H···Y hydrogen bonded system is a function of the X-Y distance. This includes O-H bonds. These authors derived empirical fits to experimental data that mathematically expressed this relationship. The second observation of interest is that the infrared absorption intensity of a unit concentration of OH in a solid is obviously not constant. Paterson (1982) confirmed that the strength of the OH absorption in the 3600 to 3000 cm−1 region was frequency dependent. From the calibrations available for various substances, he presented a single empirical calibration line that related the OH intensity to band position that could be applied as a first approximation for determining the amount of OH in a variety of substances such as silicate glasses, quartz, and various forms of water. This was the first generic calibration specifically designed for the study of hydrous components in minerals and glasses. Paterson demonstrated that the intensity of an OH band (normalized to a unit concentration of H2O) increases when the band occurs at lower wavenumbers (stronger hydrogen bonding). This trend has been used by a number of authors to estimate the OH content of various minerals. Subsequent work has shown that determinations based on Paterson’s trends are a reasonable first estimate, but that accurate determinations do require mineral-specific calibrations. Paterson’s method first assumes that if a crystal is being measured, it is in a known crystallographic orientation. To determine the concentration of hydroxyl groups in the sample, the integrated absorbance is determined by integration of the infrared spectra over the region dominated by the stretching vibrations due to O-H bonds, typically from approximately 3750 to 3000 cm–1. The integral molar absorption coefficient (I) is scaled to reflect the higher intrinsic intensities of bands at lower wavenumbers (stronger H-bonds) through the equation:

I = γ ×150 × (3780 − ν) (3)

where ν is the wavenumber and gamma (γ) is a factor to take account of the anisotropy of the crystal based on an assumption that O-H bonds are oriented in a single direction. The OH concentration is then calculated from a Beer-Lambert law relationship: ConcentrationOH =

1 A(ν ) dν (150 × γ ) ∫ (3780 − ν )

( 4)

assuming that the data are scaled for 1 cm sample thickness. Although uncertainties in this calibration were thought by Paterson to be about 30%, it has been widely adopted, partly in the hope that it would eliminate the need for more involved polarized light observations with multiple crystallographic directions. However, the studies of Libowitzky and Rossman (1997) and Bell et al. (2003) show that it can result in non-systematic underestimates of hydrogen concentrations. Examples of mineral specific calibrations that fall far from the trend are documented, particularly those that involve nominally anhydrous minerals with low concentrations of OH. As examples, the pyrope analyzed by Bell et al. (1995) departs from the Paterson trend by nearly a factor of three, the nuclear reaction analysis

Analytical Methods

5

of olivine by Bell et al. (2003) departs by more than a factor of two (Fig. 1) and the SIMS analysis of both olivine and orthopyroxene (Koga et al. 2003) show that the Paterson trend also underestimates their OH concentrations. Libowitzky and Rossman’s revision. Libowitzky and Rossman (1997) presented an updated version of the correlation of Paterson (1982). They measured polarized IR absorption data from single crystal minerals that contained stoichiometric water contents in the form of either OH or H2O. These data were used to construct a calibration curve for the intensity of the infrared absorption as a function of the band energy. Specifically, integrated molar absorption coefficient, εi (in units of cm−2 per moleH2O/liter), was evaluated as function of the mean wavenumber of the OH stretching band (in units of cm−1). The result in Figure 2 shows that an increase in the hydrogen bonding leads to a decrease in the energy of the OH stretching energy which, in turn, is associated with an increase in the intensity of absorption. The form of the correlation is εi = 246.6 × (3752 − ν) (5)



where ν is the mean wavenumber of the OH stretching band. The results in Figure 2 show that the revised calibration produces εi values about threequarters of those of Paterson (1982). Measurements of minerals with stoichiometric OH are difficult to obtain. Their OH intensities are so high that crystals must be prepared very thin (perhaps as thin as 2 μm). Such preparations are difficult to near impossible; and when successful, the determination of their thickness to a high degree of accuracy is difficult.

Figure 1. Comparison of the results of the calibration developed by Bell et al. (2003) using Nuclear Reaction Analysis and the OH analysis method of Paterson (1982), as applied to polarized (solid circle) olivine spectra. Modified after Fig. 6 of Bell et al. (2003).

Figure 2. The correlation of the integrated molar absorption coefficient of OH stretching vs. wavenumber. Circles are experimental data points for stoichiometric minerals. The correlation of Paterson (1982) is shown for comparison. This means that if all things are equal, the Paterson trend underestimates the OH content. From Libowitzky and Rossman (1997).

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In a related effort, Libowitzky (1999) evaluated correlations specific to minerals between the frequency of the O-H stretching vibration and the length of the oxygen-oxygen distance and the H···O distances in the O-H···O hydrogen bond. Effectively, the shorter these distances are, the lower becomes the energy of the O-H stretch. Because the intensity of the OH band is related to the energy of the vibration (Libowitzky and Rossman 1997), such correlations provide some degree of a predictive estimate about the intensity of an OH absorption that arises from a particular site in a crystal. Use of unoriented grains. Asimow et al. (2006) present a method that allows multiple, randomly oriented grains of a mineral to be used to determine the total absorbance. In their method, the spectra of oriented sample of the phase of interest must already exist. Then, the spectra of three different randomly oriented crystals are measured, and the orientations of the grains are determined via methods such as electron backscatter diffraction (EBSD) or from the silicate overtone bands in the infrared spectra. They demonstrated that such methods result in angular errors of typically only 6 degrees and provide a surprising good determination of the OH content of the phase. Polarizer considerations. A linear polarizer must be used in the infrared beam of conventional spectrometers to obtain the total absorbance of anisotropic crystals. Commonly, the polarizers are made of a fine, parallel wire grid deposited on an infrared-transparent substrate such as CaF2 or KRS5 (a thallium bromide iodide). These polarizers have wide acceptance angles and are readily available, but have only moderate polarization ratios. Crystal polarizers of a design similar to calcite polarizers used in the visible wavelength region are also available, but often have a narrow range of wavelengths over which they function. Lithium iodate covers a wide wavelength range and has a very high polarization ratio, but is hydroscopic and no longer readily available. Libowitzky and Rossman (1996) discussed the principles of quantitative absorbance measurements of anisotropic crystals and paid particular attention to the influence of the quality of the polarizers upon the results. First, they showed that the use of unpolarized radiation with an anisotropic crystal could not produce quantitatively accurate results. The Beer-Lambert law demands that the height of an absorption band will scale with the thickness of the sample. Figure 3 demonstrates how the spectrum taken with linearly polarized radiation follows the law. It also shows that unpolarized spectra do not scale according to the law. This means that unpolarized spectra should not be used to calibrate the infrared spectrum of OH

Figure 3. Comparison of the intensity of a carbonate overtone band in the calcite spectrum taken with well-polarized and unpolarized radiation. The experiment that used different thickness of calcite to test the Beer Lambert law shows that unpolarized spectra are not appropriate to quantitatively measure anisotropic crystals. From Libowitzky and Rossman (1997).

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7

in an anisotropic standard, and cannot be used to accurately determine the concentration of OH in an anisotropic unknown. The more highly anisotropic the sample is, the more problematic this issue will become. Libowitzky and Rossman also showed that the intensity of an absorption band of an anisotropic crystal is highly dependent upon the polarization ratio of the polarizers (Fig. 4) which means that if band heights are used to calibrate the infrared spectra, results can vary significantly from lab to lab if the appropriate in-lab standards are not available. Baselines issues. Figure 5 shows that strongly rising, non-linear baselines may be an intrinsic part of the spectrum in the OH region. These baselines commonly arise from Fe2+ and may arise from silicate overtones in thick samples. A major, subjective source of uncertainty in IR measurements of OH in minerals remains the choice of the baseline.

Figure 4. The intensity of absorbance depends on the quality of the polarizer used for the measurement. Here, the spectrum (E ⊥ c) of three bands in the calcite spectrum was obtained with a high efficiency polarizer (LiIO3), a lower efficiency wire-grid polarizer (gold wire on AgBr), and without polarization. Modified after Figure 6 of Libowitzky and Rossman (1996).

Figure 5. Infrared spectra of a clinopyroxene that show the baseline remaining after the crystal is fully dehydrated. Contributions from ferrous iron cause the rising baseline towards the long wavenumber side. From Bell et al. 1995, Figure 1.

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Comments on terminology. The terminology for spectroscopic units has not been consistent in the literature. Chemical terminology, the source of these terms, has evolved, and geoscience has had to modify some of the standard terms for anisotropic materials. Table 1 presents a compendium of terminology taken from the web site of the International Union of Pure and Applied Chemistry. In addition, the currently preferred terminology is compared to other terminology found in the literature.

Mineral specific calibrations While the generic calibrations developed by Paterson (1982) and later refined by Libowitzky and Rossman (1997) are useful first approximations, they are not necessarily accurate. There is no principle of science that demands that the infrared absorption intensity of all OH bonds be the same, or that the intensity of all OH bonds vary smoothly with the O-H···O hydrogen bond distance. Unpublished work by this author has shown that the intensity of other bonds such as C-O (carbonyl) and C-N (cyano) can vary by orders of magnitude. Thus, there is the need for mineral-specific calibrations. A variety of experimental methods, discussed in the following sections, have been used over the years to independently determine the amount of hydrous components in minerals. As is often the case in the history of development of analytical methods for trace components, early attempts suffered from large (and often Table 1. Selected terminology used in quantitative spectroscopy of minerals. Absorbance = log(I0/I) directly measured by the instrument Attenuation coefficient Analogous to the absorption coefficient, but differs from it because it accounts for the diffusion of radiation that includes absorption as well as scattering and luminescence. Formerly, it was called the extinction coefficient, a term that is now discouraged. linear absorption coefficient = Absorbance divided by the optical path length molar absorption coefficient = ε = molar absorptivity in earlier literature = linear absorption coefficient divided by the amount concentration amount concentration = molarity in prior literature commonly expressed in units of moles per liter integral molar absorption coefficient I (in units of cm-2 per molH/liter). εi (in units of cm-2 per molH2O/liter). (note that the mols of H = mols OH) Integral absorbance (not defined by IUPAC) Integrated absorbance Abstot =

1 t1

ν2

1

ν2

1

ν2

∫ Abs1dν + t2 ∫ Abs2 dν + t3 ∫ Abs3dν

ν1

Absorbancetotal Δ Integrated-Abstot

ν1

ν1

Analytical Methods

9

unrecognized) backgrounds, and the inability to separate the contributions of the hydrogen in the host phase from hydrogen contained in inclusions, cracks, and alteration products.

Thermogravimetric methods Thermogravimetric analysis (TGA) is a commonly used analytical method to determine the amount of mass lost from a sample during heating. It involves simultaneously heating and weighing a sample to produce a weight-loss vs. temperature curve. It is frequently used to determine water of hydration in minerals with more than trace quantities of water. The method has also been applied to water loss from nominally anhydrous minerals but with limited success. Early attempts to determine the H-content of garnets used the TGA method (Aines and Rossman 1984a) and coupled the results of this method with infrared spectra of the same samples. We now recognize that many of the earlier thermogravimetic methods over-estimated the water content of the NAMs due to the inclusion of contaminating water that remained trapped on the surface of the ground samples, even after the sample was “dried” by heating to over 125 °C prior to analysis. TGA was used to determine the water content of nepheline from Bancroft, Ontario (0.36 wt% H2O), and from Mt. Somma, Italy (0.17 wt% H2O) (Beran and Rossman 1989). Because these minerals have comparatively large water contents, the error introduced by the TGA method is small compared to what it may be when minerals with a few hundred ppm or less are analyzed by this method. The results of this method were also used to calibrate the infrared spectra of nepheline. While TGA analyses are conventionally conducted on ground samples, step heating experiments on slabs of single crystals used for infrared experiments demonstrate how difficult it can be to fully dehydrate a sample. Controlled heating experiments that were accompanied with infrared spectra of OH bands indicated that temperatures of about 1400 °C are needed to fully dehydrate slabs of some silicate minerals (sillimanite: Beran et al. 1989). Similar experiments with slabs of single crystal zircon indicated that OH is tightly held. Some OH persists in zircons even after the crystals are heated at 1500 °C (Woodhead et al. 1991). Ilchenko and Korzhinskaya (1993) also conducted step-heating experiments on kimberlitic zircon crystals and found that OH ions were only partially removed after heating to 1300 °C.

P2O5 cell coulometry P2O5 cell coulometry is based on the principle that water released during the thermal decomposition of a sample can react with P2O5, a non-conductor, and turns it into H3PO4, an electrical conductor. The amount of H3PO4 formed can be determined by the amount of electric current (coulombs) necessary to reverse the hydration reaction. One of the more popular commercial models used in mineral analysis was the DuPont moisture evolution analyzer (MEA). It consisted of a thermal decomposition chamber that led to a column containing a pair of closely spaced, P2O5-coated, Pt wire electrodes wound in a helical fashion. A dry nitrogen flow would carry the released water vapor into the electrodes where electrical current would flow between the wires whenever the P2O5 reacted with the water. A known mass of a stoichiometrically hydrated material was used to calibrate the system. The moisture evolution analyzer found use in some of the earlier analyses such as Wilkins and Sabine (1973) study, and the Aines and Rossman (1984) calibration of garnets. In practice, these systems had to be used regularly to prevent the P2O5 columns from going bad, and proved difficult for many users to regenerate once the columns did degrade. Because blanks with this method are typically several tens of micrograms of H2O, samples of at least a few hundred milligrams are required for the analysis of the nominally anhydrous minerals. (Aines and Rossman 1984).

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Rossman

Hydrogen extraction with uranium reduction methods Hydrogen manometry. Hydrogen manometry has long been a standard and generally reliable method to determine the water content of samples. In this method, several hundred milligrams to gram quantities of samples are weighed into a metal (Mo, or Pt) crucible, and first degassed under vacuum and low heat to drive off the adsorbed moisture. The sample crucibles are then heated with an induction furnace to liberate the bound water while under vacuum. The volatiles (H2 and H2O) are converted to just water and trapped and separated from the condensable and non-condensable gases by distillation in cryogenic traps. The water vapor is next passed over a hot furnace containing uranium metal (Bigeleisen et al. 1952) to reduce the water to molecular hydrogen. Alternatively, zinc has been used to reduce water (Michel and Villemant 2003). The hydrogen is then moved by a mercury-piston Toepler pump into a calibrated chamber in which the volume of hydrogen can be measured at a known pressure. From the PV = nRT relationship, the absolute amount of hydrogen can be determined. The system can be calibrated by known amounts of water, or by dehydration of minerals or compounds with known, stoichiometric water contents. For minerals with very low hydrogen contents such as the nominally anhydrous minerals, significant blank corrections must be applied that correct for degassing from the crucibles (Bell et al. 1995). Errors have been reported to be much less than 1% with this method (Dyar et al. 1996). Additional details of the technique can be found in Holdaway et al. (1986) This method has been used to determine the water content of minerals that are used to calibrate infrared spectra. The advantage of using this approach is that once the sample is destroyed by the hydrogen extraction procedure, its value as a calibrant remains through the calibration of the infrared spectrum which can be used to analyze additional samples of the calibrated phase that have similar spectra. Furthermore, the infrared spectrum allows reevaluation of the calibration because the original spectrum can be compared to the spectrum of other samples re-calibrated by improved methods years later. Early calibration efforts with hydrogen extraction (Aines and Rossman 1984) include a grossular with 0.18 wt% H2O, and a pyrope with 0.08 wt% [that is probably overestimated based on the more recent calibration of Bell et al. (1995) that indicate about 37 ppm H2O]; and perthite feldspar from two pegmatites (Hofmeister and Rossman 1985a,b) that had water in the 0.09 to 0.15 wt% range. More recent calibrations with lower blank contributions (Fig. 6) consist of a pyrope with 56 ppm, an enstatite with 217 ppm, and an augite with 268 ppm (Bell et al. 1995). In these experiments, large quantities of sample had to be carefully prepared, and checked to eliminate inclusions, cracks and other imperfections. The clean material was

Figure 6. Quality of hydrogen extraction determination of water in garnet and two pyroxenes using aliquots of different mass to determine the water content. Figure 4 from Bell et al. (1995).

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then crushed to less than 2 mm particles and the fraction less than 100 μm was discarded to minimize the effects of adsorbed water. Continuous flow mass spectrometry. A more recent variation of the hydrogen extraction technique uses continuous flow mass spectrometry to measure the absolute amount of hydrogen released from minerals by heating (O’Leary et al. 2006). This method is a modification of the method of Eiler and Kitchen (2001) used to determine D/H isotopic ratios of picoliter quantities of hydrogen. It requires about 1/1000 the amount of hydrogen required by conventional hydrogen manometry. Samples in the range of 50 μg to 20 mg of coarsely ground minerals are heated to release hydrous components, which are collected and converted to hydrogen by reaction with uranium (as opposed to carbon in the Eiler and Kitchen paper). The hydrogen is then detected in a mass spectrometer. The system is calibrated with a few hundred micrograms of zoisite grains of known H content. This system has been used to independently calibrate a series of garnets and pyroxenes that have been previously calibrated by conventional hydrogen extraction manometry or by nuclear methods. The linearity and agreement with previous calibrations has been excellent with samples at the few hundred-ppm H2O level and higher (Fig. 7).

Figure 7. Comparison of the water contents determined by the new micro-extraction method compared to conventional methods (O’Leary et al. 2006).

Nuclear methods for hydrogen determination A variety of nuclear reactions can be used to analyze hydrogen in solids (Lanford 1992). Some make use of nuclear reactions and others make use of nuclear scattering. Beams of ions accelerated to high energy can undergo a resonant nuclear reaction with the hydrogen ions in the target sample. Such methods are known either as Nuclear Resonant Reaction Analysis (NRRA), Nuclear Reaction Analysis (NRA) or Nuclear Profile Analysis (NPA) (when the hydrogen concentration is determined as a function of depth in the sample). The 6.42 MeV resonance of 19F with hydrogen and the 6.385 MeV resonance of 15N with hydrogen are the two that are typically used. Additional resonances of 19F at 16.44 MeV and 15N at 13.35 MeV can also be used (Xiong et al. 1987). In each of these reactions, the analysis depends upon the detection of gamma rays emitted from a heavier element that formed from transmutation of the ion beam from its reaction with hydrogen. 19

F Nuclear reaction analysis. Initially, 19F was the ion of choice for analysis of hydrogen in solids. The reaction involves the interaction of 19F with 1H to yield an 16O atom plus an alpha particle and a gamma ray. In the geological sciences, the 16.4 MeV resonance has found use for measuring hydration profiles in glass such as obsidian (Lee et al. 1974) and measurements of the H concentration in synthetic and natural quartz (Clark et al. 1978). Early work on analysis of H in garnets (Rossman 1990) also used 19F, but found that the reproducibility needed improve-

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ment. Because some accelerators can bring the 19F ion to as much as 22 MeV, significant depth profiles are possible. 15 N Nuclear reaction analysis. The most sensitive analyses of hydrogen in minerals have been made by a nuclear resonant reaction using the 15N technique (Lanford, 1978) that is based on the nuclear reaction 1H(15N,αγ)12C. In this method (Fig. 8), the hydrogen ions in the sample (the target) interact with a beam of 15N ions and are transmuted into 16O that immediately decays through alpha decay into 12C in a nuclear excited state. The 12 C has a decay path that emits a gamma ray that is detected in the analysis. The number of 12C gamma rays is proportional to the amount of hydrogen in the sample and does not depend on the chemical species of the hydrous component. A single calibration point is all that is needed to use the method for quantitative analysis of hydrogen.

Figure 8. The nuclear reaction scheme in the 15 N nuclear reaction method. The reaction of 15 N and 1H produce 16O in a nuclear excited state. A decay path of oxygen produces 12C, which comes to the ground nuclear state with the emission of a 4.44 MeV gamma ray that is the analytical signal.

The methods for mineral analysis were initially refined at Caltech and later, when the Caltech accelerator shut down, were transferred to the accelerator laboratory of the Institut für Kernphysik, Frankfurt am Main, where a beam of 15N2+ ions was delivered by a 7-MeV Van de Graaff accelerator onto a sample under high vacuum. At Frankfurt, the apparatus was specially designed and modified for the analysis of low hydrogen concentrations (to 10 ppm wt) in mineral samples. A detailed description of the experimental design can be found in the works of Endisch et al. (1993, 1994). Salient aspects include a Pb-shielded bismuth germanate (BGO) scintillation detector with an anticoincidence counting system for reduction of cosmic ray background, with the sample holder placed in an ultra-high-vacuum (10−10 mbar) chamber. The NPA method for low concentrations of H in minerals has been under development since the late 1970’s. Initially, F-19 was the ion beam of choice, but with the discovery of weak, interfering reactions, the ion beam was changed to N-15. Initially, weak nuclear reactions from carbon contamination were problematic, but improved detection methods, improved instrument vacuum and trapping of carbon compounds in the sample chamber brought them down to a manageable level (Kuhn et al. 1990). Ultimately, the layer of hydrous materials on the surface of the sample became the limiting problem, but high voltage ion sputtering was able to reduce this limitation to low levels (Maldener and Rauch 1997). An additional modification described by Maldener and Rauch allowed accurate sample positioning by Rutherford backscattering. Despite the extensive measures employed to minimize background hydrogen, a finite background or blank level may contribute to the amount of hydrogen measured. Due to the evolving methods of background reduction, the absolute background contribution to each analysis was subject to some degree of variation. One of the key calibrations for olivine was establish using this method (Fig. 9). In the most recent set of procedures, analysis of anhydrous silica glass and a silicon wafer placed the background estimate at 2 ± 2 ppm H2O. In late 2004, the accelerator at Frankfurt was decommissioned and work there on hydrogen in minerals has ceased. During the lifetime of the Frankfurt facility, the nuclear profile analysis method has been applied a variety of minerals including garnets (Rossman et al. 1988; Maldener et al. 2003), olivines (Bell et al. 2003), kyanite (Bell et al. 2004), rutile and cassiterite (Maldener et al. 2001), titanite (Hammer et al. 1996), ortho- and clinopyroxenes and zircon (Rossman et al. in prep.).

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Figure 9. The olivine calibration established with 15N nuclear reaction analysis. This calibration relates the total integrated absorption of the infrared spectrum to the water content determined by NRA. From Bell et al. 2003.

Other workers have used the NPA method for analysis of H in minerals and geological materials. Rauch et al. (1992) used the 15N method to determine the hydration of tektite glass. Semi-quantitative hydrogen concentration depth profiles were obtained on forsterite crystals by Fujimoto et al. (1993). They treated crystals under water at different pH and temperature conditions and found that high surface hydrogen concentrations developed. Under medium to high pH conditions at 25 °C, they found that the hydrogen-rich region extended less than 20 nm into the surface while at low pH conditions; it reached as deep as 200 nm. Elastic recoil detection analysis (ERDA). Methods based on the scattering of nuclei by protons are also used in the analysis of minerals. A particularly promising method is known as Elastic Recoil Detection analysis (Barbour et al. 1995; Sie et al. 1995). This method (Fig. 10) involves using 2 MeV 4He+ ion beam that is focused on the polished surface of the sample at a low angle (15°). Forward scattered 1H+ ions that come from the hydrous component in the mineral (the recoil spectrum) are detected by a silicon ERDA detector. Because the forward scattered protons loose energy as they traverse through the thickness of the sample, their energy at the

Figure 10. A diagram of a typical ERDA sample chamber where a beam of 2 MeV 4He ions are scattered at low angles by protons in the sample. Modified after Fig. 1 of Sweeney et al. 1997.

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detector is a function of the depth of interaction with the 4H+ ion. Sweeney et al. (1997) used a microbeam elastic recoil detection analysis to determine the hydrogen content of minerals. With suitable calibration, a depth profile (Fig. 11) as well as the absolute H-concentration can be obtained, in principle. In ERDA, there is a problem of H-loss due to diffusion away from the He+ beam, but it is quantifiable and the technique is readily applicable to the analysis of H in both hydrous and nominally anhydrous minerals down to the 0.04 wt% (400 ppm wt) level. Sweeney et al. state that this detection limit is potentially improvable with better protected electronics. Proton-proton scattering. Furuno et al. (2003) describe an application of a proton–proton elastic recoil coincidence spectroscopy to hydrogen analysis using a proton microbeam at an energy of 20 MeV. This method provides depth profiles of hydrogen over a thickness of 200 μm of silicate samples in a short time. A typical beam size is as small as 27 × 32 μm. The depth resolution is about 10 μm. The present work proves that the proton–proton elastic recoil coincidence spectroscopy is a promising method for measurements of hydrogen in mineral and rock samples with thickness up to 200 μm. Proton beams at energies of 20 MeV can pass through several hundred micrometers with an energy loss of only a few MeV. Protons passing through a sheet of material experience proton-proton elastic recoil. Measurement of the energy-loss distribution from the protonproton scattering events is specific for H and has a sensitivity in the ppm range (Cohen et al. 1972). A typical detection system (Fig. 12) consists of two detectors that detect scattered protons in coincidence with the recoil protons. If the detectors are the same distance from the sample, both protons arrive at detectors at the same time, but with a 90° separation. Their energy will be less than the incident beam because of energy loss that is a non-linear function of the depth of the reaction below the sample surface (Fig. 13). This method was used by Wegdén et al. (2004) with a 2.8 MeV proton beam at Lund, Sweden. They were able to get strong signals from a synthetic pyroxene with 300 ppm water. Further development of this method demonstrated the analysis of hydrogen at the 100 of ppm H2O concentration level (Fig. 14) and showed that surface hydration could be distinguished from the intrinsic bulk hydrogen content (Wegdén et al. 2005) where depth profiles exceeded 1 micrometer. Reichart et al. (2004) used a similar method to produce a three-dimensional image of the hydrogen distributions in a polycrystalline synthetic CVD diamond film and showed that the hydrogen atoms were concentrated along the grain boundaries.

Figure 11. An ERDA profile of a grossular garnet 3-232 compared to a zero-hydrogen synthetic Al2O3 blank. The grossular (GRR 1386) contains 0.17 wt% H2O as was previously determined by NMR, H-extraction and FTIR. Modified after Figure 4 of Sweeney et al. 1997.

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Figure 12. Schematic drawing of the detection system for p-p scattering. Modified after Furuno et al. (2003).

Figure 13. The results of a typical proton-proton scattering experiment on a sample containing a hydrous inclusion. Modified from Figure 4 of Furuno (2003).

Figure 14. Depth profile of the hydrogen concentration (as H) of an orthopyroxene. (25 ppm H = 223 ppm H2O). Modified after Figure 1 of Wegdén et al. (2005).

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In each of these examples of p-p scattering, the potential of the method for geologic samples was clearly demonstrated, but as was the case of the NRA in the early 1980’s, significant effort will be required before it becomes a rigorous, accurate analytical technique. Other approaches have been suggested (Wirth 1997) such as electron energy-loss spectroscopy (EELS), but have not been developed into accurate analytical methods for hydrogen in the nominally anhydrous minerals. Hopefully, geoscientists will remain associated with the nuclear physics community to bring these promising tools into the realm of a routinely useable analytical instrument.

Nuclear magnetic resonance At first, one would think that proton nuclear magnetic resonance (1H-NMR), should be an ideal method for studying H in minerals if the content of iron and other paramagnetic ions is low (less than about 0.4 wt% FeO). Although proton NMR is widely used in the chemical sciences, it has seen comparatively little application to the low concentrations of water in the nominally anhydrous minerals in part because of its low sensitivity for protons. A major challenge to the investigation of nominally anhydrous silicate minerals is overcoming or accommodating the sensitivity limits of the technique. Quantitative NMR measurements becomes difficult at H concentrations less than about 1000 ppm wt because the probe background overwhelms the sample signal unless the background is minimized through the use of pulse sequences or is somehow subtracted from the sample signal. Furthermore, the concentrations of paramagnetic transition elements are sufficiently high in most minerals that they seriously compromise or effectively eliminate the proton signal through inhomogeneous magnetic interactions. Consequently, the small amount of proton NMR conducted on minerals has been largely focused on stoichiometrically hydrous minerals and, in particular, on synthetic ones with a minimal paramagnetic component. Early NMR work on a nominally anhydrous mineral focused on the channel water in beryl where workers found the NMR signal from water in the channels and were able to conclude that the H-H vector was parallel to the c-axis (Pare and Ducros 1964; Sugitani et al. 1966; Zayarzina et al. 1969). Later work by (Carson et al. 1982) was concerned with the water in cordierite and found that the water was undergoing some kind of motion on a time scale faster than one microsecond. However, none of these studies attempted to quantitatively determine the absolute amount of water in the minerals from the NMR spectrum. Subsequent studies of beryl did distinguish between two orientations of water and determined their relative proportions (Charoy et al. 1996; Lodzinski et al. 2005). The first attempt to examine a range of nominally anhydrous minerals (Yesinowski et al. 1988) used a method known as magic angle spinning NMR. NMR spectra of solids are usually very broad due to magnetic anisotropic interactions among components of the crystals. However, high-resolution spectra can often be obtained through a method known as “magic angle” spinning NMR (MAS-NMR). In this experiment, the sample holder is rapidly spun with its axis 54.7° with respect to the applied magnetic field. If the line shape of the non-spinning sample is dominated by inhomogeneous interactions, as it often is for minerals with low hydrogen contents, magic angle spinning produces a sharp central band as well as a set of “spinning sidebands” spaced at integer multiples of the spinning frequency. Paramagnetic metal ions in the sample can complicate the NMR experiment because they introduce additional interactions with their unpaired electron spins. In addition to a number of stoichiometrically hydrous minerals, Yesinowski et al. (1988) examined microcline, quartz, and nepheline and grossular with 1H MAS NMR spectra. Although the found mostly fluid inclusions, they were able to show that different hydrous species could be distinguished but determined only their relative amounts (Fig. 15). Cho and Rossman (1993) further developed the technique for minerals and presented data on OH in grossular crystals with 0.17 to 0.31 wt% H2O. They were able to show that in low

Analytical Methods

17

water-content garnets, the mode of substitution is not dominated by the hydrogarnet substitution (H4O44−), but rather by protons in pairs (Fig. 16). Proton NMR is sensitive to just the hydrogen environment and, and is inherently quantitative. Relative amounts of various species can be determined, and, with suitable calibration, so can the absolute hydrogen content of the sample. To avoid the problem with paramagnetic components in natural samples, Kohn (1996) synthesized synthetic pyroxenes and forsterite and used NMR to study their hydrous components. He reported that they contained 0.02 to 0.24 wt% H2O. This and a subsequent report (Kohn 1998) indicated that the concentrations of hydrous components in these fine-grained materials were much higher than any earlier study suggested. Keppler and Rauch (2000) subsequently showed that polycrystalline materials have much higher water contents than the corresponding single crystal and suggested that the high water contents reported by Kohn (1996) were not representative of the true water content of the crystals. Contributions from hydrous species on grain boundaries, growth defects and submicroscopic fluid (or melt) inclusions are possible sources of these problems. Keppler and Rauch repeated an observation that this author’s group has long recognized: “measurements [of low hydrogen content] on powders are generally not reliable, no matter which analytical method is applied.” A following section discusses observations of elevated concentrations of water in mineral surfaces in more detail. An approximately universal absorption coefficient for the infrared spectra of feldspars was determined from 1H-MAS NMR spectra by Johnson and Rossman (2003). In this study, the spectra were used to determine the H concentration of three alkali feldspars and for the first time, eight plagioclase feldspars. To accurately measure structural H concentrations in

Figure 15. 1H magic angle NMR spectra at 500 MHz of (top) microcline feldspar from Lake George, Colorado and (bottom) microcline from the White Queen Mine, Pala, California. Peak A is an organic contaminant; peak B is water in fluid inclusions; peak C is a structurally bound, isolated H2O group; and peaks C* are the spinning sidebands of the structurally bound H2O group. Modified after Figure 7 of Yesinowski et al. (1988).

Figure 16. Proton MAS-NMR spectra of grossular from the Lelatema Mountains, Tanzania, with 0.17 wt% H2O showing 2 types of hydrogen. Data were obtained with 2000 scans, a 4 second delay between each scan, and a Gaussian line fit. The narrow line signal near zero KHz is from a proton either far removed from other H nuclei or is part of a mobile species within the sample. The broad band arises from pairs of protons in close proximity to each other, rather than a hydrogarnet substitution (Cho and Rossman 1993)

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samples with such low H contents (20 µm, synchrotron IR radiation was

Structure of Hydrous Species Using Polarized IR Spectroscopy

33

increasingly used as a light source in the last decade. The extreme brightness facilitates measurements with excellent signal-to-noise ratio down to the diffraction limits 2.7 Å (Emsley et al. 1981). The lower and upper limits are found at 2.4 Å (without external pressure) and beyond 3 Å (with a continuous transition to non-bonded entities). Due to the attractive force of the acceptor, the hydrogen atom is pulled away from the donor and the O-H bond is attenuated compared to a non-bonded unit. Whereas the O-H distance is approximately 0.98 Å in a free hydrous group, it is successively elongated up to 1.20 Å in the shortest hydrogen bonds (which are therefore symmetric without distinction between donor and acceptor atoms). Due to the variability in hydrogen bond distances (and forces) the strength of a hydrogen bond correlates closely with the frequency of its stretching vibration over a wide range of wavenumbers . The common regions are 3200-3750 cm−1 for weak H bonds (and non-bonded units), 1600-3200 cm−1 for strong H bonds, and 700-1600 cm−1 for very strong H bonds. 4000

O-H stretching wavenumber (cm-1)

The bond length vs. stretching frequency correlation of H bonds has been investigated theoretically by Bellamy and Owen (1969), and empirical correlation diagrams have been published since the 1950s, e.g., Nakamoto et al. (1955), Novak (1974), Mikenda (1986), Libowitzky (1999). The diagram for O···O bond lengths in Figure 2 shows the typical positive and curved trend line of the correlation. Scatter of data is caused by deviation of bonds from a straight O-H···O geometry and by influence of cations (see below). In general, it is observed that (very) strong H bonds tend to be more linear, whereas weak ones (such as those mostly observed in hydrous defects) are frequently bent. Further correlation diagrams that may be useful under certain circumstances

3500

3000

2500

2000

1500

1000 2.4

2.6

2.8

3.0

3.2

3.4

d(O···O) (Å)

Figure 2. Correlation between hydrogen bond length d(O···O) and O-H stretching frequency (wavenumbers) after Libowitzky (1999).

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are stretching frequency vs. d(H···O) by e.g., Libowitzky (1999), stretching vs. d(O-H) by Novak (1974), stretching vs. bending for OH groups (Novak 1974), hydrogen bond strength vs. ν1-ν3 splitting of the H2O molecule (Schiffer et al. 1976). All these empirical correlation diagrams have been obtained by comparison of spectroscopic data with structural data from X-ray and neutron diffraction of hydrous compounds. Thus, in the case of NAMs where only spectroscopic data are available on hydrous defects, they may give important structural information. However, as discussed below, distances obtained from correlation diagrams must be used with caution, because stoichiometric compounds with well-defined hydrogen sites may not be directly comparable with extremely low concentrations of hydrogen atoms located at locally distorted defect sites. Another correlation with hydrogen bond strength is observed in the band widths of O-H stretching bands. Weak H bonds at high wavenumbers in general reveal sharp bands with small full width at half maximum (FWHM), e.g., a few cm−1. With increasing H bond strength and decreasing wavenumber the band width increases up to several hundred cm−1 in case of very strong H bonds (Novak 1974). The reason for this behavior is the increasing anharmonicity of the vibration that correlates also with the increasing H bond strength (Szaly et al. 2002). These extremely broad bands centered at very low wavenumbers that resemble uneven background lines may be recognized in polarized spectra of a few stoichiometric hydrates with high water contents and very strong H bonds (e.g., Hammer et al. 1998), however they have never been observed in NAMs. It may be speculated that they do not exist in the form of defects or that they are simply invisible due to the peculiar background-like shape and the low concentration. In contrast, broad bands in the common O-H stretching region (~2500-3800 cm−1) have been observed in a number of NAMs, e.g., enstatite (Mierdel and Keppler 2004), ringwoodite (Smyth et al. 2003), wadsleyite (Jacobsen et al. 2005), and originate from strong H bonding or other phenomena such as structural disorder. In general, these broad bands and uneven background lines, aggravated by insufficient S/N ratio and small sample size, may affect accurate analysis of water contents in NAMs by IR spectroscopy (see Rossman 2006, this volume).

The spatial orientation of hydrous species Symmetry considerations. IR radiation traveling through a crystal never affects one individual O-H bond in a single unit cell, but rather many of them at the same time and phase. Thus, the vibrations are not independent and couple in-phase and out-of-phase in various combinations for all symmetry-equivalent entities. The rules for coupling according to symmetry are given by group theory in the form of normal mode analysis for molecules and by factor group analysis for crystals (e.g., Fadini and Schnepel 1989). As an example, the three atoms of a single H2O molecule in the gas phase (besides 3 translations and rotations in the three coordinates of space) possess three fundamental vibrations: a bending mode (ν2) above 1600 cm−1 and the symmetric and antisymmetric stretching modes (ν1 and ν3) with slightly different frequencies above 3600 cm−1. Thus, because of symmetry the vibrations do not occur independently along each O-H vector direction but in a coupled way along the vector sum and vector difference, i.e., parallel and perpendicular to the molecular axis. If more than one H2O molecule were contained in the primitive unit cell of a stoichiometric hydrate, further combinations of vibrations were possible. It is an advantage of low concentrations of hydrogen defects in NAMs that the vibrating species are diluted, and coupling of vibrations across many unit cells does not affect the vibrational energies. However, if symmetry-equivalent O-H bonds are grouped together in close vicinity within a unit cell, e.g., in the form of H2O molecules or clustered OH defects, splitting of bands by symmetry must be considered. Polarized radiation. In an optically anisotropic (non-cubic) crystal (IR) light is split into two rays with perpendicularly oriented polarization directions vibrating parallel to the main axes of the indicatrix section. In an absorption experiment, light that is already polarized parallel to

Structure of Hydrous Species Using Polarized IR Spectroscopy one of the indicatrix directions (X, Y, Z) is affected only by the component (Ax , Ay , Az) of an absorber which is parallel to this polarization direction, i.e., when the electric vector E of the light wave is parallel to (a component of) the oscillating dipole (Fig. 3). This component has a simple cosine squared relationship (Table 1) to the magnitude of total absorbance (Libowitzky and Rossman 1996). Thus, by measuring a crystal section in the two principal polarization directions the orientation of the absorber in this section is obtained. By measuring all three principal polarization directions of the indicatrix ellipsoid, the spatial orientation of the absorber is obtained. Moreover, only the sum of all three polarized component spectra yields the full magnitude of the absorber, i.e., the total absorbance (Libowitzky and Rossman 1996).

35

z

Az

A g

x

a b Ay

y Ax

Figure 3. Spatial orientation of an absorber A in an orthogonal optical axis system X, Y,

Polarizers for IR radiation are available acZ. Only components of absorption Ax, Ay, Az cording to two construction principles: (a) wire can be accessed during an IR absorption exgrid polarizers on an IR transparent material (or periment with polarized radiation, and facilitate calculation of the spatial orientation of even without a support), absorbing radiation parthe absorber and the total absorbance (after allel to the extremely fine, parallel (gold) wires, Libowitzky and Rossman 1996). polarize radiation over a wide angular range but their efficiency is usually limited to ~1:100. (b) Crystal polarizers, constructed similar to the well-known Nicol’s prisms, are made from IRtransparent but strongly birefringent material (e.g., LiIO3). They operate only in a narrow angular range, but their efficiencies may be as high as 1:105.

Total absorbance: a first step towards quantitative water analysis Due to the logarithmic relation between transmittance and absorbance (see above) only the total absorbance is proportional to the concentration of an absorber. Therefore unpolarized measurements and powder samples of optically anisotropic crystals are not recommended for quantitative measurements. Even the use of low-quality polarizers may bias results (Libowitzky and Rossman 1996). In cases where oriented single-crystals cannot be prepared, statistical analysis of polarized measurements on randomly oriented mineral grains in a rock section can be treated by comparing the measured spectra with polarized reference spectra of the same material (Asimov et al. 2006). In general, it must be emphasized that only integrated measurement of absorbance (Ai), i.e., the area of a band with properly treated background results in reasonable quantitative data. In that way the various band widths (FWHMs) and even overlapping peaks are reliably evaluated. Correct subtraction of the background line is of high importance. Though a linear background line can be chosen in many cases, problems may be encountered in the case of curved background shape, broad bands (see above) and very low band heights. Once the total absorbance and thickness of the sample have been measured, the concentration can be calculated according to Beer-Lambert’s law. Unfortunately, the molar absorption coefficient is not a unique constant for hydrogen in minerals. In contrast, it varies by orders of magnitude depending upon hydrogen bond strength and stretching wavenumber. Though the linear relation between ε and the wavenumber of the O-H stretching vibration can be used for a general water calibration trend (Libowitzky and Rossman 1997), mineral specific calibrations (in reference to other analytical methods) are preferred. A detailed review of this topic is given by Rossman (2006) in this volume.

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CONCEPTS OF STRUCTURAL MODELS FROM INFRARED DATA At the beginning of this paragraph it should be stressed that all concepts of structural incorporation models for traces of water in NAMs (whether in the form of H2O or OH− defects) have been developed from structural and crystal chemical observations of more or less hydrous minerals, where the information has been extracted from both diffraction and spectroscopy experiments in many cases. Therefore the examples at the end of this chapter contain also hydrous phases with stoichiometric hydrogen.

Charge balance and substitution Among the principles of crystal chemistry Pauling’s five rules (Pauling 1960) represent the most basic ideas on stable ionic compounds. Whereas the first rule comments on bond distances and coordination numbers resulting from the sums and ratios of effective ionic radii, respectively, the second rule comments on charge neutrality. The sum of charges arriving from the ligands at the center of a stable coordination polyhedron equals the (negative) charge of the central atom itself, referred to as the bond strength sum (e.g., Gibbs et al. 2003). In a more general way charge is compensated in the immediate surrounding (coordination sphere) of a charged particle. This principle is also employed in modern structural analysis to check the consistency of a crystal structure and to find hidden hydrogen atoms (missing charges!) in X-ray structural refinements (Brown 1981). This so-called bond valence analysis may even be applied to find preferred oxygen sites for trace hydroxyl substitution in a crystal structure. Thus, the most underbonded O atom in a structure may be considered an ideal docking site for H, e.g., O1 in wadsleyite (Smyth 1987). Equivalent ideas can be applied to the incorporation of a hydrogen defect in a host crystal structure. If a hydrogen atom (actually a H+ ion or “proton”) enters a crystal structure, its positive charge must be compensated. Or, in other words, if an O2− atom in a crystal structure is replaced by an OH− group, the missing negative charge must be compensated. An easy way to do so is to change the charge of a neighboring element with different valence states, e.g., Fe3+ + O2− + ½ H2 ↔ Fe2+ + OH− (e.g., Skogby and Rossman 1989; Koch-Müller et al. 2005). Another common coupled substitution in silicates may involve tetrahedral Si - Al exchange: Si4+ + O2− + ½ H2 ↔ Al3+ + OH− (e.g., Andrut et al. 2003). Whereas the former process involves only electronic charge transfer, the latter requires exchange of framework atoms and appears more likely to occur during crystal growth than by later diffusion processes. Another substitution mechanism which may easily occur during growth and which does not even require charge compensation is the incorporation of OH− groups for halogen atoms (F−, Cl−) such as in apatite (Baumer et al. 1985) and topaz (see example below). In general, unlike crystal growth the later gain and loss of hydrogen and charge compensating neighbor atoms require diffusion processes which are described in more detail by Ingrin (2006) in this volume. Probably the simplest way to compensate for the positive charge of an additional H atom is the simultaneous creation of a cation vacancy. This type of defect is known in olivine (Libowitzky and Beran 1995), perovskite (Beran et al. 1996) and others. Even in synthetic high-P phases, e.g., wadsleyite (Jacobsen et al. 2005), ringwoodite (Smyth et al. 2003), a clear correlation between H2O content (up to 1 wt %) and cation vacancies was established. In the hydrogarnet substitution, a cluster of four OH− groups is facilitated by a Si4+ vacancy at a tetrahedral site (see below). The latter has not only been observed as a trace defect but also as a major constituent of natural grossular garnets containing more than 1 wt% of H2O (Rossman and Aines 1991). Even if investigation of the correlation mechanisms of hydrogen defects with other substituents by chemical analysis may be an easy task for minerals with considerable H contents and characteristic trace element concentrations, it is almost impossible in cases of hydrogen trace defects (of the order of tens of wt. ppm) because the concentrations of accompanying minor and trace elements in the investigated minerals are commonly higher by orders of magnitude.

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Electrostatic considerations on defect geometry Whereas the electrostatic considerations above provide charge neutrality around a hydrous defect site, the charges of the surrounding atoms also constrain the orientation of an OH− or H2O group. Both hydrous species are polar with the negative end at the oxygen and the positive end at the hydrogen atom(s). Therefore, their orientation in a structure is strongly influenced by the attractive and repelling forces of surrounding ions. Because the oxygen atom of an OH− group is usually part of the crystal structure, it is connected to cations in its first coordination sphere. The same holds true for the H2O molecule in a number of stoichiometric hydrates. A characteristic coordination environment around the oxygen atom of a hydroxyl group is a flat trigonal pyramid with the cations at the corners of the base triangle and the OH group on top with the H atom pointing upwards, i.e., perpendicular to the base triangle, away from the other positive charges. This coordination type is an important part of the brucite sheet structure (e.g., Nagai et al. 2000). The orientation of the O-H vector perpendicular to the basal cation plane is caused only by three ions of equal charge, e.g., Mg2+, Mg2+, Fe2+. Substitution of cations by atoms with different valence, e.g., Al3+, Li+, or even by a vacancy (thus changing also the coordination number) results in considerable deviation of the O-H vector from the normal. A nice example for this type of coordination and the influence of (missing) cations and changed coordination is observed in the mica minerals (Fig. 4). In trioctahedral micas the octahedral layer builds up the regular brucite-type environment of the OH group and thus the O-H vector is aligned exactly perpendicular to the octahedal layer. In dioctahedral micas one third of the octahedral layer cations are missing. Because of this asymmetric distribution of only two positive charges around each OH group, the OH vector is strongly inclined towards the octahedral layer (Beran 2002). At the proton end, attractive forces of H bond acceptors may influence and distort the orientation of the hydrous defect and lead to elongated O-H bonds resulting in a decrease of O-H stretching frequencies (see above). A contrasting effect (i.e., increase of the stretching wavenumber beyond 3700 cm−1) is observed if cations opposite to the proton cause a compressed O-H bond, such as in amphiboles with occupied A site (Rowbotham and Farmer 1973).

a)

H1 Mg

O4 Fe Mg

b) O6 Al

Al

H1

Figure 4. Octahedral layer with OH groups in two mica minerals. (a) Trioctahedral biotite (structural data from Brigatti and Davoli 1990): Three cations (Mg, Fe2+) around the hydroxyl group cause the O-H vector to be approximately perpendicular to the layer (brucite-type coordination). (b) Dioctahedral muscovite (structural data from Rothbauer 1971): The missing third cation (dashed square) in the coordination of the hydroxyl group forces the O-H vector into a tilted direction.

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In practice, once the orientation of an O-H vector has been obtained from polarized IR absorption measurements, the probable defect site is considered best by using a three-dimensional ball-and-stick model of the host structure. Thus, electrostatic constraints (see above) can be verified and, moreover, the necessary space (see below) to host a hydrous defect can be investigated. Two-dimensional structure drawings and even animated computer plots may be helpful but they remain always limited in information (as do abstract lists of bond lengths and angles).

Space requirements: ideal and distorted models Though hydrogen is a very small and mobile ion, the possible arrangements of O-H bonds and of O-H···O hydrogen bonds impose certain space requirements, which are not available at any position in a crystal structure and so help to constrain possible sites of hydrogen incorporation. Data for hydrogen bond lengths derived from stretching vibrations (see above) further help to develop a probable model for a suitable site of hydrogen incorporation. However, as mentioned above in the course of charge balance considerations, the incorporation of hydrous species is charge-compensated by other defects such as different cations or even vacancies in the close neighborhood. Similar to major substitution in solid solution series these structural defects cause distortions of the structure (limiting the use of H bond length calculations) which may be considered in two contrasting ways (e.g., Urusov 1992; Andrut et al. 2004). VCA model. In the virtual crystal approximation (VCA) model, no structural relaxation around a site of substitution or a defect is assumed and thus the surrounding bond distances represent an arithmetic average of the substituted and unsubstituted geometries, according to the amount of substitution (Vegard’s rule). In case of trace defects this would imply an almost unchanged defect environment where calculated H bond distances could be easily applied. Because of the “averaging” and “bulk” character of diffraction methods, their results frequently seem to support the VCA model. Hard sphere model. In the hard sphere model, full relaxation of the structure around a “hard” substituent or defect is assumed. Thus the defect environment, e.g., bond distances, develops undisturbed, as if the whole structure would consist to 100% of this substitution or defect. Because of the two different environments (with and without defect/substitution) at a single site, spectroscopic methods are superior to reveal the real situation in the structure. The real situation, expressed by the degree of relaxation (Urusov 1992), is usually found between these two extremes. The example of the hydrogarnet substitution (see below) elucidates the power of IR spectroscopy to identify the true defect environment and points out the limits of theoretical H bond calculations.

Influence on band energies from cation substitution Substitution of cations by elements with different valences and even by vacancies in the vicinity of hydrous defects has been discussed above as an important mechanism to achieve charge neutrality. Another aspect of cationic substitution is that different cations in the neighborhood and coordination sphere of, for example an OH− group may affect the energy of the O-H stretching vibration by more than 50 cm−1. In turn, shifts and splitting of absorption bands may indicate different cationic surroundings of OH defects in crystal structures. Substitution of Mg by different cations (e.g., Mn, Zn, Ni, Fe2+, Fe3+) and formation of solid solution series in common silicates with the brucite-type OH coordination such as amphiboles and talc shows interesting results. The band shift to lower wavenumbers is linearly correlated with increasing electronegativity of the substituting element, i.e., from Mg to Fe2+ and further to Fe3+ (Strens 1974). Moreover, even the number of substituents can be derived from the spectra. An intermediate Mg-Fe actinolite shows four equally spaced OH stretching bands (Fig. 5), which can be correlated according to their intensities and by comparison to pure endmember tremolite to four cationic environments around the OH group (Burns and Strens 1966): MgMgMg (~3670 cm−1), MgMgFe, MgFeFe, FeFeFe (~3625 cm−1).

Structure of Hydrous Species Using Polarized IR Spectroscopy

39

Figure 5. The four O-H stretching bands in an intermediate Mg-Fe-actinolite and their assignment to different cationic environments. Modified after Burns and Strens (1966).

Discrimination among hydrous defects Single OH− group. A single hydroxyl group, i.e., a trace defect without any symmetryequivalence in its close vicinity, is characterized by a single band whose position is mainly dependent upon the strength of hydrogen bonding (see above). Nevertheless, it is commonly observed in the 3000-3750 cm−1 region. However, because this single defect may occur with different cationic environments in different parts of the crystal (e.g., MgMgMg or MgFeMg, see above) every type of environment and even a vacant site may cause a separate band. Further vibrations which are caused by an OH− group are a Me-O-H bending mode around 700-1400 cm−1 (depending also upon H bond strength) and the combination mode of the stretching and bending vibrations around 4500 cm−1. However, if a hydroxyl defect occurs only in trace concentration, the former is usually hidden by the strong vibrations of the host mineral, and the latter may be too weak to be detected. Single H2O molecule. Because of its symmetry an undistorted water molecule possesses two stretching vibrations, one symmetric and one antisymmetric mode. As for all O-H vibrations their band positions depend upon H bonding, but are frequently observed in the 3000-3750 cm−1 region, meaning that they cannot be distinguished from bands of OH− groups. Fortunately, the water molecule also has a bending mode at approximately 1600-1650 cm−1, which is a characteristic feature of this unit. Another characteristic band may be observed at ~5200 cm−1, which is the combination of stretching and bending vibrations. However, because of its weak intensity it may be invisible at trace concentrations. H3O+ group. Hydronium (if present as the only hydrous species) is identified by four vibrations with a characteristic bending mode around 1100 cm−1 (Nakamoto 1977). All these vibrations are similar to water and hydroxyl stretching and bending modes, so that they cannot be distinguished unambiguously from a combination of different H2O or H2O + OH− species. H3O2− group. This unit, which has been considered to possess a symmetric, very short hydrogen bond in its center, has only been observed in stoichiometric phases (e.g., Beran et al. 1997). These investigations confirmed the (very) strong central hydrogen bond, but indicated also the non-symmetric configuration of the bond. Thus, the unit is considered as a linked H2O plus OH− group. The observed vibrations are the stretching modes of the terminal O-H units at high wavenumbers, the stretching mode of the central (very) strong H bond at low wavenumbers, and the bending mode of an H2O unit. Because of the very broad band shape of

40

Libowitzky & Beran

the (very) strong central H bond (see above) it is unlikely that this feature is found at minor or trace concentration levels. H2 molecule. Due to its high symmetry the H2 molecule is only Raman active with a stretching vibration at ca. 4155 cm−1. Though observed by micro-Raman spectroscopy in special fluid inclusions in melts and glasses (if present as a major constituent), it is unlikely to occur commonly as a structural defect in minerals. Nevertheless, it should be stressed that, except under very peculiar circumstances (distortion of the molecular symmetry by attraction of H2 to the host structure), IR spectroscopy is not suited to detect H2 in minerals. Clusters of hydrous defects. Clusters of vibrating O-H units (in principle, an H2O molecule is a simple cluster of two O-H vectors with a common O atom) are identified by more than one vibration according to their symmetry (see above). Unfortunately, this is not an absolute necessity, as is demonstrated in the case of (OH−)4 clusters in hydrogarnet (see below), which are characterized by a single OH stretching band (Rossman and Aines 1991). However, charge balance considerations (in this case a vacancy at the Si4+ position) indicate the necessity of more than one OH− group. Hydrous inclusions. One of the pitfalls of IR spectroscopic identification of hydrous defects is that even microscopically clear, gem-quality samples may include sub-microscopic fluid inclusions that resemble true structural defects. Fluid inclusions are readily identified by their characteristic broad water bands around 3400 cm−1 and by the appearance of sharp ice bands upon freezing. However, inclusions of hydrous minerals may be very difficult to identify. This problem has been discussed by Khisina et al. (2001). In other cases, however, IR spectroscopy may be the perfect tool to identify “invisible” mineral inclusions by their characteristic fingerprints in the OH stretching region, e.g., kaolinite in kyanite (Wieczorek et al. 2004) or corundum (Beran and Rossman 2006). During heating (either by nature or by experiment) these hydrous inclusions may act as a source of hydrogen to incorporate further structural defects.

Deuteration Though the fingerprint of O-H stretching vibrations can usually be distinguished from the fundamentals and overtones of the host mineral, problematic cases need additional treatment. Because diffusion of hydrogen at elevated temperatures is rapid, the isotope deuterium (D) can be incorporated into the material in exchange for hydrogen (Ryskin 1974). Due to the dependence of vibrational frequencies on mass (see above), corresponding O-D bands are observed at lower wavenumbers, shifted by a factor of ~1.35 (depending upon H bonding and anharmonicity, the latter also causing deviation from the ideal value √2). The reduced anharmonicity of O-D stretching vibrations in comparison to O-H modes may also result in sharper peak shapes helping to deconvolute interfering absorption bands. The problem of overlapping bands may also be solved by cooling samples to liquid nitrogen temperature in a cooling stage (see above), which may lead both to reduced FWHM and variable shifts of the peaks.

EXAMPLES Vesuvianite: orientation and hydrogen bonding of hydroxyl groups Vesuvianite, ~Ca19(Mg,Fe)3(Al,Fe)10Si18O70(OH,F)8, is an ideal first example to demonstrate the power of quantitative IR data using the wavenumber vs. hydrogen bond distance correlation and polarized spectra to constrain the O-H vector orientations. Though a chemically complicated sorosilicate (Groat et al. 1992), two different hydroxyl groups can be clearly distinguished in the structure, which have also been determined by neutron diffraction (Lager et al. 1999). The latter is the reason for the selection of a hydrous mineral as example rather than a NAM.

Structure of Hydrous Species Using Polarized IR Spectroscopy

41

Figure 6 shows polarized IR absorption spectra of a (hk0) slab of tetragonal vesuvianite with the E vector of light vibrating parallel and perpendicular to the c axis, respectively. Bands of the E//c spectrum are obviously more intense than those of the E⊥c direction. There is a strong band around 3100-3200 cm−1 and a group of strong bands between 3450 and 3700 cm−1. The former is quite broad and its low wavenumber indicates a hydrogen bond distance of d(O···O) ~ 2.70 Å (Libowitzky 1999). Its intensity perpendicular to c is zero and thus indicates an O-H orientation parallel to the c axis. The latter bands are sharper and their wavenumbers indicate only weak or no hydrogen bonding. Exact evaluation of the band areas in both polarization directions confirms an O-H vector orientation of ~35° tilted from the c axis (Bellatreccia et al. 2005). The identical pleochroic behavior of all high-energy bands indicates one hydroxyl site with different cationic environment resulting in the slightly different positions of these bands. The inset of Figure 6 shows a detail of the vesuvianite structure and confirms the band assignment from above. There is actually a moderately strong hydrogen bond at H(2) with d(O10···O10) ~ 2.72 Å (Lager et al. 1999), the O(10)-H(2) vector pointing exactly parallel to the c axis. Another hydroxyl group is located at O(11)-H(1) confirming the acute angle towards the c axis and various cations in its environment. Its weak, bifurcated H bond is in good agreement with the high wavenumbers of the IR bands. Quantitative measurements of the OH content in vesuvianites by IR spectroscopy and SIMS analyses (Bellatreccia et al. 2005) confirmed the general calibration trend of Libowitzky and Rossman (1997), although it had been demonstrated to be inaccurate for a number of NAMs.

Hydrogarnet substitution - the (OH)44− cluster Garnets contain a wide range of water concentrations, starting from a few wt. ppm in mantle garnets up to several wt% in samples of the grandite (grossular-andradite) series (Beran and Libowitzky 2006). However, their optically isotropic character, their complicated chemistry

Absorbance

1.5

1.0

Vesuvianite Rotkopf, Zillertal, Austria

Y(1)

O(10) ~ 2.7 Å

X(3)

H(2) X(3) O(10)

0.5 E || c 0.0

E^c 3800

3600

3400

3200

3000

-1

Wavenumber (cm ) Figure 6. Polarized IR absorption spectra of vesuvianite (data from Kurka 2002). The strong band at ~3100 cm−1 occurs only in the c spectrum and indicates by its low wavenumber a moderately strong hydrogen bond, which is readily assigned to O(10)-H(2)···O(10) in the structure of vesuvianite (modified after Lager et al. 1999).

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due to a number of solid solution series, and a wide variety of observed OH stretching modes at rather high wavenumbers (3500-3700 cm−1) (indicating absence of hydrogen bonding) makes unambiguous identification of distinct OH− defects in silicate garnets difficult. In contrast, the high concentration of water in certain grossular garnets facilitated investigation by diffraction, analytical and spectroscopic methods, which revealed four OH− groups substituting for a SiO44− group, i.e., the (OH)44− cluster in the so-called hydrogarnet (hydrogrossular) substitution. Figure 7 shows the configuration of this cluster in comparison with a common silicate tetrahedron. With regard to the theoretical considerations on defects in crystal structures above, a number of interesting features are observed.

(SiO4)

4-

(O4H4)

4-

Figure 7. A silicate tetrahedron (left) and the hydrogarnet substitution (right) indicating the increased size of the tetrahedron, the empty Si4+ position (square) and the H atoms above the tetrahedral faces (structural data from Lager et al. 1987).

The incorporation of four protons is charge-compensated by a Si4+ vacancy. Thus, charge balance is achieved in the closest vicinity (coordination sphere) of the defect site. However, the four hydrogen atoms are not placed inside the tetrahedron (pointing towards the empty silicon site) as was proposed in an earlier paper (Sacerdoti and Passaglia 1985). Although this configuration might be considered an ideal mechanism for local charge compensation for the missing Si4+, it is not favorable due to electrostatic repulsion of the four protons in close proximity to each other. More recent papers confirm that the positions of the H atoms are, instead, rather slightly above (outside) the faces of the tetrahedron (e.g., Lager et al. 1987). Because of the missing central charge of Si4+, the size of the hydrogarnet tetrahedron is increased by ~20 % with respect to the silicate tetrahedron (Si-O ~ 1.65 Å, -O ~ 1.95 Å) in pure endmember hydrogrossular. The corresponding IR spectrum shows a single absorption band at 3660 cm−1 (Rossman and Aines 1991). This high wavenumber is in agreement with only weak or no hydrogen bonding along the edge of the tetrahedron (O···O > 3 Å). Intermediate solid solutions that contain both silicate and hydrogrossular tetrahedra are characterized by two bands at 3600 and 3660 cm−1 with different intensities. This classical two-mode behavior was interpreted by a pure hydrogrossular environment (band at 3660 cm−1) and an (OH)44− defect surrounded by silicate tetrahedra (3600 cm−1). The latter wavenumber is still in agreement with only weak hydrogen bonding in a strongly inflated tetrahedron and thus confirms the hard sphere model. If, in contrast, the VCA model (see above) were pertinent, wavenumbers at low hydrogarnet concentrations would be expected at rather low wavenumbers due to short O···O distances in a SiO44− tetrahedron with almost unchanged size. Finally, it should be emphasized that the hydrogarnet substitution is not limited to grossular garnets, but has also been observed at low concentration levels in pyrope (Beran et al. 1993; Geiger et al. 2000) and even other minerals outside the garnet group, e.g., hydrozircon (Caruba et al. 1985). Moreover, the replacement of Si4+ by a cluster of four protons has been proposed as an important hydrogen incorporation mechanism by atomistic simulations (see also Wright 2006, this volume) for e.g., olivine (Braithwaite et al. 2003) and ringwoodite (Blanchard et al. 2005). With Ti4+ substituting for Al3+ in close vicinity to the tetrahedral (vacant) site, even an incomplete cluster of [(OH)3O]5− was suggested in pyrope by Khomenko et al. (1994). The combination of the hydrogarnet cluster with moderately strong hydrogen bonding was observed in the tetragonal garnet henritermierite (Armbruster et al 2001), where the distorted octahedron around Mn3+ provides an oxygen atom acting as H bond acceptor at rather close distance. In a

Structure of Hydrous Species Using Polarized IR Spectroscopy

43

suite of non-cubic garnets of the grossular-uvarovite join, Andrut et al. (2002) observed a number of varieties of the hydrogrossular substitution related to pleochroic IR absorption bands.

Water molecules in structural cavities: beryl and cordierite The framework silicates beryl, Be3Al2Si6O18, and cordierite, Mg2Al4Si5O18, contain structural units of 6-membered rings of tetrahedra which are stacked in such a way that channels parallel to the hexagonal c axis (in beryl) and parallel to the two-fold c axis (in orthorhombic low-cordierite) are formed. Both sets of channels are lined with oxygen atoms from the tetrahedral ligands, and with a maximum width of ~5.1 Å, separated by bottlenecks of ~2.8 Å, they can be occupied by alkalis, H2O and CO2 molecules (Kolesov and Geiger 2000a,b). Vibrational spectra contain water stretching and bending modes with wavenumbers around 3550-3700 cm−1 for stretching vibrations (Aurisicchio et al. 1994) indicative of weak or no hydrogen bonding. This is consistent with water molecules contained in the wide cavities of the channels (Fig. 8). The pleochroism of H2O stretching and bending vibrations in Raman and IR spectroscopic experiments confirms two possible orientations of the H2O molecule: Type I is oriented with the molecular axis perpendicular to the channel axis, whereas type II is oriented parallel to c. It is interesting to note that, although both stretching vibrations of H2O (ν1 symmetric, ν3 antisymmetric stretching) are IR and Raman active, due to strongly different activation cross sections, only ν3 occurs at 3700 cm−1 in IR spectra with E parallel to c for type I H2O, whereas Raman spectra yield only ν1 at 3607 cm−1 (Kolesov and Geiger (2000a). Type I H2O is predominantly observed, if alkalis are absent from the channels of the structure, whereas type II is found together with alkali ions (Aurisicchio et al. 1994). Considering the polar character of the water molecule, the structural incorporation model according to Figure 8 was developed. Moreover, as suspected from IR spectra, another possible type of incorporation in the form of an OH− group attached to a large alkali ion can be derived. Studies at variable temperature showed that type I water is dynamic and “rotates” about the channel axis down to very low temperatures (Winkler 1996). A recent study by Gatta et al. (2006) shows that this “rotation” is better described by a dynamic disorder of the water molecule over 6 equivalent positions. With increasing temperature both types of water approach the gaseous state followed by dehydration without destruction of the mineral structure (Aines and Rossman 1984). H2O I

H2O II

OH

-

9.2 Å

Rb,Cs Na

Na

2.8 Å Alkali-free

5.1 Å Na-rich

Alkali-rich

Figure 8. Three types of hydrous species in the structural channels of beryl (and similarly of cordierite). H2O I occurs preferably in alkali-free channels, H2O II in alkali (Na)-rich channels, and OH− ions are considered in correlation with large alkalis (Rb, Cs). Modified after Aurisicchio et al. (1994).

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OH substitution in topaz Although topaz has been studied by IR and Raman spectroscopy in a number of investigations (Gebert and Zemann 1965; Aines and Rossman 1985; Beny and Piriou 1987; Wunder et al. 1999; Bradbury and Williams 2003), all these papers lack one or another aspect of spectral evaluation. Therefore, we decided to demonstrate a worked example on a new topaz data set that has not been published previously. Moreover, we chose topaz because it contains a single OH group (with a single IR absorption band) substituting for the fluorine atom in its structure and because its orthorhombic symmetry is ideal for geometric considerations. A clear, colorless, gem-quality topaz crystal from Spitzkoppe, Namibia with a size of 9 × 11 × 17 mm was chosen for this study. The F-rich composition of this specimen, Al2SiO4F1.85(OH)0.15, (OH)/(OH + F) = 0.075, was confirmed by the correlation of Ribbe and Rosenberg (1971) using the lattice parameters from X-ray powder diffraction (space group Pbnm, Z = 4, a = 4.652 Å, b = 8.804 Å, c = 8.390 Å, CuKα, 5% Si standard). The sample was oriented according to the excellent cleavage parallel to (001) and along the optical extinction directions (optical setting: a = X, b = Y, c = Z). Three platelets (100), (010), (001) were cut from the sample, attached with crystal bond epoxy resin to a glass plate sample holder and diamond-polished to a final thickness of ~15 µm (uncertainty due to resin layer). To retain the large size of the extremely thin sections they were not removed from the glass plate, and spectra were corrected for background absorption from glass and epoxy resin. Figure 9 shows the polarized spectra of topaz in the O-H stretching region parallel to the three main axis directions with a strong absorption band at 3649 cm−1 and Figure 10 gives the angular absorption plots of the integrated absorbance of the OH stretching band in all three crystal sections. Both figures confirm that the absorption is almost zero parallel to the b axis (Y) direction. Thus the O-H dipole must be aligned within the (010) plane. Further inspection of the (010) angular absorbance plot (Fig. 10) and comparison of the X and Z directions (Fig. 9) indicate a preferred O-H orientation along the c axis (Z) direction. The ratio of the Z:X integrated band intensities is ~2:1, and application of the cosine squared relation (Table 1) results in an angle of ~35° between the O-H dipole and the c axis direction. This result is in excellent agreement with diffraction data (Zemann et al. 1979; Parise et al. 1980; Belokoneva et al. 1993) that yield ~29°, and also with crystal chemical considerations. Figure 11 shows the environment of the OH group in the structure of topaz. The repelling forces of the two Al atoms coordinating the OH group are such that the H atom is aligned almost exactly within the Al-O-Al plane and bisects the Al-O-Al angle. The weak H bonds around the H atom are in agreement with the high wavenumber position of the OH absorption band and have only a very minor influence on the alignment of the O-H vector. Finally it should be emphasized that the OH stretching band of F-rich topaz is not a single band but rather contains another component at ~3640-3646 cm−1 (depending upon peak fit constraints). This feature has been frequently ignored in older literature on topaz with low H content, but it was definitely described and discussed in recent papers on synthetic OH-rich topaz, e.g., Wunder et al. (1999). These details of peak fitting will be presented and discussed in a separate paper (Libowitzky, in prep.).

OH incorporation in diopside Pyroxenes contain significant amounts of hydrogen with concentrations ranging from a few 10s to more than 1000 wt. ppm H2O. Thus, pyroxene may indeed be a major storage site for hydrogen in the Earth’s upper mantle (Skogby 2006, this volume, and references therein). IR spectra of clinopyroxenes (cpx), i.e., diopside-hedenbergite, augite, and omphacite, are characterized by four regions of pleochroic OH stretching bands centered at 3630-3640, 35303540, 3450-3470 and 3350-3360 cm−1 (the latter only in a number of diopsides). Two different

Structure of Hydrous Species Using Polarized IR Spectroscopy 1.2

45

Topaz

Absorbance

1.0

Figure 9 (left). Polarized IR absorption spectra of colorless topaz from Spitzkoppe, Namibia, in the O-H stretching region parallel to the three principal axis directions. Sample thickness: ~15 µm, Perkin Elmer 1760X FTIR spectrometer (ceramic light source, KBr beam splitter, TGS detector), gold wire grid polarizer (efficiency ~ 1:100), circular sample aperture: 4 mm diameter, spectral resolution: 4 cm−1, 32 scans each averaged.

0.8 0.6 0.4 Z

0.2

X Y

0.0 3750

3700

3650

3600

3550

Wavenumber (cm-1)

20

15

10

Topaz (100)

Topaz (010)

Topaz (001)

Z

Z

X

5

Y

0 0

5

10

15

15

10

5

20

0

2.38

0.98 Al

H1

O1

5

10

15

8

6

4

2

Y

0 0

2

4

6

8

Figure 10 (above). Absorbance figures depicting the pleochroic scheme of the O-H stretching band of topaz (integrated absorbance vs. sample to polarizer angle). Lack of absorbance parallel to the b axis (Y) indicates an O-H vector orientation in the (010) plane. The anisotropic absorbance in (010) indicates an O-H orientation closer to the c axis (Z direction).

2.23

2.40

X

0

2.29

Al

c a

Si

b

Figure 11 (left). The environment of the hydroxyl group in the structure of topaz. The view was chosen in such a way that the OH group in the center of the picture and the two coordinating Al atoms (broken circles, connected by bold broken lines) are in the plane of projection. Broken lines indicate weak H···O bonds, the numbers give distances in Å. Structural data from Zemann et al. (1979).

Libowitzky & Beran

46

types of pleochroic behavior can be distinguished. Bands in the 3630-3640 cm−1 region are α- and β-polarized (group I bands), the lower energetic bands are γ-polarized (group II bands). Compare to Figure 12 in this chapter, and Table 1 in Skogby (2006), this volume. The two types of pleochroic bands suggest that at least two types of OH positions exist simultaneously in the diopside structure (Beran 1976; Ingrin et al. 1989; Skogby and Rossman 1989; Skogby et al. 1990). The position and pleochroism of the absorption bands are similar for different cpx samples, but the absolute intensities vary strongly, e.g., spectra of omphacites with jadeite-rich compositions show a strong γ-polarized absorption band at 3460-3470 cm−1, whereas those with diopside-rich compositions reveal a strong α-polarized band at 3620 cm−1 (Smyth et al. 1991). As an example for OH defect characterization by polarized IR spectroscopy, the study of a hydrothermally formed, gem-quality diopside crystal from Rotkopf, Tyrol, Austria, is given below (Andrut et al. 2003). The extremely strong pleochroism of the high-energy group I band in (010) sections at 3647 cm−1 (Fig. 12) suggests a strong preferred orientation of the OH dipole approximately parallel to the α index of refraction, i.e., the direction of the long diagonal of the unit cell projection parallel to [010] (Fig. 13). The moderate pleochroism of this band in (100) with a stronger component parallel to [010] (equivalent to nβ) indicates a strong deviation of the OH vector direction from an alignment within the (010) plane. These results confirm the model proposed by Beran (1976) that OH defects partially replace the O2 “zigzag” oxygen atoms pointing to the O3 oxygen atom of a neighboring silicate chain (Fig. 13). O2 is coordinated by 1 Mg, 1 Ca and 1 Si, thus forming the top of a flat slightly distorted trigonal pyramid, being an ideal candidate for a partial OH replacement. This replacement mode also occurs in a 1100 °C temperaturetreated crystal and evidently represents a very stable OH defect position. Another model of OH defect incorporation on O2 sites (with similar O-H vector orientation) can also be derived under the assumption of a vacant M1 site, resulting in a coordination of the OH defect by Ca and Si. Owing to the pleochroism of the low-energy band doublet at 3464 and 3359 cm−1 in (010) (Fig. 12), the OH dipole direction must be oriented roughly parallel to the γ index of refraction, i.e., the direction of the short diagonal in the (010) section of the unit cell. In addition, a slight deviation from the (010) plane is indicated. An OH dipole direction that is in agreement with the observed pleochroic behavior can be provided under the assumption of M2 vacancies. OH defects coordinated by 1 Mg and 1 Si are generated by a partial replacement of O2 oxygen atoms with an orientation pointing strongly above the Ca vacancy site. The separation of the low-energy bands is explained by a replacement of the coordinating Mg by Fe or Si by Al.

-1

Linear absorption coefficient (cm )

3.0 2.5

group I group II

2.0 1.5 1.0

Figure 12. Polarized OH absorption spectra of lightgreen diopside from Rotkopf, Zillertal, Tyrol, Austria, measured on (100) and (010) plates (modified after Andrut et al. 2003).

a in (010)

b in (100)

0.5 g in (010)

0.0

3800

3700

3600

3500

3400 -1

Wavenumber (cm )

3300

Structure of Hydrous Species Using Polarized IR Spectroscopy OH defects in perovskite

The IR spectrum of OH containing CaTiO3 perovskite consists of two bands with maxima centered at 3394 and 3326 cm−1 (Fig. 14). From the weak pleochroism of the bands in (001) and the more distinct pleochroism in (110), with a stronger component of absorption perpendicular to [001], an OH direction roughly pointing along [110] with the O2 oxygen atoms acting as donor is deduced (inset of Fig. 14). Using the hydrogen bond length vs. stretching frequency correlation of Libowitzky (1999), excellent agreement between the expected (calculated) H bond lengths and the actual O···O distances in the structure (2.75-2.78 Å) is obtained. The O-H vector orientation is facilitated only by the presence of a vacant Ca site (Beran et al. 1996, inset of Fig. 14). The OH defect positions coordinated by two Ti and one Ca atoms correspond to those proposed by Meade et al. (1994) in synthetic high-pressure MgSiO3 perovskite, where OH bands occur at 3483 and 3423 cm−1. Similar to CaTiO3 perovskite, the assumption of a vacant Mg position seems necessary from geometric and electrostatic considerations.

c

g a O3

Ca/Mg

Si

O1

O2

Ca/Mg

O3

O3 Si

O1

O1 O3

O3

a

Ca/Mg

O2

O2

Ca/Mg

a/2

Figure 13. Part of the diopside structure projected parallel to [010] with OH defects at O2 pointing to O3 of a neighboring silicate chain (modified after Beran 1999). Dark grey atoms, labels and tetrahedron belong to the silicate chain behind the light grey units. 0.10 b

Ca

Ti a

0.08

Absorbance

(Mg,Fe)SiO3 perovskite is most likely the major mineral phase in the Earth’s lower mantle and its role as a storage site for hydrogen has been recently discussed in review articles by Bolfan-Casanova (2005) and Ohtani (2005). On the other hand, though being a rare mineral, natural CaTiO3 perovskite forms in various geological environments, including kimberlitic rocks, and shows a wide range of compositions (Hu et al. 1992). Based on Paterson’s (1982) calibration, Beran et al. (1996) reported about 70 wt ppm H2O in perovskite of metasomatic origin.

47

[Ca] O2

0.06

0.04

0.02 3600

3400

3200

3000

2800

-1

Wavenumber (cm ) Figure 14. IR absorption spectra of perovskite. The inset shows the structure of perovskite. The two arrows, deviating from the plane of projection, indicate possible O-H vectors close to a vacant Ca site (modified after Beran et al. 1996).

In contrast, high-P,T experiments performed by Bolfan-Casanova et al. (2000) in the MgO-SiO2-H2O system did not detect any OH in MgSiO3 perovskite. Whereas MgSiO3 akimotoite, coexisting with MgSiO3 perovskite, synthesized at 24 GPa and 1600°C dissolved significant amounts of water (see below), perovskite was essentially dry.

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The conflicting experimental findings of Bolfan-Casanova et al. (2000) and Meade et al. (1994) might be related to the different synthesis conditions of the two studies. In the former study the sample was held at P and T for 3.5 hours, in the latter study for only several minutes. Therefore, any defects in MgSiO3 perovskite may have been annealed out in the Bolfan-Casanova et al. (2000) study whereas they were retained in the latter. Therefore, the presence of vacancies at the Mg site (see above) may be a key factor whether MgSiO3 perovskite incorporates OH groups in its crystal structure or not (Ross et al. 2003). Moreover, under natural conditions, the lower mantle contains elements, such as Fe and Al, which have not been involved in the synthesis and which may facilitate OH incorporation in perovskite (Bolfan-Casanova 2005).

OH traces in corundum The presence of OH in corundum as accessory mineral of mantle rocks is rather speculative. IR spectra of a 320 µm thick corundum from a South African eclogite assemblage showed no indication for the presence of OH (Rossman and Smyth 1990). Natural ruby and sapphire samples from crustal origin showed extremely weak absorption bands at 3310, 3230, and 3185 cm−1 (Beran 1991). Smith et al. (1995) confirmed OH bands also in sapphires from Southern Vietnam. In a comprehensive IR spectroscopic study of about 150 corundum samples from worldwide localities, Beran and Rossman (2006) established the presence of OH defects in corundum of crustal occurrences, however at concentration levels around only 0.5 wt ppm H2O or even lower. On the other hand, knowledge of possible OH defect incorporation mechanisms in this hexagonally close-packed mineral structure has a definite geophysical interest due to its close relation to MgSiO3 akimotoite in the high-PT regions of the Earth’s interior. Therefore, a number of synthetic samples have been studied in the past. OH groups in hydrothermally grown corundum were originally recognized by Belt (1967). A polarized IR spectroscopic study of a suite of Verneuil-grown corundum crystals (Beran 1991) revealed that variously colored samples show a distinct variability in the region of the OH fundamental vibration. Narrow strongly polarized OH bands with varying intensities are centered at 3310, 3230, and 3185 cm−1 (see above). Additional weak bands at 3290 cm−1 occur in (V Cr Fe Ti)-doped “alexandrite” sapphires, weak bands at 3280 and 3160 cm−1 appear in colorless corundum. Due to the strong polarization with maximum absorption perpendicular to the c axis and the deduced orientation of the OH dipoles perpendicular to c, Beran (1991) proposed a model where, under the assumption of vacant Al sites, OH defects are coordinated by two Al atoms, forming groups of face-sharing [Al2(OH)O8] double octahedra. According to Moon and Phillips (1991) the OH defects appear to be correlated to vacant Al sites as well as to the presence of Ti4+. In addition, two types of OH absorption bands were reported for hydrothermally treated synthetic sapphires by Kronenberg et al. (2000). The first type of bands observed at 3308, 3293, 3278, 3231, 3208, 3183, and 3163 cm−1 is characterized by narrow bands and strong pleochroism, the second type consists of a broad isotropic band centered at 3400 cm−1, resembling closely the OH bands of hydrothermally grown quartz crystals. Polarized IR spectra of synthetic high-P MgSiO3 akimotoite (Bolfan-Casanova et al. 2002), consist of five pleochroic OH absorption bands—three sharp strong bands at 3390, 3320, and 3300 cm−1 and two weak bands at 3260 and 3050 cm−1. Based on Paterson’s (1982) calibration the H2O content was calculated to 350 wt ppm. The bands at 3320 and 3300 cm−1 are strongly polarized perpendicular to the c axis. Similar to corundum, under the assumption of Mg vacancies, the pleochroic behavior is consistent with OH groups oriented nearly parallel to the plane of the shared face between two SiO6 octahedra. The two OH bands of corundum (at 3310 and 3230 cm−1) with the same pleochroic behavior occur at lower frequencies compared to those of MgSiO3 akimotoite. The band at 3390 cm−1 has maximum intensity

Structure of Hydrous Species Using Polarized IR Spectroscopy

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parallel to the c axis and is therefore consistent with OH groups pointing into a tetrahedral void of the close-packed oxygen sublattice. This band, however, has no analogue in the IR spectrum of corundum.

ACKNOWLEDGMENTS The authors wish to thank H. Keppler and J. Smyth for the invitation to contribute to the present MSA volume. S.D. Jacobsen, H. Keppler, and an anonymous referee helped to improve the quality of the manuscript. The topics of this paper were partly sponsored by the European Commission, Human Potential-Research Training Network, HPRN-CT-2000-0056.

REFERENCES Aines RD, Rossman GR (1984) The high temperature behavior of water and carbon dioxide in cordierite and beryl. Am Mineral 69:319-327 Aines RD, Rossman GR (1985) The high temperature behavior of trace hydrous components in silicate minerals. Am Mineral 70:1169-1179 Andrut M, Wildner M, Beran A (2002) The crystal chemistry of birefringent natural uvarovites. Part IV. OH defect incorporation mechanisms in non-cubic garnets derived from polarized IR spectroscopy. Eur J Mineral 14:1019-1026 Andrut M, Brandstätter F, Beran A (2003) Trage hydrogen zoning in diopside. Mineral Petrol 78:231-241 Andrut M, Wildner M, Rudowicz CZ (2004) Optical absorption spectroscopy in geosciences. Part II: Quantitative aspects of crystal fields. In: Spectroscopic Methods in Mineralogy. Beran A, Libowitzky E (eds) EMU Notes in Mineralogy, vol 6. Eötvös University Press, p 145-188 Asimov PD, Stein LC, Mosenfelder JL, Rossman GR (2006) Quantitative polarized infrared analysis of trace OH in populations of randomly oriented mineral grains. Am Mineral 91:278-284 Armbruster T, Kohler T, Libowitzky E, Friedrich A, Miletich R, Kunz M, Medenbach O, Gutzmer J (2001) Structure, compressibility, hydrogen bonding, and dehydration of the tetragonal Mn3+ hydrogarnet, henritermierite. Am Mineral 86:147-158 Aurisicchio C, Grubessi O, Zecchini P (1994) Infrared spectroscopy and crystal chemistry of the beryl group. Can Mineral 32:55-68 Baumer A, Ganteaume M, Klee WE (1985) Determination of OH ions in hydroxyfluorapatites by infrared spectroscopy. Bull Minéral 108:145-152 Bellamy LJ, Owen AJ (1969) A simple relationship between the infrared stretching frequencies and the hydrogen bond distances in crystals. Spectrochim Acta A25:329-333 Bellatreccia F, Della Ventura G, Ottolini L, Libowitzky E, Beran A (2005) The quantitative analysis of OH in vesuvianite: a polarized FTIR and SIMS study. Phys Chem Minerals 32:65-76 Belokoneva EL, Smirnitskaya YuYa, Tsirel’son VG (1993) Electron density distribution in topaz Al2(SiO4)(F,OH)2 as a result of accurate X-ray diffraction study. Zhurnal Neorganicheskoi Khimii 38:1346-1350 Belt FR (1967) Hydrothermal ruby: Infrared spectra and X-ray topography. J Appl Phys 38:2688-2689 Beny JM, Piriou B (1987) Vibrational spectra of single-crystal topaz. Phys Chem Minerals 15:148-154 Beran A (1976) Messung des Ultrarot-Pleochroismus von Mineralen. XIV. Der Pleochroismus der OHStreckfrequenz in Diopsid. Tschermaks Min Petr Mitt 23:79-85 Beran A (1991) Trace hydrogen in Verneuil-grown corundum and its colour varieties - an IR spectroscopic study. Eur J Mineral 3:971-975 Beran A (1999) Contribution of IR spectroscopy to the problem of water in the Earth’s mantle. In: Microscopic Properties and Processes in Minerals. Wright K, Catlow R (eds) NATO Science Series. Kluwer Acad. Publishers, p 523-538 Beran A (2002) Infrared spectroscopy of micas. Rev Mineral Geochem 46:351-369 Beran A, Libowitzky E (2006) Water in natural mantle minerals II: olivine, garnet and accessory minerals. Rev Mineral Geochem 62:169-191 Beran A, Rossman GR (2006) OH in naturally occurring corundum. Eur J Mineral 18, in press Beran A, Langer K, Andrut M (1993) Single crystal infrared spectra in the range of OH fundamentals of paragenetic garnet, omphacite and kyanite in an eklogitic mantle xenolith. Mineral Petrol 48:257-268 Beran A, Libowitzky E, Armbruster T (1996) A single-crystal infrared spectroscopic and X-ray diffraction study of untwinned San Benito perovskite containing OH groups. Can Mineral 34:803-809 Beran A, Giester G, Libowitzky E (1997) The hydrogen bond system in natrochalcite-type compounds - An FTIR spectroscopic study of the H3O2− unit. Mineral Petrol 61:223-235

50

Libowitzky & Beran

Bolfan-Casanova N (2005) Water in the Earth’s mantle. Mineral Mag 69:229-257 Bolfan-Casanova N, Keppler H, Rubie DC (2000) Water partitioning between nominally anhydrous minerals in the MgO-SiO2-H2O system up to 24 GPa: implications for the distribution of water in the Earth’s mantle. Earth Planet Sci Lett 182:209-221 Bolfan-Casanova N, Keppler H, Rubie DC (2002) Hydroxyl in MgSiO3 akimotoite: A polarized and highpressure IR study. Am Mineral 87:603-608 Blanchard M, Wright K, Gale JD (2005) A computer simulation study of OH defects in Mg2SiO4 and Mg2GeO4 spinels. Phys Chem Minerals 32:585-593 Bradbury SE, Williams Q (2003) Contrasting bonding behavior of two hydroxyl-bearing metamorphic minerals under pressure: Clinozoisite and topaz. Am Mineral 88:1460-1470 Braithwaite JS, Wright K, Catlow CRA (2003) A theoretical study of the energetics and IR frequencies of hydroxyl defects in forsterite. J Geophys Res 108(B6):2284 Brigatti MF, Davoli P (1990) Crystal structure refinement of 1M plutonic biotites. Am Mineral 75:305-313 Brown ID (1981) The bond-valence method: An empirical approach to chemical structure and bonding. In: Structure and Bonding in Crystals, vol II. O’Keeffe M, Navrotsky A (eds), Academic Press, p 1-30 Burns RG, Strens RGJ (1966) Infrared study of the hydroxyl bands in clinoamphiboles. Science 153:890-892 Buse K, Breer S, Peithmann K, Kapphan S, Gao M, Krätzig E (1997) Origin of thermal fixing in photorefractive lithium niobate crystals. Phys Rev B56:1225-1235 Caruba R, Baumer A, Ganteaume M, Iacconi P (1985) An experimental study of hydroxyl groups and water in synthetic and natural zircons: a model of the metamict state. Am Mineral 70:1224-1231 Dixon JE, Leist L, Langmuir C, Schilling J-G (2002) Recycled dehydrated lithosphere observed in plumeinfluenced mid-ocean-ridge basalt. Nature 420:385-389 Dobson DP, Meredith PG, Boon SA (2002) Simulation of subduction zone seismicity by dehydration of serpentine. Science 298:1407-1410 Dumas P, Tobin MJ (2003) A bright source for infrared microspectroscopy: synchrotron radiation. Spectr Eur 15/6:17-23 Emsley J, Jones D, Lucas J (1981) Detecting and measuring strong hydrogen bonds: recent developments. Rev Inorg Chem 3:105-140 Fadini A, Schnepel F-M (1989) Vibrational spectroscopy. Ellis Horwood Limited Gatta GD, Nestola F, Bromiley GD, Mattauch S (2006) The real topological configuration of the extraframework content in alkali-poor beryl: A multi-methodological study. Am Mineral 91:29-34 Geiger, CA (2004) An introduction to spectroscopic methods in the mineral sciences and geochemistry. In: Spectroscopic Methods in Mineralogy. Beran A, Libowitzky E (eds) EMU Notes in Mineralogy, vol 6. Eötvös University Press, p 1-42 Geiger CA, Stahl A, Rossman GR (2000) Single-crystal IR- and UV/VIS- spectroscopic measurements on transition-metal-bearing pyrope: the incorporation of hydroxide in garnet. Eur J Mineral 12:259-271 Gibbs GV, Rosso KM, Cox DF, Boisen Jr MB (2003) A physical basis for Pauling’s definition of bond strength. Phys Chem Minerals 30:317-320 Griffiths PR, de Haseth JA (1986) Fourier Transform Infrared Spectrometry (Chem Anal 83). John Wiley & Sons Griggs D (1967) Hydrolytic weakening of quartz and other silicates. Geophys J R Astr Soc, 14:19-31 Groat LA, Hawthorne FC, Ercit TS (1992) The chemistry of vesuvianite. Can Mineral 30:19-48 Gebert W, Zemann J (1965) Messung des Ultrarot-Pleochroismus von Mineralen. III. Der Pleochroismus der OH-Streckfrequenz in Topas. N Jb Miner Mh 1965:380-384 Hall D, Wood MK (1985): A molecular-packing analysis of the crystal structures of ice. Acta Cryst B41:169172 Hammer VMF, Libowitzky E, Rossman GR (1998) Single-crystal IR spectroscopy of very strong hydrogen bonds in pectolite, NaCa2[Si3O8(OH)], and serandite, NaMn2[Si3O8(OH)]. Am Mineral 83:569-576 Hirth G, Kohlstedt DL (1996) Water in the oceanic upper mantle: Implications for rheology, melt extraction and the evolution of the lithosphere. Earth Planet Sci Lett 144:93-108 Hu Meisheng, Wenk H-R, Sinitsyna D (1992) Microstructures in natural perovskites. Am Mineral 77:359-373 Ingrin J, Blanchard M (2006) Diffusion of hydrogen in minerals. Rev Mineral Geochem 62:291-320 Ingrin J, Latrous K, Doukhan J-C, Doukhan N (1989) Water in diopside: an electron microscopy and infrared spectroscopy study. Eur J Mineral 1:327-431 Jacobsen SD (2006) Effect of water on the equation of state of nominally anhydrous minerals. Rev Mineral Geochem 62:321-342 Jacobsen SD, van der Lee S (Eds) (2006) Earth’s Deep Water Cycle. Am Geophys Union Monogr Ser, in press Jacobsen SD, Demouchy S, Frost DJ, Boffa Ballaran T, Kung J (2005) A systematic study of OH in hydrous wadsleyite from polarized FTIR spectroscopy and single-crystal X-ray diffraction: Oxygen sites for hydrogen storage in Earth’s interior. Am Mineral 90:61-70

Structure of Hydrous Species Using Polarized IR Spectroscopy

51

Jung H, Green HW (2004) Experimental faulting of serpentinite during dehydration: Implications for earthquakes, seismic low-velocity zones, and anomalous hypocenter distributions in subduction zones. Int Geol Rev 46:1089-1102 Karato S (1990) The role of hydrogen in the electrical conductivity of the upper mantle. Nature 347:272-273 Khisina NR, Wirth R, Andrut M, Ukhanov AV (2001) Extrinsic and intrinsic mode of hydrogen occurrence in natural olivines: FTIR and TEM investigation. Phys Chem Minerals 28:291-301 Khomenko VM, Langer K, Beran A, Koch-Müller M, Fehr T (1994): Titanium substitution and OH-bearing defects in hydrothermally grown pyrope crystals. Phys Chem Minerals 20:483-488 Kleppe AK, Jephcoat AP, Welch MD (2003) The effect of pressure upon hydrogen bonding in chlorite: A Raman spectroscopic study of clinochlore to 26.5 GPa. Am Mineral 88:567-573 Koch-Müller M, Abs-Wurmbach I, Rhede D, Kahlenberg V, Matsyuk, S (2005) Hydration and dehydration of omphacite. Berichte der Deutschen Mineralogischen Ges, Beih z Eur J Mineral 17, No 1:69 Kohlstedt DL, Mackwell SJ (1998) Diffusion of hydrogen and intrinsic point defects in olivine. Z Phys Chem 207:147-162 Kolesov AB, Geiger CA (2000a) A single-crystal Raman study of the orientation and vibrational states of molecular water in beryl. Phys Chem Minerals 27:557-564 Kolesov AB, Geiger CA (2000b) Cordierite II: The role of CO2 and H2O. Am Mineral 85:1265-1274 Kronenberg AK, Castaing J, Mitchell TE, Kirby SH (2000) Hydrogen in α-Al2O3 and water weakening of sapphire and alumina ceramics between 600 and 1000 °C – I. Infrared characterization of defects. Acta Mater 48:1481-1494 Kurka A (2002) IR-spektroskopische Untersuchungen im OH-Streckschwingungsbereich von Vesuvian und Gossular. MSc thesis, University of Vienna, Austria Lager GA, Armbruster T, Faber F (1987) Neutron and X-ray diffraction study of hydrogarnet, Ca3Al2(O4H4)3. Am Mineral 72:756-765 Lager GA, Qianyen X, Ross FK, Rossman GR, Armbruster T, Rotella FJ, Schultz AJ (1999) Hydrogen-atom positions in P4/nnc vesuvianite. Can Mineral 37:763-768 Libowitzky E (1999) Correlation of O-H stretching frequencies and O-H···O hydrogen bond lengths in minerals. Mh Chemie 130:1047-1059 Libowitzky E, Beran A (1995) OH-defects in forsterite. Phys Chem Minerals 22:387-392 Libowitzky E, Beran A (2004) IR spectroscopic characterisation of hydrous species in minerals. In: Spectroscopic methods in mineralogy. Beran A, Libowitzky E (eds) EMU Notes in Mineralogy, Vol 6. Eötvös University Press, p 227-279 Libowitzky E, Giester G (2003) Washing soda (natron), Na2CO3 · 10H2O, revised: Crystal structures at low and ambient temperatures. Mineral Petrol 77:177-195 Libowitzky E, Rossman GR (1996) Principles of quantitative absorbance measurements in anisotropic crystals. Phys Chem Minerals 23:319-327 Libowitzky E, Rossman GR (1997) An IR absorption calibration for water in minerals. Am Mineral 82:11111115 Mackwell SJ, Kohlstedt DL, Paterson MS (1985) The role of water in the deformation of olivine single crystals. J Geophys Res 90B:11319-11333 Meade C, Reffner JA, Ito E (1994) Synchrotron infrared absorbance measurements of hydrogen in MgSiO3 perovskite. Science 264:1558-1560 van der Meijde M, Marone F, Giardini D, van der Lee S (2003) Seismic evidence for water deep in Earth’s upper mantle. Science 300:1556-1558 Mierdel K, Keppler H (2004) The temperature dependence of water solubility in enstatite. Contr Mineral Petrol 148:305-311 Mikenda W (1986) Stretching frequency versus bond distance correlation of O-D(H)···Y (Y = N, O, S, Se, Cl, Br, I) hydrogen bonds in solid hydrates. J Mol Struct 147:1-15 Moon AR, Phillips MR (1991) Defect clustering in H,Ti:α-Al2O3. J Phys Chem Solids 52:1087-1099 Nagai T, Hattori T, Yamanaka T (2000) Compression mechanism of brucite: An investigation by structural refinement under pressure. Am Mineral 85:760-764 Nakamoto K (1977) Infrared and Raman spectra of inorganic and coordination compounds. Wiley, New York Nakamoto K, Margoshes M, Rundle RE (1955) Stretching frequencies as a function of distances in hydrogen bonds. J Am Chem Soc 77:6480-6486 Novak A (1974) Hydrogen bonding in solids: Correlation of spectroscopic and crystallographic data. Struct Bond 18:177-216 Ohtani E (2005) Water in the mantle. Elements 1:25-30 Parise JB, Cuff C, Moore FH (1980) A neutron diffraction study of topaz: evidence for lower symmetry. Min Mag 43:943-944 Paterson MS (1982) The determination of hydroxyl by infrared absorption in quartz, silicate glasses and similar materials. Bull Minéral 105:20-29

52

Libowitzky & Beran

Pauling L (1960) The Nature of the Chemical Bond. Cornell Univ. Press Ribbe PH, Rosenberg PE (1971) Optical and X-ray determinative methods for fluorine in topaz. Am Mineral 56:1812-1821 Ross NL, Gibbs GV, Rosso KM (2003) Potential docking sites and positions of hydrogen in high-pressure silicates. Am Mineral 88:1452-1459 Rossman GR (1988) Vibrational spectroscopy of hydrous components. Rev Mineral 18:193-206 Rossman GR (2006) Analytical methods for measuring water in nominally anhydrous minerals. Rev Mineral Geochem 62:1-28 Rossman GR, Aines RD (1991) The hydrous components in garnets: Grossular-hydrogrossular. Am Mineral 76: 1153-1164 Rossman GR, Smyth JR (1990) Hydroxyl contents of accessory minerals in mantle eclogites and related rocks. Am Mineral 75: 775-780 Rothbauer R (1971) Untersuchung eines 2M1-Muskovits mit Neutronenstrahlen. N Jb Mineral Mh 1971:143154 Rowbotham G, Farmer VC (1973) The effect of “A” site occupancy upon hydroxyl stretching frequency in clinoamphiboles. Contrib Mineral Petrol 38:147-149 Ryskin YI (1974) The vibrations of protons in minerals: hydroxyl, water and ammonium. In: The Infrared Spectra of Minerals. Farmer VC (ed) Mineralogical Society, p 137 -181 Sacerdoti M, Passaglia E (1985) The crystal structure of katoite and implications within the hydrogrossular group of minerals. Bull Minéral 108:1-8 Schiffer J, Intenzo M, Hayward P, Calabrese C (1976) The FS correlation and a systematic analysis of the stretching absorptions of water in the condensed state. J Chem Phys 64:3014-3020 Serratosa JM, Bradley WF (1958) Determination of the orientation of OH bond axes in layer silicates by infrared absorption. J Phys Chem 62:1164-1167 Skogby H (2006) Water in natural mantle minerals I: pyroxenes. Rev Mineral Geochem 62:155-167 Skogby H, Rossman GR (1989) OH− in pyroxene: An experimental study of incorporation mechanisms and stability. Am Mineral 74:1059-1069 Skogby H, Bell DR, Rossman GR (1990) Hydroxide in pyroxene: Variations in the natural environment. Am Mineral 75:764-774 Smith CP, Kammerling RC, Keller AS, Peretti A, Scarratt KV, Khoa ND, Repetto S (1995) Sapphires from southern Vietnam. Gems Gemology 31:168-186 Smyth JR (1987) β-Mg2SiO4: A potential host for water in the mantle? Am Mineral 72:1051-1055 Smyth JR, Frost DJ (2002) The effect of water on the 410-km discontinuity: An experimental study. Geophys Res Lett 29:10.1029/2001GL014418 Smyth JR, Bell DR, Rossman GR (1991) Incorporation of hydroxyl in upper-mantle clinopyroxenes. Nature 351:732-735 Smyth JR, Holl CM, Frost DJ, Jacobsen SD, Langenhorst F, McCammon CA (2003) Structural systematics of hydrous ringwoodite and water in Earth’s interior. Am Mineral 88:1402-1407 Strens RGJ (1974) The common chain, ribbon, and ring silicates. In: The Infrared Spectra of Minerals. Farmer VC (ed) Mineralogical Society, p 305-330 Szalay V, Kovács L, Wöhlecke M, Libowitzky E (2002) Stretching potential and equilibrium length of the OH bond in solids. Chem Phys Lett 354:56-61 Thompson AB (1992) Water in the Earth’s upper mantle. Nature 358:295-302 Tillmanns E, Zemann J (1965) Messung des Ultrarot-Pleochroismus von Mineralen. I. Der Pleochroismus der OH-Streckfrequenz in Azurit. N Jb Miner Mh 1965:228-231 Tsuboi M (1950) On the position of the hydrogen atoms in the crystal structure of muscovite, as revealed by the infra-red absorption study. Bull Chem Soc Japan 23:83-88 Urusov VS (1992) A geometric model of deviations from Vegard’s rule. J Solid State Chem 98:223-236 Wieczorek A, Libowitzky E, Beran A (2004) A model for the OH defect incorporation in kyanite based on polarised IR spectroscopic investigations. Schweiz Min Petr Mitt 84:333-343 Wilkins RWT, Sabine W (1973) Water content of some nominally anhydrous silicates. Am Mineral 58:508-516 Winkler B (1996) The dynamics of H2O in minerals. Phys Chem Minerals 23:310-318 Wright K (2006) Atomistic models of OH defects in nominally anhydrous minerals. Rev Mineral Geochem 62:67-83 Wunder B, Andrut M, Wirth R (1999) High-pressure synthesis and properties of OH-rich topaz. Eur J Mineral 11:803-813 Zemann J, Zobetz, E, Heger G, Völlenkle H (1979) Strukturbestimmung eines OH-reichen Topases. Anz Österr Akad Wiss, Math-Naturwiss Kl 116:145-147

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Reviews in Mineralogy & Geochemistry Vol. 62, pp. 53-66, 2006 Copyright © Mineralogical Society of America

Structural Studies of OH in Nominally Anhydrous Minerals Using NMR Simon C. Kohn Department of Earth Sciences University of Bristol Bristol, BS8 1RJ, United Kingdom e-mail: [email protected]

INTRODUCTION Nuclear magnetic resonance (NMR) is a technique that is used very widely throughout science and medicine. There are approaching 20 journals devoted exclusively to NMR and magnetic resonance imaging, and well in excess of 10,000 papers published annually which involve this family of techniques. It has found extraordinarily diverse applications, and its applications to inorganic solids such as minerals represent a very small part of NMR as a whole. There had been occasional NMR studies of minerals since the discovery of the NMR effect in 1946, but until the early 1980s, the large width of NMR resonances in the solid state precluded widespread application of the technique. The most important factor in the application of NMR to minerals was the development of magic angle spinning (MAS), a technique for narrowing lines in solid state NMR. More recently the development of more complex sample spinning arrangements and multiple pulse methods together with the availability of ever higher magnetic fields opens up more and more possibilities for NMR in mineralogy and geochemistry. NMR is far too diverse and complex for the whole subject to be covered here, so this review will be tightly focused on aspects which relate to understanding the structural role of water in nominally anhydrous minerals. Two review papers (Kirkpatrick 1988; Stebbins 1988) in the “Reviews in Mineralogy” volume on Spectroscopic Methods in Mineralogy and Geology provide an excellent starting point for Earth scientists wishing to learn more about NMR in general. More recent reviews aimed at Earth scientists include those by Fechtelkord (2004) and Kohn (2004), more general reviews of NMR of inorganic solids include Engelhardt and Michel (1987) and MacKenzie and Smith (2002). The web page maintained by J.P. Hornak (http://www.cis.rit.edu/htbooks/nmr/) is also an extremely good resource. NMR is a multinuclear technique, which means that a sample containing several different NMR active nuclei can be studied independently using NMR of each of the nuclei. Thus, for example, the zeolite natrolite with a composition Na2[Al2Si3O10]·2H2O could be studied using 23Na, 27Al, 29Si and 1H NMR, and each would provide different and complementary information. If suitably isotopically enriched samples were available, 17O and 2H spectra would provide additional distinct information. However, if we are interested in the structural aspects of H dissolved in nominally anhydrous minerals the situation is rather different. In this case the concentration of water is usually too low to have a significant influence on the local environments of the major components. For example the H/Si ratio in enstatite containing 200 ppm H2O is 2.2 × 10−3. 1H NMR is therefore the most promising NMR technique for NAMs except in particular circumstances which will be described later. There are a number of problems that have limited the use of NMR in studying NAMs. In this chapter the principles of solid state NMR will be briefly outlined, then both the benefits and problems of 1H MAS NMR applied to the problem of water in NAMs will be discussed. A review of published studies on 1529-6466/06/0062-0003$05.00

DOI: 10.2138/rmg.2006.62.3

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Kohn

1 H MAS NMR will follow and finally other NMR methods, such as static 1H NMR and use of other NMR active nuclei, and prospects for future developments will be discussed.

PRINCIPLES OF SOLID STATE NMR The references given in the introduction give a much more complete introduction to NMR spectroscopy, so the following description should only be considered to be a highly simplified outline of the most important concepts which are relevant for NMR studies of NAMs. Some nuclei (depending on the numbers of protons and neutrons in the nucleus) have a quantized property known as spin. The NMR effect is based on the nuclear spin, and can be conceptualized by either classical or quantum mechanical descriptions. In the quantum mechanical view the spin energy levels become non-degenerate in an external magnetic field, and the difference between the energy levels is given by ∆E =

µB0 I

(1)

where B0 is the applied magnetic field and I is the nuclear spin. The nuclear magnetic moment, μ, is given by µ=

γ hI 2π

(2 )

where γ is the magnetogyric ratio (a constant which is specific for each nucleus) and h is Planck’s constant. The energy differences, ΔE, are typically in the radio frequency range for the magnetic fields used in NMR spectroscopy, so transitions between the energy levels can be stimulated by irradiating with the appropriate tuned RF frequency. The relaxation of the nuclear spins back to equilibrium is then measured by emission of the same RF frequency. In the classical description, the nuclei are considered to behave like tiny bar magnets in a magnetic field. At equilibrium, the torque exerted on the magnetic moments causes them to precess around the direction of the B0 magnetic field with a frequency known as the Larmor frequency; this leads to a net bulk magnetization along the same direction as the applied B0 magnetic field. Applying a relatively small additional B1 magnetic field, by irradiation with a radio frequency field oscillating at close to the Larmor frequency, applies an additional torque to this bulk magnetization. The result is that the bulk magnetization moves away from the direction of the B0 field. When the B1 field is turned off, two things happen i) local differences in the B0 field mean that individual spins experience slightly different local fields. Hence when the magnetization is in the transverse plane phase coherence is lost as the differing Larmor precession frequencies cause the spins to fan out. The timescale of this process is known as the spin-spin relaxation rate, with a time constant T2. ii) the bulk magnetization is built up to re-attain equilibrium. This process is known as spin-lattice relaxation, with a time constant T1. Spin lattice relaxation is usually much slower than spin-spin relaxation in solids, i.e., T1 > T2. NMR of liquids usually gives narrow resonances such that chemically inequivalent sites in a molecule can be identified. The resonances are narrow, because the rate of the tumbling motion of the molecules is fast compared with the strength of all the line broadening interactions, so anisotropy in the NMR interactions, and coupling between nuclei are averaged. In solids, however, there are a range of interactions that cause line broadening, and obscure most of the useful information in the spectra. Most NMR measurements that have been reported in the mineralogical literature have, therefore, used the magic angle spinning (MAS) technique. In this technique the sample is packed into a ceramic rotor (typically 2-7 mm in diameter), and a polymer or ceramic cap is inserted into the end of the rotor to contain the

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sample. The rotor is then placed into an NMR probe and physically spun at very high speeds using a compressed gas system. The angle between the axis of rotation of the sample and the direction of the B0 field is set to be 54.7°, the magic angle. At this angle, the term (3cos2θ−1), which appears in many of the line broadening functions, becomes zero. Magic angle spinning at sufficiently high speeds narrows most of the possible broadening interactions. A single NMR resonance can potentially provide information from its position (frequency or chemical shift), width and intensity. The nature of the information contained in these three parameters will vary from nucleus to nucleus, but the case of 1H MAS NMR is reviewed below.

Positions of 1H MAS NMR resonances

Isotropic chemical shift (ppm)

One of the most useful features of NMR is that small differences in the chemical environment around a nucleus result in a slightly modified magnetic field, and hence the position of a 1H NMR peak. These differences in resonance frequency are expressed as a chemical shift with units of ppm and reflect the chemical environment of the hydrogen. 1H chemical shifts can be measured most easily using fast magic angle spinning. In nominally anhydrous silicate minerals one would expect H to always be strongly bonded to an adjacent oxygen, and several studies on hydroxyl containing minerals and other materials have shown that there is a strong correlation between 1H chemical shift and O-H distance, rOH (Brunner and 18 Sternberg 1998). Hydroxyl groups in silicates may be oriented towards other oxygens in the 16 structure and form hydrogen bonds of variable 14 strength. It has been shown that the O-H..O distance (rO..O), which reflects the strength 12 of a hydrogen bond, is also correlated with 10 both rOH and 1H chemical shift (Eckert et al. 1988). The correlation between rO..O and 1H 8 shift shown in Figure 1 is particularly useful as 6 rO..O is measurable from X-ray diffraction data, whereas rOH requires neutron diffraction data. If 4 several peaks are observable in a spectrum, rO..O can be deduced for each hydrogen environment. 2 These correlations between 1H NMR shift 0 and structure are analogous to those for O-H stretching frequency (Libowitzky 1999), so in 2.4 2.5 2.6 2.7 2.8 2.8 3.0 3.1 principle NMR can give the same structural d(O-H..O)/angstroms information as FTIR spectroscopy, even though 1 Figure 1. H NMR chemical shift as a function the latter is much more frequently used. The of O-H...O distance for various crystalline key to extracting this structural information is compounds (after Eckert et al. 1988). that the width of the resonances must be small 1 compared with the shift range of H in NAMs.

Widths of 1H MAS NMR resonances Magic angle spinning dramatically reduces the widths of 1H NMR spectra of solids because it averages homonuclear and heteronuclear dipolar coupling between protons and also chemical shift anisotropy. Nonetheless a variety of mechanisms are involved in determining the linewidth of a 1H MAS NMR resonance, and these are important in interpreting the spectra.

i) Chemical shift dispersion. This is the range of chemical shifts present, and reflects the actual distribution of H environments in the sample. The linewidth in ppm resulting from chemical shift dispersion is independent of magnetic field.

Kohn

56 ii) Residual

iii) Dipolar coupling to quadrupolar

2000 1800

Full width at half maximum (Hz)

homonuclear dipolar broadening. If protons are close together in the structure of a material, magic angle spinning at conventional speeds of 5-15 kHz is insufficient to completely remove the dipolar coupling. This effect can be seen in Figure 2 where the residual linewidth decreases as a function of spinning speed for two minerals with high H-density. For NAMs, dipolar coupling should be completely removed by magic angle spinning unless hydrogens are clustered, for example in water molecules, as pairs of H charge balancing a divalent cation vacancy or as a hydrogarnet substitution. If dipolar coupling is not completely removed, spinning sidebands will be observed.

1600

Datolite

1400 1200 1000

Pyrophyllite

800 600 400 200 1

2

6 4 5 3 Spinning speed (Hz)

7

8

Figure 2. The full width at half maximum (FWHM) of 1H MAS spectra of datolite and pyrophyllite as a function of spinning speed (after Yesinowski et al. 1988). The decrease in width with increasing spinning speed for these minerals indicates that the homonuclear dipolar coupling is incompletely averaged by the spinning speeds that were available at that time.

nuclei. Quadrupolar nuclei (those with I > ½) are subject to additional line broadening mechanisms compared with dipolar nuclei, and these are not completely averaged by MAS. However, quadrupolar interactions are reduced relative to the chemical shift interaction at higher magnetic fields. Protons which undergo strong dipolar coupling to quadrupolar nuclei can be broadened, but in this case the linewidth (in Hz) will also decrease with increasing magnetic field.

iv) Paramagnetic samples. Coupling between the nuclear dipole and the much larger

dipole of unpaired electrons in paramagnetic samples (such as iron-bearing, natural mantle minerals) can cause enormous broadening of NMR signals. Although there are some circumstances where NMR of paramagnetic samples is possible, NMR of natural iron-bearing samples does not generally give useable signals.

v) Motional narrowing. Protons that undergo rapid isotropic motion (such as those in macroscopic fluid inclusions) will give very narrow lines. Even certain restricted motions, such as rotation of water molecules about a single axis can be effective in narrowing lines.

Intensity of 1H MAS NMR resonances In favorable circumstances, and if the NMR experiment is performed using the correct conditions, the area of an NMR resonance is proportional to the number of resonating nuclei. Thus if a 1H spectrum consists of several peaks, the relative areas of the peaks can be used directly to determine the relative abundances of the different H environments which give rise to the different peaks. Furthermore if the area of the resonance for a standard with a known concentration of H is measured, a comparison with the area of the resonance for an unknown will yield the absolute number of H nuclei in the sample, as long as the masses of the standard and sample are taken into account. Obviously the standard that is used should be very well characterized, as any uncertainty in the water concentration of the standard will be reflected in the uncertainty in the water concentrations of the unknowns. The standard should also have a similar magnetic

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susceptibility to the samples, and have a relatively low water concentration if possible. These two factors are not thought to be critical compared with the other difficulties outlined below, but more work on the effect of using different standard materials should probably be undertaken. In practice there are a number of reasons why the observed NMR signal may not be fully quantitative. i) Resonance may be broadened, by residual dipolar coupling or interaction with paramagnetic centers, to the point where it is unobservable. A broad line in an NMR spectrum (frequency domain) corresponds to a rapidly decaying signal in the time domain. Even with fast digitization, the signal of a broad line may only correspond to a few data points, and these may be obscured by ringdown of the probe, which occurs for a few μs after the pulse. ii) The relaxation delay between pulses may not be sufficiently long. To obtain quantitative spectra it is crucial to allow a sufficiently long time between pulses to allow complete spin-lattice relaxation for all the nuclei. Materials with high concentrations of hydrogen usually have quite short values of T1, and since the signal is extremely large, only a few repetitions of the pulse and acquisition cycle are required. In the case of dilute hydrogen, T1 can be much longer, and since many more cycles are required to get an acceptable signal, it can be difficult, but nonetheless crucial to ensure that the relaxation delay is sufficiently long. This can be done either by a rigorous T1 determination, or simply by increasing the relaxation delay (while keeping other parameters constant) until the signal intensity per pulse becomes constant. iii) The B1 field may not be homogeneous. The B1 field is the magnetic field generated by applying the pulse of RF to the NMR coil. Over a long distance this cannot be homogeneous, especially at the ends of the coil, and the best that can be aimed for is that it is homogeneous over the sample volume. The true homogeneity can be tested by checking whether all parts of the sample have the same 90° pulse length, but for practical purposes the effect of B1 inhomogeneity on signal size can be tested by quantifying spectra collected with vary amounts of sample in a rotor. Figure 3 shows the integrated signal (after background subtraction) plotted as a function of sample mass for the specialized 1H MAS NMR probe used in our laboratory. This probe is based around a DOTY Si3N4 stator, and gives an excellent linear relationship between sample mass and integrated signal. 350

Figure 3. The integrated area of 1H MAS NMR spectra of a hydrous glass standard plotted as a function of sample mass within a 5 mm MAS rotor. The linear plot illustrates the quantitative nature of NMR, and that for this NMR probe the B1 field is effectively constant over the sample volume.

Area of spectrum (arbitrary units)

300 250 200 150 100 50 0 0

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Mass of sample (mg)

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Kohn APPLICATION OF 1H MAS NMR TO NOMINALLY ANHYDROUS MINERALS

Attractive features of 1H MAS NMR for studies of NAMs There are a number of characteristics of 1H NMR in particular, which, in principle, can be exploited to determine both the concentration and environment of H in NAMs. 1H has the highest magnetogyric ratio of all nuclei and is nearly 100% abundant. Furthermore, it has a spin of ½, and is therefore generally free of all the complications and line-broadening experienced by quadrupolar nuclei. Therefore, of all the NMR active nuclei, it is the most suitable for measurement at low concentrations, and should be a valuable tool for studying the concentration and environment of low concentrations of H in NAMs. Concentrations of less than 1 ppm H2O can be detected in favorable circumstances. As described earlier, NMR is an intrinsically quantitative technique, so as long as one suitable standard is available, no further calibration for different bulk compositions is required. This a major advantage over FTIR. Of course, NMR is also an element specific technique, so all the intensity in a 1H NMR spectrum corresponds to H in the sample. This contrasts with vibrational spectroscopy, where a peak in a spectrum could be result from an OH vibration, or a combination of other vibrations of the structure. NMR will also provide information on all dissolved H, unlike FTIR, which is only sensitive to O-H species. This point could be important if it turns out that other species such as Si-H or other hydrides, organic molecules or H2 are significant for the H budget in NAMs, as proposed by Freund et al. (2002). The chemical shift for gaseous H2 is 4.45 ppm (Raynes 1977) and H2 intercalated into different materials shows a significant range of shifts, so molecular H2 would not necessarily be easy to distinguish from H2O molecules or OH groups on the basis of chemical shift. There is a well known correlation between 1H MAS NMR shift and structure and information on H-H distances and clustering can be obtained by exploiting the dipole-dipole interactions between H nuclei, so abundant structural information is accessible.

Problems and difficulties in applying 1H MAS NMR to NAMs Despite all the potentially attractive features of 1H NMR for studying NAMs, there have been very few published studies because of the practical difficulties of working with the small signal from the low concentration of H. The specific problems are

i) The NMR rotor, the caps on the rotor and the stator (the assembly in which the

rotor is contained) can be made of H-containing materials or contain adsorbed water. This problem can be minimized by drying all components before use, by using caps made from Kel-f (a fluorine based polymer) and an NMR probe with a stator containing minimal H containing materials, and by avoiding the use of porous materials in the stator (as they are difficult to dry). Even if all these precautions are taken, it is impossible to completely eliminate background signals. A spectrum of an empty NMR rotor should therefore be taken, and this spectrum should be subtracted from all sample and standard spectra prior to further processing and integration. An example from the work of Johnson and Rossman (2003) is shown in Figure 4. The bottom spectrum (A) is that of an empty rotor, while spectrum (B) is for anhydrous labradorite powder, and (C) is an uncorrected spectrum of an OH bearing sample of andesine. The top spectrum (D) is the final background subtracted data, obtained by subtracting (B) from (C), and even with careful background subtraction there are still artifacts in the spectrum around 1.5 ppm. The difficulty of performing an effective background subtraction is compounded for experimental equipment with a large proton background and samples with small signals. In the case of samples containing paramagnetic ions a normal background subtraction is not applicable, because the paramagnetic susceptibility of the sample influences the local B0 and broadens the background (M. Fechtelkord, personal communication).

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Figure 4. Illustration of the procedure for subtracting a background signal from the 1H MAS NMR spectrum of a nominally anhydrous mineral (after Johnson and Rossman 2003).

ii) The signal from H dissolved in NAMs can easily be confused with signals from

small concentrations of contaminant H-bearing phases such as melt inclusions, fluid inclusions and hydrous minerals as well as adsorbed water on the surfaces of grains and trapped water on grain boundaries. In our experience it is this problem which is the most limiting for NMR studies of NAMs. It is hard to obtain sufficiently large samples, either natural or experimentally produced, which are free of hydrous impurities. It is well known that narrow peaks from surface contamination by organic species have a shift of around 1-2 ppm, and thus overlap with the expected shift from non hydrogen bonded OH. Great care has therefore to be taken in interpreting any features in this region of the spectrum.

iii) The spin lattice relaxation time, T1, for 1H can be long. This can be a particular problem for materials where protons are distant from each other and for chemical pure systems where there are no paramagnetic cations to help relaxation.

iv) NMR studies are generally restricted to samples that are free of significant

concentrations of paramagnetic ions. Even if all the other problems can be overcome, 1 H NMR is probably only suitable for studies of iron-free synthetic analogues of mantle minerals rather than natural mantle mineral samples.

1

H MAS NMR studies of orthopyroxene

A 1H MAS NMR spectrum of synthetic enstatite prepared at 1.5 GPa and 1050 °C was presented by Kohn (1996). The spectrum consisted of a broad feature, with narrower peaks at 7.9 and 5.9 ppm, together with peaks which were interpreted as either fluid inclusions or organic surface contaminants (Fig. 5). The shift and relative intensities of the two narrow peaks are entirely consistent with FTIR spectra of pure enstatite which has peaks at 3062 and 3360 cm−1 (e.g., Rauch and Keppler 2002; Stalder and Skogby 2002; Grant et al. 2006). The absolute concentrations of water in the sample was calculated to be 870 ppm H2O. However, it was noted that of the broad components could be related to hydrous mineral or glass impurities or water molecules at grain boundaries and that if only the narrow components are considered the solubility in enstatite would be 240 ppm H2O. Keppler and Rauch (2000)

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suggested that the solubilities in the study of Kohn (1996) could be overestimated because of hydrous inclusions, growth defects and surface water. Subsequent studies in our laboratory (Najorka and Kohn, in prep) suggest that the lower figure of 240 ppm is close to the true solubility and that the broad features in the spectra result mainly from a hydrous phase formed upon quenching the coexisting aqueous fluid at the end of the synthesis run. 1

H MAS NMR studies of clinopyroxene

Three clinopyroxene samples in the system diopside-CaTs were studied by Kohn (1996). These samples had much higher dissolved water concentrations than the enstatite sample, with the total spectral areas corresponding to 2615-4100 ppm H2O. If only the narrow parts of the spectra are considered to be dissolved water the solubilities are in the range 1160-2430 ppm for synthesis conditions of 1.5 GPa and 1000-1150 °C. These values are comparable with water concentrations in other aluminous Figure 5. The 1H MAS NMR centrebands of clinopyroxenes determined using FTIR experimentally synthesised forsterite (Fo-10), (Skogby et al. 1990; Smyth et al. 1991; enstatite (En-2) and three clinopyroxenes on the Di-CaTs join (Dicat1-9, Dicat2-9 and Dicat3-1) Bromiley and Keppler 2004). The spectra (Kohn 1996). The prominent peak at 4.7 ppm is have different shapes, with the main peak due to water in fluid inclusions and those at 1.3 becoming broader and more asymmetric and 0.8 ppm are probably due to contamination by with increasing CaTs component (Fig. 5). organic compounds. The spectra also have prominent spinning sidebands (Fig. 6), suggesting that some of the H has strong H-H dipolar coupling because of clustering of hydrogen in species such as hydrogarnet substitution or included water molecules. 1

H MAS NMR studies of olivine

Kohn (1996) presented a 1H MAS NMR spectrum of a synthetic forsterite sample. This spectrum contained a broad resonance at 4.3 ppm, together with a small feature at 6.9 ppm and an intense peak at 1 ppm (Fig. 5). Although peaks from surface contamination are known to be near 1 ppm the peak in the olivine sample was anomalously large, so it was interpreted as being a possible peak for dissolved OH in the forsterite. The total area of the spectrum corresponded to 1790 ppm and even the narrow part alone give a value of 560 ppm. These values were much higher than expected, and should be treated with caution. A much improved spectrum of forsterite synthesised at 2.1 GPa and 1100 °C, and in equilibrium with a small amount of enstatite to buffer silica activity, was subsequently presented in a published abstract (Kohn 1999). This spectrum contained a broad resonance at 1 ppm (in addition to the background signal), and a distinct peak at 6.7 ppm, corresponding to relatively strongly hydrogen bonded OH. Although there is still some ambiguity in the intensity of features around 1 ppm, because of adsorbed organic contaminants, this spectrum does not have the very broad feature that dominates the earlier spectrum. Quantification suggests a solubility of 400-500 ppm H2O. This is still higher than the solubility determined on different samples using FTIR, although it is closer to solubili-

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ties calculated with the FTIR calibration of Bell et al. (2003) than those calculated using alternative calibrations. More work needs to be done to resolve this important issue, although one possibility is that elevated nonequilibrium concentrations of OH can be incorporated in forsterite under certain crystal growth conditions (Lemaire et al. 2004). The peak at 6.7 ppm is an interesting feature of the spectrum. The correlation shown in Figure 1 would suggest that this corresponds to an O-H…O distance of 0.284 nm (Eckert et al. 1988), which in turn would be predicted to give an O-H stretching frequency around 3450 cm−1 (Libowitzky 1999). However, there is a large spread in the data on which the Libowitzky (1999) correlation is based (at this distance), and the corresponding stretching frequency could be as low as 3200 cm−1. The peak at 6.7 ppm is therefore consistent with the low frequency OH stretching peaks which have been observed in the FTIR spectra of forsterite when synthesized under conditions of high silica activity (Lemaire et al. 2004; Grant et al. 2006). 1

H MAS NMR studies of garnet

Figure 6. The same 1H MAS NMR spectra as shown in Figure 5, but displayed over a wider frequency range to show the spinning sidebands.

A sample of grossular garnet that contained 3.6 wt% H2O was included in the early 1H MAS NMR study of hydrous minerals by Yesinowski et al. (1988). The MAS spectrum was broader than would be expected based on the average H density, suggesting either that proton-proton dipolar coupling is larger than would be expected based on a homogeneous distribution of H in the sample or that an alternative line-broadening mechanism was operating. It was suggested that the width was too large to be explained by chemical shift dispersion, and that the width was not related to paramagnetic impurities because the width of the spinning sideband envelope did not increase significantly with increasing field. It was therefore concluded that the width was due to proton-proton dipolar coupling, and hence that protons are clustered within the structure. This conclusion was not unexpected as there is known to be a solid solution series between grossular and hydrogrossular and the latter contains the clustered (OH)4 defect. 1

H MAS NMR studies of SiO2 polymorphs

Yesinowski et al. (1988) reported a 1H MAS NMR spectrum for a natural quartz sample. The only peaks that were observed were attributed to the organic contaminant at 1.5 ppm and fluid inclusions at 4.7 ppm. No features for dissolved hydroxyl were observed. No NMR study of hydrogen in coesite or stishovite has yet been published. 1

H MAS NMR studies of feldspars and other aluminosilicate framework minerals

The earliest application of 1H MAS NMR to NAMs was the study of feldspars by Yesinowski et al. (1988). Three different resonances were observed, termed A, B and C. Resonance A was a narrow peak at 1.5 ppm, and was attributed to contamination from organic species because its intensity was reduced by refluxing the samples in CCl4, then packing the sample into

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hexane-washed rotors using gloves. Peak B was assigned to fluid inclusions because it was very narrow even at slow spinning speeds, had no spinning sidebands, and had a shift close to that for liquid H2O. In contrast peak C had an extensive, field-independent, spinning sideband pattern that was similar to that for water molecules in analcite. This resonance was therefore attributed to isolated water molecules which could be experiencing 180° flips or librational motion, but for which any motion was strongly anisotropic. The spin-lattice relaxation time, T1, was also measured for A, B and C, and found to be much longer for C than A or B, consistent with the assignments of A and B to surface or included components which cannot effectively interact with the dissolved C component. The Yesinowski et al. (1988) study also included the ammonium feldspar, buddingtonite. 1H MAS NMR of this sample gave a narrow peak at 6.8 ppm with moderate sideband intensities. This shift is consistent with ammonium in other compounds and the narrow spectrum indicates that the strong dipolar coupling between the protons of the NH4 group must be partially averaged by molecular motion. More recently Johnson and Rossman (2003) have used both FTIR and 1H MAS NMR to characterize the hydrous components in feldspars and to calibrate the infrared absorption by both OH and H2O. The NMR data showed that microcline samples contained structural water molecules at a level of 1000-1400 ppm H2O, whereas a sanidine sample contained 170 ppm H2O as hydroxyl. Plagioclase samples were also studied and dissolved OH concentrations were reported, but difficulties were encountered in unambiguously assigning intensity to dissolved or adsorbed OH. As in previous studies, the spectra were also complicated by the presence of organic molecules at the surface and the presence of fluid inclusions. Xia et al. (2000) also used 1H MAS NMR to determine the water concentrations of several anorthoclase megacrysts from Cenozoic alkalic basalts from China. The total water contents of three of the samples are 405 ppm, 915 ppm and 365 ppm, but more work would be needed to determine the distribution of water of different types in these samples. 1

H MAS NMR studies of wadsleyite

Wadsleyite was produced unintentionally in an NMR study of high-pressure hydrous silicates (Phillips et al. 1997). In one sample of superhydrous phase B, a significant proportion of the sample was actually either phase B or wadsleyite. This enabled the spectra of both superhydrous phase B and wadsleyite to be obtained. The 1H MAS spectrum of wadsleyite consisted of a narrow peak (indicating isolated hydroxyl groups) at 1.5 ppm. The intensity of the signal, when coupled to the proportions of the phases was used to calculate a water concentration in the wadsleyite of 0.2 wt%. Kohn et al. (2002) measured the 1H MAS NMR spectrum of a wadsleyite sample with a much higher water concentration of 1.5%. This sample had a complex and asymmetric spectrum that was fitted with six peaks in the range 1.4-11.0 ppm (Fig. 7). The 1H MAS NMR data were consistent with FTIR spectra measured on the same samples, and quantification of both data sets suggested that the Libowitzky and Rossman (1997) calibration of IR intensities works well. The resolution between the sites was inferior to that for the FTIR spectra for the same sample. One could speculate that the linewidth results from residual dipolar coupling between protons, but subsequent studies at higher magnetic fields and with faster magic spinning (unpublished data) did not show a significant improvement in resolution, so the controls on linewidths in this system are not yet clear.

NON-SPINNING 1H NMR EXPERIMENTS In an earlier section it was explained that a variety of interactions cause broadening of NMR lines in the solid state. Many of these interactions are averaged by magic angle spinning, leaving the spectrum as a measure of the distribution of chemical shifts in the sample. MAS enables much useful information to be obtained, and by reducing linewidths, also increases sensitivity. This reduction in sensitivity is a major drawback in applying static NMR to NAMs,

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Figure 7. The 1H MAS NMR centerbands of experimentally synthesised hydrous wadsleyite. (a) is the experimental spectrum, (b) is a fit the spectrum with the Lorentzian peaks shown in (c), and (d) is the residual. This NMR spectrum is less well resolved than the FTIR spectrum of the same sample, so this fit is not intended to be unique, but does provide a basis for comparison (see Kohn et al. 2002 for details).

a b c d 25

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and all the problems of accidental contamination by H-bearing compounds or materials are, if anything, more severe for static measurements, so application of static NMR to NAMs is of limited applicability. However, some of the information which is lost under MAS is useful in its own right, and a variety of strategies can, in principle, be employed to exploit this information. The simplest option is to perform static (non-spinning) experiments on powdered samples. In this case characteristic lineshapes are obtained which result from summing the contributions of all directional interactions over all orientations. The classic case is that for pairs of protons, such as immobile water molecules in a structure. The strong dipolar interaction between the two protons leads to a characteristic broad doublet feature, known as the Pake doublet. In this case, the strength of dipole-dipole coupling can be measured, and the H-H distance calculated (e.g., Phillips et al. 1997). If large single crystals are available, additional information can be obtained by orienting the crystal, and collecting spectra with the crystal fixed at varying angles to the magnetic field. The three dimensional nature of the different interactions can be explored in this way. A more sophisticated method of obtaining data on H environments using static samples was applied to garnet samples by Cho and Rossman (1993). Grossular samples with a water concentration of 0.2-0.3 wt% H2O were studied using non-spinning NMR experiments, and the data were compared with data for hydrogrossular. A multiple quantum technique was used to deduce that one sample contained mostly pairs of H, whereas the other sample contained clusters of two and four protons. As expected the hydrogrossular contained clusters of four protons. In addition Cho and Rossman (1993) analyzed the static lineshapes in detail and obtained an average interproton distance of 0.169 nm in grossular, compared with 0.216 nm in hydrogrossular.

STUDIES OF OTHER NUCLEI IN NAMS 1 H is the most sensitive nucleus for NMR, so it is the obvious candidate for NMR studies of the dissolution of water in NAMs. 2H could potentially be used for deuterated samples, but 2H is very much less sensitive than 1H and it is a quadrupolar nucleus (i.e., I > ½) so is subject to additional line-broadening mechanisms. Hydrous glasses with water concentrations of 1.6-4.8

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wt% have been studied with 2H (Eckert et al. 1987) and speciation and dynamic information were obtained. It is unlikely that 2H NMR of NAMs will be successful in the near future. In most cases NAMs will not contain enough H to have a significant effect on the average environment of any of the other atoms in the mineral. There are a few exceptions however, and in these cases NMR studies of other nuclei may be useful. Two important examples are the high pressure polymorphs of Mg2SiO4, wadsleyite and ringwoodite. These phases, although nominally anhydrous, can contain up to several wt% H2O, and there will therefore be significant differences in the environments of 29Si and 17O environments between hydrous and anhydrous samples. The relevant 17O NMR data for hydrous wadsleyite and ringwoodite have not been published, but comparison with 17O NMR for dry samples (Ashbrook et al. 2005) could provide an important constraint on the dissolution mechanism. In some cases the molar concentration of H becomes significant compared to that of other important minor components in NAMs. A recent study of hydrous and anhydrous aluminous orthopyroxene (Kohn et al. 2005) used 27Al NMR to test the proposition (Rauch and Keppler 2002) that dissolution of H is coupled to tetrahedral Al3+, with a consequent increase in the tetrahedral Al: octahedral Al ratio. Rauch and Keppler (2002) suggested that this ratio could potentially be used as the basis of a geohygrometer, even for mantle xenoliths which had dehydrated on ascent. In contrast, the 27Al NMR data of Kohn et al. (2005) suggest that the tetrahedral Al: octahedral Al ratio is the same in both dry and hydrous samples. It should be noted, however, that 27Al is a quadrupolar nucleus with I = 5/2, so interpretation of 27Al MAS NMR spectra is not straightforward and requires an understanding of the nature of the quadrupole interaction (MacKenzie and Smith 2002). Figure 8 shows 27Al spectra for dry and hydrous orthopyroxenes, together with a subtraction that emphasizes the difference between them. Although the [Al]4:[Al]6 is constant the wet sample contains new Al environments which are not present in the dry sample. The concentrations of these new sites imply that each dissolved H modifies the environment of one tetrahedral and one octahedral Al, consistent with protonation of O21 and O22 sites (Kohn et al. 2005).

Figure 8. 27Al MAS NMR spectra of dry and hydrous aluminous enstatite, obtained at a magnetic field of 18.8 T. The ratio of tetrahedral Al : octahedral Al is (within error) the same for the two samples, and after correction for the magnitude of the quadrupole coupling constants, is very close to 50:50 (Kohn et al. 2005). The difference spectrum shows that new Al[4] and Al[6] sites are present in the wet sample, and this information can be used to make deductions about the mechanism of water incorporation. [Used with permission of Elsevier from Kohn et al. (2005) Earth Planet. Sci. Lett., Vol. 238, Fig. 1, p. 342-350.]

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PROSPECTS FOR FUTURE DEVELOPMENT OF NMR FOR STUDIES OF NAMS In the context of the whole body of research on nominally anhydrous minerals, NMR data have played a relatively minor part. This is certainly because the measurements of 1H are not easy, and special precautions have to be made to overcome the problems of background H signal. More work on obtaining H-free probe materials, drying and purging etc., could potentially reduce backgrounds, but it is debatable whether the improvement would justify the required effort. Various pulse sequences have been used for background suppression in NMR e.g., (Cory and Ritchey 1988; Chen et al. 2004); these have not yet been applied widely to NAMs. The availability of higher magnetic fields could potentially increase the intrinsically small signal for 1H NMR of NAMs. However, the improvement would not be dramatic, and faster spinning would also be required in order to avoid problems resulting from increased chemical shift anisotropy. For some samples, increasing the field could increase the resolution of 1H MAS spectra, depending on the line broadening mechanisms. Studies of quadrupolar nuclei, such as 27Al, 17O and 23Na are more likely to be successful at high fields, and some progress in studying dissolution mechanisms via the effect of dissolved H on other nuclei should be expected. The use of cross-polarization from 1H to other nuclei (e.g., Farnan et al. 1987) could have a major effect on the sensitivity of NMR of other nuclei to hydrous species. If hydrogen is clustered in NAMs, the H-H distances could potentially be accessed using double quantum MAS techniques (e.g., Schnell and Spiess 2001). In summary, despite several partially successful attempts to introduce NMR as a major tool for studying the dissolution of water in NAMs, the full potential of the technique has not yet been realized. The problem of paramagnetic cations in most naturally occurring NAMs will remain a major limitation of NMR, so its major contribution is likely to be in simplified synthetic systems. As with most spectroscopic techniques, the most crucial factor in obtaining high quality 1H NMR spectra is probably sample selection. Samples should be entirely free of fluid or melt inclusions or hydrous phases, and should be only coarsely crushed to minimize the surface area. The full potential of NMR in this field may ultimately be limited more by experimental methodologies for producing such clean samples than by the methods of obtaining NMR spectra.

ACKNOWLEDGMENTS I would like to thank NERC for funding, Michael Fechtelkord, Mark Smith and an anonymous reviewer for helpful comments on the manuscript and Hans Keppler for his efforts in editing this volume.

REFERENCES Ashbrook SE, Berry AJ, Hibberson WO, Steuernagel S, Wimperis S (2005) High-resolution 17O MAS NMR spectroscopy of forsterite (α-Mg2SiO4) wadsleyite (β-Mg2SiO4), and ringwoodite (γ-Mg2SiO4). Am Mineral 90:1861-1870 Bell DR, Rossman GR, Maldener J, Endisch D, Rauch F (2003) Hydroxide in olivine: A quantitative determination of the absolute amount and calibration of the IR spectrum. J Geophys Res-Solid Earth 108(B2) Art. No. 2105 Bromiley GD, Keppler H (2004) An experimental investigation of hydroxyl solubility in jadeite and Na-rich clinopyroxenes. Contrib Mineral Petrol 147:189-200 Brunner E, Sternberg U (1998) Solid-state NMR investigations on the nature of hydrogen bonds. Prog Nucl Magn Reson Spectrosc 32:21-57 Chen Q, Hou SS, Schmidt-Rohr K (2004) A simple scheme for probehead background suppression in one-pulse 1 H NMR. Solid State Nucl Magn Reson 26:11-15

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Cho H, Rossman GR (1993) Single crystal NMR studies of low-concentration hydrous species in minerals grossular garnet. Am Mineral 78:1149-1164 Cory DG, Ritchey WM (1988) Suppression of signals from the probe in Bloch decay spectra. J Magn Reson 80:128-132 Eckert H, Yesinowski JP, Silver LA, Stolper EM (1988) Water in silicate glasses - quantitation and structural studies by 1H solid echo and MAS NMR methods. J Phys Chem 92:2055-2064 Eckert H, Yesinowski JP, Stolper EM, Stanton TR, Holloway J (1987) The state of water in rhyolitic glasses a deuterium NMR study. J Non-Cryst Solids 93:93-114 Engelhardt G, Michel D (1987) High-Resolution Solid-State NMR of Silicates and Zeolites. Wiley Farnan I, Kohn SC, Dupree R (1987) A study of the structural role of water in hydrous silica glass using crosspolarization Magic Angle Spinning NMR. Geochim Cosmochim Acta 51:2869-2873 Fechtelkord M (2004) Solid state NMR spectroscopy as supporting method in Rietveld refinements of rockforming minerals: New developments and examples. EMU Notes Mineral 6:421-463 Freund F, Dickinson JT, Cash M (2002) Hydrogen in rocks: an energy source for deep microbial communities. Astrobiology 2:83-92 Grant KJ, Kohn SC, Brooker RA (2006) Solubility and partitioning of water in synthetic forsterite and enstatite in the system MgO-SiO2-H2O±Al2O3. Contrib Mineral Petrol 151:651-664 Johnson EA, Rossman GR (2003) The concentration and speciation of hydrogen in feldspars using FTIR and 1H MAS NMR spectroscopy. Am Mineral 88:901-911 Keppler H, Rauch M (2000) Water solubility in nominally anhydrous minerals measured by FTIR and 1H MAS NMR: the effect of sample preparation. Phys Chem Mineral 27:371-376 Kirkpatrick RJ (1988) MAS NMR spectroscopy of minerals and glasses. Rev Mineral 18:341-403 Kohn SC (1996) Solubility of H2O in nominally anhydrous mantle minerals using 1H MAS NMR. Am Mineral 81:1523-1526 Kohn SC (1999) Partitioning of water between nominally anhydrous minerals in the upper mantle. In: Processes and Consequences of Deep Subduction. Vol 99/7. Mysen B, Rubie D, Ulmer P, Walter M (eds) Terra Nostra, Alfred Wegener Stifung, p 58-59 Kohn SC (2004) NMR studies of silicate glasses. EMU Notes Mineral 6:399-419 Kohn SC, Brooker RA, Frost DJ, Slesinger AE, Wood BJ (2002) Ordering of hydroxyl defects in hydrous wadsleyite (β-Mg2SiO4). Am Mineral 87:293-301 Kohn SC, Roome BM, Smith ME, Howes AP (2005) Testing a potential mantle geohygrometer; the effect of dissolved water on the intracrystalline partitioning of Al in orthopyroxene. Earth Planet Sci Lett 238: 342-350 Lemaire C, Kohn SC, Brooker RA (2004) The effect of silica activity on the incorporation mechanisms of water in synthetic forsterite: a polarised infrared spectroscopic study. Contrib Mineral Petrol 147:48-57 Libowitzky E (1999) Correlation of O-H stretching frequencies and O-H..O hydrogen bond lengths in minerals. Monatshefte für Chemie 130:1047-1059 Libowitzky E, and Rossman GR (1997) An IR absorption calibration for water in minerals. Am Mineral 82: 1111-1115 MacKenzie KJD, Smith ME (2002) Multinuclear Solid State Nuclear Magnetic Resonance of Inorganic Materials. Pergamon Phillips BL, Burnley PC, Worminghaus K, Navrotsky A (1997) 29Si and 1H NMR spectroscopy of high-pressure hydrous magnesium silicates. Phys Chem Mineral 24:179-190 Rauch M, Keppler H (2002) Water solubility in orthopyroxene. Contrib Mineral Petrol 143:525-536 Raynes WT (1977) Theoretical and physical aspects of nuclear shielding. NMR Spectrosc Period Rep Chem Soc 7:1-25 Schnell I, Spiess HW (2001) High-resolution 1H NMR spectroscopy in the solid state: very fast sample rotation and multiple quantum coherences. J Magn Reson 151:153-227 Skogby H, Bell DR, Rossman GR (1990) Hydroxide in pyroxene - variations in the natural environment. Am Mineral 75:764-774 Smyth JR, Bell DR, Rossman GR (1991) Incorporation of hydroxyl in upper-mantle clinopyroxenes. Nature 351:732-735 Stalder R, Skogby H (2002) Hydrogen incorporation in enstatite. Eur J Mineral 14:1139-1144 Stebbins JF (1988) NMR spectroscopy and dynamic processes in mineralogy and geochemistry. Rev Mineral 18:405-429 Xia QK, Pan YJ, Chen DG, Kohn S, Zhi XC, Guo LH, Cheng H, Wu YB (2000) Structural water in anorthoclase megacrysts from alkalic basalts: FTIR and NMR study. Acta Petrologica Sinica 16:485-491 Yesinowski JP, Eckert H, Rossman GR (1988) Characterization of hydrous species in minerals by high-speed 1 H MAS NMR. J Am Chem Soc110:1367-1375

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Reviews in Mineralogy & Geochemistry Vol. 62, pp. 67-83, 2006 Copyright © Mineralogical Society of America

Atomistic Models of OH Defects in Nominally Anhydrous Minerals Kate Wright Nanochemistry Research Institute Curtin University of Technology GPO Box 1987 Perth, Western Australia 6845, Australia e-mail: [email protected]

INTRODUCTION The Earth’s upper mantle may contain substantial amounts of water dissolved in nominally anhydrous minerals (NAMs) such as the Mg2SiO4 polymorphs, pyroxenes and garnets. This water, incorporated into the crystal lattice as hydrogen defects, can have a profound influence on the physical properties of the mantle, even when present at low concentrations. An understanding of these defects at the atomic level is therefore of fundamental importance for the development of models of the evolution and dynamics of the Earth’s mantle. The incorporation of hydrogen and its influence on the properties of NAMs has been an active area of research for almost three decades. High pressure synthesis of hydrous phases, analyzed using a range of spectroscopic techniques (see Kohn 2006; Libowitzky and Beran 2006; Rossman 2006), have yielded a wealth of information that allow us to determine the concentration of hydrogen that can be accommodated by various NAMs, and provide information on the mechanisms of uptake. However, these data are often complex, and difficult to interpret unambiguously. Computer simulation methods can offer real insights at the atomic level, often not accessible by experiment, and provide an alternative way to explore hydrogen defects in minerals. The past 20 years have seen an explosion in the use of computational modeling to study a range of phenomena in minerals. These include the high-pressure behavior of mantle (Oganov and Price 2005) and core (Vocadlo et al. 2003) phases, diffusion (Walker et al. 2003), dislocation structures (Walker et al. 2005), and mineral surface reactivity (Kerisit et al. 2005). A broad introduction to the methods and applications to the geosciences is given in the recent MSA volume edited by Cygan and Kubicki (2001). In this chapter we explore the contribution of computational methods to the development of atomistic models of hydrogen defects in NAMs of the Earth’s upper mantle and transition zone. We begin with an introduction to defects in solids, followed by a brief overview of the computational methods used to model them, and then review the results of studies on the most important upper mantle NAMs.

POINT DEFECTS IN MINERALS All crystalline materials contain imperfections, or defects, that disrupt their long-range ordering. These can be point defects, line defects (dislocations) or planar defects (stacking faults, crystallographic shear planes), as well as interfaces such as grain boundaries and twin planes. A whole range of different defect types can be present in the same crystal, depending on the conditions of formation and subsequent history, and can interact in a variety of ways. 1529-6466/06/0062-0004$05.00

DOI: 10.2138/rmg.2006.62.4

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Point defects are the simplest type of imperfection that can occur in a crystal, and involve the removal, inclusion or replacement of an atom or ion at specific sites in the crystal. They are important, since they are the means by which atomic migration takes place and can influence color, electrical conductivity and reactivity. In general, we define three types of point defect: vacancies, where an ion is removed from its normal lattice site; interstitials, where an ion is present at a non-lattice site; and impurities, where dopants are present either at lattice or interstitial sites (Fig. 1). In ionic and semi-ionic crystals, these point Figure 1. Schematic representation of point defects defects will typically be charged species in a crystal of composition MX. and so must occur in balanced defect populations to maintain charge neutrality. Within pure crystals, defects made up of balanced populations of cations and anions are termed Schottky defects, while those composed of a vacancy and interstitial of the same species are known as Frenkel defects. In the strictest sense, Schottky disorder requires that charge neutrality and stoichiometry be maintained, however, the term is fairly loosely applied in the literature and we will use Schottky to refer to any charge neutral group of vacancies. The formation energy of a Schottky defect (ESch) is the sum of the individual vacancy energies (Ex) plus the lattice energy (U) of the phase removed: ESch = EV1 + EV 2 + .......EVn + U

(1)

For a Frenkel defect, the energy is simply the sum of the corresponding vacancy and interstitial energies. Point defects occur in all crystals at temperatures above 0 K and, in pure crystals, there will be a finite population of these intrinsic defects in thermodynamic equilibrium with the system. The change in free energy (∆G) associated with the introduction of a defect is expressed in the usual way as: ∆G = ∆H − T ∆S

(2)

∆H is the enthalpy, associated with changes in nearest neighbor interactions, and ∆S the entropy increase due to the introduction of disorder into an otherwise perfect crystal. The entropy term includes vibrational disorder in the atoms around the defect as well as configurational terms related to way in which the defects are distributed within the crystal. The equilibrium concentration of defects in a stoichiometric material of composition MX (e.g., MgO, NaCl) can be approximated by the following: ndef = Ne−∆H / 2 RT

(3)

where N is the number of sites, ∆H is the enthalpy required to form the defect, R the universal gas constant, and T temperature. The full derivation of this formula can be found in Putnis (1992). Populations of intrinsic defects are generally small; for a typical alkali halide at room temperature less than 1 site per million will be vacant (Tilley 1987). Normally, we would expect one type of point defect to dominate, and this will be the one with the lowest value of ∆H. Generally speaking, Schottky defects tend to dominate in close packed solids, while Frenkel defects are more common in framework and layered structure materials.

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In many crystalline materials, non-intrinsic vacancies and interstitials can be created in response to the presence of impurities and are thus termed extrinsic. These impurities may occur as neutral species (e.g., Mn2+ replacing Mg2+), or as charged species (e.g., Al3+ replacing Si4+) that must be charge balanced by another impurity (coupled substitution) or by an accompanying vacancy or interstitial. Since our defects are all charged species, we might expect them to interact strongly and form discrete defect clusters with significant binding energies. When defects form, the ions around them must relax to accommodate the new configuration, and in some cases, the energy will be lower for a cluster than for the same isolated defects. At this stage it’s a good idea to introduce the notation used for point defects, so that they can be easily identified and written down. The Kröger-Vink (Kröger 1972) notation is widely used to describe point defects. Vacancies (V) are defined in terms of species (subscript) and charge (superscript), where the charge may be neutral (x), positive (•) or negative ( / ). For // example, VMe describes a metal (Me) vacancy with an effective 2− charge. A Me2+ interstitial is x written Mei•• , while a neutral impurity (A) at a Me site is given as AMe . In the case of hydrogen, − 2− we are generally referring to an (OH) group replacing an O , which is written (OH)•O .

THEORETICAL BACKGROUND This section presents a brief introduction to the different computational approaches used for the study of defects. Technical details of the different theories are not included since there are many excellent texts available that cover computational methods (e.g., Foresman and Frisch 1997; Leach 2001; Gale and Rohl 2003; Griffiths 2005) and readers are referred to these for a more detailed and rigorous treatment. Computer simulation methods aim to determine the energy of a solid as a function of the interaction of all particles within that system, with varying degrees of approximation depending on the level of theory used. Simulations can be static, where the system is essentially at 0 K, or dynamic, where the free energy is calculated by molecular dynamics (MD) or lattice dynamics techniques. This can be obtained either by quantum mechanics or by atomistic techniques, based on classical mechanics, in which the details of the electronic structure are subsumed into a series of effective interactions that depend only on the nuclear positions.

Quantum mechanical methods Within quantum mechanical (QM) theory, both the electrons and nuclei are explicitly considered and their interactions are generally calculated using either Hartree-Fock (HF) or density functional theory (DFT). In both cases the Born-Oppenheimer approximation, that the motion of electrons can be separated from that of the nuclei, is assumed to hold true. However, the difficulty arises when trying to calculate interactions between electrons, since the potential experienced by one electron depends on the position of all other electrons in the system. These exchange interactions, between electrons of like spin, and correlation interactions, between electrons of opposite spin, are treated in different ways depending on the approach used. HF theory calculates the exchange energy explicitly but ignores correlation, although post-HF techniques, such as Moller-Plesset (Leininger et al. 2000) and Coupled Cluster theory, can overcome this to some extent at the price of significantly increased computational cost. DFT, while being in principle an exact theory, in practice has to approximate both the exchange and correlation potentials, using either the Local Density Approximation (LDA) or Generalized Gradient Approximation (GGA). Hybrid functionals, such as B3LYP (Becke 1993), that combine GGA or LDA with exact HF exchange, are also available. The wave function is generally described by a linear combination of basis functions that can be atom centered Gaussian type functions, or plane waves. In many cases, only the valence electrons need

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to be explicitly considered, as it is these that are responsible for bonding. The electrostatic potential due to the frozen core electrons and the nucleus are commonly represented by a pseudopotential, and can lead to substantial computational savings, particularly when used in conjunction with plane waves. Quantum mechanical calculations, are by their very nature, the most accurate and reliable approach, although they require significant computational resources. Early studies using these techniques were limited to the use of clusters of atoms representing a solid, or very small unit cells. Recent developments in both hardware and software mean that it is now possible to calculate the properties of complex mineral phases using these methods and thus their use is increasing. DFT is by far most commonly used technique within the Earth Sciences, with particular success being enjoyed using the planewave, pseudopotential codes such as CASTEP (Segall et al. 2002) and VASP (Kresse and Furthmuller 1996a,b). DFT does, however, have its limitations; band gaps are typically underestimated by about 50%, and while LDA overestimates binding, GGA underestimates it leading to calculated cell parameters that are normally 1-2% too large. In addition, long-range van de Waals interactions are not well described, so that layered structures can prove difficult to model accurately.

Classical methods Classical atomistic, or molecular mechanics (MM), simulation techniques employ interatomic potential functions to describe the total energy of the system in terms of atomic positions. These potentials include long-range electrostatic effects as well as short-range interactions produced by the overlap of nearest neighbor electron clouds. Terms to describe oxygen ion polarizibility and directionality of bonding are also available. The effective potential parameters are derived either by fitting to experimental data (structure, elastic constants, dielectric constants, etc.), or by fitting to potential energy surfaces generated by high level QM calculations. The equilibrium positions of the ions are then evaluated by minimizing the lattice energy until all forces acting on the crystal are close or equal to zero. The majority of defect calculations carried out using these methods are performed at 0 K and 0 GPa and so the energies obtained are enthalpies rather than free energies of defect formation, although free energies can be obtained by the use of lattice dynamical techniques. A comprehensive overview of interatomic potential methods can be found in Gale and Rohl (2003). Interactions between closed shell ionic species are well modeled by standard two-body potentials of the Born-Meyer type but bonded molecules, such as (OH), need to be treated differently. The hydroxyl molecule is generally described using a Morse potential of the form: 2   − α(r −r ) UijMorse = D  1 − e ij 0  − 1   

( 4)

where D corresponds to the dissociation energy of the bond, r0 is the equilibrium bond length and α, in combination with D, is related to the vibrational frequency of the stretching mode. Traditionally, studies of hydrogen defects in minerals use the Morse potential parameters originally derived from HF calculations on NaOH (Saul et al. 1985). These have been used extensively in the study of hydroxyl in zeolites (e.g., Schröder et al. 1992), and in a whole range of other minerals including muscovite (Collins and Catlow 1990) and goethite (Steele 2002). In the Saul et al. model, polarizibility of the hydroxyl oxygen ion is not included, and the ions have fractional charges (qO = −1.412, qH = +0.412) such that the (OH) unit has −1 overall charge. Despite its success in calculating defect structures and energies, the Saul et al. potential does not give O-H vibrational frequencies that agree with experiment. This is a reflection on the fact that HF calculations were used in the original fitting procedure. Gatzemeier and Wright (2006) modified the α parameter by fitting it to O-H stretching frequencies in forsterite obtained using QM/MM methods (Braithwaite et al. 2003). Other,

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more sophisticated, Morse potentials have been developed for OH, such as that by Baram and Parker (1996). The model explicitly includes oxygen polarizibility via the use of a shell model (Dick and Overhauser 1956), where the ion is divided into a core containing all of the mass, and a shell, coupled by an harmonic spring. Other potential forms used to model water and hydroxyl groups include simple LennardJones type models as well as more sophisticated potentials, such as that of Stillinger-David (Stillinger and David 1980) that allow the dissociation of the water molecule. However, these models have not been used in the study of hydrogen in NAMs and so will not be discussed further.

Treatment of defects The choice of theoretical method to use for defect calculations depends on the level of accuracy required and the actual quantity to be calculated. This, along with the limitations imposed by available computational resources, determine which level of theory to use. Classical MM methods have been extremely successful at predicting defect behavior in a whole range of solids from complex ionic materials, such as zeolites (Schröder et al. 1992), clays (Cygan et al. 2004) and carbonates (Austen et al. 2005), to battery materials (Islam et al. 2005) and semi-conductors (Wright and Gale 2004). They need minimal computer time and memory so that large numbers of possible configurations can be easily sampled. However, the quality of the results will only be as good as the potential parameters available for the system under consideration. Interatomic potential methods do have their limitations, as they are generally unable to model bond breaking and bond forming processes, although there are some exceptions to this such as the reactive empirical bond order (REBO) type models (e.g., Brenner et al. 2002). Equally importantly, they cannot be used to assess the influence of defects on those properties that explicitly depend on the electronic structure, such as JahnTeller distortions associated with transition metal ions. QM methods usually give a much more accurate description of a system, but use far more resources. There are essentially two approaches to calculating the structure and energetics of defects in solids, namely the supercell (SC) and cluster methods. The supercell approach, illustrated in Figure 2, has the defect within in a large supercell and the system is modeled in any code, QM (e.g., CASTEP, VASP) or classical (e.g., GULP, PARAPOCS), using periodic boundary conditions. The cell should be large enough that the defect does not interact with those in the periodic images as this will introduce an additional component into the total energy obtained. Defect-defect interactions can be corrected for in terms of electrostatic multipoles, and, in the case of charged defects, a neutralizing background needs to be applied (Leslie and Gillan 1985). Although this first consideration is not an issue for force field calculations, which can handle cells containing thousands of atoms, it can be a problem when performing QM calculations. For periodic DFT codes (PDFT) such as CASTEP and VASP, the CPU time required to run a calculation goes up in Figure 2. Illustration of a supercell used in defect a non-linear fashion as the number of atoms calculations. The cell defined by a dark line is increases. With the advances in parallel the unit cell which is periodically repeated in the supercell. computing, and increases in efficiency, large-

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scale simulations of cells containing hundreds of atoms are possible, although most calculations use much less than this. The cluster method involves cutting out a fragment of a crystal that has the defect at its centre, and embedding it in some representation of the bulk material (Fig. 3). The embedding approach is most commonly implemented in MM calculations, where the polarization caused by introducing the defect is handled using the formulation of Mott and Littleton (Mott and Littleton, 1938). In this approach the crystal is divided into two concentric spherical regions (Fig. 3). In region 1, which contains the defect at its centre, an explicit atomistic Figure 3. The embedded cluster has a central area containing simulation is carried out to adjust the defect that is embedded in a representation of the bulk the coordinates of all ions within material. the region until they are at positions at which no net forces act on them, i.e. they are relaxed around the defect. In region 2, the effects of the defect are relatively weak and the relaxation can be calculated essentially as the polarization response to the effective charge of the defect. In practice an interface region between regions 1 and 2, referred to as 2A, is normally used. The resulting defect energy is a measure of the perturbation by the defect of the static lattice energy of the crystal. As with supercells, size matters, and region 1 should be large enough that the defect energy is converged with respect to region size. The Mott-Littleton method is implemented in codes such as GULP (Gale and Rohl 2003). Although the cluster approach works well within classical calculations, its use in quantum mechanical simulations is more problematic, primarily due to edge effects and the limited number of atoms that can be included in the cluster. Hybrid, so called QM/MM embedded cluster methods (see for example Braithwaite et al. 2002) overcome these problems, by having a quantum region surrounded by a classical one. The central region contains the defect of interest and is treated at the quantum mechanical level of theory, using either HF or DFT. This QM region is normally terminated with atoms described by effective core pseudopotentials and is embedded in a large (>50 Å) array of point charges which represent the potential due to the bulk crystal that acts on the embedded cluster. Between the two, is a sphere of classical atoms that are described by interatomic potentials. The embedded cluster approach overcomes many of the problems associated with studying charged defects using periodic supercell methods, including the problem of the energy term produced by defect – defect interactions.

OH DEFECTS IN MANTLE SILICATES Water can be incorporated into NAMs via a number of different mechanisms, depending on the chemistry and defect structure of the host mineral. Equation (5) describes the formation of a hydrogarnet defect, where two molecules of water interact with a silicon ion to form a silicon vacancy charge balanced by four (OH) groups, and a unit of SiO2: 2H 2O + SiSix + 4OOx → [ VSi⋅ 4(OH)O ] + SiO2 x

(5)

Atomistic Models of OH in NAM’s

73

The energy of the above reaction is found by summing together the energies of the different terms, including the self-energy of the water molecule, calculated using the same level of theory. In some MM calculations, this self-energy is substituted by a proton transfer term representing the energy of the H2O + O2− → 2(OH)− reaction (see Wright et al. 1994 for details). Other reaction pathways (Eqns. 6-8) involve the creation of other vacancies, or reactions with impurities and their energetics are calculated in a similar manner. In Equation (6), water // is incorporated via the formation of VMe  , two (OH) groups and a unit of metal oxide. x H 2O + MeMe + 2OOx → [ VMe⋅ 2(OH)O ] + MeO x

(6)

Similar reactions can occur for metal cations of different charges, with corresponding numbers of (OH) and different oxide products. Reactions, such as those in Equations (7) and (8), involve interactions of water with impurity cations and vacancies on the oxygen sub-lattice. Within clinopyroxenes, Al Si/ and Na /Me substitutions can be charge balanced by the inclusion of (OH) as: / H 2O + 2 AlSi + 2OOx + VO•• → 2 [ AlSi⋅ (OH)O ] + OOx

( 7)

/ H 2O + 2 Na Me + VO•• + OOx → 2 [ Na Me⋅ (OH)O ]

(8)

x

x

The identity of the charge compensating defects in NAMs had been the subject of considerable debate in the literature, as has the extent to which these defects bind with the hydrogen. This is the sort of problem that can readily be addressed by computational methods, as the relative stabilities of known defect configurations can be assessed by calculating their formation energies, both bound and unbound. In addition, the O-H stretching frequency can be determined for each configuration and compared with experiment. In this way, the results obtained from the calculations can be used to both constrain models, and to aid in the interpretation of experimental data. In the following sections we consider the literature on hydrogen defects in the major mineral phases of the Earth’s upper mantle and transition zone; i.e., the Mg2SiO4 polymorphs and the clinopyroxenes diopside and jadeite. Hydrogen defects in a number of other important nominally anhydrous minerals, have been studied using computational methods including garnets (Wright et al. 1994; Nobes et al. 2000), quartz (Lin et al. 1994; Purton et al. 1992) and feldspar (Wright et al. 1996) although these will not be covered here.

The Mg2SiO4 polymorphs Forsterite. There is considerable experimental evidence for the presence of hydrogen in all three of the Mg2SiO4 polymorphs, as discussed in various chapters in this volume. Of the three, forsterite is by far the most well studied both experimentally (Kohlstedt et al. 1996; Matveev et al. 2001; Demouchy and Mackwell 2003) and computationally (see Table 1 for references). Olivine [(Fe,Mg)2SiO4] is the dominant mineral in the Earth’s upper mantle and thus will exert a major control on the rheological behavior. Natural samples show levels of hydrogen in the range 1 to 140 ppm H2O, where hydrogen is expected to be incorporated into the olivine lattice in association with both silicon and magnesium vacancies (e.g., Kohlstedt et al. 1996; Kohn 1996). Concentrations of OH in natural samples appear to show some correlation with the geological setting and composition suggesting that P/T history, as well as local stoichiometry, can affect the uptake of hydrogen. There is some evidence (Bell and Rossman 1992) suggesting that iron rich olivines contain a greater proportion of hydrogen than those with low iron content; however, this relationship has not been quantified in any way. Calculations, based on both QM and classical methods (Table 1), have been used to investigate the relative stability of hydrogen defects at different positions in the forsterite

Wright

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Table 1. Summary of calculations carried out on the Mg2SiO4 polymorphs. All calculations with the exception of those marked with *, were carried out at 0 K and 0 GPa. Mineral

Method

Reference

Forsterite

P-DFT P-DFT QM/MM MM MM MM

Haiber et al. (1997) Brodholt and Refson (2000) Braithwaite et al. (2003) Wright and Catlow (1994) de Leeuw et al. (2000) Walker et al. (2006)

Wadsleyite

P-DFT MM MM MM

Haiber et al. (1997)* Wright and Catlow (1996) Parker et al. (2004)* Walker et al. (2006)

Ringwoodite

MM P-DFT

Blanchard et al. (2005) Haiber et al. (1997)

lattice and to calculate energies of the reactions in Equations (5) and (6). Forsterite has an orthorhombic unit cell with isolated SiO4 tetrahedra separated by magnesium ions octahedrally coordinated by oxygens. There are two symmetry inequivalent magnesium positions, and three different oxygen sites, as shown in Figure 4. Looking at the structure of forsterite, we can identify a number of possible environments for hydrogen: (i) interstitial hydrogen bound to any one of the three oxygen sites but isolated from any cation vacancies; (ii) hydrogen bound to oxygen adjacent to either M1 or M2 vacancies; and (iii) hydrogen bound to oxygen adjacent to silicon vacancies. Of the three oxygen sites, all calculations agree that O3 is the most easily protonated, and that the M1 vacancy has a lower formation energy that M2. The most favorable defect configuration involving magnesium vacancies has hydrogen bound to two O2 oxygens around the vacant Mg1 site, as shown in Figure 5a. The third possibility, of hydrogen surrounding a silicon vacancy is shown in Figure 5b. Calculated binding energies (Brodholt and Refson 2000; Braithwaite et al. 2003; Walker et al. 2006) for [VMg⋅2(OH) ]x and [VSi⋅4(OH) ]x and associated clusters, given in Table 2, are ° ° substantial, and hence there is a strong driving force for hydroxyl groups to combine with cation vacancies. Haiber et al. (1997) have suggested that under mantle conditions entropy would cancel out any vacancy-hydrogen binding and therefore only isolated interstitial hydrogen defects would be present. However, the magnitude of the binding energies is sufficient to overcome the activation energy for hydrogen diffusion in olivine, estimated as 130 kJ·mol−1 (Mackwell and Kohlstedt 1990), so that cation vacancies will act as a ‘sink’ for unassociated hydroxyl species. Brodholt and Refson (2000) suggest that reactions with water could actively promote the formation of metal vacancies, particularly silicon vacancies, leading to a form of hydrolytic weakening. Calculated energies for the dissolution reactions in Equations (5) and (6) are given in Table 3 and show considerable variation depending on the methodology used. The P-DFT and QM/MM calculations compare well, with both methods showing that reactions of water with silicon vacancies will be exothermic. The error on the QM/MM values comes from uncertainties in the value for the lattice energy of oxide phases produced on formation of the defect, which by necessity had to be calculated using a periodic QM code. Calculated values from MM calculations are much higher in energy than either of the QM values, although reaction with

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Figure 4. The unit cell of forsterite (Mg2SiO4) viewed along the [100] direction.

(a)

(b)

Figure 5. Structure of hydrogen defect complexes in forsterite produced from the data of Braithwaite et al. (2003). (a) [VMg⋅2(OH) ]x cluster, and (b) hydrogarnet ° defect [VSi⋅4(OH) ]x cluster. °

Table 2. Defect binding energies in forsterite calculated using periodic DFT (Brodholt and Refson 2000) and QM/MM DFT (Braithwaite et al. 2003). Binding energy (kJ·mol−1) Reaction

1

// / VMg + H•I → H Mg

2

/ H Mg + H•I → ( 2H )

3

VSi////

4

( 3H )

5

VSi////

6

( 3H )

/ // + H Mg → ( 4H ) + VMg

7

( 3H )

// + VMg

8

( 4H )

+ H•I / Si

/// → HSi

+ H•I

/ + H Mg / Si / Si x Si

X Mg

→ ( 4H )

X Si

/// → HSi

// + VMg X Si

→ ( 2H )

X Si

/ + H Mg

/ x + H Mg → ( 3H ) + (2H)Mg / Si

P-DFT

QM/MM DFT

−239

−245

−157

−203

−546

−519

−155

−202

−3.18 0.84 117 −2.89

Wright

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Table 3. Comparison of solution reaction energies in forsterite calculated using different techniques. P-DFT, Brodholt and Refson (2000), QM/MM, Braithwaite et al. (2003), MM Walker et al. (2006). Equation #

Reaction energy per OH (kJ·mol−1) P-DFT

QM/MM

MM

(5)

−23

−7 (± 30)

43

(6)

58

5 (± 30)

145

water leading to formation of the hydrogarnet defect is still most favorable. These energy differences most likely arise form the overestimation of polarization effects induced by the defect. All of the studies discussed above considered only defects in otherwise pure, perfect crystals of forsterite. However, other defects such as dislocations and grain boundaries will be sites for high concentrations of point defects and thus could be sinks for hydrogen. De Leeuw et al. (2000) investigated the formation of [VMg⋅2(OH)]x defect complexes in the bulk and along {010} tilt grain boundaries of forsterite. Their calculations showed this process to be more favorable along the grain boundary by over 100 kJ·mol −1 compared to the bulk. Infra-red (IR) spectroscopy has been used to measure the concentration of water in both natural olivine and synthetic forsterite. Polarized IR in particular can provide information on defect structures, as individual O-H vectors can be resolved. However, these analyses are complex with multiple peaks in the O-H stretching region, particularly in the case of mantle derived olivines that can make them difficult to resolve. Therefore the unambiguous assignment of specific frequencies to any one defect can be problematic. In general, two distinct groups of frequencies can be identified in the IR spectra of olivine, designated as group 1 for higher bands (3450 – 3650 cm−1), and group 2 for those in the lower frequency range of 3200-3450 (Bai and Kohlstedt 1993). IR frequencies can be calculated, within the harmonic approximation, from the second derivatives of the total energy with respect to each bond length. Braithwaite et al. (2003) identified two distinct bands analogous to those seen experimentally; the first, at around 3200 cm−1, was associated with the [VMg⋅2(OH) ]x defect cluster, and the second, higher band ° (3266-3478 cm  −1) with the hydrogarnet [VSi⋅4(OH) ]x cluster. IR frequencies calculated using ° interatomic potential methods are consistently around 300 wavenumbers higher than those from the QM/MM embedded clusters (Walker et al. 2006). The assignment of OH bands to defects made by Braitewaite et al. (2003) is supported by the recent experimental work of Lemaire et al. (2004), who synthesized hydrous forsterite over a range of silica activity conditions. These authors found that samples with low silica activity (Si vacancies dominant) exhibited OH bands at 3620-3450 cm−1, while the high silica activity samples (Mg vacancies dominant) showed bands at 3160, 3220 and 3600 cm−1. Comparison of experimental and computational results is more difficult in the case of olivine, where the oxidation state of Fe and its influence on defect chemistry must be considered. Studies by Matveev et al. (2001) on synthetic olivine also attribute group 1 bands to OH at Si vacancies and group 2 to OH at Mg vacancies. However, these authors note that the signature of OH in mantle derived olivines is quite different to that of synthetic samples, which has been interpreted by Berry et al. (2005) as being due to the presence of trace elements. Wadsleyite. Wadsleyite is believed to be the dominant mineral in the upper part of Earth’s transition zone. The wadsleyite structure (Fig. 6) is based on a nearly perfect cubic closepacking of oxygen atoms with silicon atoms in tetrahedral sites. The structure is orthorhombic

Atomistic Models of OH in NAM’s

77

Figure 6. The unit cell of wadsleyite (Mg2SiO4) viewed along [001].

(space group Imma), but it also has a monoclinic polymorph (space group I2/m) recently identified by Smyth et al. (1997). Wadsleyite differs from forsterite in that it contains Si2O7 groups, so that there are four distinct oxygen sites as well as three magnesium sites. Smyth (1987), using analyses of bond strengths and electrostatic site potential, predicted that O1 would be a very favorable site to attach a proton, charge-balanced by metal vacancies. If all O1 sites were protonated then wadsleyite would be able to incorporate up to 3 wt% H2O (Smyth 1994) making it an enormous reservoir for water in the Earth’s mantle. A later study by Ross et al. (2003) carried out a similar investigation using the Laplacian of the electron density, as calculated from HF theory, to determine possible sites for protonation in wadsleyite and a number of other high pressure silicates. Their results also suggested that O1 would be the most likely site for protonation. However, Downs (1989), suggested that the O2 site could also be important. Calculations using P-DFT (Haiber et al. 1997) and interatomic potentials (Wright and Catlow 1996; Parker et al. 2004; Walker et al. 2006) methods confirm that O1 is the most favorable site for protonation, followed by O4, O3, and lastly by O2. The O1 site is quite isolated from silicon, and from other oxygens and is thus underbonded. Mulliken population analysis (Haiber et al. 1997) shows that O1 has an anomalously low charge that becomes more like normal oxygen when protonated. Magnesium vacancies, silicon vacancies and iron impurities have all been investigated with interatomic potential methods as possible charge compensating defects in wadsleyite. Mg1 has the lowest vacancy energy of the three magnesium positions, followed by Mg3 and Mg2, with Mg1 being approximately 95 kJ·mol−1 more stable than Mg2 (Walker et al. 2006). The solution reaction energies of wadsleyite and water are summarized in Table 4. The values of Parker et al. (2004) are for free energies calculated at simulated mantle conditions (1900 K and 15 GPa). This is in contrast to the other MM calculations discussed here, where it is the enthalpy that is calculated for 0 K and 0 GPa, so that the two are not directly comparable. However, the calculations all show that reactions leading to formation of the defect complex at the Mg3 site are most favorable. Table 4. Calculated solution reaction energies in wadsleyite. Equation # (5) (6)

Reaction energy per OH (kJ·mol−1) >5001 1

60

662

5883

2

1063

32

[1] Wright and Catlow (1996), [2] Parker et al. (2004), [3] Walker et al. 2006.

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A further reaction considered by Wright and Catlow (1996) involved the incorporation of hydrogen via the reduction of ferric iron as: x H 2O + 2 Fe•Mg + 2 OOx → 2 ( OH )O + 2 Fe Mg + •

1 O2( g ) 2

( 9)

Here, Fe3+ is present as an impurity replacing magnesium at a magnesium site and reacts with a water molecule leading to the formation of two OH groups and Fe2+. This reaction was predicted to be exothermic, and therefore a highly favorable mechanism for water incorporation in wadsleyite. Haiber et al. (1997) carried out QM molecular dynamics calculations on the optimized O1-H defect, predicting an O-H stretching frequency of 3180 cm−1 (± 30 cm−1), at a simulated temperature of 400 K. The MM calculations of Walker et al. (2006) give a value of 3506 cm−1 for the same defect at a temperature of 0 K, the difference between the two values being similar to that found between MM and QM/MM calculation on forsterite. Walker et al. (2006) also calculated OH frequencies for defects associated with magnesium and silicon vacancies and found that, as with forsterite, lower frequency vibrations were related to OH at magnesium vacancies, and higher ones to OH associated with silicon vacancies. These results are broadly supported by the experimental work of Jacobsen et al. (2005) and with that of Kohn et al. (2002). Ringwoodite. The final polymorph in this family, ringwoodite, γ-Mg2SiO4, is the major constituent of the mantle between 520 and 660 km depth, i.e., lower part of the transition zone. It has a cubic spinel structure, as illustrated on Figure 6. The oxygens in this structure are close-packed with silicon in tetrahedral sites (isolated) and magnesium in octahedral sites (Fig. 7). Some disorder over the cation sites is considered likely. High-pressure experiments have shown that this mineral can incorporate up to 2.7 wt% H2O in its structure in the form of OH groups (Kohlstedt et al. 1996; Bolfan-Casanova et al. 2000). To date, only two computational studies of OH in ringwoodite have appeared in the Figure 7. The unit cell of ringwoodite literature. The DFT study of Haiber et al. (1987) (Mg2SiO4), which has the spinel structure. quotes a protonation energy for the oxygen that is comparable to the O2 and O3 values for forsterite, but less favorable than any of the oxygen sites in wadsleyite. No further information on the spinel structure is given. The second study is based on interatomic potentials (Blanchard et al. 2005) and investigates a range of different hydrogen defect positions. In addition to OH associated with magnesium and silicon vacancies, the authors consider the influence of iron (Eqn. 9) and of cation disorder. This last reaction is described by (Blanchard et al. 2005): x

x // H 2O + Mg Mg + SiSix + 3OOx →  MgSi ( OH )2  + VMg + VO•• + SiO2

(10)

The calculated defect energies indicate that the most favorable mechanisms for hydrogen incorporation are coupled with reduction of iron (Eqn. 9) or with the creation of silicon vacancies. As with the other Mg2SiO4 polymorphs, binding energies between OH and cation vacancies are large and negative, so that isolated hydroxyls are not expected to occur in significant amounts. The solution reaction energies for Equations (5), (6), (9) and (10) calculated by Blanchard et al. (2005) are shown in Table 5 and indicate that substantial amounts of water could be incorporated in ringwoodite via reaction with ferric iron, and with silicon and magnesium vacancies.

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Summary of Mg2SiO4 polymorphs. InteTable 5. Calculated solution reaction grating the information obtained on hydrogen energies per OH in ringwoodite from defects by computational methods, some clear Blanchard et al. (2005). trends emerge. Firstly, it seems that the majority of hydrogen defects in forsterite, wadsleyite Equation # Energy (kJ·mol−1) and ringwoodite will be closely bound to cation vacancies and that the formation of isolated (5) −297 OH groups is not expected to be an important (6) −626 mechanism for the uptake of water in these (9) −680 minerals. In pure, iron free forsterite and ringwoodite, hydrogen can be easily incorporated (10) 170 through the formation of silicon defects. It has been suggested (Brodholt and Refson 2000) that reactions with water will drive the creation of silicon vacancies in olivine. A mechanism involving magnesium vacancies will be important in wadsleyite, especially the Mg3 site, and also in ringwoodite. Hydrogen incorporation via redox reactions with iron is exothermic in wadsleyite and ringwoodite and thus energetically favored. Grain boundaries and dislocations in forsterite are likely to contain high concentrations of hydrogen compared to the bulk thus deformation history could influence the ability of minerals to incorporate water. Finally, OHvibrational spectra for forsterite and wadsleyite indicate that higher frequencies (~3500-3800) are associated with silicon defect complexes, and low frequency vibrations (~3000-3300) with hydrogen at magnesium defects. In general, the results show excellent qualitative agreement between the different methods. Indeed, where results are comparable, as in the case of forsterite (Table 3), agreement is good quantitatively as well. However, it is worth pointing out that the majority of the above calculations have been performed at pressures of 0 GPa and thus the conclusions based on the results are only strictly valid for low pressure regimes.

Pyroxene Concentrations of OH in natural clinopyroxene samples range from 100-1300 ppm H2O (Skogby et al. 1990) with diopside, augite and omphacite showing the highest water concentrations of all NAMs, greater than olivine and pyrope garnet, which contain only trace amounts. Diopside (CaMgSi2O6) and jadeite (NaAlSi2O6) are monoclinic pyroxenes with space group C2/c. Each has two distinct cation sites; M1 is a regular 6-fold octahedral site, while M2, is distorted such that it becomes 8-coordinate, and three distinct oxygen sites. In diopside, the M1 site is generally occupied by Mg, and by Al in jadeite, while the M2 site is occupied by Ca and Na in diopside and jadeite respectively. There are also three different oxygen sites as shown on Figure 8. Gatzemeier and Wright (2006) have recently published the first study of hydrogen defects in clinopyroxenes. They use MM methods with interatomic potentials similar to those used by Wright and Catlow (1996) and Walker et al. (2006). Their calculations of intrinsic Schottky energies indicate that cation vacancies in both diopside and jadeite are more favored on the M2 than the M1 site, while the lowest energy oxygen vacancies are at different sites in the two phases—O2 in diopside and O3 sites in jadeite—although O2 is most easily

Figure 8. The unit cell of diopside and jadeite viewed along [001].

Wright

80

protonated in both. Reported solution reaction energies for various defect complexes are given in Table 6. In the pure phases hydrogen is most easily incorporated via the formation of [VSi(OH)4]x hydrogarnet type defects. When components of the two phases are mixed, then solution energies can become exothermic. The substitution of Al for Si and Na for Ca or Mg in diopside, provides favorable routes for hydrogen incorporation. In jadeite, Al rich compositions, with Al at Si sites, and the presence of both Ca and Mg at Al sites, also favor hydrogen incorporation, with exothermic values of solution energy. Thus the amount of water present in these minerals in the Earth’s upper mantle will vary with composition. Analysis of IR frequencies associated with O-H stretching at specific defect clusters in diopside and jadeite (Gatzemeier and Wright 2006) give hydrogen-oxygen bond lengths in good agreement with those correlated by Libowitzky (1999). Comparison of experimental and calculated IR frequencies were problematic, partly due to the complexity of experimental spectra, but also due to possible deficiencies in the ability of the model to accurately describe the O-H stretching frequency. Table 6. Calculated solution reaction energies per OH in diopside and jadeite from Gatzemeier and Wright (2006).

Equation #

Defect complex

Energy (kJ·mol−1)

Diopside (5) (6) (6) (7) (8) (8)

[VSi⋅4(OH) ] ° [VCa⋅2(OH) ]x ° [VMg⋅2(OH) ]x ° [AlSi⋅(OH) ]x °x [NaCa(OH) ] ° [NaMg⋅(OH) ]x ° x

Defect complex

Energy (kJ·mol−1)

Jadeite 87 227 203 −281 −105 −103

[VSi⋅4(OH) ]x ° [VAl⋅3(OH) ]x ° [VNa⋅2(OH) ]x ° [AlSi⋅(OH) ]x ° x [MgAl⋅(OH) ] ° [CaAl⋅(OH) ]x °

61 142 338 −335 −307 −277

General remarks and future directions This review has concentrated on studies of the more important minerals in the Earth’s upper mantle and transition zone that contain hydrogen defects. For forsterite and wadsleyite there is good general agreement between the different computational methods used and with experiment. Thus a comprehensive atomistic model for OH in these minerals is emerging. For all of the phases considered, it appears that: (a) hydrogen is closely bound with metal vacancies and with charged impurities; (b) the magnitude of binding energies is such that water could facilitate the formation of point defects; and (c) the concentration of hydrogen is closely linked to chemistry and to the populations of other defects, such as dislocation, present in NAMs. However, the effects of pressure on defect energetics and on calculated frequencies must be considered before the results can be applied realistically to minerals under mantle conditions. Clearly simulations will continue to play a critical role in understanding the incorporation of water into NAMs and its influence on their properties. Of particular importance is the calculation of IR frequencies, and of the interaction of hydrogen with extended defects such as dislocations. The accurate prediction of IR frequencies associated with specific defect complexes can only be obtained using QM methods, and with advances in hardware and software, such calculations on minerals with large unit cells now become possible. Modeling dislocations in complex ionic materials is still in its infancy however, and requires a simulation

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cell containing many thousands of ions (Walker et al. 2005), making it more suited to MM approaches. Therefore both QM and MM techniques will continue to provide insights into OH defect behavior in nominally anhydrous minerals at the atomic level.

ACKNOWLEDGMENTS The bulk of the work reviewed here was carried out in the UK and funded by the Natural Environmental Research Council, the Engineering and Physical Sciences Research Council, the Royal Society, and the European Union. The author is grateful for funding from all of these agencies during the last 15 years and would also like to acknowledge very fruitful collaborations with Richard Catlow over this time. In addition, thanks to input from former students and post-docs Andrew Walker, Alex Gatzemeier, Marc Blanchard and Spencer Braithwaite. Finally, thanks to Julian Gale for discussions and comments on the manuscript.

REFERENCES Austen KF, Wright K, Slater B, Gale JD (2005) The interaction of dolomite surfaces with impurities: A computer simulation study. Phys Chem Chem Phys 7:4150-4156 Bai Q, Kohlstedt DL (1993) Effects of chemical environment on the solubility and incorporation mechanism for hydrogen in olivine. Phys Chem Minerals 19:460-471 Baram PS, Parker SC (1996) Atomistic simulation of hydroxide ions in inorganic solids. Phil Mag B 73:49-58 Becke AD (1993) A missing of Hartree-Fock and local density-functional theories. J Chem Phys 98:1372-1377 Bell DR, Rossman GR (1992) Water in the Earth’s mantle: the role of nominally anhydrous minerals. Science 255:1391-1397. Berry AJ, Hermann J, O’Neill HSC, Foran GJ (2005) Fingerprinting the water site in mantle olivine. Geology 33:869-872 Blanchard M, Gale JD, Wright K (2005) A computer simulations of OH defects in Mg2SiO4 and Mg2GeO4 spinels. Phys Chem Minerals 32(8-9):585-593 Bolfan-Casanova N, Keppler H, Rubie DC (2000) Water partitioning between nominally anhydrous minerals in the MgO-SiO2-H2O system up to 24 GPa: implications for the distribution of water in the Earth’s mantle. Earth Plan Sci Lett 182:209-221 Braithwaite JS, Sushko PV, Wright K, Catlow CRA (2002) Hydrogen defects in Forsterite: A test case for the embedded cluster method. J Chem Phys 116(6):2628-2635 Braithwaite JS, Wright K, Catlow CRA (2003) A theoretical study of the energetics and IR frequencies of hydroxyl defects in forsterite. J Geophys Res Solid Earth 108(B6) article 2284 Brenner DW, Shenderova OA, Harrison JA, Stuart SJ, Ni B, Sinnott SB (2002) A second-generation reactive empirical bond order (REBO) potential energy expression for hydrocarbons. J Phys Condens Matter 14: 783-802 Brodholt JP, Refson K (2000) An ab initio study of hydrogen in forsterite and a possible mechanism for hydrolytic weakening. J Geophys Res 105:18977-18982 Collins R, Catlow CRA (1992) Computer simulation of the structure and cohesive properties of micas. Am Mineral 77:1172-1181 Cygan RT, Kubicki JD (eds) (2001) Molecular Modeling Theory: Applications to the Geosciences. Reviews in Mineralogy and Geochemistry 42. Mineralogical Society of America. Cygan RT, Guggenheim S, van Groos AFK (2004) Molecular models for the intercalation of methane hydrate complexes in montmorillonite clay. J Phys Chem B 108:15141-15149 de Leeuw NH, Parker SC, Catlow CRA, Price GD (2000) Proton-containing defects at forsterite {010} tilt grain boundaries and stepped surfaces. Am Mineral 85:1143-1154 de Leeuw NH, Parker SC (2001) Density functional theory calculations of proton-containing defects in forsterite. Rad Effects Defects Solids 154:255-259 Demouchy S, Mackwell SL (2003) Water diffusion in synthetic iron-free forsterite. Phys Chem Minerals 30: 486-494 Dick BG Jr, Overhauser AW (1958) Theory of the dielectric constants of alkali halide crystals. Phys Rev 112: 90-103 Downs JW (1989) Possible sites for protonation in β-Mg2SiO4 from an experimentally derived electrostatic potential. Am Mineral 74:1124-1129 Foresman JB, Frisch A (1996) Exploring Chemistry with Electronic Structure Methods. Gaussian Inc.

82

Wright

Gale JD, Rohl AL (2003) The General Utility Lattice Program (GULP). Molecular Simulation 29:291-341 Gatzemeier A, Wright K (2006) Computer modeling of hydrogen defects in the clinopyroxenes diopside and jadeite. Phys Chem Minerals 33:115-125 Griffiths DJ (2005) Introduction to Quantum Mechanics. Pearson Prentice Hall Haiber M, Ballone P, Parrinello M (1977) Structure and dynamics of protonated Mg2SiO4: An ab initio molecular dynamics study. Am Mineral 82:913-922 Islam MS, Driscoll DJ, Fisher CAJ, Slater PR (2005) Atomic-scale investigations of defects, dopants and lithium transport in the LiFePO4 olivine type battery material. Chem Mater 17:5085-5092 Jacobsen SD, Demouchy S, Frost DJ, Boffa-Ballaran T, Kung J (2005) A systematic study of OH in hydrous wadsleyite from polarized FTIR spectroscopy and single-crystal X-ray diffraction: oxygen sites for hydrogen storage in the Earth’s interior. Am Mineral 90:61-70 Kerisit S, Marmier A, Parker SC (2005) Ab initio surface phase diagram of the {10-14} calcite surface. J Phys Chem B 109(39):18211-18213 Kohlstedt DL, Keppler H, Ruby DC (1996) Solubility of water in the α, β and γ phases of (Mg,Fe)2SiO4. Contrib Mineral Petrol 123:345-357 Kohn SC (1996) Solubility of H2O in nominally anhydrous minerals using 1H MAS NMR. Am Mineral 81: 1523-1526 Kohn SC (2006) Structural studies of OH in nominally anhydrous minerals using NMR. Rev Mineral Geochem 62:53-66 Kohn SC, Brooker RA, Frost DJ, Slesinger AE, Wood BJ (2002) Ordering of hydroxyl defects in hydrous wadsleyite (beta-Mg2SiO4). Am Mineral 87:293-301 Kresse G, Furthmuller J (1996a) Efficient iterative schemes for ab initio total energy calculations using a planewave basis set. Phys Rev B 54:11169-11186 Kresse G, Furthmuller J (1996b) Efficiency of ab initio total energy calculations for metals and semiconductors using a planewave basis set. Comput Mater Sci 6:15-50 Kröger FA (1972) The Chemistry of Imperfect Crystals. North-Holland Leach AR (2001) Molecular Modeling Principles and Applications. Prentice Hall Leininger ML, Allen WD, Schaefer HF, Sherrill CD (2000) Is Moller-Plesset perturbation theory a convergent ab initio method? J Chem Phys 112(21):9213-9222 Lemaire C, Kohn SC, Brooker RA (2004) The effect of silica activity on the incorporation mechanisms of water in synthetic forsterite: a polarized infrared spectroscopic study. Contrib Mineral Petrol 147:48-57 Leslie M, Gillan MJ (1985) The energy and elastic dipole tensor of defects in ionic crystals calculated by the supercell method. J Phys Chem 18:973-982 Libowitzky E (1999) Correlation of O–H stretching frequencies and O–H···O hydrogen bond length in minerals. Monatshefte für Chemie 130:1047-1059 Libowitzky E, Beran A (2006) The structure of hydrous species in nominally anhydrous minerals: information from polarized IR spectroscopy. Rev Mineral Geochem 62:29-52 Lin JS, Payne MC, Heine V, McConnell JDC (1994) Ab initio calculations of (OH)4 defects in alpha-quartz. Phys Chem Minerals 21:150-155 Mackwell SL, Kohlstedt DL (1990) Diffusion of hydrogen in olivines: Implications for water in the mantle. J Geophys Res 95:5079-5088 Matveev S, O’Neill HSC, Ballhaus C, Taylor WR, Green DH (2001) Effect of silica activity on OH− IR spectra of olivine: implications for low-αSiO2 mantle metasomatism. J Petrol 42:721-729 Mott NF, Littleton MJ (1938) Conduction in polar crystals. I. Electrolytic conduction in solid salts. Trans Faraday Soc 34:485 Nobes RH, Akhmatskaya EV, Milman V, White JA, Winkler B, Pickard CJ (2000) An ab initio study of hydrogarnets. Am Mineral 85:1706-1715 Oganov AR, Price GD (2005) Ab initio thermodynamics of MgSiO3 perovskite at high pressure and temperature. J Chem Phys 122(12) art. No. 124501 Parker SC, Cooke DJ, Kerisit S, Marmier AS, Taylor SL, Taylor SN (2004) From HADES to PARADISE – atomistic simulations of defects in minerals. J Phys Cond Matter 16:S2735-S2749 Purton J, Jones R, Heggie M, Oberg S, Catlow CRA (1992) Pseudopotential calculations of the alpha-quartz structure and hydrogarnet defect. Phys Chem Minerals 18:389-392 Putnis A (1992) Introduction to Mineral Sciences. Cambridge University Press Ross NL, Gibbs GV, Rosso KM (2003) Potential docking sites and positions of hydrogen in high-pressure silicates Am Mineral 88(10):1452-1459 Rossman GR (2006) Analytical methods for measuring water in nominally anhydrous minerals. Rev Mineral Geochem 62:1-28 Saul P, Catlow CRA, Kendrick J (1985) Theoretical studies of protons in sodium hydroxide. Philos Mag B 51: 107-117

Atomistic Models of OH in NAM’s

83

Schröder KP, Sauer J, Leslie M, Catlow CRA, Thomas JM (1992) Bridging hydroxyl groups in zeolitic catalysts: a computer simulation study. Chem Phys Lett 188:320-325 Segall MD, Lindan PJD, Probert MJ, Pickard CJ, Hasnip PJ, Clarke SJ, Payne MC (2002) First-principles simulation: ideas, illustrations and the CASTEP code. J Phys Cond Matter 14:2717-2743 Skogby H, Bell DR, Rossman GR (1990) Hydroxide in pyroxene: variations in the natural environment. Am Mineral 75:764–774 Smyth JR (1987) β-Mg2SiO4: a potential host for water in the mantle? Am Mineral 72:1051-1055 Smyth JR (1994) A crystallographic model for hydrous wadsleyite (β-Mg2SiO4): an ocean in the Earth’s interior? Am Mineral 79:1021-1024 Smyth JR, Kawamoto T, Jacobson SD, Swope RJ, Hervig RL, Holloway JR (1997) Crystal structure of monoclinic hydrous wadsleyite [β-(Mg,Fe)2SiO4]. Am Mineral 82:270-275 Steele HM, Wright K, Hillier IH (2002) Modeling the adsorption of uranyl on the surface of goethite. Geochimica Cosmochim Acta 66:1305-1310 Stillinger FH, David CW (1980) Study of the water octamer using the polarization model of molecular interactions. J Chem Phys 73:3384-3389 Tilley RJD (1987) Defect Crystal Chemistry and its Applications. Blackie and Sons Ltd. Vocadlo L, Alfe D, Gillan MJ, Price GD (2003) The properties of iron under core conditions from first principles calculations. Phys Earth Planet Ints 140:101-125 Walker AM, Wright K, Slater B (2003) A computational study of oxygen diffusion in olivine. Phys Chem Minerals 30:536-545 Walker AM, Slater B, Gale JD, Wright K (2005) Atomic scale modeling of the cores of dislocations in complex materials part 2: applications. Phys Chem Chem Phys 7(17):3235-3242 Walker AM, Demouchy S, Wright K (2006) Computer simulation of hydroxyl groups in α- and β-Mg2SiO4. Euro J Mineral (in press) Wright K, Freer R, Catlow CRA (1994) Energetics and structure of the hydrogarnet defect in grossular: A computer simulation study. Phys Chem Minerals 20:500-504 Wright K, Catlow CRA (1994) A computer simulation study of OH defects in olivine. Phys Chem Minerals 20: 515-518 Wright K, Catlow CRA (1996) Calculations on the energetics of water dissolution in wadsleyite. Phys Chem Minerals 23:38-41 Wright K, Catlow CRA, Freer R (1996) Water related defects and oxygen diffusion in albite. Contrib Mineral Petrol 125:161-166 Wright K, Cygan RT, Slater B (2002) Impurities and non-stoichiometry in the bulk and on the (10-14) surface of dolomite. Geochimica Cosmochim Acta 66:2541-2546 Wright K, Gale JD (2004) Interatomic potentials for the simulation of the zinc-blende and Wurtzite forms of ZnS and CdS: Bulk structure, properties and phase stability. Phys Rev B 70:03521

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Reviews in Mineralogy & Geochemistry Vol. 62, pp. 85-115, 2006 Copyright © Mineralogical Society of America

Hydrogen in High Pressure Silicate and Oxide Mineral Structures Joseph R. Smyth Department of Geological Sciences University of Colorado Boulder, Colorado, 80309, U.S.A. e-mail: [email protected]

INTRODUCTION Earth is the water planet. Liquid water covers more than 70% of the surface and dominates all surface processes, geological, meteorological, and biological. However the hydrosphere composes only about 0.025% of the planet’s mass, so that small amounts of H incorporated into the oxygen minerals of the interior may constitute the majority of Earth’s total water. The Earth is thought to be generally similar in composition to the chondrite meteorites which average about 0.10% by weight H2O. So if the Earth were strictly chondritic in its H content, about 75% of that H as water would have either been tied up in the minerals of the interior or lost to space. Understanding how H behaves at the atomic scale in these materials will help us to understand how the Earth balances and retains its water and may help us to understand how water planets develop and how common they might be. In addition to the surface processes, water also controls the processes of the interior. Water dramatically reduces the melting temperature of rocks controlling igneous processes. Even trace amounts of hydrogen have a major effect on some physical properties such as deformation strength and electrical conductivity (Karato 1990). The nominally anhydrous minerals of the Earth’s interior are capable of incorporating many times the amount of water in the hydrosphere, and these phases would need to be saturated before stoichiometrically hydrous minerals could be stable. Hydrogen in amounts reported in olivine, wadsleyite, and ringwoodite by Kohlstedt et al. (1996) as recalibrated by Bell et al. (2003), if present in the Earth, would constitute a significant fraction of the total water budget of the planet. The amounts that can be incorporated into the nominally anhydrous minerals of the Transition Zone (410-660 km depth) may constitute the largest reservoir of water in the planet and may have controlled the chemical evolution and interior processes of the planet. Hirschmann et al. (2005) have estimated the storage capacities of the various mineral reservoirs in the mantle. These volumes of water imply that there may be a deep water cycle in the Earth whereby some of the water in subducted slabs may be returned to the large deep interior reservoir and then be released in mid-ocean ridge basalts so that the amount of water in the Earth’s oceans would represent a dynamic balance between these processes. This process would depend on the ability of the nominally anhydrous phases of the upper mantle to incorporate the water released by the breakdown of the hydrous phases on increasing pressure and temperature with subduction.

Geochemistry of H Hydrogen is the most abundant element in the cosmos, and the geochemical behavior of hydrogen is unlike that of any other element. Because the proton does not behave like other cations in the crystal, it is generally inappropriate to treat H as an incompatible element 1529-6466/06/0062-0005$05.00

DOI: 10.2138/rmg.2006.62.5

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or to compare its compatibility with other cations. In the highly reducing conditions of the condensing solar nebula, H was primarily atmophile, as the diatomic gas H2, but also as methane, ammonia, and water. However, in the Earth’s crust and mantle, hydrogen in its ionic state, H+, is strongly lithophile. It substitutes readily in silicates and other oxygen minerals in both trace and stoichiometric amounts. Because H does not occupy a normal cation site in a mineral, it does not have an effective ionic radius that controls its geochemical behavior. Its compatibility is therefore not systematic as other trace cations are, but strongly dependent on temperature, pressure, and the chemical activity of possible charge-balancing cations. The chalcophile nature of H is not well known, and its substitution in sulfide minerals in trace amounts is difficult to measure and poorly studied. H2S is an abundant volatile in mafic to silicic volcanic systems, and there are a few OH-bearing sulfide minerals such as tochilinite [6(Fe0.9S)·5(Fe,Mg)(OH)2)] (Beard 2000), but I was unable to identify a single H-bearing sulfide mineral that does not also contain oxygen. Under reducing conditions, neutral H is highly soluble in metallic liquids and forms solid metal hydrides. However, very little is known about H partitioning between silicate and metallic liquids, and the amount of H in the core is unknown as is its effect on liquid metal densities under conditions of the core. The objectives of this review are to examine the various structural substitution mechanisms whereby H enters major high pressure silicate and oxide minerals in stoichiometric amounts and then use this information to look at H substitution in nominally anhydrous minerals of the Earth’s mantle.

Crystal Chemistry of H Because oxygen is the only anionic species of significant abundance in the crust and mantle, we think of hydrogen and water as synonymous. At low pressure, water can enter silicates either as molecular water or as hydroxyl, or both. In low-density silicates such as zeolites and clays, the water molecules are located in large cavities or interlayer sites and freely flow into and out of the crystals. In other low-temperature minerals such as gypsum, the molecular water is tightly bound structurally and does not exchange. Hydrogen also enters low temperature and pressure minerals as structural hydroxyl. At higher pressures, hydrogen occurs in the solid minerals of the mantle in several forms, but generally does not exchange. It can be present as discrete, structurally bound water molecules as in lawsonite, K-cymrite, or 10 Å phase, but most often it is present as hydroxyl, OH−. The hydroxyl can be stoichiometric, part of the nominal mineral formula, or it can be a minor constituent, where the hydrogen may substitute ionically for other cations in the structure. In nominally hydrous silicates, the hydroxyl rarely bonds directly to the Si cation. This is also true of other small, high-fieldstrength cations such as B, C, P, and S6+. The proton position is difficult to locate by X-ray diffraction, but neutron single-crystal or powder diffraction can give proton positions with high precision. Additionally, the protonated oxygen is relatively easy to identify from X-ray data by a simple Pauling bond strength calculation. When H enters nominally anhydrous minerals, the proton does not occupy the normal cation position, but attaches to one or more of the oxygens. The usual proton to oxygen nucleus distance (0.95 to 1.2 Å) is less than the nominal oxygen radius (1.32 to 1.4 Å). The oxygen atoms that can be protonated in a stoichiometric, fully occupied structure are those that are most underbonded. The degree of underbonding can be calculated on the basis of Pauling bond strength or a Madelung site potential calculation. Pauling bond strength at the oxygen is calculated as the sum of the bond strengths (nominal cation valence divided by coordination number) around an oxygen atom. The Madelung site potential (Smyth 1987, 1989) is the nominal valence charge divided by distance and summed to convergence. These methods may identify the oxygen most likely to be protonated if there are several non-equivalent oxygen positions in a structure, but does not identify the proton location. Libowitzky (1999) reports a correlation of O-H-O distance with O-H stretching frequency. This has been used together

Hydrogen in High Pressure Silicate & Oxide Mineral Strutures

87

with polarization vectors to deduce proton positions in nominally anhydrous structures such as wadsleyite (Kohn et al. 2002) and akimotoite (Bolfan-Casanova et al. 2002). Ross et al. (2003) propose a computational method to identify non-bonding electron-pairs on oxygens in order to locate potential docking sites for protons in high pressure silicates. Extensive protonation of an oxygen site in a nominally anhydrous mineral generally requires a charge balancing substitution or a cation vacancy. Cation vacancies normally result in a significant expansion of the vacant coordination polyhedron, however such vacant polyhedra are typically large and highly compressible (Jacobsen 2006). Tetrahedral cation vacancies charge-balanced by protons are well documented. This is the so-called hydrogarnet substitution because the H4O4 tetrahedron can completely replace the silicate tetrahedron in hydrogarnets (Lager et al. 2005). The H4O4 tetrahedron is larger than the silicate tetrahedron with the Si-O distance in silicate garnets being 1.60 to 1.64 Å whereas the equivalent distance (4-O) in hydrogarnet is over 2.0 Å. This means that pressure will inhibit this substitution mechanism so that garnets from natural high pressure (2-5 GPa) environments generally contain less than about 50 ppmw H2O (Bell and Rossman 1992). It may be possible that this substitution mechanism may again become viable at pressures above about 7 GPa, as it has been proposed to be present in hydrous coesite above this pressure (Koch-Mueller et al. 2003). Oxygen-oxygen edges are typically 2.6 to 2.8 Å for tetrahedral silicon, whereas edges of Mg octahedra typically are 2.8 to 3.0 Å. These distances have been used to infer proton positions from infrared spectra based on the calibration of Libowizky (1999) (e.g., Kohn et al. 2002), however the correlation curve is quite flat in this region and cation vacancy may result in local distortion of the coordination polyhedra. Octahedral cation vacancy charge-balanced by protons appears to be more common at pressures of the upper mantle. Protonated octahedral cation vacancy appears to become a very significant substitution mechanism in olivine (Smyth et al. 2006a) and wadsleyite (Smyth et al. 1997; Ross et al. 2003; Jacobsen et al. 2005). Wadsleyite (β-Mg2SiO4) can contain more than 3 wt% H2O (Inoue et al. 1995; Kohlstedt et al. 1996), where the charge balance mechanism is octahedral site vacancy, principally at M3 (Smyth et al. 1997; Kohn et al. 2003). Even trace hydration (10 to 1000 ppmw H2O) can have very large effects of physical properties of nominally anhydrous phases such as mechanical strength (Kavner 2003), effective viscosity (Karato et al. 1986), and electrical conductivity (Karato 1990; Huang et al. 2005). Minor hydration (1000 to 10000 ppmw H2O) can have a major effect on density, compressibility (Smyth et al. 2003; 2004), seismic velocity (Jacobsen et al. 2005; 2006), and pressure-temperature conditions of phase transitions (Wood 1995; Smyth and Frost 2002). In order to understand the crystal chemistry of H at high pressure, it is necessary to first look briefly at the nominally hydrous phases on the Earth’s mantle and then to examine the mechanisms for minor and trace substitution of H in the nominally anhydrous silicates and oxides that compose the mantle. The dense hydrous magnesium and aluminum silicate phases covered here are listed in Table 1 along with formulae, cell parameters and calculated densities. The dense anhydrous magnesium and aluminum silicate phases covered here are listed in Table 2.

Nominally Hydrous High-Pressure Silicate Phases Compositions of the dense hydrous magnesium silicate (DHMS) phases can be displayed in the magnesia-silica-brucite (MgO-SiO2-Mg(OH)2) ternary (Fig. 1). Along the anhydrous edge fall periclase (MgO), anhydrous phase B, forsterite and its polymorphs (wadsleyite and ringwoodite), enstatite and its polymorphs (akimotoite, and perovskite-type MgSiO3) and quartz and its polymorphs (coesite, and stishovite). On the brucite-forsterite join lie phase A and the humites (norbergite, chondrodite, humite and clinohumite). Near the bruciteanhydrous phase B join, lie phase B and super-hydrous phase B (Fig. 1).

Formula

Mg(OH)2 Mg3Si2O5(OH)4 Mg3Si4O10(OH)2 Mg5Al2Si3O10(OH)8

KAl2AlSi3O10(OH)2 KMg3AlSi3O10(OH)2 KMgAlSi4O10(OH)2

(Mg,Fe)7Si8O22(OH)2 Ca2Mg5Si8O22(OH)2

CaAl2Si2O7(OH)2·H2O

Ca2Al3Si3O12(OH) Ca2Al3Si3O12(OH) Ca2FeAl2Si3O12(OH)

Mg9Si4O16(OH)2 Mg5Si2O8(OH)2

Mg7Si2O8(OH)6 Mg12Si4O19(OH)2 Mg10Si3O14(OH)4 MgSi2O4(OH)2 Mg2SiO2(OH)4

AlSiO3(OH) Al3Si2O7(OH)3 Al2SiO4(OH)2

KAlSi3O8 H2O

Mineral

Brucite Serpentine Talc Chlorite

Mica Group Muscovite 2M1 Phlogopite1M Phengite 2M1

Amphibole Group Cummingtonite Tremolite

Lawsonite

Epidote Zoisite Clinozoisite Epidote

Humite Group Clinohumite Chondrodite

Phase A Phase B Suphyd. Phase B Phase D Phase E

Phase Egg Phase Pi Topaz-OH

K-Cymrite

296.356

120.074 300.137 180.063

456.398 742.096 619.403 178.499 176.739

621.162 339.744

454.366 454.366 483.232

314.243

843.953 812.419

398.317 417.290 396.753

58.327 277.137 379.294 555.838

F.W. (g)

7.1441 6.0885 4.7203

7.868 10.588 5.089 4.745 2.967

4.741 4.7459

16.188 8.861 8.888

8.795

9.5220 9.863

5.192 5.308 5.205

3.142 5.332 5.290 5.327

a (Å)

P6/mmm 5.3348

P21/n P1 Pbnm

P63 P21/c Pnnm P 31m R 3m

P21/b P21/b

Pnma P21/m P21/m

Ccmm

C2/m C2/m

C2/c C2/m C2/c

P 3 m1 P31m C1 C2/m

S.G.

5.3348

4.3346 7.2832 8.9207

7.868 14.097 13.968 4.775 2.967

10.275 10.3480

5.550 5.583 5.628

5.847

18.1833 18.048

9.015 9.190 9.037

3.142 5.332 9.173 9.227

b (Å)

7.7057

6.9525 7.7234 8.4189

9.577 10.073 8.696 4.345 13.886

13.704 7.9002

10.034 10.141 10.152

13.142

5.3184 5.285

20.046 10.166 19.886

4.766 7.223 9.460 14.327

c (Å)

90

90 115.71 90

90 90 90 90 90

100.1 108.70

90 90 90

90

90 90

90 90 90

90 90 90.46 90

α (°)

90

98.40 88.85 90

90 104.1 90 90 90

90 90

90 115.46 115.38

90

102.020 104.79

95.74 100.10 95.62

90 90 98.68 96.81

β (°)

2 4 2 1 1

2 2

4 2 2

4

2 2

4 2 4

1 1 2 2

Z

120

1

90 4 92.89 2 90 4

120 90 90 120 120

90 90

90 90 90

90

90 90

90 90 90

120 120 90.09 90

γ (°)

114.372

32.066 92.795 53.371

154.569 219.530 186.122 59.287 73.613

201.006 116.094

135.719 136.388 138.147

101.745

271.634 273.88

140.546 147.012 140.133

24.54 107.24 136.63 210.540

MVol (cm3)

Table 1. Physical properties of nominally hydrous high pressure silicate phases.

2.591

3.744 3.234 3.373

2.952 3.380 3.327 3.010 2.401

3.089 2.926

3.347 3.331 3.497

3.088

2.879 2.966

2.834 2.838 2.831

2.377 2.584 2.776 2.639

Fasshauer et al. (1997)

Schmidt et al. (1998) Wunder et al. (1993) Northrup et al. (1994)

Kagi et al. (2000) Finger et al. (1993) Pacalo and Parise (1992) Yang et al. (1997) Shieh et al. (2000)

Ross and Crichton (2001) Ross and Crichton (2001)

Grevel et al. (2000) Pawley et al. (1996) Gabe et al. (1973)

Baur (1978)

Yang et al. (1998) Hawthorne et al. (1976)

Rothbauer (1971) Hazen and Burnham (1973) Smyth et al. (2000)

Zigan and Rothbauer (1967) Mellini (1982) Perdikatsis and Burzlaff (1981) Smyth et al. 1997

ρ STP References (g/cm3)

88 Smyth

140.709 140.709 140.709 203.779

864.789

SiO2 SiO2 SiO2

Mg2SiO4 Fe2SiO4

Mg2SiO4 Mg2SiO4 Mg2SiO4 Fe2SiO4

Mg14Si5O24

Mg3Al2Si3O12 Fe3Al2Si3O12 Mg3(MgSi)Si3O12

Mg2Si2O6 Fe2Si2O6 Mg2Si2O6 NaAlSi2O6 CaMgSi2O6

Mg2Si2O6 Al2SiO5

MgSiO3 MgSiO3 ZrSiO4 CaTiSiO5

Quartz Coesite Stishovite

Olivine Forsterite Fayalite

Wadsleyite Wadsleyite II Ringwoodite γ-Fe2SiO4

Anhyd. Phase B

Garnets Pyrope Almandine Majorite

Pyroxenes Orthoenstatite Orthoferrosilite Clinoenstatite Jadeite Diopside

Akimotoite Kyanite

Perovskite Post-perovskite* Zircon Titanite

4.211 4.311 4.7589 4.5845

a (Å)

4.775 2.65 6.6042 7.069

4.728 7.126

R3 P1

Pbnm Cmcm I41/amd P21/a

18.227 18.427 9.618 9.423 9.750

11.452 11.531 11.501

5.908

5.711 5.6884 8.092 8.234

4.753 4.820

Pbca Pbca P21/c C2/c C2/c

Ia 3 d Ia 3 d I41/a

Pmcb

Imma Imma Fd 3 m Fd 3 m

Pbnm Pbnm

P3121 4.914 C2/c 7.137 P42/mnm 4.179

Fm 3 m Fm 3 m R 3c P42/mnm

S.G.

* decompressed from Murakami et al. (2004) assuming K = 300 GPa and K′ = 5

100.397 100.397 183.304 196.063

200.795 162.047

200.795 263.865 200.795 202.139 232.330

403.153 497.758 401.590

140.709 203.779

60.086 60.086 60.086

40.311 71.846 101.961 79.899

MgO FeO Al2O3 TiO2

Periclase Wüstite Corundum Rutile

F.W. (g)

Formula

Mineral

4.929 8.69 6.6042 8.722

4.728 7.852

8.819 9.076 8.815 8.564 8.926

11.452 11.531 11.501

14.241

11.467 28.924 8.092 8.234

10.190 10.479

4.914 12.370 4.179

4.211 4.311 4.7589 4.5845

b (Å)

6.908 6.59 5.9796 6.566

13.559 5.575

5.179 5.237 5.176 5.223 5.251

11.452 11.531 11.480

10.069

8.256 8.2382 8.092 8.234

5.978 6.087

5.405 7.174 2.665

4.211 4.311 12.9912 2.9533

c (Å)

90 90 90 90

90 89.99

90 90 90 90 90

90 90 90

90

90 90 90 90

90 90

90 90 90

90 90 90 90

α (°)

90 90 90 113.86

90 101.11

90 90 108.37 107.56 105.90

90 90 90

90

90 90 90 90

90 90

90 119 90

90 90 90 90

β (°)

90 90 90 90

120 106.03

90 90 90 90 90

90 90 90

90

90 90 90 90

90 90

120 90 90

90 90 120 90

γ (°) 11.242 12.060 25.577 18.693

MVol (cm3)

4 4 4 4

3 4

8 8 4 4 4

8 8 8

2

8 20 8 8

4 4

24.482 22.89 39.264 55.739

52.700 44.219

62.666 65.930 62.704 60.498 66.167

113.056 115.412 114.305

255.081

40.699 40.812 39.886 42.023

43.596 46.283

3 22.685 16 20.573 2 14.014

4 4 6 2

Z

Table 2. Physical properties of nominally anhydrous high pressure silicate phases.

4.100 4.386 4.668 3.517

3.810 3.664

3.204 4.002 3.202 3.341 3.511

3.565 4.312 3.513

3.390

3.457 3.447 3.527 4.848

3.227 4.402

2.648 2.920 4.287

3.585 5.956 3.986 4.274

Horiuchi et al. (1987) Murakami et al. (2004) Hazen and Finger (1976) Kek et al. (1995)

Horiuchi et al. (1982) Winter and Ghose (1979)

Yang and Ghose (1995) Sueno et al. (1976) Pannhorst (1984) Cameron et al. (1973) Cameron et al. (1973)

Armbruster et al. (1992) Armbruster et al. (1992) Angel et al. (1989)

Hazen et al. (1992)

Finger et al. (1993) Smyth et al. (2005) Smyth et al. (2004) Yagi et al. (1974)

Smyth et al. (2006) Fujino et al. (1981)

Kihara (1990) Smyth et al. (1987) Ross et al. (1990)

Hazen (1976) Hazen (1981) Newnham and DeHaan (1962) Shintani et al. (1975)

ρ STP References (g/cm3)

Hydrogen in High Pressure Silicate & Oxide Mineral Strutures 89

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90

Figure 1. The compositions of the dense, hydrous and anhydrous magnesium silicate phases displayed on the MgO-SiO2-Mg(OH)2 ternary. W denotes the field for wadsleyite, olivine, and ringwoodite; aB is anhydrous phase B; B is phase B, sB is superhydrous phase B; CH clinohumite; H humite, Chd chondrodite; Nb norbergite; A phase A; E phase E; and D phase D.

Brucite Brucite, Mg(OH)2 is the first phase discussed among the nominally hydrous minerals of the mantle. Although brucite is not a silicate, it forms a prominent structural component in many silicate minerals. Because of its very high water content, more than 30% by weight or about 75% water by volume, it is not a likely mantle mineral. Brucite forms the most hydrous end member in our systems and is a common ingredient in starting compositions to experimentally produce hydrous high pressure phases. This component, with some Al substitution also occurs in the chlorite structure. The brucite structure (Fig. 2) is trigonal, P 3 m1, and consists of tri-octahedral layers of Mg(OH)6 octahedra parallel to (001). All oxygens in the structure are equivalent and protonated, so that each oxygen is bonded to three Mg atoms and one proton. The layers are bonded by relatively weak hydroxyl bonds giving the mineral its perfect basal cleavage. Gibbsite, Al(OH)3, is isostructural with brucite except that one third of the octahedra are vacant.

Serpentine Serpentine, ideally Mg3Si2O5(OH)4, is a major alteration phase in ultramafic rocks. It is stable at ambient pressure and to depths of roughly 250 km in a cool, subducting slab (Kawamoto et al. 1996; Schmidt and Poli 1998). It contains roughly 13% H2O by weight which corresponds to more than 30% by volume. The structure consists of a tri-octahedral brucite-like layer attached to a single pure-silica tetrahedral layer (Fig. 3). The structure has several stacking polytypes, but most are similar in composition and density. In its absestiform habit known as chrysotile, the sheets are rolled into tubes, so that the actual space groups and structure are not well defined. Several stacking polytypes have been described with differing degrees of order each having a density of about 2.58 g/cm3. Lizardite 1H is trigonal P31m (Table 1) (Mellini 1987). The well crystallized massive form is known as antigorite, the space group is triclinic, P1. There are no Si-OH bonds in the structure, so that each oxygen is bonded either to three Mg and one Si or to three Mg and one proton.

Talc Talc, Mg3Si4O10(OH)2, is also a major alteration phase of mafic rocks. It contains more silica than serpentine and may occur in more siliceous rock compositions than serpentine. The water content is a bit less than 5% by weight. Its triclinic, C 1 structure (Fig. 4) is that of a T-O-T layer silicate like mica, but without interlayer cations. The bonding between layers is just the weak Van der Waals bonds resulting in a very soft and easily deformable structure. There are no Si-OH bonds in the structure so that one sixth of the oxygen atoms are protonated and bonded to one proton and three Mg atoms. The remaining oxygen atoms are each bonded to one Si and three Mg atoms.

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Figure 2. The structure of brucite (Mg(OH)2) is trigonal, P 3 m1. All oxygen atoms are equivalent and bonded to one H and three Mg atoms. The Mg octahedra are arranged in a sheet parallel to (001). The sheets are H-bonded together giving the mineral its perfect basal cleavage.

Figure 3. The simple trigonal structure of the serpentine mineral lizardite, Mg3Si2O5(OH)4, is P31m. The octahedral Mg atoms are arranged in a trioctahedral sheet as in brucite. All nonsilicate oxygens are protonated.

Figure 4. The structure of talc, Mg3Si4O10(OH)2, is triclinic, C1. The octahedral Mg atoms are arranged in a trioctahedral sheet as in brucite and serpentine except tat there are tetrahedral sheets on both sides of the octahedral sheet. Again, all non-silicate oxygens are protonated.

True micas The true micas (Fig. 5) have a T-O-T layer like that of talc, but one fourth of the Si cations are replaced by Al and charge balanced by an interlayer alkali cation, dominantly K. Like talc, the micas contain 4.5 to 5% H2O by weight. In muscovite, KAl2AlSi3O10(OH)2, the octahedral layer is dioctahedral with two Al cations, whereas in biotite and phlogopite, KMg3AlSi3O10(OH)2, it is trioctahedral with three divalent cations, Mg or Fe, per formula unit. The phengite substitution into the dioctahedral micas puts additional silicon into the tetrahedral layer in place of Al which is charge-balanced by Mg in the dioctahedral layer. This substitution is stabilized by pressure, and high-silica phengites have been synthesized at pressures as high as 11 GPa (Domanik and Holloway 1996; Smyth et al. 2000). Phengite is stable in a mafic composition to over 300 km depth if K is present and temperatures are low as in a subducting slab. The micas exist in several polytypes, that is, different stacking sequences,

92

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Figure 5. The structure of phlogopite 1M, KMg3AlSi3O10(OH)2 is monoclinic C2/m. One third of the Si atoms in the tetrahedral layer are replaced by Al and charge-balanced by the interlayer K atoms (gray sphere). Muscovite, KMg3AlSi3O10(OH)2, is similar except that one third of the octahedra are vacant and the rest replaced by Al. There are several distinct stacking arrangements called polytypes.

predominantly 2M1 (C2/c) and 3T (P3112) in dioctahedral micas, and 1M (C2/m) and 2M1 in trioctahedral micas. The different polytypes commonly coexist in natural samples and are so close in physical properties that separate stability fields for the different polytypes have not been documented. Again, there are no Si-OH bonds and protons coordinate the non-silicate oxygens in the octahedral layer. The 10 Å phase, Mg3Si4O10(OH)·H2O, is a mica-like dense hydrous magnesium silicate phase that occurs at 3-5 GPa as a breakdown product of serpentine and chlorite (Yamamoto and Akimoto 1977). It is structurally similar to talc and phlogopite, but has neutral molecular water in the inter-layer (Fumagalli et al. 2001; Comodi et al. 2005). It is likely to be an important host phase for H in subducting hydrated lithosphere (Fumagalli and Poli 2005).

Chlorite Chlorite, Mg3AlSi3O10(OH)2·Mg2Al(OH)6, is another low pressure alteration phase of mafic and ultramafic rocks. Like talc, it is stable to about 100 km depth, but is distinct from talc in its Al-content. The structure (Fig. 6) is monoclinic C2/m or triclinic, C 1, and consists of trioctahedral talc-like layer, but with one fourth of the tetrahedral sites occupied by Al instead of Si, giving the layer a net negative charge. Instead of an interlayer cation as in micas, there is a trioctahedral brucite-like layer with one third of the octahedra occupied by Al instead of Mg giving the layer a net positive charge. In the brucite-like layer, all of the oxygen atoms are protonated, whereas in the talc-like later one-sixth of the oxygens are protonated. Again, there are no Si-OH bonds in the structure. Chlorite, like serpentine, contains about 13% H2O by weight.

Amphiboles The amphiboles A0-1X7Y8O22(OH)2, are complex hydrous chain silicate minerals of high grade metamorphic and igneous rocks in which A is an alkali cation, X an octahedral divalent or trivalent cation, and Y is tetrahedral Si or Al. The structure (Fig. 7) is based on a double tetrahedral chain parallel to c. Again, there are no Si-OH bonds and all non-silicate oxygens (one in 12) are protonated. Amphiboles are stable in subducting lithosphere to about 3 GPa (Kawamoto et al. 1996; Schmidt and Poli 1998), so they are not expected to be major hosts for H in the sub-lithospheric mantle. Amphibole-like double chain defects are relatively common in pyroxenes at low pressures and so may be a water-carrying defect in mantle pyroxenes.

Lawsonite Lawsonite, CaAl2Si2O7(OH)2·H2O, contains molecular water as well as hydroxyl. The structure (Fig. 8) is orthorhombic, Ccmm, and is a sorosilicate with Si2O7 groups, Al in octahedral coordination, and Ca in 8-coordination. Again, there are no Si-OH bonds, and the two of the non-silicate oxygen atoms coordinating Al are hydroxyls, and two of the oxygens coordinating Ca are water molecules. Lawsonite is a common hydrous alteration product of mafic igneous rocks, replacing calcic plagioclase feldspar. The total water content of lawsonite

Hydrogen in High Pressure Silicate & Oxide Mineral Strutures

93

Figure 6. The structure of chlorite, Mg3AlSi3O10(OH)2 Mg2Al(OH)6 is monoclinic, C2/m, or triclinic, C1. There are talc-like layers interspersed with brucitelike layers. One third of the Si atoms are replaced by Al giving the talc layer a net negative charge, and one third of the Mg atoms in the brucite-like layer are replaced by Al to give that layer a net positive charge. There are no Si-O-H bonds and all non silicate oxygens are protonated.

Figure 7. The structure of the amphibole tremolite, Ca2Mg5Si8O22(OH)2, is monoclinic C2/m. The only non-silicate oxygen is protonated (black).

Figure 8. The structure of lawsonite, CaAl2Si2O7(OH)2·H2O is orthorhombic, Ccmm. Lawsonite is a sorosilicate containing Si2O7 groups. The structure contains molecular water as well as hydroxyl. The Ca atom (gray sphere) is 8-coordinated, whereas the Al is octahedral, and the Si tetrahedral.

is high at about 11.5% by weight, and it is stable to relatively high pressures (~10 GPa) and low temperatures (Pawley 1994). Despite its high water content, it is about 10% denser than anorthite, and relatively incompressible with an isothermal bulk modulus of 122 GPa (Boffa-Balaran and Angel 2003). Being stable to depths of 300 km in the crustal portion of a subducting slab, lawsonite may act as a major conduit for water in the crustal portion of the slab to depths approaching those of the transition zone.

Epidote The epidote group comprises epidote (Ca2(Al,Fe)3Si3O12(OH)), zoisite, and clinozoisite (Ca2Al3Si3O12(OH)). Epidote is a very common metamorphic alteration product of mafic igneous rocks, whereas zoisite and clinozoisite are more restricted in composition and occurrence to aluminous and peraluminous rocks. The pressure stability ranges from less than 0.1 GPa to near 7 GPa (Poli and Schmidt 2004). There are also Mn-rich varieties (piemontite), and rare-earth-rich (allanite) varieties as well as several more named chemical variants (Franz and Liebscher 2004). The structure (Fig. 9) is monoclinic, P21/m (Z = 2), and has both isolated SiO4 tetrahedra as well as Si2O7 groups, so it is classed as a sorosilicate. Zoisite is orthorhombic, Pnma, with a nearly identical structure, but twice the unit cell volume (Z = 4). Most of the iron

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Figure 9. The structure of epidote, CaAl2FeSi3O12(OH), and clinozoisite, CaAl3Si3O12(OH), is monoclinic P21/m. The only non-silicate oxygen has a proton (black).

is ferric, and epidote has all of its trivalent cations in octahedral coordination, so it is also denser than anorthite. The O10 position is a non-silicate oxygen and is protonated. There is another non-silicate oxygen, O4, which is bonded to three trivalent metal octahedra. This oxygen is not protonated directly but shares a longer hydrogen bond to the proton on O10.

Humite The humite group comprises norbergite (Mg3SiO4(F,OH)2), chondrodite (Mg5Si2O8 (F,OH)2), humite (Mg7Si3O12(F,OH)2), and clinohumite (Mg9Si4O16(F,OH)2). Humites are relatively rare components of hydrothermally altered ultramafic rocks. They also occur in other silica-undersaturated rocks such as skarns and carbonatites. Natural humites almost always contain more F than hydroxyl. All of the humites lie on the join forsterite-brucite (Fig. 1). The formulas can be thought of as n·(Mg2SiO4)·Mg(OH)2 where n is one for norbergite, two for chondrodite, three for humite, and four for clinohumite, so that all have a higher (Mg+Fe)/Si ratio than does olivine. Because of this they are not thought to be major hydrous components of the mantle, which is generally considered to have a lower (Mg+Fe)/Si ratio than olivine. The F-free pure Mg clinohumite and chondrodite are stable to pressures greater than 14 GPa and temperatures greater than 1250 °C, but are not known to coexist with enstatite. The structures of chondrodite and clinohumite are illustrated in Figures 10 and 11. Although humite and norbergite have not been reported from pressure higher than about 3 GPa, hydroxy-chondrodite and hydroxy-clinohumite are stable at pressures and temperatures well into the transition zone (Yamamoto and Akimoto 1977; Burnley and Navrotsky 1996; Wunder 1998).

Clinohumite Clinohumite (Mg9Si4O16(OH)2) can coexist with chondrodite or with olivine at high pressure but not with phase A or enstatite (Fig. 1). The structure (Fig. 10) is monoclinic, P21/b (a-unique). The odd setting of the space group is chosen to preserve the olivine axial relation (Table 1). The c-axis is greater than that of chondrodite by approximately 6 Å, and the α-angle is reduced to about 100°. Hydroxy-clinohumite has the problem of protonating two identical oxygens symmetrically disposed about the inversion (Friedrich et al. 2001). But again, there are no Si-OH bonds and all non-silicate oxygens are protonated in pure hydroxy-clinohumite. Berry and James (2001) report a second partially occupied deuteron position in pure hydroxyl clinohumite located on the hydroxyl oxygen approximately 180° away from the position near the inversion on the O9-O9 edge.

Chondrite Chondrodite (Mg5Si2O8(OH)2) can coexist with phase A or with hydroxy-clinohumite at pressures to about 14 GPa. Its structure (Fig 11) resembles olivine with a and b axes nearly the same as olivine, but c different and the space group is monoclinic, P21/b. The O5 is the only non-silicate oxygen and is protonated. As with clinohumite, there is a problem with protonation

Hydrogen in High Pressure Silicate & Oxide Mineral Strutures

Figure 10. The structure of clinohumite, Mg9Si4O16(OH)2, is monoclinic P21/b. The odd setting of the space group is chosen to maintain the structural relation to olivine.

95

Figure 11. The structure of chondrodite, Mg5Si2O8(OH)2, is monoclinic P21/b. The odd setting of the space group is chosen to maintain the structural relation to olivine.

of every O5 in that this position is close to the inversion center and putting the proton on the O5-O5 edge would put the protons too close to each other. For synthetic deuterated chondrodite (Mg5Si2O8(OD)2), Lager et al. (2001) identified a second partially occupied deuteron position located approximately 180° away from the primary deuteron position on O5. Again, there are no Si-OH bonds and all non-silicate oxygens are protonated in pure hydroxy-chondrodite.

Phase A Phase A (Mg7Si2O8(OH)6) is stable under very hydrous conditions at pressures of 3 to about 8 GPa and temperatures of 550 to about 1250 °C (Yamamoto and Akimoto 1977). The structure (Fig. 12) is hexagonal, P63, and consists of slightly distorted closepacked layers of oxygen atoms and hydroxyl groups repeating along the caxis in an ABCB sequence (Horiuchi et al. 1979). This contrasts with the hexagonal close-packed sequence of Figure 12. The structure of Phase A, Mg7Si2O8(OH)6, ABAB in olivine and the humites. is acentric hexagonal, P63. Mg occupies one special position on the 3-fold axis (M3) and two general positions, M1 and M2. Si occupies two special positions, one each on the 3-fold and on the 63 axes, so that there is one in each layer of cations. All tetrahedra point in the same direction along c, so that the structure is acentric. The O2 and O4 oxygen sites are hydroxyls (Kagi et al. 2000), so all non-silicate oxygens are protonated and there are no Si-OH bonds in the structure. The density is relatively low (2.95 g/cm3) consistent with its high water content (~12% by weight) and limited pressure stability range. Phase A is a possible phase in the mantle as a breakdown product of serpentine, and may coexist with brucite or chondrodite, but probably not with olivine (Luth 1995).

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96 Phase B

Phase B (Mg12Si4O19(OH)2) along with superhydrous phase B (SHyB) and anhydrous phase B (AHyB), contains Si in both octahedral and tetrahedral coordination and has a Mg/Si ration greater than two (Finger et al. 1989). Phase B is stable under pressure and temperature conditions of the Transition Zone. The density is greater that that of forsterite, but less than that of wadsleyite or ringwoodite, despite the presence of octahedral silicon. The structure (Fig. 13) is monoclinic, P21/c, and all atoms except M1 and M3 are in general positions. There are four Si sites, three of which are tetrahedral and one octahedral. There are 13 distinct Mg octahedral sites and 21 distinct oxygen sites of which two are hydroxyls. All non-silicate oxygens are protonated and there are no Si-OH bonds in the structure.

Superhydrous Phase B Superhydrous Phase B (Mg10Si3O14(OH)4) (SHyB) is similar to phase B in having both octahedral and tetrahedral silicon, a stability range within the transition zone, and a Mg/Si ratio greater than two. The density is slightly less than that of phase B. The structure (Fig. 14) is orthorhombic, Pnnm (Pacalo and Parise 1992), and half of the Si atoms are octahedral and half tetrahedral. There are four distinct Mg octahedra and six oxygen sites. Although all non-silicate oxygens are protonated and there are no Si-OH bonds in the structure, one of the silicate oxygens is under-bonded (O3) and one is over bonded (O6) which leads to the distortions of the coordination polyhedra. As with the other B-phases, its Mg/Si ratio is greater than two, so it is not a likely phase in an enstatite or majorite bearing mantle assemblage. Koch-Müller et al. (2005) report polymorphic inversion in superhydrous phase B with an ordered low temperature polymorph having symmetry Pnn2. The reduction in symmetry with ordering causes splitting of the Mg positions, but not the Si positions.

Phase D Phase D (MgSi2O4(OH)2) is stable into the lower mantle at pressures of 17 to 50 GPa (Frost and Fei 1999) and has both Mg and Si in octahedral coordination (Fig. 15). The structure is

Figure 13. The structure of Phase B, Mg12Si4O19(OH)2, is monoclinic, P21/c.

Figure 14. The structure of superhydrous Phase B (ShyB), Mg10Si3O14(OH)4, is orthorhombic, Pnnm.

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highly disordered with variable Mg/Si ratios and water contents ranging from 10 to 18% by weight (Yang et al. 1997). The density of the ideal trigonal structure (P 31m) is about 3.01 g/cm3. All oxygens are equivalent and bonded to two Si, one Mg, and one proton, although only about one third of the proton positions can be occupied. Although the density is only about 75% that of the lower-mantle anhydrous assemblage, phase D is the likely host phase for H in the lower mantle.

Phase E

Figure 15. The structure of Phase D, MgSi2O4(OH)2, is trigonal, P 31m.

Phase E (Mg2SiO2(OH)4) is a highly disordered structure with Si in tetrahedral and Mg in octahedral coordination. The structure (Fig. 16) is trigonal R 3 m with variable Mg/Si ratio and H content (Kudoh et al. 1993). In the structure the M2 site occurs in an octahedral site adjacent to the Si tetrahedron, so that either one or the other can be occupied but not both. There is no long range order and charge balance is made up by protonation. The structure occurs in very hydrous compositions as a breakdown product of serpentine at pressures of 13 to 17 GPa and temperatures of 800 to 1300 °C (Kanzaki 1991).

Phase Pi Phase Pi (Al3Si2O7(OH)3) is so called because was formerly thought to be the poorly described synthetic mineral piezotite (Coes 1962). The mineral has been synthesized at low temperatures and moderate pressures (500-650 °C and 4-5.5 GPa) (Wunder et al. 1993) in the hydrous aluminosilicate system. The structure (Fig. 17), is acentric triclinic, P1, with Al in octahedral and Si in tetrahedral coordination (Daniels and Wunder 1993, 1996). There are 20 distinct oxygen atoms in the unit cell, of which six should be hydroxyls if the formula is correct. Four of the oxygens (O9, O10, O19, O20) are bonded to just two Al atoms and so

Figure 16. The structure of Phase E, Mg2SiO2(OH)4 is trigonal R 3 m. The structure is highly disordered. The silicate layer (dark) can have tetrahedral voids occupied by Si or octahedral voids occupied by Mg.

Figure 17. The structure of phase Pi (Al3Si2O7(OH)3) is acentric, triclinic P1. Although the proton positions have not been determined for this phase, the oxygen atoms shown as black spheres are protonated nonsilicate oxygens. The remaining two underbonded oxygens shown as white spheres are likely protonated silicate oxygen atoms.

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are certainly hydroxyls. The remaining oxygens all bond to Si. Of these, O4 and O14 bond to one Si and one Al, and so are also underbonded. They each have very long Al-O distances so are apparently hydroxyls, but unusual in that they may be protonated silicate oxygens.

Topaz-OH Topaz-OH (Al2SiO4(OH)2) also occurs in the hydrous aluminosilicate system at temperatures of 600-1000 °C and pressures up to about 12 GPa (Pawley 1994; Schmidt et al. 1998; Wunder et al. 1999). The structure (Fig. 18) is orthorhombic, Pbnm, with Al in octahedral and Si in tetrahedral coordination. The structure is relatively dense (3.37 g/cm3), more dense than phase Pi, but less dense than phase Egg or kyanite. Curiously, it is significantly less dense than fluoro-topaz. This may be because the protons are disordered over two distinct positions (Northrup et al. 1994).

Figure 18. The structure of topaz-OH, Al2SiO4(OH)2 is orthorhombic Pbnm. Despite its stability to quite high pressures (~12 GPa) it is significantly less dense at zero pressure than fluorotopaz.

Phase Egg Phase Egg AlSiO3(OH) is named after the first author to describe the phase (Eggleton 1978) and has a 1:1 Al:Si ratio. It occurs at pressure ranges into the transition zone at 11-18 GPa and temperatures of 700-1300 °C as a high pressure breakdown product of hydroxyl-topaz. The structure was solved and proton positions located to high precision by neutron powder diffraction (Schmidt et al. 1998). The structure is monoclinic, P21/n, and has both Si and Al in octahedral coordination (Fig. 19). There are four distinct oxygens in the structure. The O1 and O2 oxygens are bonded to two Si and one Al positions, whereas O4 is the hydroxyl, but it is also bonded to two Al and one Si, as is O3. The long H bond extends to O3. With octahedral silica and a single hydroxyl, the structure is relatively incompressible with a bulk modulus of 157 GPa and K′ of 6.5 (Vanpeteghem et al. 2003).

Figure 19. The structure of Phase Egg, AlSiO3(OH), is monoclinic P21/n. It has both Al (light) and Si (dark) in octahedral coordination.

K-cymrite K-cymrite (KAlSi3O8·H2O) occurs as a hydration product of sanidine at pressures above 3GPa and temperatures of 350-750 °C. It is isostructural with cymrite (BaAl2Si2O8·H2O) and has a layered structure of a double tetrahedral sheet (Fig. 20) with molecular water within the layer and K between the sheets (Fasshauer et al. 1997). The symmetry is P6/mmm so that the tetrahedral Al and Si are disordered over the sheet. There are bridging

Figure 20. The structure of K-cymrite (KAlSi3O8·H2O) is hexagonal P63/mmm and is composed of a double hexagonal layer of disordered Al-Si tetrahedra. K atoms (black) form the interlayer, and molecular water (gray) is in the tetrahedral layer.

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oxygens, non-bridging silicate oxygens, as well as molecular water, which lies within the tetrahedral layer (Fig. 20). The proton positions have not been determined, but are likely to be locally determined by the Al occupancy of the nearest tetrahedra. None of the silicate oxygens are protonated. K-cymrite is slightly denser than sanidine (Table 1).

Nominally Anhydrous High-Pressure Silicate and Oxide Phases Periclase-wüstite Periclase-wüstite ((Mg,Fe)O) is isometric, Fm 3 m, with the rock-salt structure (Fig. 21). Pure MgO is stable at low pressures and not known to undergo any high pressure phase transformations, whereas wüstite (FeO) is known to undergo a rhombohedral distortion of this structure at pressures above 20 GPa (Shu et al. 1998; Jacobsen et al. 2005). At low to modest pressures the structure can accommodate significant ferric iron in tetrahedral voids associated with octahedral vacancies. The oxygen site potentials of the nominally anhydrous mantle phases are given in Table 3. Periclase and wüstite have some of the shallowest oxygen potentials of any mantle minerals, which make these phases likely hosts for H if charge balance can be achieved. Murakami et al. (2002) report up to 2000 ppmw H2O in (Mg,Fe)O ferro-periclase synthesized from a hydrous peridotite composition at 25.5 GPa and 1650 °C. However their FTIR spectra show pleochroism unexpected for a cubic phase raising the possibility of an included hydrous phase of power symmetry. Bolfan Casanova et al. (2000) report only about 2 ppmw H2O in periclase at 24 GPa and 1500 °C in a pure MgO-SiO2-H2O system. Bolfan-Casanova et al. (2003) also report very low H contents in ferro-periclase up to 10 Gpa, so it appears that H2O solubility in pure MgO and in ferro-periclase of possible lower mantle composition is quite limited.

Figure 21. The structure of periclase (MgO) and wüstite (FeO) is the cubic rock salt structure, Fm 3 m.

Corundum Corundum (Al2O3) (Fig. 22) is rhombohedral, R 3 c, and isostructural with hematite (Fe2O3), eskolaite (Cr2O3), karelianite (V2O3), and synthetic Ti2O3. Ilmenite (FeTiO3) and akimotoite (MgSiO3) are also closely related structures with subgroup symmetry, R3. Natural corundum has not been reported with appreciable H contents, but it is not a common mineral in high pressure assemblages. It occurs in high grade peraluminous rocks with zoisite or in peraluminous eclogites. Rossman and Smyth (1990) report no observable OH stretch features in the FTIR spectrum of a natural corundum from a

Figure 22 The structure of corundum (Al2O3) and hematite (Fe2O3) is trigonal, R 3 c.

Mg2SiO4

Wadsleyite II

O O

Mg2SiO4

Wadsleyite

Mg3Al2Si3O12 Fe3Al2Si3O12

Fe2SiO4

Fayalite

Garnets Pyrope Almandine

Mg2SiO4

Mg2SiO4 Fe2SiO4 Mg14Si5O24

O1 O2 O3 O1 O2 O3 O1 O2 O3 O4 O1 O2 O3 O4 O5 O6 O7 O8 O O O1 O2 O3 O4 O5 O6 O7 O8 O9

SiO2

Stishovite Olivine Forsterite

Ringwoodite γ-Fe2SiO4 Anhyd. Phase B

O O O O O O1 O2 O3 O4 O5 O

MgO FeO Al2O3 TiO2 SiO2 SiO2

Periclase Wüstite Corundum Rutile Quartz Coesite

Site

Formula

Mineral

4Mg,1Al,1Si 4Fe,1Al,1Si

3Mg,1Si 3Mg,1Si 3Mg,1Si 3Fe,1Si 3Fe,1Si 3Fe,1Si 5Mg 1Mg,2Si 3Mg,1Si 3Mg,1Si 3Mg,1Si 5Mg 3Mg,1Si 3Mg,1Si 3Mg,1Si 3Mg,1Si 1Mg,2Si 3Mg,1Si 3Mg,1Si 3Fe,1Si 3Mg,1SiIV 4Mg,1SiVI 3Mg,1SiIV 6Mg 3Mg,1SiIV 3Mg,1SiIV 3Mg,1SiIV 4Mg,1SiVI 3Mg,1SiIV

6Mg 6Fe 4Al 3Ti 2Si 2Si 2Si 2Si 2Si 2Si 3Si

Coordination

27.06 27.08

27.69 27.53 26.35 27.38 27.20 26.12 21.28 30.94 26.78 26.97 26.83 20.03 26.41 26.68 26.71 27.34 30.51 27.50 26.57 26.28 27.09 25.62 26.32 22.70 28.77 27.84 26.66 25.09 26.91

23.92 23.36 26.40 25.94 30.82 29.07 29.64 30.67 30.41 30.97 28.61

Potential (V)

MgSiO3 MgSiO3 ZrSiO4 CaTiSiO5

Perovskite Post-perovskite* Zircon Titanite

Mg2Si2O6 Al2SiO5

CaMgSi2O6

Diopside Akimotoite Kyanite

NaAlSi2O6

Mg2Si2O6

MgAlAlSiO6

Mg2Si2O6

Formula

Jadeite

Clinoenstatite

Mg-Tschermaks

Pyroxenes Orthoenstatite

Mineral O1a O2a O3a O1b O2b O3b O1a O2a O3a O1b O2b O3b O1a O2a O3a O1b O2b O3b O1 O2 O3 O1 O2 O3 O O1 O2 O3 O4 O5 O6 O7 O8 O9 O10 O1 O2 O1 O2 O O1 O2a O2b O3a O3b

Site 3Mg,1Si 2Mg,1Si 1Mg,2Si 3Mg,1Si 2Mg,1Si 1Mg,2Si 1Mg,2Al,1Si 1Mg,1Al,1Si 1Mg,2Si 3Al,1Mg 2Al,1Mg 2Al 3Mg,1Si 2Mg,1Si 1Mg,2Si 3Mg,1Si 2Mg,1Si 1Mg,2Si 1Na,2Al,1Si 1Na,1Al,1Si 2Na,2Si 1Ca,2Mg,1Si 1Ca,1Mg,1Si 2Ca,2Si 2Mg, 2SiVI 2Al,1Si 4Al 2Al,1Si 2Al,1Si 2Al,1Si 4Al 2Al,1Si 2Al,1Si 2Al,1Si 2Al,1Si 2Mg,2SiVI 3Mg,2SiVI 2Mg,2SiVI 3Mg,2SiVI 2Zr,1Si 1Ca,2Ti 1Ca,1Ti,1Si 1Ca,1Ti,1Si 2Ca,1Ti,1Si 2Ca,1Ti,1Si

Coordination

Table 3. Cation coordinations and electrostatic potentials of oxygen sites in high pressure silicate and oxide phases.

26.22 26.38 30.89 26.39 26.48 30.57 32.09 30.25 32.49 24.59 22.39 21.99 26.33 26.34 30.90 26.39 26.34 30.59 27.54 27.15 30.34 25.53 25.78 30.74 27.38 28.60 25.91 27.73 28.01 28.60 25.81 27.70 27.97 28.28 28.36 26.91 26.86 27.76 26.77 31.49 24.89 26.86 26.96 26.98 26.87

Potential (V)

100 Smyth

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high pressure corundum-kyanite eclogite. There is a single oxygen site in the structure. The site potential is significantly deeper than that of periclase but might allow minor protonation if charge balance can be achieved. However, significant protonation of the isostructural akimotoite (MgSiO3) does occur as discussed below.

Coesite Coesite (SiO2) is the high pressure polymorph of SiO2 stable between about 3 and 8 GPa. The structure (Fig. 23) is a relatively dense tetrahedral framework with monoclinic C2/c symmetry. Natural coesite is normally quite pure SiO2 with only trace levels of other elements. All oxygens are bridging oxygens bonded only to two Si atoms. There are five distinct oxygen sites in the structure, all with deep potentials similar to quartz (Table 3). Of these O1 has the shallowest potential and the most likely one to be protonated if there were a small amount of B or Al substitution in the tetrahedra. Rossman and Smyth (1990) report no observable OH in a natural coesite in a relatively hydrous coesite-kyanite eclogite. Koch-Mueller et al. (2001) and Mosenfelder (2000) however report up to 200 ppmw H2O in coesite synthesized at pressures of 7.5 GPa and 1100 °C, but undetectable amounts in coesite synthesized at pressures below 5 GPa. Koch-Müller et al. (2003) report that the major substitution mechanism in coesite is by the hydrogarnettype (H4O4) with relatively minor amounts of H being associated with B and Al substitution. In a low-symmetry tetrahedral framework structure such as coesite, any Si vacancy would result in protonation of the terminating oxygens, but there would be nothing to constrain these oxygens to maintain a tetrahedral configuration, as there is in garnet. Koch-Müller et al. (2001) propose several possible proton locations for coesite on the oxygens coordinating a vacant Si2 position consistent with O-H dipoles observed in polarized infrared spectra. They further suggest that vacancy at Si1 is unlikely because of difficulty in accounting for the pleochroism of one of the major O-H vibrations.

Stishovite and rutile

Figure 23. The structure of coesite, SiO2, is monoclinic C2/c. All oxygens are bridging oxygens bonded to two tetrahedral Si atoms. Trace hydration of this structure has only been observed in samples quenched from pressures above 5 GPa.

Figure 24. The structure of stishovite (SiO2) and rutile (TiO2) is tetragonal, P42/mnm. All oxygens are equivalent. Protonation of these compounds can accompany trivalent ion substitution in the octahedra. Proton positions determined by neutron single crystal diffraction for rutile (Swope et al. 1995) are illustrated.

Stishovite (SiO2) and rutile (TiO2) (Fig. 24) are isostructural and both may incorporate considerably more H than either coesite or quartz. The stishovite structure is tetragonal P42/mnm with all cations in octahedral coordination, and the octahedra share edges in the c-direction. All oxygens in the structure are equivalent, and protonation of the oxygens can accompany Al for Si substitution in the octahedra (Smyth et al.

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1995). The oxygen site potential is substantially lower than those of quartz or coesite (Table 3). Bolfan-Casanova et al. (2000) report up to 72 ppmw H2O in stishovite in an Al-free system. Vlassopoulos et al. (1993) report up to 8000 ppmw H2O in natural rutile containing minor amounts of trivalent cations (Cr, Fe, V, Al). Principal rutile absorptions in the OH range are at 3290 and 3365 cm−1 (Rossman and Smyth 1990; Vlassopoulos et al. 1993) and are strongly polarized normal to the c-axis. Swope et al. (1995) report a proton position on the shared octahedral edge for hydrous rutile at x/a = 0.4176; y/b = .5033, and z/c = 0, based on neutron single crystal diffraction of a natural sample. This position is consistent with the strong IR pleochroism and is illustrated in Figure 24.

Pyroxenes Pyroxenes of major importance to mantle dynamics include enstatite (Mg2Si2O6), diopside (CaMgSi2O6), and jadeite (NaAlSi2O6), which are all significant components of the upper mantle. For a recent review of pyroxene structures at temperature and pressure see Yang and Prewitt (2000). Enstatite is an orthopyroxene, orthorhombic, Pbca (Fig. 25), at pressures to about 7 GPa, whereas enstatite quenched from higher pressures is monoclinic P21/c. Clinoenstatite transforms to majorite garnet at about 15 GPa, in a pyrolite composition and gradually dissolves into the garnet phase through the upper Transition Zone. Mantle peridotites and lherzolites contain up to about 15 modal percent clinopyroxene that is typically a Cr-diopside with very minor amounts of Na, Al or Fe3+. In eclogites, however, diopside and jadeite form a complete crystalline solution known as omphacite, which is monoclinic C2/c at high temperatures. Omphacite composes 50% or more of eclogites that form from subducting basalt at pressures of 3 to 13 GPa. Eclogites are quite distinct from peridotites and lherzolites, so that rocks of intermediate composition are virtually unknown among rocks of high pressure origin.

Figure 25. The structure of ortho-

enstatite, Mg2Si2O6, is orthorhombic Orthoenstatite can be a major host for water in the Pbca. This view down c with ashallow (lithospheric) upper mantle. Rauch and Keppler vertical, shows the alternating layers (2002) report that the solubility of H2O in enstatite of T1 and T2 tetrahedra. The likely increases to a maximum of about 850 ppmw at 1100 °C sites of protonation are on the O2b at 7.5 GPa and decreases slightly at higher pressures in and O1b oxygens (spheres) with the O-H vectors lying in the b-c plane. the clinoenstatite field. In pure Mg enstatite, the strongest OH absorptions in the infrared spectra are polarized parallel to c. However, Al has a dramatic effect on the water solubility and on the FTIR spectra of orthoenstatite, especially at pressures of 1 to 2 GPa at which Al substitution in the tetrahedral site can be extensive (Mierdel et al. 2006). In aluminous enstatite, H2O solubilities can approach 9000 ppmw at 900 °C and 1.5 GPa. In these enstatites, the O-H polarizations are strongest perpendicular to c (Mierdel et al. 2006).

Orthoenstatite is orthorhombic, Pbca, with two distinct tetrahedral sites, T1 and T2, arranged is separate layers of tetrahedral chains (Fig. 25). Al enters the structure as a coupled substitution where the Al is in both an M1 octahedron and one of the tetrahedral sites. Tetrahedral Al is known to strongly order in the structure with a very strong preference for T2 (Takeda 1973). There are six distinct oxygen sites in the structure, O1a, O1b, O2a, O2b, O3a,

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and O3b, with the ‘a’ oxygens in the T1 chains and the ‘b’ oxygens in the T2 chains. The O3 atoms are the bridging oxygens in the chains. Electrostatic site potentials for the oxygens for pure Mg orthoenstatite are given in Table 3, and the O2b has the shallowest potential and is therefore the most likely site for protonation. Structure refinement of a hydrous, aluminous orthopyroxene shows up to 5% cation vacancy at M2 with nearly equal amounts of Al substitution in both M1 and T2 sites, based on chemical analysis and volumes of coordination polyhedra (Smyth et al. 2006b). Also reported in Table 3 is an oxygen site potential calculation for a hypothetical fully “Mg-Tschermaks” orthoenstatite of composition MgAlAlSiO6, fully ordered with all tetrahedral Al in T2. In this structure both O2b and O1b are substantially underbonded and likely sites for protonation. The O3b oxygen is also underbonded, but Al-avoidance would not allow Al in T2 to exceed 50%, so O3b is not as likely to protonate as O2b or O1b. It appears then that the major hydrous components are MgAlAlSiO6 and H2AlAlSiO6 (“hydro-Tschermaks”), with a cation vacancy at M2 and protons on the O1b-O2b edges of the vacant M2 polyhedron, consistent with the observed O-H polarization in the a-b plane. This substitution mechanism achieves a net volume reduction of the unit cell, and nearly 1% H2O by weight (Mierdel et al. 2006), but because it requires tetrahedral Al, H solubility decreases sharply with increasing pressure. The O2b and O1b oxygen sites are indicated by spheres in Figure 25. This “hydro-Tschermaks” substitution appears to be strongly abetted by the ordering of tetrahedral Al in T2 which can only happen in the Pbca structure. At pressures near the 410 km discontinuity, enstatite is monoclinic, P21/c, after quenching to low temperature, but C2/c at relevant mantle temperatures. The solubility of H is much less than that in aluminous orthopyroxene at lower crustal pressure, so that clinoenstatite in equilibrium with forsterite containing >8000 ppmw H2O contains less than 1000 ppmw (Smyth et al. 2006a) and somewhat less (~650 ppmw) in equilibrium with wadsleyite (Bolfan-Casanova et al. 2000). The principal substitution mechanism appears to be divalent cation vacancies, principally at M2. Natural omphacites can contain up to about 3000 ppmw H2O (Katayama and Nakashima 2003; Smyth et al. 1991). Bromiley and Keppler (2004) experimentally investigated water solubility in jadeite and found a maximum H2O content of about 450 ppmw at 2 GPa, but dramatically higher solubilities in more complex solid solutions. Natural omphacites are very complex chemically containing 10% or more of up to eight chemical end members (Smyth 1980), but crystallographically relatively simple, having space group C2/c at mantle conditions of temperature and pressure. The hydrous component referred to as Ca-Eskola pyroxene Ca0.50.5AlSi2O6, may be better described as HAlSi2O6. Crystal structure refinements of natural H-rich omphacites indicate significant M2 site vacancy (Smyth 1980). Textural evidence of kyanite and garnet exsolution from omphacite suggests that H2O solubility in these pyroxenes may approach 1% by weight (Smyth et al. 1991). Bromiley et al. (2004) have experimentally hydrated natural Cr-diopside crystals at 1100 °C and pressures of 1.5 to 4 GPa. They report up to about 450 ppmw at 1.5GPa and infer proton positions on the O2-O1 and O2-O3 edges of the M2 polyhedron based on polarizations of the O-H vector in the a-b plane, which are similar to those reported for orthopyroxene by Mierdel et al. (2006).

Akimotoite Akimotoite (MgSiO3) is the ilmenite-type polymorph of enstatite stable at pressures of the lower transition zone (18-22 GPa). The structure is trigonal R3 and has alternating layers of Si and Mg octahedra (Fig. 26). Bolfan-Casanova et al. (2000) report up to about 450 ppmw H2O in pure Mg akimotoite at 21 GPa and 1500 °C coexisting with stishovite and melt. BolfanCasanova et al. (2000, 2002) report strongly pleochroic FTIR spectra for the O-H stretching vibration in this phase with strong absorptions at 3390 cm−1 parallel to c and 3320 and 3300 cm−1 perpendicular to c. Based on the polarizations and the relation of frequency to O-H-O

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Figure 26. The structure of akimotoite (ilmenite-type MgSiO3) is trigonal R3 and closely related to that of corundum.

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Figure 27. the structure of garnet is cubic Ia 3 d . All oxygen atoms are identical and the tetrahedra and octahedra form a corner-sharing framework structure.

distance (Libowitzky 1999), they deduce two proton positions, both likely associated with Mg vacancies. Inasmuch as the structure is essentially isostructural with corundum, possible Al substitution for octahedral Si might have a significant impact on the H solubility in this phase.

Garnet Garnet (X3Y2Z3O12) (Fig. 27) is isometric, Ia 3 d, with Si (Z) in tetrahedral coordination forming a framework by sharing oxygens with Al (Y) in octahedral coordination. Interstitial to the framework is the dodecahedral divalent cation site, which may be occupied by Mg, Fe, or Ca (X). In this high-symmetry structure, all oxygens are equivalent and in a general position. At pressures of the transition zone, garnet can accept equal amounts of Si and Mg into the octahedral site in place of a trivalent cation. The Mg3(MgSi)2Si3O12 (MgSiO3) end-member is majorite. Majorite quenches to tetragonal, I41/a, by ordering of Mg and Si in the octahedral site, although it is likely disordered Ia 3 d at mantle conditions (Angel et al. 1989). Hydrogen is accommodated in the garnet structure by Si vacancies so that the terminating octahedral oxygens are protonated. The tetrahedral site has 4 point symmetry, so symmetry constrains the oxygens to maintain a tetrahedral configuration, but the distance from the 4 point position to the oxygen increases from about 1.63 Å for the occupied site to about 1.95 Å for the vacant site (Lager and von Dreele 1996). This means that pressure inhibits the substitution so that garnets from high pressure environments generally contain less than 50 ppmw H2O (Bell and Rossman 1992). Lager et al. (1987) and Lager and von Dreele (1996) report deuteron positions for a deuterated hydrogarnet (Ca3Al2D12O12) on the edges of the vacant tetrahedra based on neutron single crystal diffraction.

Olivine Olivine ((Mg,Fe)2SiO4) is generally believed to be the most abundant phase in the upper mantle from the Moho to 410 km discontinuity. Natural olivines as reviewed in the current volume (Beran and Libowitzky 2006) contain up to about 400 ppm by weight (ppmw) H2O, but typically less than 100 ppmw (Bell et al. 2004). Olivine synthesized at high pressures and quenched can contain much more H. Kohlstedt et al. (1996) report up to 1510 ppmw in olivine equilibrated at 1100 °C and 12 GPa. Recalculating this amount based on Bell et al. (2003) one

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gets about 4000 ppmw (Hirschmann et al. 2005). Mosenfelder et al. (2006) report up to 6400 ppm H2O in olivine quenched from 12 GPa and 1100 °C. Smyth et al. (2006a) report up to 8900 ppmw in olivine synthesized at 1250 °C and 12 GPa in equilibrium with either enstatite or clinohumite, but decreasing at higher temperatures with the onset of melting. Water contents approaching one per cent by weight would make olivine a major host for water in the upper mantle. The olivine structure (Fig. 28) is orthorhombic, Pbnm, with two distinct octahedra, M1 and M2, and one silicate tetrahedron. There are three distinct oxygen sites in the structure, with O1 and O2 lying on the mirror, and O3 being in a general position. All oxygens are bonded to three Mg and one Si atom (Table 3) and site potentials range from 26.3 V for O3 to 27.7 V for O1. Smyth et al. (2006a) report that the major H substitution mechanism in olivine is protonation of the O1-O2 edges of vacant M1 octahedra. The proton position suggested by Smyth et al. (2006a) at x/a = 0.95; y/b = 0.04; z/c = 0.25 is illustrated in Figure 28. They further report a volume of hydration at ambient conditions: V = 290.107 + 5.5×10−5 *cH2O Å3



where V is cell volume in Å3, and H2O is the ppm by weight H2O as determined from the calibration of Bell et al. (2003).

Wadsleyite Wadsleyite is the first high pressure polymorph of Mg2SiO4, and the olivine-wadsleyite transition at about 13 GPa is thought be responsible for the 410 km discontinuity. The wadsleyite structure (Fig. 29) is usually orthorhombic, Imma, with three distinct divalent metal octahedra, M1, M2 and M3. The structure is similar to that of spinelloid III in the Nialuminosilicate system (Ma and Sahl 1975). Unlike olivine which is based on a hexagonal close-packed array of oxygens, wadsleyite and the other spinels and spinelloids are based on a cubic close-packed oxygen array. Unlike olivine and ringwoodite, wadsleyite is a sorosilicate

Figure 28. The structure of forsterite, Mg2SiO4, and fayalite Fe2SiO4, is orthorhombic Pbnm. Hydration appears to be compensated by octahedral cation vacancies principally at M1. The proton position inferred from polarized FTIR spectroscopy on the O1-O2 shared edge of the M1 octahedron is illustrated.

Figure 29. The structure of wadsleyite, (Mg,Fe)2SiO4, is orthorhombic Imma. Hydrous wadsleyite may deviate slightly from orthorhombic symmetry as monoclinic, I2/m, due to ordered cation vacancies in M3 in violation of the mirror perpendicular to a. The structure has a non-silicate oxygen which is readily protonated. Charge balance is maintained by Mg vacancies at M3.

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with Si2O7 groups, a bridging oxygen (O2) and a non-silicate oxygen (O1). Smyth (1987) calculated oxygen site potentials and predicted that the under-bonded non-silicate oxygen would be a potential site for protonation. Wadsleyites with up to 3% by weight H2O have been reported (Inoue et al. 1995). The major hydrogen substitution mechanism appears to be protonation of the vacant M3 octahedral edges and ordering of the vacancies so that hydrous wadsleyites with more than about 1% H2O are monoclinic, I2/m (a subgroup of Imma). Beta angles up to 90.4° have been reported (Smyth et al. 1997; Jacobsen et al. 2005). Wadsleyite shows a significant zero-pressure volume expansion that is similar in magnitude to that of olivine. Holl (2006) reports the volume expansion as: V = 538.64 + 9.4 × 10−5 *cH2O Å3 Hydrous wadsleyite shows a strong O-H stretching absorption at about 3325 cm−1 which shows minimal pleochroism. A potential proton location on the O1-O4 edge of a vacant M3 octahedron at about x/a = 0.11; y/b = 0.20; z/c = 0.36 would be consistent with the observed frequency and pleochroism of this polarization and is illustrated in Figure 29. The complexity of the infrared absorption spectrum, however, indicates that there are multiple possible proton locations in the structure (Kohn et al. 2002).

Wadsleyite II Wadsleyite II is isostructural with spinelloid IV (Smyth and Kawamoto 1997; Smyth et al. 2005). It has only been reported from long-duration hydrous peridotite composition runs at 17.5 to 18 GPa, between the wadsleyite and ringwoodite fields. It is a well-ordered phase with a- and c-axes similar to wadsleyite but with a b-axis 2.5 times that of wadsleyite at about 30 Å. The structure is very difficult to distinguish from wadsleyite by powder diffraction or by Raman spectroscopy. The structure (Fig. 30) contains both isolated SiO4 tetrahedra as well as Si2O7 groups in three distinct tetrahedral sites. It also contains six distinct octahedral sites and eight distinct oxygens, of which O2 is a non-silicate oxygen and a potential protonation site. Analogous to wadsleyite, a possible proton location would be near the O2-O4 edge of the M6 octahedron or the O2-O5 edge of the M5 octahedron. Wadsleyite II in the high pressure peridotite system is only known with about 2.8 wt% H2O, whereas spinelloid IV in the Ni aluminosilicate system is thought to be anhydrous (Akaogi et al. 1982; Horioka et al. 1981).

Ringwoodite Ringwoodite is the true spinel polymorph of forsterite and is stable as the dominant phase in a pyrolite composition mantle from about 525 to 670 km depth. The ringwoodite to perovskite plus periclase transition is thought to be responsible for the 670 km discontinuity. The structure (Fig. 31) is cubic, Fd 3 m with octahedral Mg and tetrahedral Si. Kohlstedt et al. (1996) report up to about 2.4 wt% H2O in ringwoodite. The FTIR spectrum shows a

Figure 30. The structure of wadsleyite II, (Mg,Fe)2SiO4, is orthorhombic Imma. This structure, like wadsleyite is a spinelloid, but contains both isolated SiO4 groups as well as Si2O7 groups.

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Figure 31. The structure of ringwoodite is a true spinel and is cubic, Fd 3 m. Si is in tetrahedral (dark) and Mg in octahedral (light) coordination. All oxygens are equivalent and bonded to one Si and three Mg atoms. There are no bridging or nonsilicate oxygens. Hydration is compensated by octahedral site vacancies.

broad absorption feature in the range 2600 to 3600 cm−1 (Smyth et al. 2003; Keppler and Smyth 2005). Although there is no IR pleochroism in the cubic system, the OH does appear to be structural because OH concentration computed from the FTIR spectrum correlates with a zero-pressure unit cell volume increase (Smyth et al. 2003) that is similar in magnitude to those observed for forsterite and wadsleyite cited above. Peaks in the spectra correlate with protonation of both the octahedral and tetrahedral edges (Libowitzky 1999) and crystal structure refinements indicate both octahedral and tetrahedral vacancies (Kudoh et al. 2000; Smyth et al. 2003).

Anhydrous phase B Anhydrous phase B (Mg14Si5O24) lies on the anhydrous edge of the DHMS ternary between forsterite and periclase. As with the other B-phases, anhydrous phase B has Mg/Si ratio greater than two, and so is not expected to coexist with either enstatite or majorite. It is therefore not expected to be a significant phase in the transition zone. The structure (Fig. 32) is orthorhombic, Pmcb (Hazen et al. 1992) and has Si in both octahedral and tetrahedral coordination. Little is known about its trace H content, but its oxygen sites are all electrostatically balanced according to Pauling bond strength sums, bonded to either three octahedral Mg and a tetrahedral Si, six Mg, or four Mg and one octahedral Si. Of these, the O4 is the non-silicate oxygen, has the lowest electrostatic potential and is thus a potential protonation site (Table 3). The density (3.39 g/cm3) lies between that of forsterite and periclase, but less than either wadsleyite or ringwoodite, despite its octahedral silicon.

Kyanite Kyanite (Al2SiO5) is triclinic P1, with Al in octahedral and Si in tetrahedral coordination. There are ten distinct oxygen sites in the structure (Fig. 33) most of which are bonded to two octahedral Al and one tetrahedral Si. The O2 and O6 positions are non-silicate oxygens and

Figure 32. The structure of anhydrous Phase B (AnHB), Mg14Si5O24, is orthorhombic Pmcb, and had Si in both octahedral and tetrahedral coordination.

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bonded to only four Al atoms (Table 3). These are potential hydration sites if charge balance can be achieved by divalent cation substitution for Al. Although Beran and Goetzinger (1987) and Rossman and Smyth (1990) report relatively large amounts of OH in kyanite up to about 4000 ppmw H2O, Bell et al. (2004) report a new calibration for kyanite, greatly reducing this amount and reporting a maximum H2O content for kyanite of about 230 ppmw.

Perovskite Perovskite-type (Mg,Fe)SiO3 is believed to be the major phase in the lower mantle, so small amounts of H in this phase can have a large effect on the total water budget of the planet. The strucFigure 33. The structure of kyanite, Al2SiO5, ture (Fig. 34) is orthorhombic, Pbnm, with Mg in is triclinic P1. eight coordination, Si in octahedral coordination, and two distinct oxygen sites. Both oxygen sites have relatively deep electrostatic potentials near 27 V (Table 3). The structure is dense (4.1 g/cm3). Meade et al. (1994) report only minor amounts of H in MgSiO3 perovskite. Bolfan-Casanova et al. (2000) report no detectable H by FTIR spectroscopy in pure MgSiO3 perovskite in equilibrium with hydrous akimotoite in an Al-free composition, however Higo et al. (2001) report up to 500 ppmw H2O by SIMS analysis of similar samples. Murakami et al. (2002) report up to 2000 ppmw H2O in (Mg,Fe)SiO3 perovskite synthesized at 25.5 GPa and 1600 °C in an Al-bearing peridotite composition. Litasov et al. (2003) observed only Figure 34. The structure of perovskite-type about 100 ppm in pure MgSiO3 perovskite, but MgSiO3 is orthorhombic Pbnm. 1400 to 1800 ppmw H2O in Al and Fe bearing perovskites in a hydrous peridotite system. None of the FTIR spectra of silicate perovskites in pure MgSiO3 or MgSiO3-Al2O3 systems show sharp absorption bands so there has been some disagreement as to whether these features represent structurally bound hydroxyl (Bolfan-Casanova et al. 2003; Litasov et al. 2003). Perovskite samples synthesized in chemically complex systems show a consistent but broad OH absorption feature at about 3397 cm−1, but variable other features. It appears that while H2O solubility in pure MgSiO3 perovskite is likely negligible, perovskite crystallized from more chemically complex systems may incorporate significant amounts of water, but in reports of higher water contents, the possibility of hydrous inclusions within the perovskite cannot be ruled out. Perovskite-type CaSiO3 is believed to be a minor phase in the lower mantle. Although it is isostructural with MgSiO3 perovskite (orthorhombic, Pbnm), is appears to form a separate phase in lower mantle synthesis experiments. Murakami et al. (2002) report up to 4000 ppmw H2O in CaSiO3 perovskite synthesized at 25.5 GPa and 1600 °C. This phase does not appear to be quenchable so interpretation of FTIR spectra on quenched material is difficult.

Post-perovskite Post-perovskite (MgSiO3) is a new structure type reported for MgSiO3 at pressures of the lower-most lower mantle near the core-mantle boundary (Murakami et al. 2004). It is

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Figure 35. The structure of post-perovskite-type MgSiO3 is orthorhombic Cmcm.

postulated that the perovskite to post perovskite transition may account for the discontinuity that defines the D′′ layer near 2600 km depth. The structure (Fig. 35) is orthorhombic, Cmcm, and has edge-sharing silicate octahedra forming chains parallel to a, which are corner-linked to form sheets in the a-c plane. The sheets are linked together with 8-coodinated Mg atoms to form a strongly anisotropic structure. There are two distinct oxygen sites in the structure. Of these, O1 is slightly underbonded, being coordinated to two Si and two Mg atoms, whereas O2 is slightly overbonded to two Si and three Mg. However the potentials are rather similar to those of MgSiO3-perovskite (Table 3).

Zircon Zircon (ZrSiO4) is a primary accessory phase in nearly all igneous rocks, and a major host phase for minor U, Th, and rare earth elements in the Earth. Though nominally anhydrous, nonmetamict zircons of mantle origin can contain up to about 100 ppmw H2O (Woodhead et al. 1991; Nasdala et al. 2001). This minor hydration is consistent with the very deep potential of the oxygen site (Table 3), and probably requires trivalent cation substitution for Zr. Additionally, metamict zircons, which have experienced radiation damage from the decay of U and Th, may contain much more H2O, more than 16% by weight H2O (Woodhead et al. 1991). The structure (Hazen and Finger 1979) is illustrated in Figure 36 and has Si in tetrahedral and Zr in eight-coordination. All O atoms are equivalent and bonded to tetrahedral Si so there are no non-silicate oxygens. Woodhead et al. (1991) report that strong absorption features at 3385 cm−1 perpendicular to c, and a weaker feature at 3420 cm−1 parallel to c, are associated with an occupied tetrahedron and trivalent cation substitution for Zr. However, if the proton is located on an O-O polyhedral edge, the only edge of the Zr polyhedron that does not have a component Figure 36. The structure of zircon, ZrSiO4, is in the c-direction is the edge shared with the tetragonal, I41/amd. In this c-axis projection, tetrahedron. This would be consistent with the the Zr is seen as eight-coordinated dipyramids suggestion of Nasdala et al. (2001) that hydra(light) and the Si (dark) is tetrahedral. All tion also appears to occur by the hydro-garnet oxygens are equivalent and bonded to two Zr and one Si. substitution involving tetrahedral vacancy.

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110 Titanite Titanite (CaTiSiO5), like zircon, is a very common primary accessory phase in igneous rocks. The structure (Fig. 37) is monoclinic, P21/a (b-unique) and has Ca in eight-coordination with Ti in octahedral and Si in tetrahedral coordination. Although it is nominally anhydrous, it can accommodate substantial amounts of both OH and F with Al substitution for Ti. There is one non-silicate oxygen in the structure (O1) which is bonded to one Ca and two Ti atoms. It is under-bonded in the Pauling sense, and its electrostatic site potential is 24.9 V which makes it the obvious candidate for protonation to accommodate Al or Fe3+ in the octahedron.

Figure 37. The structure of titanite, CaTiSiO5, is monoclinic, P21/a.

Conclusions The structure of the nominally hydrous and anhydrous phases that compose the Earth’s mantle have been reviewed and compared. Among the nominally hydrous high-pressure silicate phases, we have examples of molecular water in lawsonite and K-cymrite. We also see that for hydroxyl-bearing silicates, the hydroxyls are in general, non-silicate oxygens. We see no examples of a proton on tetrahedral silicate oxygens. There are a few examples of protonated tetrahedral silicate oxygens in nature such as in the pyroxenoids, pectolite (NaHCa2Si3O9) and serandite (NaHMn2Si3O9). In these structures the chains are so strongly kinked that two of the non-bridging oxygens approach so closely that there is a H-bond between the two (Jacobsen et al. 2000). We also see a few examples of Si-OH bonds for octahedral silica, as in the very high pressure phases D and Egg. This is consistent with the octahedral Si-O bond being longer and weaker than the tetrahedral Si-O bond. Among the nominally anhydrous phases we see that the phases that have only bridging tetrahedral silicate oxygens are able to accommodate the least amount of H, whereas phases containing non-silicate oxygens are readily hydrated. The minerals containing octahedral silica can accept up to several thousand ppmw H2O if Al is present to substitute for octahedral silica.

ACKNOWLEDGMENT The author thanks U.S. National Science Foundation for grant NSF-EAR 03-36611, the Bayerisches Geoinstitut Visitors Program, and the Alexander von Humboldt Foundation. The author also thanks H. Keppler, T. Boffa-Balaran and P. Comodi for constructive, thorough, and competent reviews.

REFERENCES Akaogi M, Akimoto S, Horioka H, Takahashi K, Horiuchi H (1982) The system NiAl2O4-Ni2SiO4 at high pressures ad high temperatures: spinelloids with spinel-related structures. J Solid State Chem 44:257-267 Angel RJ, Finger LW, Hazen RM, Kanzaki M, Weidner DJ, Liebermann RC, Veblen DR (1989) Structure and twinning of single-crystal MgSiO3 garnet synthesized at 17 GPa and 1800 C. Am Mineral 74:509-512 Armbruster T, Geiger CA, Lager G (1992) Single crystal X-ray study of synthetic pyrope almandine garnets at 100 and 293 K. Am Mineral 77:512-521

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Baur WH (1978) Crystal structure refinement of lawsonite. Am Mineral 63:311-315 Beard JS (2000) Occurrence and composition of tochilinite and related minerals in Site 1068 serpentinites. In: Proceedings of the Ocean Drilling Program, Scientific Results, Return to Iberia. Beslier M-O, Whitmarsh RB, Wallace PJ, Girardeau J (eds) OPD, 173:1–9 Bell DR, Rossman GR (1992) Water in the Earth’s mantle: Role of nominally anhydrous minerals. Science 255: 1392-1396 Bell DR, Rossman GR, Moore RO (2004) Abundance and partitioning of OH in a high-pressure magmatic system: megacrysts from the Monastery Kimberlite, South Africa. J Petrol 45:1539-1564 Bell DR, Rossman GR, Maldener J, Endisch D, Rauch F (2004) Hydroxide in kyanite: A quantitative determination of the absolute amount and calibration for the IR spectrum. Am Mineral 89:998-1003 Bell DR, Rossman GR, Maldener J, Endisch D, Rauch F (2003) Hydroxide in olivine: a quantitative determination of the absolute amount and calibration of the IR spectrum. J Geophys Res 108 doi:10.1029/ 2001JB000679 Beran A, Goetzinger MA (1987) Quantitative IR spectroscopic determination of structural OH groups in kyanites. Mineral Petrol 36:41-49 Bercovici D, Karato S-I (2003) Whole-mantle convection and the transition-zone water filter. Nature 425:3944 Beran A, Libowitzky E (2006) Water in natural mantle minerals II: olivine, garnet and accessory minerals. Rev Mineral Geochem 62:169-191 Berry A, James M (2001) Refinement of hydrogen positions in synthetic hydroxyl clinohumite by powder neutron diffraction. Am Mineral 86:181-184 Boffa-Balaran T, Angel RJ (2003) Equation of state and high-pressure phase transitions in lawsonite. Eur J Min 15:241-246 Bolfan-Casanova N, Keppler H, Rubie DC (2002) Hydroxyl in MgSiO3 akimotoite: a polarized and high pressure IR study. Am Mineral 87:603-608 Bolfan-Casanova N, Mackwell S, Keppler H, McCammon CA, Rubie DC (2002) Pressure dependence of H solubility in magnesiowustite up to 25 GPa: Implications for the storage of water in the Earth’s lower mantle. Geophys Res Lett 29:1029-1032 Bolfan-Casanova N, Keppler H, Rubie DC (2003) Water partitioning at the 660 km discontinuity and evidence for very low water solubility in magnesium silicate perovskite. Geophys Res Lett 30:1905-1908 Bolfan-Casanova N, Keppler H, Rubie DC (2000) Water partitioning between nominally anhydrous minerals in the MgO-SiO2-H2O system up to 24 GPa: implications for the distribution of water in the Earth’s mantle. Earth Planet Sci Lett 182:209-221 Bromiley GD, Keppler H (2004) An experimental investigation of hydroxyl solubility in jadeite and Na-rich clinopyroxenes. Contrib Mineral Petrol 147:189–200 Bromiley GD, Keppler H, McCammon C, Bromiley FA, Jacobsen SD (2004) Hydrogen solubility and speciation in natural, gem-quality chromian diopside. Am Mineral 89:941–949 Burnley PC, Navrotsky A (1996) Synthesis of high-pressure hydrous magnesium silicates: observations and analysis. Am Mineral 81:317–326 Cameron M, Sueno S, Prewitt CT, Papike JJ (1973) High temperature crystal chemistry of acmite, diopside, hedenbergite, jadeite, spodumene, and ureyite. Am Mineral 58:594-618 Coes L (1962) Synthesis of minerals at high pressures. In: Modern Very High Pressure Techniques. Wentorf R (ed) Butterworths, p 137-150 Comodi P, Fumagalli P, Nazzareni S, Zanazzi PF (2005) The 10 Å phase: Crystal structure from single-crystal X-ray data. Am Mineral 90:1012-1016 Daniels P, Wunder B (1993) Crystal structure of tri-aluminium tri-hydroxo disilicate, Al3Si2O7(OH)3. Z Kristallogr 206:103-105 Daniels P, Wunder B (1996) Crystal structure of phase Pi, Al3Si2O7(OH)3. Eur J Min 8 1283 Domanik K, Holloway JR (1996): The stability and composition of phengitic muscovite and associated phases from 5.5 to 11 GPa: Implications for deeply subducted sediments. Geochim. Cosmochim Acta 64:41334150 Eggleton RA, Boland JN, Ringwood AE (1978) High-pressure synthesis of a new aluminum silicate: Al5Si5O17OH. Geochem J 12:191–194 Fasshauer DW, Chatterjee ND, Marler B (1997) Synthesis, structure, thermodynamic properties and stability relations of K-cymrite. Phys Chem Minerals 24:455-462 Finger LW, Hazen RM, Zhang J, Ko J, Navrotsky A (1993) The effect of Fe on the crystal structure of wadsleyite β-(Mg1-xFex)2SiO4, 0.00 < x < 0.40. Phys Chem Minerals 19:361-368 Finger LW, Ko J, Hazen RM, Gasparik T, Hemley RJ, Prewitt CT, Weidner DJ (1989) Crystal chemistry of Phase B and an anhydrous analogue: Implication for water storage in the mantle. Nature 341:140-142 Franz G, Liebscher A (2004) Physical and chemical properties of the epidote minerals – an introduction. Rev Mineral Geochem 56:1-81

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Friedrich A, Lager GA, Kunz M, Chakoumakos BC, Smyth JR, Schultz AJ (2001) Temperature-dependent single-crystal neutron diffraction study of natural chondrodite and clinohumite. Am Mineral 86:981-989 Frost DJ, Fei Y (1999) Static compression of the hydrous magnesium silicate phase D to 30 GPa at room temperature. Phys Chem Minerals 26:415-418 Fujino K, Sasaki S, Takeuchi Y, Sadanaga R (1981) X-ray determinations of electron distributions in forsterite, fayalite, and tephroite. Acta Crystallogr B 37:513-518 Fumagalli P, Poli S (2005) Experimentally determined phase relations in hydrous peridotites to 6.5 GPa and their consequences on the dynamics of subduction zones. J Petrol 46:555–578 Fumagalli P, Stixrude L, Poli S, Snyder D (2001) The 10 Å phase: a high-pressure expandable sheet silicate stable during subduction of hydrated lithosphere. Earth Planet Sci Lett 186:125-141 Gabe EJ, Portheine JC, Whitlow SH (1973) A reinvestigation of the epidote structure: confirmation of the iron location. Am Mineral 58:212-253 Grevel K-D, Nowlan, EU, Fasshauer DW, Burchard M (2000) X-ray diffraction analysis of lawsonite and zoisite at high pressures and temperatures. Am Mineral 85:206-216 Hawthorne FC, Grundy HD (1976) The crystal chemistry of the amphiboles IV. X-ray and neutron refinements of the crystal structure of tremolite. Can Mineral 14:334-345 Hazen RM (1976) Effects of temperature and pressure on the cell dimension and X-ray temperature factors of periclase. Am Mineral 61:266-271 Hazen RM (1981) Systematic variation of bulk modulus of wüstite with stoichiometry. Carnegie Inst Wash Yearb 80:277-280 Hazen RM, Burnham CW (1973) The crystal structures of one-layer muscovite and annite. Am Mineral 58: 889-900 Hazen RM, Finger LW (1979) Crystal structure and compressibility of zircon at high pressure. Am Mineral 64: 196-201 Hazen RM, Finger LW, Ko J (1992) Crystal chemistry of Fe-bearing anhydrous phase B: Implications for transition zone mineralogy. Am Mineral 77:217-220 Higo Y, Inoue T, Irifune T, Yurimoto H (2001) Effect of water on the spinel-postspinel transformation in Mg2SiO4. Geophys Res Lett 28:3505-3508 Hirschmann MM, Aubaud C, Withers AC (2005) Storage capacity of H2O in nominally anhydrous minerals in the upper mantle. Earth Planet Sci Lett 236:167-181 Holl CM (2006) Effects of hydration on the structure and compression of wadsleyite: Implications for water in the Earth’s interior. Ph.D. Dissertation Geological Sciences, University of Colorado, Boulder. Horioka K, Takahashi K, Morimoto N, Horiuchi H, Akaogi M, Akimoto S (1981) Structure of nickel aluminosilicate (Phase IV): A high-pressure phase related to spinel. Acta Crystallogr B37:635-638 Horiuchi H, Morimoto N, Yamamoto K, Akimoto S (1979) Crystal structure of 2Mg2SiO4 · 3Mg(OH)2, a new high-pressure structure type. Am Mineral 64:593-598 Horiuchi H, Hirano M, Ito E, Matsui Y (1982) MgSiO3 (ilmenite-type): Single crystal X-ray diffraction study. Am Mineral 67:788-793 Horiuchi H, Ito E, Weidner D (1987) Perovskite-type MgSiO3: Single-crystal X-ray diffraction study. Am Mineral 72:357-360 Huang X, Xu Y, Karato SI (2005) Water content of the transition zone form electrical conductivity of wadsleyite and ringwoodite. Nature 434:746-749 Inoue T, Yurimoto H, Kudoh Y (1995) Hydrous modified spinel, Mg1.75SiH0.5O4: a new water reservoir in the mantle transition region. Geophys Res Lett 22:117-120 Jacobsen SD (2006) Effect of water on the equation of state of nominally anhydrous minerals. Rev Mineral Geochem 62:321-342 Jacobsen SD, Smyth JR (2006) Effect of water on the sound velocities of ringwoodite in the transition zone. Earth’s Deep Water Cycle, AGU Monograph Series (in press) Jacobsen SD, Lin JF, Angel RJ, Shen G, Prakapenka VB, Dera P, Mao HK, and Hemley RJ (2005) Single-crystal synchrotron X-ray diffraction study of wüstite and magnesiowüstite at lower-mantle pressures. J Synchr Rad 12:577-583 Jacobsen SD, Demouchy S, Frost DJ, Boffa-Ballaran T, Kung J (2005) A systematic study of OH in hydrous wadsleyite from polarized FTIR spectroscopy and single-crystal X-ray diffraction: Oxygen sites for hydrogen storage in the Earth’s interior. Am Mineral 90:61-70 Jacobsen SD, Smyth JR, Spetzler H, Holl CM, Frost DJ, (2004) Sound velocities and elastic constants of ironbearing hydrous ringwoodite. Phys Earth Planet Int 143-144:47-56 Jacobsen SD, Smyth JR, Swope RJ, Sheldon RI (2000) Two proton positions in the very strong H-bond of serandite [Na,Mn2Si3O8OH]. Am Mineral 85:745-752 Libowitzky E, Beran A (2006) The structure of hydrous species in nominally anhydrous minerals: information from polarized IR spectroscopy. Rev Mineral Geochem 62:29-52

Hydrogen in High Pressure Silicate & Oxide Mineral Strutures

113

Kagi H, Parise JB, Cho HM, Rossman GR, Loveday JS (2000) Hydrogen bonding interactions in phase A [Mg7Si2O8(OH)6] at ambient and high pressure. Phys Chem Minerals 27:225–233 Kanzaki M (1991) Stability of hydrous magnesium silicates in the mantle transition zone. Phys Earth Planet Int 66:307–312 Karato S (1990) The role of hydrogen in the electrical conductivity of the upper mantle. Nature 347:272-273 Karato S, Paterson MS, Fitzgerald D (1986) Rheology of synthetic olivine aggregates: Influence of grain size and water. J Geophys Res 91:8151-8176 Kavner A (2003) Elasticity and strength of hydrous ringwoodite at high pressure. Earth Planet Sci Lett 214: 645-654 Katayama I, Nakashima S (2003) Hydroxyl in clinopyroxene from the deep subducted crust: Evidence for H2O transport into the mantle. Am Mineral 88:229-234 Kawamoto T, Hervig RL, Holloway JR (1996) Experimental evidence for a hydrous transition zone in the early Earth’s mantle. Earth Planet Sci Lett 142:587-592 Kek S, Aroyo M, Bismayer U, Schmidt C, Eichhorn K, Krane HG (1997) The two-step phase transition of titanite, CaTiSiO5: a synchrotron radiation study. Zeit Kristallogr 212:9–19 Keppler H, Smyth JR (2005) Optical and near infrared spectra of ringwoodite to 21.5 GPa. Am Mineral 90: 1209-1214 Kihara K (1990) An X-ray study of the temperature dependence of the quartz structure. Eur J Mineral 2:63-77 Koch-Mueller M, Dera P, Fei Y, Hellwig H, Liu Z, Van Orman J, Wirth R (2005) Polymorphic phase transition in superhydrous phase B. Phys Chem Minerals 32:349-361 Koch-Müller M, Dera P, Fei Y, Reno B, Sobolev N, Hauri E, Wysoczanski R (2003) OH- in synthetic and natural coesite. Am Mineral 88:1436-1445 Koch-Müller M, Fei Y, Hauri E, Liu Z (2001) Location and quantitative analysis of OH in coesite. Phys Chem Minerals 28:693-705 Kohlstedt DL, Keppler H, Rubie DC (1996) Solubility of water in the α, β, and γ phases of (Mg,Fe)2SiO4. Contr Mineral Petrol 123:345-357 Kohn SC, Brooker RA, Frost DJ, Slesinger AE, Wood BJ (2002) Ordering of hydroxyl defects in hydrous wadsleyite (β-Mg2SiO4). Am Mineral 87:293-301 Kudoh Y, Kuribayashi T, Mizohata H, Ohtani E (2000) Structure and cation disorder of hydrous ringwoodite, γ-Mg1.89Si0.97H0.34O4. Phys Chem Minerals 27:474-479 Kudoh Y, Finger LW, Hazen RM, Prewitt CT, Kanzaki M, Veblen DR (1993) Phase E: a high pressure hydrous silicate with unique crystal chemistry. Phys Chem Minerals 19:357-360 Lager GA, Von Dreele RB (1996) Neutron powder diffraction study of hydrogarnet to 9.0 GPa. Am Mineral 81: 1097-1104 Lager GA, Marshall WG, Liu Z, Downs RT (2005) Re-examination of the hydrogarnet structure at high pressure using neutron powder diffraction and infrared spectroscopy. Am Mineral 90:639-644 Lager GA, Ulmer P, Miletich R, Marshall WG (2001) O-D...O bond geometry in OD-chondrodite. Am Mineral 86:176-180 Lager GA, Armbruster T, Faber J (1987) Neutron and X-ray diffraction study of hydrogarnet Ca3A12(O4H4)3. Am Mineral 72:756–765 Libowitzky E (1999) Correlation of O-H stretching frequencies and O-H. . .O hydrogen bond lengths in minerals. Monatsh. Chemie 130:1047-1059 Litasov K, Ohtani E, Langenhorst F, Yurimoto H, Kubo T, Kondo T (2003) Water solubility in Mg-perovskites and wate storage capacity in the lower mantle. Earth Planet Sci Lett 211:189-203 Luth RW (1995) Is phase A relevant to the Earth’s mantle? Geochim Cosmochim Acta 59:679-682 Ma C-B, Sahl K (1975) Nickel aluminosilcate, phase III. Acta Crystallogr B31:2142-2143 Meade C, Reffner JA, Ito E (1994) Synchrotron infrared absorbance measurements of hydrogen in MgSiO3 perovskite. Science 264:1558–1560 Mellini M (1982) The crystal structure of lizardite 1T: Hydrogen bonds and polytypism. Am Mineral 67:587598 Mierdel K, Keppler H (2004) The temperature dependence of water solubility in enstatite. Contrib Mineral Petrol 148:305–311 Mierdel K, Kepler H, Smyth JR, Langenhorst F (2006) The origin of the Earth’s asthenosphere. Nature (in review) Mosenfelder JL (2000) Pressure dependence of hydroxyl solubility in coesite. Phys Chem Minerals 27:610617 Mosenfelder JL, Deligne NI, Asimow PD, Rossman GR (2006) Hydrogen incorporation into olivine at 2-12 GPa. Am Mineral 91:285-294 Murakami M, Hirose K, Yurimoto H, Nakashima S, Takafuji N (2002) Water in the Earth’s lower mantle. Science 295:1885-1887

114

Smyth

Murakami M, Hirose K, Kawamura K, Sata N, Ohishi Y (2004) Post-perovskite phase transition in MgSiO3. Science 304:855-858 Nasdala L, Beran A, Libowitzky E, Wolf D (2001) The incorporation of hydroxyl groups and molecular water in natural zircon (ZrSiO4). Am J Sci 301:831-857 Newnham RE, deHaan YM (1962) Refinement of the alpha-Al2O3, Ti2O3, V2O3, and Cr2O3 structures. Z Kristallogr 117:235-237 Northrup PA, Leinenweber K, Parise JB (1994) The location of H in the high-pressure synthetic Al2SiO4(OH)2 topaz analogue. Am Mineral 79:401-404 Pacalo REG, Parise JB (1992) Crystal structure of superhydrous B, a hydrous magnesium silicate synthesized at 1400 °C and 20 GPa. Am Mineral 77:681-684 Pannhorst W (1984) High temperature crystal structure refinements of low clinoenstatite up to 700 °C. N Jahrb Mineral Abh 150:219-228 Pawley AR (1994) The pressure and temperature stability limits of lawsonite: implications for H2O recycling in subduction zones. Contr Mineral Petrol 118:99-108 Pawley AR, Redfern SAT, Holland TJB (1996) Volume behavior of hydrous minerals at high pressure and temperature: 1. Thermal expansion of lawsonite, zoisite, clinozoisite, and diaspore. Am Mineral 81:335340 Perdikatsis B, Burzlaff H (1981) Strukturverfeinerung am Talk Mg3[(OH)2Si4O10]. Z Kristalogr 156:177-186 Poli S, Schmidt MW (2004) Experimental subsolidus studies of epidote minerals. Rev Mineral Geochem 56: 171-195 Rauch M, Keppler H (2002) Water solubility in orthopyroxene. Contrib Mineral Petrol 143:525-536 Ross NL, Crichton WA (2001) Compression of synthetic hydroxyl clinohumite [Mg9Si4O16(OH)2] and hydroxyl chondrodite [Mg5Si2O8(OH)2]. Am Mineral 86:990-996 Ross NL, Gibbs GV, Rosso KM (2003) Potential docking sites and positions of hydrogen in high-pressure silicates. Am Mineral 88:1452-1459 Ross NL, Shu JF, Hazen RM, Gasparik T (1990) High-pressure crystal chemistry of stishovite. Am Mineral 75: 739-747 Rossman GR, Smyth JR (1990) Hydroxyl contents of accessory minerals in mantle eclogites and related rocks. Am Mineral 75:775-780 Rothbauer R. (1971) Untersuchung eines 2M1-Muskovits mit Neutronenstrahlung. Neues J Mineral Monats 1971:143-154 Schilling FR, Sinogeikin SV, Bass JD (2003) Single-crystal elastic properties of lawsonite and their variation with temperature. Phys Earth Planet Inter 136:107-118 Schmidt MW, Poli S (1998) Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation. Earth Planet Sci Lett 163:361-179 Schmidt MW, Finger LW, Angel RJ, Dinnebier RE (1998) Synthesis, crystal structure and phase relations of AlSiO3OH, a high pressure hydrous phase. Am Mineral 83:881-888 Shieh SR, Mao HK, Konzett J, Hemley RJ (2000) In-situ high pressure X-ray diffraction of phase E to 15 GPa. Am Mineral 85:765-769 Shintani H, Sato S, Saito Y (1975) Electron density distribution in rutile crystals. Acta Crystallogr B31:19811982 Shu J, Mao HK, Hu J, Fei Y, Hemley RJ (1998) Single-crystal X-ray diffraction of wüstite to 30 GPa hydrostatic pressure. N Jahrb Mineral Abh 172:309–323 Skogby H (2006) Water in natural mantle minerals I: pyroxenes. Rev Mineral Geochem 62:155-167 Smyth JR (1980) Cation vacancies and the crystal chemistry of breakdown reactions in kimberlitic omphacites. Am Mineral 65:1185-1191 Smyth JR (1987) β-Mg2SiO4: a potential host for water in the mantle? Am Mineral 72:1051-1055 Smyth JR (1989) Electrostatic characterization of oxygen sites in minerals. Geochim Cosmochim Acta 53: 1101-1110 Smyth JR, Artioli G, Smith JV, Kvick A (1987) Crystal structure of coesite a high pressure form of SiO2 at 15 and 298 K from single crystal neutron and X-ray diffraction data: test of bonding models. J Phys Chem 91:988-992 Smyth JR, Dyar MD, May HH, Bricker OP, and Acker HA (1997) Crystal structure refinement and Mössbauer spectroscopy of an ordered triclinic chlorite. Clays Clay Mineral 45:544-550 Smyth JR, Frost DJ (2002) The effect of water on the 410-km discontinuity: an experimental study. Geophys Res Lett 29:1485, doi:10.1029/2001GL014418 Smyth JR, Kawamoto T (1997) Wadsleyite II: A new high pressure hydrous phase in the peridotite-H2O system. Earth Planet Sci Lett 146:E9-E16 Smyth JR, Frost DJ, Nestola F, Holl CM, Bromiley G (2006a) Olivine hydration in the deep upper mantle: effects of temperature and silica activity. Geophys Res Lett 33, doi:2006GL026194 (in press)

Hydrogen in High Pressure Silicate & Oxide Mineral Strutures

115

Smyth JR, Holl CM, Frost DJ, Jacobsen SD, (2004) High pressure crystal chemistry of hydrous ringwoodite and water in the Earth’s interior. Phys Earth Planet Int 143-144:271-278 Smyth JR, Holl CM, Frost DJ, Jacobsen SD, Langenhorst F, McCammon CA (2003) Structural systematics of hydrous ringwoodite and water in the Earth’s interior. Am Mineral 88:1402-1407 Smyth JR, Holl CM, Langenhorst F, Laustsen HMS, Rossman GR, Kleppe A, McCammon CA, Kawamoto T, van Aken PA (2005) Crystal chemistry of wadsleyite II and water in the Earth’s interior. Phys Chem Minerals 31:691-705 Smyth JR, Jacobsen SD, Swope RJ, Angel RJ, Arlt T, Domanik K, Holloway JR (2000) Crystal structures and compressibilities of synthetic 2M1 and 3T phengite micas. Eur J Mineral 12:955-963 Smyth JR, Kawamoto T, Jacobsen SD, Swope RJ, Hervig RL, Holloway JR (1997) Crystal structure of monoclinic hydrous wadsleyite. Am Mineral 82:270-275 Smyth JR, Mierdel K, Keppler H, Langenhorst F, Dubrovinsky L, Nestola F (2006b) Crystal chemistry of hydration in aluminous orthopyroxene. American Mineralogist (in prep.) Smyth JR, Rossman GR, Bell DR (1991) Incorporation of hydroxyl in upper mantle clino-pyroxenes. Nature 351:732-735 Smyth JR, Swope RJ, Pawley AJ (1995) H in rutile-type compounds: II. Crystal chemistry of Al substitution in H-bearing stishovite. Am Mineral 80:454-456 Sueno S, Cameron M, Prewitt CT (1976) Orthoferrosilite: High temperature crystal chemistry. Am Mineral 61: 38-53 Swope RJ, Smyth JR, Larson AC (1995) H in rutile-type compounds: I. Single-crystal neutron and X-ray diffraction study of H in rutile. Am Mineral 80:448-453 Takeda H (1973) Tetrahedral sizes of orthopyroxenes and silicon-aluminum ordering. Am Mineral 58:10961097 Vlassopoulos D, Rossman GR, Haggerty SE (1993) Coupled substitution of H and minor elements in rutile and the implications of high OH contents in Nb- and Cr-rich rutile from the upper mantle. Am Mineral 78: 1181-1191 Vanpeteghem CB, Ohtani E, Kondo T, Takemura K, Kikegawa T (2003) Compressibility of phase Egg AlSiO3OH: Equation of state and role of water at high pressure. Am Mineral 88:1408-1411 Winter JK, Ghose S (1979) Thermal expansion and high temperature crystal chemistry of the Al2SiO5 polymorphs. Am Mineral 64:573-586 Wood BJ (1995) The effect of H2O on the 410-kilometer seismic discontinuity. Science 268:74-76 Woodhead JA, Rossman GR, Thomas AP (1991) Hydrous species in zircon. Am Mineral 76:1533-1546 Wunder B (1998) Equilibrium experiments in the system MgO-SiO2-H2O (MSH): stability fields of clinohumiteOH [Mg9Si4O16(OH)2], chondrodite-OH [Mg5Si2O8(OH)2] and phase A [Mg7Si2O16(OH)6]. Contr Mineral Petrol 132:111-120 Wunder B, Andrut M, Wirth R (1999) High pressure synthesis and properties of OH-rich topaz. Eur J Mineral 11:803-813 Wunder B, Medenbach O, Daniels P, Schreyer W (1995) First synthesis of the hydroxyl end-member of humite Mg7Si3O12(OH)2. Am Mineral 80:638-640 Wunder B, Medenbach O, Krause W, Schreyer W (1993)Synthesis, properties and stability of Al3Si2O7(OH)3 (phase Pi), a hydrous high-pressure phase in the system Al2O3-SiO2-H2O (ASH). Eur J Mineral 5:637649 Yamamoto K, Akimoto S (1977) The system MgO-SiO2-H2O at high pressure and temperatures – stability field for hydroxyl-chondrodite, hydroxyl-clinohumite, and 10 Å-phase. Am J Sci 277:288-312 Yagi T, Marumo F, Akimoto S (1974) Crystal structures of spinel polymorphs of Fe2SiO4 and Ni2SiO4. Am Mineral 59:486-490 Yang H, Prewitt CT (2000) Chain and layer silicates at high temperatures and pressures. Rev Mineral Geochem 41:211-256 Yang H, Prewitt CT, Frost DJ (1997) Crystal structure of the dense hydrous magnesium silicate, phase D. Am Mineral 82:651-654 Yang H, Hazen RM, Prewitt CT, Finger LW, Lu R, Hemley RJ (1998) High-pressure single-crystal X-ray diffraction and infrared spectroscopic studies of C2/m-P21/m phase transition in cummingtonite. Am Mineral 83:288-299 Yang H, Ghose S (1995) High temperature single crystal X-ray diffraction study of the ortho-proto phase transition in enstatite Mg2Si2O6. Phys Chem Minerals 22:300-310 Zigan F, Rothbauer R (1967) Neutronenbeugungsmessung am Brucit. N Jahrb Mineral Monats 137-143

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Reviews in Mineralogy & Geochemistry Vol. 62, pp. 117-154, 2006 Copyright © Mineralogical Society of America

Water in Nominally Anhydrous Crustal Minerals: Speciation, Concentration, and Geologic Significance Elizabeth A. Johnson* Department of Earth and Space Sciences University of California, Los Angeles Los Angeles, California, 90095, U.S.A. e-mail: [email protected] (*present address: Dept. of Geology & Environmental Sciences, James Madison Univ., Harrisonburg, VA, 22807)

INTRODUCTION Importance of nominally anhydrous minerals in the crust Why should we be interested in trace hydrous species in nominally anhydrous minerals in the Earth’s crust? After all, hydrous minerals dominate the pedosphere and are abundant to fairly common trace minerals in many metamorphic and igneous crustal rocks. On the other hand, the most abundant minerals in the crust—feldspars, quartz, pyroxenes, and garnet—are all nominally anhydrous. They are present even in systems with low total volatiles or fluid contents, or environments with low water activities where hydrous minerals are unstable. These nominally anhydrous minerals provide an opportunity to expand the extent of our knowledge of fluid composition and water activity, as well as the influence of water on physical properties and geochemical signatures of rocks. One advantage to investigations of the crustal component of the lithosphere is that many parts of the crust (especially the continental crust) are available for direct study in outcrops at the surface of the Earth. This allows the nominally anhydrous mineral and its hydrous species to be placed into the context of the hand sample, the outcrop, and even the regional geology.

Scope and goals of this chapter It would be unrealistic to try to cover every water-bearing mineral in the Earth’s crust in this chapter. I have limited my discussion to minerals that do not require hydrous species to complete their stoichiometry, and those for which research has been completed on natural crustal samples. These minerals are: quartz, the feldspars, nepheline, pyroxenes, garnets (except pyrope), kyanite, andalusite, sillimanite, rutile, cassiterite, zircon, titanite, cordierite, and beryl. This selection of minerals restricts the discussion primarily to the continental crust below about 3 km depth. Some references to eclogitic and mantle-wedge minerals are included for completeness. This is a fairly new field of study, and as such, the goal of this chapter is to give an overview of the work that has been done, and more importantly, provide directions for future work. The chapter begins with an overview of the types of hydrous species and the range of absolute concentrations for each mineral or mineral family. The second section provides examples of applications of these measurements to problems of geologic interest. It is assumed that the reader is familiar with the compositions and general structure and crystal chemistry of these minerals. It is also helpful to have a general understanding of absolute OH concentration measurement techniques, and a reading knowledge of polarized infrared spectra of hydrous species in minerals. Overviews of these topics may be found in 1529-6466/06/0062-0006$05.00

DOI: 10.2138/rmg.2006.62.6

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Rossman (1988, 2006); Libowitzky and Beran (2006); Smyth (2006). Hydrogen abundance measurements discussed in this chapter are generally obtained using manometric or infrared spectroscopic methods. A summary of infrared spectroscopic calibrations for common mineral species is given in Table 2 of Rossman (2006). The reader should consult the reference of interest for detailed information about the absorption coefficient used in a particular study.

HYDROUS SPECIES AND CONCENTRATIONS IN CRUSTAL MINERALS Quartz and coesite Quartz, a common crustal mineral, contains structural OH groups, macroscopic fluid inclusions, and nanoscale “fluid inclusions” or water clusters (especially seen in synthetic quartz). Previous studies have compiled detailed summaries of the hydrous species in natural and synthetic quartz, chert, opal, and chalcedony (Aines and Rossman 1984a; Rossman 1988). The infrared spectrum of OH in a natural quartz crystal from Brazil is shown in Figure 1. Diffusion and electrolytic exchange experiments in natural and synthetic α-quartz have established that these sharp bands are due to OH groups associated with other H+ or monovalent cations including Li+, Na+, K+, Cu+, and Ag+, and hydroxyl associated with Al3+ (Kats 1962; Aines and Rossman 1984a; Rovetta et al. 1986; Miyoshi et al. 2005). This structural OH is most commonly found in large, clear, undeformed quartz crystals from high-temperature pegmatites as well as synthetic quartz, although some structural OH bands may occur in spectra of other low-temperature forms of quartz (such as amethyst) (Aines and Rossman 1984a; Kronenberg and Wolf 1990). The OH bands in quartz have been calibrated (Chakraborty and Lehmann 1976) and the reported range of OH concentrations is 1 Ga (Hawkesworth et al. 1983; Cohen et al. 1984; Kinny et al 1989; 1994). There are also examples where the timing of metasomatism in xenoliths carried by alkali basalts cannot be separated from the magmatic event that brought the samples to the surface (Menzies and Murthy 1980). The fluids or melts that caused the chemical changes are only rarely found as crystallized melts or glasses in mineral inclusions (Schrauder and Navon 1994). In some composite xenoliths and alpine massifs metasomatism can be observed in the wall rocks adjacent to veins and dikes (Jones et al. 1982; Boyd 1990; Woodland et al. 1996). Small degree melts that crystallize in the mantle and expel fluids into the wall rocks are implicated in many metasomatic events. Some alpine peridotite massifs, however, seem to have experienced phases of pervasive metasomatism over large regions, with no apparent relationship to veining (Zanetti et al. 1999; Scambelluri et al. 2006) Metasomatism can occasionally be directly attributed to subduction zones processes, where the slab provides both fluids and a source of incompatible elements (Brandon and Draper 1996; Zanetti et al 1999; McInnes et al. 2001; Scambelluri et al. 2006). Although many metasomatized regions of the subcontinental lithosphere have not been at convergent margins for over a billion years, it is highly likely that components added by ancient subduction could be remobilized at a later date as a result of heating or decompression. In many instances metasomatism is more directly attributed to the infiltration and crystallization of small degree melts that migrate from the asthenosphere as a result of plume activity or decompression due to rifting. Menzies et al

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(1987) proposed that in general terms the chemical changes that are induced in the lithosphere occur because it forms a mechanical barrier between asthenospheric melts and the surface. A number of geochemical and laboratory based studies have addressed the nature of metasomatic agents in the mantle (Ryabchikov and Boettcher 1980; Schneider and Eggler 1986; McNeil and Edgar 1987; Bodinier et al. 1988; Gregoire et al. 2003) but it is often difficult to categorically attribute natural metasomatic assemblages to specific fluid or melt compositions. H2O-rich fluids are often implicated as metasomatic agents but the conditions where they can exist in the mantle are constrained to relatively low temperatures (e.g.,  85 and are generally Ti-poor. They occur with a wide range of fertility from harzburgite to lherzolite and detailed studies often reveal complex histories involving melt extraction, metasomatism and reaction with the host magma during transport to the surface. Where hydrous minerals occur they are generally the calcic amphibole pargasite, (often replacing spinel), and phlogopite mica. Group II (Frey and Prinz 1978) or Al-augite rocks (Wilshire and Shervais 1975) occur as clinopyroxene dominated veins or layers in xenoliths. They are Ti-rich, Cr-poor and have Mg# 8 GPa, K-richterite stability is very close to that of the pure phase. Only the thermal maxima is reduced by approximately 100 °C compared to the pure phase stability determined by Trønnes (2002). The vast majority of mantle peridotite rocks, on the other hand, are subalkaline (i.e., Na2O+K2O)/Al2O3 > 1). K-richterite is only stable in subalkaline lherzolitic bulk compositions above 6-7 GPa as a result of the reaction: 0.5K2Mg6Al2Si6O20(OH)4 + CaMgSi2O6 + NaAlSi2O6 + Mg2Si2O6 = phlogopite

in cpx

opx

KNaCaMg5Si8O22(OH)2 + Mg3Al2Si3O12 (7) K-richterite

garnet

Frost

260

Pressure (GPa)

16

Di+St +Cen+fl

5

Di+Cen +Wad+X+fl

12

8

1

K-richterite

Di+Cen +Wad+fl 4

2

4 3

Di+En+fl

K-richterite stable or unstable in peridotite

0 900 1000 1100 1200 1300 1400 1500 1600

Temperature ( C) o

Figure 11. The stability field of pure KNaCaMg5Si8O22(OH)2 K-richterite as bracketed by Trønnes (2002) is shown by curve (A). All of the indicated named products are with respect to curve (A) with Wad being wadite-structured K2Si4O9, X is Phase X, en and Cen are enstatite and clinoenstatite, Di is diopside and St is stishovite. CurvesFigure (B) 11and (C) are stability fields of K-richterite determined by Foley (1991) and Gilbert and Briggs (1974), respectively. The closed and open symbols indicate the presence and absence of K-richterite in a synthetic KNCMASH subalkaline peridotite assemblage, as determined by Konzett and Ulmer (1999) and Konzett and Fei (2000). In these experiments the coexisting assemblage always contained olivine, garnet, clinopyroxene and enstatite or clinoenstatite. The breakdown products at high pressure also contain Phase X and below curve (D) K-richterite breaks down to a phlogopite-bearing assemblage. Curve (E) shows the high pressure stability of KK-richterite K2CaMg5Si8O22(OH)2 as determined by Inoue et al. (1998).

Equation (7), the Na present equivalent of Equation (6) proposed by Sudo and Tatsumi (1990), demonstrates that the breakdown of minor amounts of phlogopite in the presence of pyroxenes containing some jadeite component can produce the observed natural mantle Krichterite without any fluid release. The identical K/OH-ratio of phlogopite and K-richterite (but not KK-richterite) is the fundamental requirement for such a fluid-free reaction (Konzett and Ulmer 1999). Equation (7) and the general lack of K-richterite in garnet-bearing mantle xenoliths (e.g., Erlank et al. 1987), indicate that the majority of such xenoliths equilibrated at depths shallower than 200 km.

Experimental studies on the stability of potential high pressure hydrous mantle minerals Whereas most mantle xenoliths originate from the upper 200 km of the mantle or up to approximately 7 GPa, experimental studies at higher pressures have identified a number of other hydrous minerals that are potentially stable in the deeper parts of the mantle, although mainly in subduction zones. These phases generally don’t have mineral names and are simply referred to by letters e.g., A, B, superhydrous B, D, E and X. Their stability in subduction zones is covered in this volume by Kawamoto (2006). The criterion for evaluating their presence in the ambient mantle is their compatibility with typical mantle minerals and their high temperature stability,

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which can only be assessed through experimental studies. Here I only consider phases with upper thermal stability limits that are close to an average mantle adiabat.

Phase X The enigmatically named Phase X has been observed in several studies as a high-pressure product of the decomposition of K-richterite. Phase X has a variable composition with reported K2O contents of between 10 and 19 wt%. In the KCMSH system Inoue et al. (1998) reported Phase X with the approximate composition K4Mg8Si8O25(OH)2 whereas Trønnes (2002) reported the composition K3.7Mg7.4Al0.6Si8O25(OH)2 in the KMASH system. In addition to Phase X with the formula K1.54Mg1.93Si1.89O7H1.04, Yang et al. (2001) synthesized and solved the structures of sodic Phase X, Na1.16K0.01Mg1.93Al0.14Si1.89O7H1.04, and the anhydrous end members K1.85Mg2.06Si2.01O7 and Na1.78Mg1.93Al0.13Si2.02O7. Phase X is composed of layers of brucite-like MgO6 octahedra linked by Si2O7 tetrahedral dimers and K cations (Yang et al. 2001; Mancini et al. 2002). Yang et al. (2001) proposed the general formula A2−xM2Si2O7Hx where A can be K and/or Na, M can be Mg or Al and x = 0-1. An increase in the K content of Phase X is therefore coupled to a decrease in the H content. The only measurement of the H2O content of Phase X, performed using SIMS, yielded a value of 1.7±0.1 wt% H2O (Inoue et al. 1998), which is significantly below the theoretical maximum of 3.51 wt%. No studies have been performed on the stability of any pure Phase X composition; however, Konzett and Fei (2000) have examined the stability of Phase X in a subalkaline KNCMASH analogue peridotite composition. Phase X coexists with a typical mantle assemblage of olivine/wadsleyite, clinopyroxene and garnet between 14 and 20 GPa and at temperatures up to 1600 °C. Phase X, therefore, has the highest thermal stability of any yet investigated nominally hydrous silicate. As shown in Figure 12 the high temperature stability of Phase X is for the main part undetermined. Konzett and Fei (2000) showed that the reactions that produce Phase X from K-richterite in a mantle peridotite composition release fluid because the K/H ratio of Phase X is higher than that of K-richterite. These results also show that the K content and K/Na ratio of Phase X both increase with pressure, which implies a decrease in the H2O content of Phase X with pressure. The change in K/Na ratio occurs as Na is partitioned into coexisting garnet with increasing pressure. Between 20 and 22 GPa Phase X breaks down to an assemblage containing K-hollandite (KAlSi3O8).

K-Hollandite bearing

Phase X out

Pressure (GPa)

K

nd lla -ho

Phase X bearing

K-richterite bearing

Temperature (oC)

in ite

Figure 12. The closed and open symbols indicate the presence and absence of phase-X in a synthetic KNCMASH subalkaline peridotite assemblage, as determined Konzett and Fei (2000). In these experiments the coexisting assemblage was that expected for a peridotite composition at the indicated conditions, i.e., olivine or high-pressure polymorphs, garnet and Ca-perovskite at and above 20 GPa. Filled rectangles show conditions where Luth (1997) observed Phase X in a KCMASH bulk composition. At high pressures Phase X breaksdown to an assemblage containing K-hollandite (KAlSi3O8).

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Humite and dense hydrous magnesium silicate phases A number of high pressure experimental studies have shown that the humite minerals chondrodite and clinohumite and the dense hydrous magnesium silicate phases A, superhydrous B, D and E can coexist with ultramafic assemblages at various conditions above 6 GPa and below 1200 °C (Kanzaki 1991; Kawamoto et al. 1995; Ohtani et al. 1995; Frost and Fei 1998; Irifune et al. 1998). The stability fields of these phases are significantly below reasonable average mantle adiabats and they are therefore only expected to be stable in the cooler regions of subduction zones, provided that significant H2O is in fact present within such regions at pressures above 6 GPa. Although the stability fields have been examined in natural systems (Luth 1995; Kawamoto et al. 1995; Frost 1999; Kawamoto 2004) there remains some question as to whether the strong partitioning of some element by a particular hydrous phase may cause some increase in thermal stability. In addition the large amounts of H2O added in some bulk compositions may result in the breakdown of hydrous phases at a lower temperature than we might expect in the mantle as a result of excessive melting. Experiments in relatively low-H2O bulk compositions show, however, that the presence of Al and Fe in phases A and E, superhydrous phases B and D has a limited effect on stability relations in comparison to the MSH system (Luth 1995; Frost 1999). Humite minerals have a preference for Ti and F. Titanian clinohumite is a common accessory mineral in metamorphosed ultrabasic rocks and occurs in serpentinites and kimberlites (López Sánchez-Vizcaíno et al. 2005). The stability of titanian clinohumite is below 1000 °C at 8 GPa although the pure fluorine clinohumite end-member is stable to over 1400 °C at 3 GPa (Weiss 1997; Ulmer and Trommsdorf 1999). The experiments of Kawamoto (2004) contained Ti and showed clinohumite and chondrodite stability to be limited to below 1100 °C at 11 GPa. Phase D is the highest-pressure dense hydrous magnesium silicate and its stability in the lower mantle is ultimately controlled by the reaction, MgSi2O4(OH)2 Phase D

+

MgO periclase

=

2MgSiO3 MgSi-perovskite

+

H2O (8) Liquid

The slope of this reaction is not clear however. From a Schreinemakers analysis of the existing experimental data Komabayashi et al. (2004) reported a negative Clapyron slope at approximately 25 GPa with a maximum thermal stability for phase D of 1100°C. Laser heated diamond cell experiments of Shieh et al. (1998) indicate that this reaction leads to the breakdown of phase D at 44 GPa at temperatures between 1000 and 1400 °C, which would be consistent with a convex shape of the reaction boundary of Equation (8), like many other dehydration reactions. Phase D may therefore be stable at temperatures higher than 1100  C at pressures between 25 and 44 GPa but it is probably unlikely that these temperatures approach that of the mantle adiabat. In natural systems phase D contains significant amounts of Al and ferric and ferrous Fe but not in quantities higher than coexisting silicate perovskite, so they have little effect on the thermal stability of phase D (Frost 1999; Frost unpublished data).

The stability of hydrous phases in ultramafic lithosphere and the convecting mantle In considering the significance of hydrous minerals in the mantle it is not only of interest to define stability fields, but it is also important to assess the proportion of hydrous minerals that may exist at particular conditions, identify how much of the mantle’s water budget they may account for and examine further factors, such as H2O activity, that may affect their stability. Changing redox conditions as a function of depth in the upper mantle and transition zone may also control fluid speciation and H2O activity, which, in turn, may affect the stability of the hydrous phases.

Stability of Hydrous Mantle Phases

263

Figure 13 shows the stability fields of the major mantle hydrous phases derived from the previously described experimental studies. The experimental data employed are from studies where hydrous phases formed in equilibrium with typical ultramafic mantle assemblages at H2O undersaturated conditions. A mantle adiabat with a potential temperature of 1600 K (i.e., the temperature at the surface when extrapolated through the melting region) is shown with branching geotherms for Achean cratonic and oceanic lithosphere. A water saturated peridotite solidus interpolated from the data of Mysen and Boettcher (1975) to 4 GPa and Kamamoto (2004) >4 GPa is shown. The solidus is not followed into the region of dense hydrous magnesium silicate stability because huge amounts of H2O are required to produce a melt at these conditions and the solvus between fluid and melt may anyway disappear. Figure 13 indicates that the only hydrous mineral to be stable along an average mantle adiabat (AMA) is Phase X, which could be present in the mantle between depths of 400 and 600 km. The Archean lithospheric geotherm (ACL), which branches off the average mantle adiabat at temperatures approaching 1400 °C, misses the stability field of K-richterite but enters the phlogopite stability field at pressures of approximately 6.8 GPa at 1280 °C. In Figure 13 the data on phlogopite and K-richterite are taken from experiments in Fe-free systems (Konzett and Ulmer 1999; Konzett and Fei 2000). Preliminary experiments seem to indicate that Fe destabilizes these hydrous phases further (Konzett and Ulmer 1999) and the extent of the stability fields in Figure 13 may, therefore, be slightly overestimated. It is important to reiterate that in nature K-richterite occurs in mantle xenoliths of peralkline rocks where the K-richterite stability field extends to much lower pressures (Konzett et al. 1997) than in normal subalkaline

700

D

SB

Phase X out

DHMS

600

Phase X

E 400

A Phlo

300

K-richterite AMA

gopit e ou t

K-richterite in

ACL Phlogopite

Depth (km)

Pressure (GPa)

500

200

OL

100

Na-amphibole

Temperature (oC) Figure 13

Figure 13. Stability fields of hydrous minerals in mantle of peridotite composition at H2O-undersaturated conditions. Data are combined from Figures 5,9,11 and 12. The grey shaded region shows where the dense hydrous magnesium silicate phases A, E, super hydrous phase B (sB) and D are stable from Kawamoto (2005). Thick grey curves show the peridotite solidus under H2O saturated and dry conditions. An average mantle adiabat (AMA) and geotherms for Archean cratonic lithosphere (ACL) and 100 million year old oceanic lithosphere (OL) are shown by thin black lines.

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peridotite compositions that are depicted in Figure 13. It seems that the only environment where K-richterite could exist in subalkaline mantle rocks is in a subduction zone. The oceanic lithosphere geotherm (OL) passes into the stability field of phlogopite and pargasitic amphibole below 3 GPa. The proportion of hydrous minerals that form along typical geotherms will depend on the Na and K content of the mantle, which in turn will depend on the degree of depletion and metasomatism. Using a primitive mantle composition as a benchmark, it is possible to appreciate the degree of metasomatic enrichment necessary for significant hydrous minerals to form in the lithosphere. If we consider a typical primitive mantle Na2O content of 0.3 wt% then from the experiments of Niida and Green (1999) we can calculate that along an oceanic geotherm at approximately 70 km the lherzolitic assemblage could contain 9 wt% pargasitic amphibole which would accommodate approximately 1500 ppm H2O, assuming stoichiometric amphibole OH contents. At only 50 km this rises to approximately 25 wt% pargasite which would host 4000 ppm H2O in the bulk. Significant amounts of amphibole can, therefore, form in lithospheric mantle with typical Na contents, mostly at the expense of clinopyroxene, by adding relatively small amounts of H2O alone. Primitive mantle K contents, on the other hand, are generally 10 times lower than corresponding Na contents. Therefore, along an Archean craton lithospheric geotherm at approximately 150 km depth 0.03 wt% K2O in the bulk rock will allow a maximum of just 0.2 wt% phlogopite to form, which will host 90 ppm H2O, using data from Konzett and Ulmer (1999). In comparison, metasomatized garnet phlogopite peridotite rocks (GPP) reported by Erlank et al. (1987) have average bulk K2O contents of 0.16% which would result in 1.4 wt% phlogopite forming at 150 km with a bulk H2O content of 600 ppm. Erlank et al. (1987) classified GPP rocks as the least metasomatized, whereas PKP rocks, which are considered to be the most metasomatized, have average K2O contents of approximately 1%. The presence of significant phlogopite in some mantle xenoliths means, therefore, that there are processes that occur in the mantle that strongly concentrate K while having much smaller effects on other major elements and in particular Na. One possibility is that such high K-bearing liquids are produced by the breakdown of the white mica phengite in subducting lithosphere (Schmidt et al. 2004). If this is the only explanation then all K-rich metasomatism of the lithosphere must be related to subduction. Another possibility is that high-pressure metasomatic fluids or melts are K-rich because Na becomes compatible in clinopyroxene during melting at high pressures and low temperatures (Blundy et al. 1995). Clinopyroxene/melt partition coefficients for K, on the other hand, are normally 2 orders of magnitude below those of Na. The compositions of low-fraction hydrous melts or fluids at pressures above 3 GPa are poorly constrained but further study may provide important insights into metasomatic agents in the lithosphere. Phase X is the only hydrous mineral that could be stable along an average mantle adiabat in the convecting mantle, at least given the available experimental data that extend to lower mantle conditions (>660 km). Assuming a primitive mantle bulk composition we can calculate how much Phase X could form at the top of the transition zone (410 km) and how much of the convecting mantle’s water budget it could account for at these depths. Using the data of Konzett and Fei (2000) who used a K2O enriched KLB-1 peridotite composition the K2O content of Phase X expected in the transition zone is 14.5 wt% and the stoichiometric H2O content is approximately 3 wt%. If the bulk rock contains 0.03 wt% K2O then 0.1 wt% Phase X can form with the proportion of H2O hosted by Phase X being just 30 ppm. In addition to the low K content of the primitive mantle, at these conditions high Ca-clinopyroxene also contains as much as 0.1% K2O and given the uncertainty on some of the values it is quite possible that all K2O and H2O may be accommodated by the nominally anhydrous assemblage. It is of course also possible that K is inhomogeneously distributed in the convecting mantle in a similar way to that found in the lithosphere, resulting in regions with higher proportions of Phase X. If the bulk of the transition zone has a primitive mantle composition, however, then the formation of these regions must leave the remaining transition zone depleted in K and the amount of H2O stored

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by Phase X over the bulk of the transition zone cannot be much greater than 30 ppm. It seems clear, therefore, that nominally anhydrous minerals and melts, or possibly fluids at reducing conditions, must host the majority of hydrogen stored in the ambient convecting mantle. As shown in Figure 4 and explained previously for pargasitic amphibole, hydrous phases display the highest thermal stability at fluid-absent conditions. Figure 13 should, therefore, depict the maximum thermal stability within ultramafic bulk compositions with respect to water activity. Lower H2O activities or H2O-saturated conditions should lead to lower hydrous mineral stability fields. One of the problems of relating experimental studies of hydrous mineral stability to natural mantle mineral assemblages is that we generally have only circumstantial evidence for the nature of the metasomatic or igneous melt/fluid phase from which the minerals formed. The activity of H2O at the conditions of formation are, therefore, poorly constrained. For this reason the previously described methodology of Popp et al. (1995) to determine H2O activity through the use of Equation (2) is particularly attractive. Above subduction zones for example where high concentrations of H2O may enter the mantle wedge the maximum thermal stability of hydrous minerals such as pargasite or phlogopite may be closer to that of H2O-saturated conditions, shown for pargasite in Figures 4 and 5, which may be a few hundred degrees below those in Figure 13. Another poorly constrained factor is that the oxygen fugacity of the mantle may decrease with depth causing C-O-H fluids to become richer in CH4 and lowering the activity of H2O (Woermann and Rosenhauer 1985; Wood et al. 1990). Several studies have argued for a lowering of mantle ƒo2 with depth as a result of the pressure effect on the ferric-ferrous equilibria that likely define mantle ƒo2 and due to changes in the solubility of ferric iron in major mantle minerals (Wood et al. 1990; Gudmundsson and Wood 1995; O’Neill et al. 1993; Ballhaus and Frost 1994; Frost et al. 2004). Several oxygen thermobarometry studies on garnet peridotite xenoliths have observed a decrease in ƒo2 with depth from values of around FMQ-1 (one log unit below the fayalite-magnetite-quartz oxygen buffer) close to the spinel peridotite field at 80 km depth down to FMQ-4 at approximately 200 km (McCammon et al. 2001; Woodland and Koch 2003; McCammon and Kopylova 2004). O’Neill et al. (1993) argued that oxygen fugacities in the transition zone may be close to the iron-wüstite buffer (IW i.e., ~FMQ-5). At these conditions a C-O-H fluid may contain over 50% CH4 and up to 5% H2 although values are uncertain as equations of states for reduced gas phases are poorly constrained at these conditions (Holloway 1987; Belonoshko and Saxena 1992). H2 contents may also increase depending on the activity of carbon and components such as H2S could also be relevant. Although there is very little experimental data on their behavior, reduced fluid phases may be more mobile in the mantle as their components likely have lower solubilities in minerals and melts and the solubilities of silicate components in these fluids may be low. As previously discussed, Taylor and Green (1988) observed an increase in the fluid-saturated peridotite solidus between 1.0 and 3.5 GPa at low ƒo2 (~FMQ-4) where CH4 became a major fluid component. This occurred because CH4 lowered the H2O activity in the fluid, which lowered the H2O solubility in the coexisting silicate melt. The presence of a reduced fluid phase with a low H2O activity in the mantle may affect hydrous phase stability and may also lower the solubility of hydroxyl in nominally anhydrous phases. The mobility and low density of a reduced fluid phase in the deeper convecting mantle may help to redistribute hydrogen and might even tend to focus H2O in the upper more oxidized regions of the upper mantle.

Acknowledgments I am tremendously grateful to Jurgen Konzett, Reidar Trønnes and Alan Woodland for lengthy discussions and for making numerous comments on an earlier version of the manuscript. I also appreciate the comments and corrections of Hans Keppler and John Winter.

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Adam J, Green TH, Sie SH (1993) Proton microprobe determined partitioning of Rb, Sr, Ba, Y, Nb and Ta between experimentally produced amphiboles and silicate melts with variable F content. Chem Geol 109: 29-49 Adam J, Green TH, Sie SH, Ryan CG (1997) Trace element partitioning between aqueous fluids, silicate melts and minerals. Eur J Mineral 9:569-584 Arai S, Matsukage K, Isobe E, Vysotskiy S (1997) Concentration of incompatible elements in oceanic mantle: effect of melt/wall interaction in stagnant or failed melt conduits within peridotite. Geochim Cosmochim Acta 61:671-675 Agrinier P, Mevel C, Bosch D, Javoy M (1993) Metasomatic hydrous fluids in amphibole peridotites from Zabargad island (Red Sea). Earth Planet Sci Lett 120:187-205 Baker MB, Wyllie PJ (1992) High-pressure apatite solubility in carbonate-rich liquids: implications for mantle metasomatism. Geochim Cosmochim Acta 56:3409-3422 Bailey DK (1982) Mantle metasomatism-continuing chemical change within the Earth. Nature 296:525-530 Ballhaus C, Frost BR (1994) The generation of oxidized CO2-bearing basaltic melts from reduced CH4-bearing upper mantle sources. Geochim Cosmochim Acta 58:4931-4940 Belonoshko AB, Saxena SK (1992) A unified equation of state for fluids of C-H-O-N-S-Ar composition and their mixtures up to very high-temperatures and pressures. Geochim Cosmochim Acta 56:3611-3626 Bureau H, Keppler H (1999) Complete miscibility between silicate melts and hydrous fluids in the upper mantle: experimental evidence and geochemical implications. Earth Planet Sci Lett 165:187-196 Blundy JD, Falloon TJ, Wood BJ, Dalton JA (1995) Sodium partitioning between clinopyroxene and silicate melts. J Geophys Res 100:15501-15515 Bodinier J-L, Dupuy C, Dostal J (1988) Geochemistry and petrogenesis of Eastern Pyrenean peridotites. Geochim Cosmochim Acta 52:2893-2907 Bose K, Ganguly J (1995) Experimental and theoretical studies of the stabilities of talc, antigorite and phase A at high pressures with applications to subduction processes. Earth Planet Sci Lett 136:109-121 Boyd FR (1959) Hydrothermal investigations of amphiboles. In: Researches in Geochemistry Vol 1. Abelson PH (ed) Wiley, p 377-396 Boyd FR, Mertzman SA (1987) Composition and structure of the Kaapvaal lithosphere, South Africa. In: Magmatic Processes: Physiochemical Principles. Mysen B (ed) Geochemical Society Special Publication 1, p 13-24 Boyd FR (1990) Mantle metasomatism: evidence from a MARID-harzburgite compound xenolith. Carnegie Institution Washington Yearbook 90:18-23 Brandon AD, Draper DS (1996) Constraints on the origin of the oxidation state of mantle overlying subduction zones: An example from Simcoe, Washington, USA. Geochim Cosmochim Acta 60:1739-1749 Bryndzia LT, Wood BJ (1990) Oxygen thermobarometry of abyssal spinel peridotites: the redox state and the C-O-H volatile composition of the Earth’s sub-oceanic upper mantle. Am J Sci 290:1093-1116 Bureau H, Keppler H (1999) Complete miscibility between silicate melts and hydrous fluids in the upper mantle: experimental evidence and geochemical implications. Earth Planet Sci Lett 165:187-196 Chazot G, Menzies MA, Harte B (1996) Determination of partition coefficients between apatite, clinopyroxene, amphibole, and melt in natural spinel lherzolites from Yemen: implications for wet melting of the lithospheric mantle. Geochim Cosmochim Acta 60:423-437 Cohen RE, O’Nions RK, Dawson JB (1984) Isotope geochemistry of xenoliths from East Africa: implications for development of mantle reservoirs and their interaction. Earth Planet Sci Lett 68:209-220 Conceicao RV, Green DH (2004) Derivation of potassic (shoshonitic) magmas by decompression melting of phlogopite+pargasite lherzolite. Lithos 72:209-229 Dawson JB, Smith JV (1977) The MARID (mica-amphibole-rutile-ilmenite-diopside) suite of xenoliths in kimberlite. Geochim Cosmochim Acta 41:309-323 Dawson JB, Smith JV (1982) Upper-mantle amphiboles: a review. Mineral Mag 45:35-46 Dawson JB (1984) Contrasting types of upper-mantle metasomatism? In: Kimberlites II: The mantle and CrustMantle Relationships. Kornprobst J (ed) Elsevier, pp 289-294 Dyar MD, Mackwell SJ, McGuire AV, Cross, LR, Robertson JD (1993) Crystal chemistry of Fe3+ and H+ in mantle kaersutite: implications for mantle metasomatism. Am Mineral 78:968-979 Edgar AD, Vukadinovic D (1992) Implications of experimental petrology to the evolution of ultrapotassic rocks. Lithos 28:205-220 Eggler DH (1978) The effect of CO2 upon partial melting of peridotite in the system Na2O-CaO-Al2O3-MgOSiO2-CO2 to 35 kb, with an analysis of melting in a peridotite-H2O-CO2 system. Am J Sci 278:305-343 Elkins-Tanton LT, Grove TL (2003) Evidence for deep melting of hydrous metasomatized mantle: Pliocene high-potassium magmas from the Sierra Nevadas. J Geophys Res 108:2350, doi:10.1029/2002JB002168

Stability of Hydrous Mantle Phases

267

Erlank AJ, Waters FG, Hawkesworth CJ, Haggerty SE, Allsopp HL, Rickard RS, Menzies M (1987) Evidence for mantle metasomatism in peridotite nodules from the Kimberley pipes, South Africa. In: Mantle Metasomatism. Menzies MA, Hawkesworth CJ (eds). Academic Press, p 313-361 Fabries J, Lorand J-P, Guiraud M (2001) Petrogenesis of the amphibole-rich veins from the Lherz orogenic lherzolite massif (Eastern Pyrenees, France): a case study for the origin of the orthopyroxene-bearing amphibole pyroxenites in the lithospheric mantle. Contrib Mineral Petrol 140:383-403 Field SW, Haggerty, SE, Erlank AJ (1989) Subcontinental metasomatism in the region of Jagersfontein, South Africa. In: Kimberlites and Related Rocks. Geol Soc Aust Spec Publ 14. Ross J (ed). Geological Society of Australia, p 771-784 Fleet ME (2003) Micas. The Geological Society Foley S (1991) High-pressure stability of the fluor- and hydroxyl-endmembers of pargasite and K-richterite. Geochim Cosmochim Acta 55:2689-2694 Foley S (1992) Vein-plus-wall-rock melting mechanisms in the lithosphere and the origin of potassic alkaline magmas. Lithos 28:435-453 Frey FA, Prinz M (1978) Ultramafic inclusions from San Carlos, Arizona: petrological and geochemical data bearing on their petrogenesis. Earth Planet Sci Lett 38:129-176 Frost DJ, Fei Y (1998). Stability of phase D at high pressure and high temperature. J Geophys Res 103B:74637474 Frost DJ (1999) The stability of dense hydrous magnesium silicates in Earth’s transition zone and lower mantle. In Mantle Petrology: Field Observations and High-Pressure Experimentation. Fei Y, Bertka CM, Mysen BO(ed) Geochemical Society, Special Publication No. 6, p 283-295 Frost DJ, Liebske C, Langenhorst F, McCammon CA, Tronnes RG, Rubie DC (2004) Experimental evidence for the existence of iron-rich metal in the Earth’s lower mantle. Nature 428:409-412 Gilbert MC, Briggs DF (1974) Comparison of the stabilities of OH- and F-potassic richterites- a preliminary report. Trans Am Geophys Union 55:480-481 Gilbert MC, Helz RT, Popp RK, Spear FS (1982) Experimental studies of amphibole stability. Rev Mineral 9B: 229-353 Green DH (1973) Experimental melting studies on a model upper mantle composition at high pressure under water-saturated and water-undersaturated conditions. Earth Planet Sci Lett 19:37-53 Gregoire M, Lorand JP, O’Reilly SY, Cottin JY (2000) Armalcolite-bearing, Ti-rich metasomatic assemblages in harzburgitic xenoliths from the Kerguelen Islands: Implications for the oceanic mantle budget of high-field strength elements. Geochim Cosmochim Acta 64:673-694 Gregoire M, Bell DR, Le Roex AP (2003) Garnet lherzolites from the Kaapvaal craton (South Africa): trace element evidence for a metasomatic history. J Petrol 44:629-657 Gudmundsson G, Wood BJ (1995) Experimental tests of garnet peridotite oxygen barometry. Contrib Mineral Petrol 119:56-67 Halliday AN, Lee D-C, Tommasini S, Davies GR, Paslick CR, Fitton JG, James DE (1995) Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic mantle. Earth Planet Sci Lett 133: 379-395 Harte B (1993) Mantle peridotites and processes- the kimberlite sample. In: Continental Basalts and Mantle Xenoliths. Hawkesworth CJ, Norry MJ (eds) Shiva, p 46-91 Hauri EH, Shimizu N, Dieu JJ, Hart SR (1993) Evidence for hotspot-related carbonatite metasomatism in the oceanic upper-mantle. Nature 365:221-227 Hawkesworth CJ, Erlank AJ, Marsh JS, Menzies MA, Van Calsteren P (1983) Evolution of the continental lithosphere: evidence from volcanics and xenoliths in southern Africa. In: Continental Basalts and Mantle Xenoliths. Hawkesworth MJ, Norry MJ (eds). Shiva, p 111-138 Hoal KEO, Hoal BG, Erlank AJ, Shimizu N (1994) Metasomatism of the mantle lithosphere recorded by rare earth elements in garnets. Earth Planet Sci Lett 126:303-313 Holloway JR (1973) The system pargasite H2O-CO2:a model for melting of a hydrous mineral with a mixedvolatile fluid. I. Experimental results to 8 kbar. Geochim Cosmochim Acta 37:651-666 Holloway JR (1987) Igneous fluids. Rev Mineral 17:211-233 Hübner SJ, Papike JJ (1970) Synthesis and crystal chemistry of sodium-potassium richterite (Na,K)NaCaMg5 Si8O22(OH,F)2): a model for amphiboles. Am Mineral 55:1973-1992 Huckenholz HG, Gilbert MC, Kunzmann T (1992) Stability and phase relations of calcic amphiboles crystallized from magnesio-hastingsite compositions in the 1 to 45 kbar pressure range. N Jahr Miner Abh 164:229-268 Inoue T, Irifune T, Yurimoto H, Miyagi I (1998) Decomposition of K-amphibole at high pressures and implications for subduction zone volcanism. Phys Earth Planet Int 107:221-231 Ionov DA, Hofmann AW (1995) Nb-Ta-rich mantle amphiboles and micas: implications for subduction-related metasomatic trace element fractionations. Earth Planet Sci Lett 131:341-356 Ionov D (1998) Trace element composition of mantle-derived carbonates and coexisting phases in peridotite xenoliths from alkaline basalts. J Petrol 39:1931-1941

268

Frost

Irifune T, Kubo T, Isshiki M, Yamasaki Y (1998) Phase transformations in serpentine and transportation of water into the lower mantle. Geophys Res Lett 22:117-120 Jenkins DM (1983) Stability and composition relations of calcic amphiboles in ultramafic rocks. Contrib Mineral Petrol 83:375-384 Jones AP, Smith JV, Dawson JB (1982) Mantle metasomatism in 14 veined peridotites from Bultfontein Mine, South Africa. J Geol 90:435-453 Jones AP (1989) Upper mantle enrichment by kimberlitic or carbonatitic magmatism. In: Carbonates: Genesis and Evolutions. Bell K (ed) Unwin Hyman, p 448-463 Kanzaki M (1991) Stability of hydrous magnesium silicates in the mantle transition zone. Phys Earth Planet Inter 66:307–312 Kawamoto T, Leinenweber K, Hervig RL, Holloway JR (1995) Stability of hydrous minerals in H2O-saturated KLB-1 peridotite up to 15 GPa. In: Volatiles in the Earth and Solar System. Farley KA (ed) American Institute of Physics, p 229-239 Kawamoto T (2004) Hydrous phase stability and partial melt chemistry in H2O-saturated KLB-1 peridotite up to the uppermost lower mantle. Phys Earth Planet Int 143-144:387-395 Kawamoto T (2006) Hydrous phases and water transport in the subducting slab. Rev Mineral Geochem 62: 273-289 King PL, Hervig RL, Holloway JR, Vennemann TW, Righter K (1999) Oxy-substitution and dehydrogenation in mantle-derived amphibole megacrysts. Geochim Cosmochim Acta 63:3635-3651 Kinny PD, Compston JW, Bristow JW, Williams IS (1989) Archaean mantle xenocrysts in a Permian kimberlite: two generations of kimberlitic zircon in Jwaneng DK2, southern Botswana. In: Kimberlites and related rocks. Geol Soc Aust Spec Publ 14. Ross J (ed). Geological Society of Australia, p 833-842 Kinny PD, Meyer HOA (1994) Zircon from the mantle: a new way to date old diamonds. J Geol 102:475-481 Komabayashi T, Omori S, Maruyama S (2004) Petrogenetic grid in the system MgO-SiO2-H2O up to 30 GPa, 1600 degrees C: Applications to hydrous peridotite subducting into the Earth’s deep interior. J Geophys Res 109:Art. No. B03206 Konzett J, Sweeney RJ, Thompson AB, Ulmer P (1997) Potassium amphibole stability in the upper mantle: an experimental study in a peralkaline KNCMASH system to 8.5 GPa. J Petrol 38:537-568 Konzett J, Ulmer P (1999) The stability of hydrous potassic phases in lherzolitic mantle- an experimental study to 9.5 GPa in simplified and natural bulk compositions. J Petrol 40:629-652 Konzett J, Fei Y (2000) Transport and storage of potassium in the Earth’s upper mantle and transition zone: an experimental study to 23 GPa in simplified and natural bulk compositions. J Petrol 41:583-603 Kushiro I (1970) Stability of amphibole and phlogopite in the upper mantle. Carnegie Inst Wash Year Book 68: 245-247 Kushiro I (1987) A petrological model of the mantle wedge and lower crust in the Japanese island arcs. In: Magmatic Processes: Physiochemical Principles. Geochem Soc Spec Publ 1. Mysen BO (ed) The Geochemical Society, p 165-181 Lambert IB, Wyllie PJ (1968) Stability of hornblende and a model for the low velocity zone. Nature 215:12401241 LaTourrette T, Hervig RL, Holloway JR (1995) Trace element partitioning between amphibole, phlogopite and basanite melt. Earth Planet Sci Lett 135:13-50 Leake BE, Woolley AR, Arps CES, Birch WD, Gilbert MC, Grice JD, Hawthorne FC, Kato A, Kisch HJ, Krivovichev VG, Linthout K, Laird J, Mandarino JA, Maresch WV, Nickel EH, Rock NMS, Schumacher JC, Smith DC, Stephenson NCN, Ungaretti L, Whittaker EJW, Guo YZ (1997) Nomenclature of amphiboles: Report of the subcommittee on amphiboles of the International Mineralogical Association, commission on new minerals and mineral names. Am Mineral 82:1019-1037 López Sánchez-Vizcaíno V, Trommsdorff V, Gómez-Pugnaire MT, Garrido CJ, Müntener O, Connolly JAD (2005) Petrology of titanian clinohumite and olivine at the high-pressure breakdown of antigorite serpentinite to chlorite harzburgite (Almirez Massif, S. Spain). Contrib Mineral Petrol 149:627-646 Luth RW (1995) Is phase A relevant to the Earth’s mantle? Geochim Cosmochim Acta 59:679-682 Luth RW (1997) Experimental study of the system phlogopite-diopside from 3.5 to 17 GPa. Am Mineral 82: 1198-1209 Lykins RW, Jenkins DM (1992) Experimental determination of pargasite stability relations in the presence of orthopyroxene. Contrib Mineral Petrol 112:405-413 McCammon CA, Griffin WL, Shee SR, O’Neill HSC (2001) Oxidation during metasomatism in ultramafic xenoliths from Wesselton kimberlite, South Africa. Implications for the survival of diamond. Contrib Mineral Petrol 141:287-296. McCammon C, Kopylova MG (2004) A redox profile of the Slave mantle and oxygen fugacity control in the cratonic mantle. Contrib Mineral Petrol 148:55-68 McInnes BIA, Gregoire M, Binns RA, Herzig PM, Hannington MD (2001) Hydrous metasomatism of oceanic sub-arc mantle, Lihir, Papua New Guinea: petrology and geochemistry of fluid-metasomatized mantle wedge xenoliths. Earth Planet Sci Lett 188:169-183

Stability of Hydrous Mantle Phases

269

McNeil AM, Edgar AD (1987) Sodium-rich metasomatism in the upper mantle: implications of experiments on the pyrolite-Na2O-rich fluid system at 950 °C, 20 kbar. Geochim Cosmochim Acta 51:2285-2294 Mancini F, Harlow GE, Cahill C (2002) The crystal structure and cation ordering of phase-X-(K1−x−n)2(Mg1−n [Al,Cr]n)2Si2O7H2x: A potential K- and H-bearing phase in the mantle. Am Mineral 87:302-306 Matson DW, Muenow DW, Garcia MO (1986) Volatile contents of phlogopite micas from South African kimberlite. Contrib Mineral Petrol 93:399-408 Mengel K, Green DH (1989) Stability of amphibole and phlogopite in metasomatized peridotite under watersaturated and water-undersaturated conditions. In: Kimberlites and related rocks. Geol Soc Aust Spec Publ 14. Ross J (ed) Geological Society of Australia, p 571-581 Menzies MA, Murthy VR (1980) Nd and Sr isotope geochemistry isotope geochemistry of hydrous mantle nodules and their host alkali basalts: implications for local heterogeneities in metasomatically veined mantle. Earth Planet Sci Lett 46:323-334 Menzies MA, Rogers N, Tindle A, Hawkesworth CJ (1987) Metasomatic enrichment processes in lithospheric peridotites, an effect of asthenospheric-lithospheric interaction. In: Mantle Metasomatism. Menzies MA, Hawkesworth CJ (eds) Academic Press, p 313-361 Merrill RB, Wyllie PJ (1975) Kaersutite and kaersutite eclogite from Kakanui, New Zealand- water-excess and water-deficient melting to 30 kilobars. Geol Soc Am Bull 86:555-570 Michael PJ (1988) The concentration, behavior and storage of H2O in the sub-oceanic upper mantle: implications for mantle metasomatism. Geochim Cosmochim Acta 52:555-566 Millhollen GL, Irving AJ, Wyllie PJ (1974) Melting interval of peridotite with 5.7 percent water to 30 kilobars. J Geol 82:575-587 Miyagi I, Matsubaya O, Nakashima S (1998) Change in D/H ratio, water content and color during dehydration of hornblende. Geochem J 32:33-48 Modreski PJ, Boettcher AL (1972) Stability of phlogopite + enstatite at high pressure: A model for micas in the interior of the Earth. Am J Sci 272:852-869 Murayama JK, Nakai S, Kato M, Kumazawa M. (1986) A dense polymorph of Ca3(PO4)2:a high pressure phase of apatite decomposition and its geochemical significance. Phys Earth Planet Int 44:293-303 Mysen BO, Boettcher AL (1975) Melting of a hydrous mantle: I Phase relations of natural peridotite at high pressures and temperatures with controlled activities of water, carbon dioxide, and hydrogen. J Petrol 16: 520-548 Mysen BO, Virgo D, Popp RK, Bertka CM (1998) The role of H2O in Martian magmatic systems. Am Mineral 83:942-946 Niida K, Green DH (1999) Stability and chemical composition of pargasitic amphibole in MORB pyrolite under upper mantle conditions. Contrib Mineral Petrol 135:18-40 Obata M (1980) The Ronda peridotite: garnet-, spinel- and plagioclase-lherzolite facies and the P-T trajectories of a high-temperature mantle intrusion. J Petrol 21:533-572 Obata M, Morten L (1987) Transformation of spinel lherzolite to garnet lherzolite in ultramafic lenses of the Austridic crystalline complex, northern Italy. J Petrol 28:599-623 O’Hara MJ, Mercy ELP (1963) Petrology and petrogenesis of some garnetiferous peridotites. Trans Royal Soc Ed 65:251-314 O’Neill HStC, Rubie DC, Canil, D, Geiger CA, Ross CR, Seifert F, Woodland AB (1993) Ferric iron in the upper mantle and in transition zone assemblages: Implications for relative oxygen fugacities in the mantle. Geophys Monograph 74 IUGG 14:73-89 O’Reilly SY, Griffin WL (2000) Apatite in the mantle: implications for metasomatic processes and high heat production in Phanerozoic mantle. Lithos 53:217-232 O’Reilly SY, Griffin WL (1988) Mantle metasomatism beneath western Victoria, Australia: I. Metasomatic processes in Cr-diopside lherzolites. Geochim Cosmochim Acta 52:433-447 Ohtani, E, Shibata T, Kubo T, Kato T (1995) Stability of hydrous phases in the transition zone and the upper most part of the lower mantle. Geophys Res Lett 22:2553-2556 Popp RK, Virgo D, Yoder HS, Hoering TC, Phillips MW (1995) An experimental study of phase equilibria and Fe oxy-component in kaersutitic amphibole: implications for the ƒH2 and aH2O in the upper mantle. Am Mineral 80:534-548 Righter K, Dyar MD, Delaney JS, Vennemann TW, Hervig RL, King PL (2002) Correlations of octahedral cations with OH−, O2−, Cl− and F− in biotite from volcanic rocks and xenoliths. Am Mineral 87:142-153 Robinson P, Spear FS, Schumacher JC (1982) Phase relations of metamorphic amphiboles: natural occurrence and theory. Rev Mineral 9B:1-227 Roden MF, Murthy R (1985) Mantle metasomatism. Ann Rev Earth Planet Sci 13:269-296 Ryabchikov ID, Boettcher AL (1980) Experimental evidence at high pressure for potassic metasomatism in the mantle of the Earth. Am Mineral 65:915-919 Sato K, Katsura T, Ito E (1997) Phase relations of natural phlogopite with and without enstatite up to 8 GPa: implication for mantle metasomatism. Earth Planet Sci Lett 146:511-526

270

Frost

Scambelluri M, Hermann J, Morten L, Rampone E (2006) Melt- versus fluid-induced metasomatism in spinel to garnet wedge peridotites (Ulten Zone, Eastern Italian Alps): clues from trace element and Li abundances. Contrib Mineral Petrol 151:372-394 Schmidt MW, Poli S (1998) Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation. Earth Planet Sci Lett 163:361-379 Schmidt MW, Vielzeuf D, Auzanneau E (2004) Melting and dissolution of subducting crust at high pressures: the key role of white mica. Earth Planet Sci Lett 228:65-84 Schneider ME, Eggler DH (1986) Fluids in equilibrium with peridotite minerals: implications for mantle metasomatism. Geochim Cosmochim Acta 50:711-724 Schrauder M, Navon O (1994) Hydrous and carbonatitic mantle fluids in fibrous diamonds from Jwaneng Botswana. Geochim Cosmochim Acta 58:761-771 Seyler M, Mattson PH (1989) Petrology and thermal evolution of the Tinaquillo peridotite (Venezuela). Journal Geophys Res 94:7629-7660 Shaw CSJ, Eyzaguirre J (2000) Origin of megacrysts in the mafic alkaline lavas of the West Eifel volcanic field, Germany. Lithos 50:75-95 Shen AH, Keppler H (1997) Direct observation of complete miscibility in the albite-H2O system. Nature 385: 710-712 Shieh SR, Mao HK, Hemley RJ, Ming LC (1998) Decomposition of phase D in the lower mantle and the fate of dense hydrous silicates in subducting slabs. Earth Planet Sci Lett 159:13-23 Shimizu N (1975) Rare earth elements in garnets and clinopyroxenes from garnet lherzolite nodules in kimberlites. Earth Planet Sci Lett 25:26-32 Sweeney RJ, Thompson AB, Ulmer P (1993) Phase relations of a natural MARID composition and implications for MARID genesis, lithospheric melting and mantle metasomatism. Contrib Mineral Petrol 115:225-241 Sudo A, Tatsumi Y (1990) Phlogopite and K-amphibole in the upper mantle: implication for magma genesis in subduction zones. Geophys Res Lett 17:29-32 Tatsumi Y, Hamilton DL, Nesbitt RW (1986) Chemical characteristics of fluid phase released from a subducted lithosphere and origin of arc magmas: evidence from high pressure experiments and natural rocks. J Volcanol Geotherm Res 29:293-310 Taylor WR, Green DH (1988) Measurement of reduced peridotite-C-O-H solidus and implications for redox melting of the mantle. Nature 332:349-352 Thibault Y, Edgar AD, Lloyd FE (1992) Experimental investigation of melts from a carbonated phlogopite lherzolite: implications for metasomatism in the continental lithospheric mantle. Am Mineral 77:784-794 Tiepolo M, Vannucci R, Oberti R, Foley S, Bottazzi P, Zanetti A (2000) Nb and Ta incorportation and fractionation in titanian pargasite and kaersutite: crystal-chemical constraints and implications for natural systems. Earth Planet Sci Lett 176:185-201 Thompson AB (1992) Water in the Earth’s upper mantle. Nature 358:295-302 Thompson JB, Laird J, Thompson AB (1981) Reactions in amphibolite, greenschist and Blueschist. J Petrol 23: 1-27 Trønnes RG, Edgar AD, Arima M (1985) A high pressure-high temperature study of TiO2 solubility in Mg-rich phlogopite: implications to phlogopite chemistry. Geochim Cosmochim Acta 49:2323-2329 Trønnes RG (2002) Stability range and decomposition of potassic richterite and phlogopite end members at 5-15 GPa. Mineral Petrol 74:129-148 Ulmer P, Trommsdorff V (1999) Phase relations of hydrous mantle subducted to 300 km. In Mantle Petrology: Field Observations and High-pressure Experimentation. Fei Y, Bertka CM, Mysen BO (ed) Geochemical Society, Special Publication No. 6, p 259-281 Ulmer P (2001) Partial melting in the mantle wedge - the role of H2O in the genesis of mantle-derived ‘arcrelated’ magmas. Phys Earth Planet Int 127:215-232 van Achterbergh E, Griffin WL, Stiefenhofer J (2001) Metasomatism in mantle xenoliths from the Letlhakane kimberlites: estimation of element fluxes. Contrib Mineral Petrol 141:397-414 Wagner C, Deloule E, Mokhtari A (1996) Richterite-bearing peridotites and MARID-type inclusions in lavas from North Eastern Morocco: mineralogy and D/H isotopic studies. Contrib Mineral Petrol 124:406-421 Wallace ME, Green DH (1988) An experimental determination of primary carbonatite magma composition. Nature 335:343-346 Wallace ME, Green DH (1991) The effect of bulk rock composition on the stability of amphibole in the upper mantle: implications for solidus positions and mantle metasomatism. Mineral Petrol 44:1-19 Wallace PJ (1998) Water and partial melting in mantle plumes: Inferences from the dissolved H2O concentrations of Hawaiian basaltic magmas. Geophys Res Lett 25:3639-3642 Waters FG (1987) A suggested origin of MARID xenoliths in kimberlites by high pressure crystallization of an ultrapotassic rock such as lamproite. Contrib Mineral Petrol 95:523-533 Watson BE (1980) Apatite and phosphorus in mantle source regions: an experimental study of apatite/melt equilibria at pressures to 25 kbar. Earth Planet Sci Lett 51:322-335

Stability of Hydrous Mantle Phases

271

Weiss M (1997) Clinohumites: a field and experimental study. ETH Dissertation No. 12202, 168 pp. Wilkinson JFG, Le Maitre RW (1987) Upper mantle amphiboles and micas and TiO2, K2O and P2O5 abundances and 100Mg/(Mg+Fe2+) ratios of common basalts and andesites: implications for modal mantle metasomatism and undepleted mantle compositions. J Petrol 28:37-73 Wilshire HG, Shervais JW (1975) Al-augite and Cr-diopside ultramafic xenoliths in basaltic rocks from Western United States. Phys Chem Earth 9:257-276 Witt-Eickschen G, Seck HA, Reys CH (1993) Multiple enrichment processes and their relationships in the subcrustal lithosphere beneath the Eifel (Germany). J Petrol 34:1-22 Witt-Eickschen G, Kramm U (1998) Evidence for the multiple stage evolution of the subcontinental lithospheric mantle beneath the Eifel (Germany) from pyroxenite and composite pyroxenite/peridotite xenoliths. Contrib Mineral Petrol 131:258-272 Woermann E, Rosenhauer M (1985) Fluid phases and the redox state of the earths mantle. Fortschr Mineral 63: 263-349 Wood BJ, Bryndzia LT, Johnson KE (1990) Mantle oxidation state and its relationship to tectonic environment and fluid speciation. Science 248:337-345 Woodland AB, Kornprobst J, McPherson E, Bodinier J-L, Menzies MA (1996) Metasomatic interactions in the lithospheric mantle: petrologic evidence from the Lherz massif, French Pyrenees. Chem Geol 134:83-112 Woodland AB, Koch M (2003) Variation in oxygen fugacity with depth in the upper mantle beneath the Kaapvaal craton, South Africa. Earth Planet Sci Lett 214:295-310 Wyllie PJ (1978) Mantle fluid compositions buffered in peridotite-CO2-H2O by carbonates, amphibole, and phlogopite. J Geol 86:687-713 Yang HX, Konzett J, Prewitt CT (2001) Crystal structure of Phase X, a high pressure alkali-rich hydrous silicate and its anhydrous equivalent. Am Mineral 86:1483-1488 Yang H-J, Frey FA, Clague DA (2003) Constraints on the source components of lavas forming the Hawaiian North Arch and the Honolulu volcanics. J Petrol 44:603-627 Yaxley GM, Green DH, Kamenetsky V (1998) Carbonatite metasomatism in the southeastern Australian lithosphere. J Petrol 39:1917-1930 Yoder HS, Kushiro I (1969) Melting of a hydrous phase: phlogopite. Am J Sci 267A: 558-582 Young ED, Virgo D, Popp RK (1997) Eliminating closure in mineral formulae with specific application to amphiboles. Am Mineral 82:790-806 Zanetti A, Mazzucchelli M, Rivalenti G, Vannucci R (1999) The Finero phlogopite-peridotite massif: an example of subduction-related metasomatism. Contrib Mineral Petrol 134:107-122

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Reviews in Mineralogy & Geochemistry Vol. 62, pp. 273-289, 2006 Copyright © Mineralogical Society of America

Hydrous Phases and Water Transport in the Subducting Slab Tatsuhiko Kawamoto Institute for Geothermal Sciences Graduate School of Science Kyoto University Beppu 874-0903, Japan e-mail: [email protected]

Introduction Arc volcanoes are typically located 90-180 km above the surface of downgoing slabs, as shown by Wadati-Benioff deep seismic foci (Gill 1981; Tatsumi 1989). The intimate relationship between the dip angles of the subducting slab and the locations of volcanic arcs indicates that subduction zone magmatism is triggered by material input from the subducting slab (Tatsumi and Eggins 1995). The slab-derived components are thought to be aqueous fluids or H2O-rich partial melts of subducted oceanic crust. Therefore, knowledge of the stability of hydrous phases and the chemical and physical properties of aqueous fluids in downgoing slabs is essential to understand the material transport in subduction zones. In this section, I will review the stability of hydrous phases in downgoing peridotite, basalt and sediment systems, and the chemical and the wetting properties of aqueous fluids. Recent experimental studies indicate that 3-4 GPa, equivalent to 90-120 km depth, is a key pressure, where (1) the chemical compositions of silicate components dissolved in aqueous fluids equilibrated with mantle minerals approach the composition of mantle peridotite itself (Stalder et al. 2001; Mibe et al. 2002; Kawamoto et al. 2004), (2) the dihedral angle between olivine and aqueous fluids starts becoming smaller than 60° (Watson et al. 1990; Mibe et al. 1998, 1999), and (3) the immisciblity gap between peridotitic melts and aqueous fluids disappears and consequently hydrous minerals liberate supercritical aqueous fluids (Mibe et al. 2004a, 2006). The similarity between these pressures and the depths of downgoing slab underneath volcanic fronts, where the maximum numbers of volcanoes are formed, 124 ± 38 km (Gill 1981) or 112 ± 19 km (Tatsumi 1986), suggests that subduction zone magmatism can be triggered by the input of supercritical fluids from the downgoing peridotite and basalt.

LOW-PRESSURE HYDROUS MINERALS AND HIGH-PRESSURE HYDROUS PHASES Many hydrous crystalline phases are stable in peridotite, basalt and sediment systems over a wide range of pressure. Their chemical formulae and H2O contents are summarized together with those of nominally anhydrous minerals in Table 1. Some hydrous phases have been found only in high-pressure and high-temperature experimental products and have not yet been found in nature: dense hydrous magnesium silicates (DHMS) or alphabet phases (Ringwood and Major 1967), phase Egg (Eggleton et al. 1978), phase Pi (Wunder et al. 1993a), topaz-OH (Wunder et al. 1993b), and δ-AlOOH (Suzuki et al. 2000). Although phase D, F, and G were originally suggested as different phases, these phases seem to be identical (Frost 1999; Ohtani et al. 2001). The chemical compositions of DHMS are plotted in Figure 1 with the estimated 1529-6466/06/0062-0012$05.00

DOI: 10.2138/rmg.2006.62.12

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Kawamoto Figure 1

H 2O

1 GPa

3 GPa 5 GPa Br

A Nor sB B

MgO

10 GPa

8 GPa 10 Å

E

AhyB Ol

Atg Chn Hywd Chm En

D Talc Ant

SiO2

Figure 1. Compositions of hydrous minerals and dense magnesium hydrous silicates stable in peridotite system plotted with compositions of silicates dissolved into aqueous fluids coexisting with forsterite and enstatite at 1100 °C at 1-10 GPa estimated by Zhang and Frantz (2000) and Mibe et al. (2002) in the MgOSiO2-H2O system. Phase D, E, antigorite, and 10 Å phase are non-stoichiometric phases. Humite is located between chondrodite and clinohumite. Abbreviations are in Table 1.

chemistry of aqueous fluids equilibrated with forsterite + enstatite in the MgO-SiO2-H2O system (Mibe et al. 2002). The hydrous crystalline phases can be divided into three major groups with respect to their stability range (Fig. 2): (1) low-pressure hydrous minerals such as chlorite (clinochlore), talc, and amphibole (the relevant end members are listed in Table 1), which are commonly observed in metamorphic rocks, (2) high-pressure hydrous phases such as DHMS (Fig. 1), K-richterite, topaz-OH, and phase Egg, and (3) middle-pressure hydrous minerals such as phlogopite, antigorite, Mg-sursassite and 10 Å phase in peridotite, lawsonite in basalt, and phengite in sediment. The last group is stable between 5 and 7 GPa, and may be important for delivering H2O from low-pressure hydrous minerals to high-pressure hydrous phases (Fig. 2). Liu (1987) recognized that phase A, a DHMS, can accommodate much more water than amphibole or phlogopite. Therefore he emphasized the important reaction forsterite + H2O = phase A + enstatite, and he described this reaction boundary as a “water-line,” implying that a region deeper than the water-line can be a H2O reservoir in the mantle. In Figure 2, the water-line is shown by the low-pressure stability of DHMS. Kawamoto et al. (1996) identified the presence of a “choke point” in a down going slab. A choke point represents a pressure and temperature condition along a PT path where low-pressure and middle-pressure hydrous minerals get dehydrated at certain pressure conditions and cannot deliver H2O to high-pressure hydrous phases (Fig. 2). The choke point curve, the curve connecting the array of choke points, represents the high-pressure and high-temperature stability limit of the lowpressure and middle-pressure hydrous minerals. In the MgO-SiO2-H2O system, the invariant point composed of antigorite, phase A, enstatite, forsterite and H2O represents the lowest temperature and highest pressure of the choke point. In recent literature, this point is at around 6.2 GPa and 620 °C (Iwamori 2004), and at 5.1 GPa and 550 °C (Komabayashi et al. 2005). In Figure 2, based on the KLB-1 peridotite data, TiO2 stabilizes chondrodite and clinohumite. Therefore, in the peridotite systems, the PT conditions where antigorite meets chondrodite and clinohumite represent the lowest temperature and highest pressure choke point. In the MgO-

Hydrous Phases & Water Transport in the Subducting Slab

Kawamoto 400 Figure 2

600

HT

+ Ol 4 Atg 6

LT

10 12 14

1000 1200 1400 1600

Chl Par +Opx

Lws

ol-fl Present Mantle Adiabat

Phl + Cpx

Phe

K-ric

100

200

gt-fl

E + Chm E + Chm Top +A Eg E

300

Ol Ol + Wd

400

Depth (km)

Pressure (GPa)

8 Chm + Chn

Temperature (°C)

800

Wet solidus

2 Talc

275

Wd

16 18 20 22 24 26

Wd + Rg 500

D + sB

Rg Hy-wd D + sB

Hy-rg Rg Pv + Mw sB

600

700

Figure 2. Pressure and temperature diagram showing stability of hydrous minerals/phases in peridotite (Kawamoto 2004a) with some hydrous phases in basalt/sediment systems. The wet solidus is from Kawamoto and Holloway (1997). Since a second critical endpoint between peridotite melt and aqueous fluids is located at around 3.8 GPa (Mibe et al. 2004a; 2006), the wet solidus is drawn by dashed line at pressures higher than 4 GPa. Stability of lawsonite in basalt is indicated by solid dots and stability boundaries among phengite, topaz-OH, and phase Egg is drawn by open dots, respectively. The stabilities of Par, Chl, Talc, Atg, Phl, K-rich, Lws, Top, Eg are after Schmidt and Poli (1998), Pawley (2003), Ulmer and Trommsdorff (1995), Sudo and Tatsumi (1990), and Ono (1998); phase boundaries among Ol, Ol + Wd, Wd, and Wd + Rg - (Mg0.9Fe0.1)2SiO4 and Rg - (Mg0.9Fe0.1)2SiO4 and Mg-perovskite (Mg-Pv) + magnesium wüstite (Mw) in dry conditions are after Katsura and Ito (1989), and Ito and Takahashi (1989), respectively. The phase boundary of Hy- wd and Hy-rg (dashed line) is at higher pressure than under dry conditions. The 60° isopleths of the dihedral angle in garnet-garnet-fluid (gt-fl) and olivine-olivine-fluid (ol-fl) are also shown (thick gray line). The data of the dihedral angle are compiled in Figure 5. HT and LT represent PT paths of high-temperature and low-temperature subducting slab surface, respectively (Peacock and Wang 1999). Abbreviations are in Table 1.

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Table 1. Formula of hydrous minerals/phases and nominally anhydrous minerals in metamorphic basalts, metamorphic sediments, and peridotite (after Wunder and Schreyer 1992, Pawley and Wood 1995, Mysen et al. 1998, Ono 1999, Forneris and Holloway 2003). Name

Symbols

Formula

(Amphibole groups) Tremolite Pargasite Barroisite Glaucophane K-richterite

Trm Par Bar Gln K-ric

Ca2Mg5Si8O22(OH)2 Na2Ca3Mg8FeAl3Si13O44(OH)4 NaCaMg3Al2Si7AlO22(OH)2 Na2Mg3Al2Si8O22(OH)2 K1.9Ca1.1Mg5Si7.9Al0.1O22(OH)2

(Peridotite system) Chlorite Talc Serpentine Antigorite Clinohumite Humite Chondrodite Norbergite Phase A Brucite Phase B Superhydrous B Anhydrous B Phase E Phase D/F/G Anthophyllite Talc 10 Å phase Mg-sursassite Hydrous wadsleyite Hydrous ringwoodite

Chl Tlc Serp Atg Chm Hm Chn Nor A Br B sB AhyB E D Ant Talc 10 Å MgS Hy-wd Hy-rg

(Mg5Al)(AlSi3)O10(OH)8 Mg6Si8O20(OH)4 Mg3Si2O5(OH)4 Mg48Si34O85(OH)62 Mg9Si4O16(OH)2, Ti0.5Mg8.5Si4O17(OH) Mg7Si3O12(OH)2 Mg5Si2O8(OH)2 , Ti0.5Mg4.5Si2O9(OH) Mg3SiO4(OH)2 Mg7Si2O8(OH)6 Mg(OH)2 Mg24Si8O38(OH)4 Mg10Si3O14(OH)4 Mg14Si5O24 Mg2.27Si1.26H2.4O6 MgSi2H2O6 Mg7Si8O22(OH)2 Mg3Si4O10(OH)2 Mg3Si4O10(OH)2 xH2O Mg5Al5Si6O21(OH)7 Mg1.75SiO4(OH)0.5 Mg1.75SiO4(OH)0.5

(Basalt and sediment systems) Zoisite/clinozoisite Staurolite Apatite Sphene Phlogopite Phase Egg Topaz-OH Phase Pi Lawsonite Chloritoid Phengite δ-AlOOH

Zo / Czo Sta Ap Spn Phl Eg Top Pi Lws Cld Phe δ-Al

Ca2Al3Si3O12(OH) (Mg,Fe)2(Al,Fe)9Si4O22(O,OH)2 Ca5(PO4)3(OH,F,Cl) CaTiSiO4(O,OH,F) KMg2Si3AlO10(OH)2 AlSiO3(OH) Al2SiO4(OH)2 Al3Si2O7(OH)3 CaAl2Si2O7(OH)2 H2O (Mg, Fe)2(Al,Fe)4Si2O10(OH)4 K(Al2-xMgx)(Si3+xAl1-x)O10(OH,F)2 AlOOH

(Nominally anhydrous minerals) Olivine/Wadsleyite/Ringwoodite Clinopyroxene Ca-perovskite Orthopyroxene/ Majorite/ Akimotoite/ Perovskite Quartz/ Coesite/ Stishovite Spinel Garnet

Ol / Wd / Rg Cpx Ca-pv Opx/ Mj / Ak / Pv Qz / Coe / St Sp Gt

Mg2SiO4 (Na,Ca)(Mg,Al)Si2O6 CaSiO3 MgSiO3 SiO2 MgAl2O4 (Fe,Mg,Ca)3Al2Si3O12

wt% H2O 2.2 2.2 2.3 2.3 2.1 13 4.8 13 12.3 2.9 - 1.4 3.75 5.3 - 2.6 9.0 11.8 30.9 2.4 1.6 11.4 10.1 2.3 4.75 7.6 - 13 7.2 3.3 3.3 2 2 1.8 1.5 4.8 7.5 10.0 9.0 11.5 8 4.6 15

Hydrous Phases & Water Transport in the Subducting Slab

277

Al2O3-SiO2-H2O system, Mg-sursassite (Gottschalk et al. 2000), which was previously called MgMgAl-pumpellyite (for example, Domanik and Holloway 1996), is stabilized at higher temperature than this invariant point (Fig. 3; Bromiley and Pawley 2003), and its presence therefore increases the temperature of the choke point. The transition zone (410-660 km depth) is also characterized by the high H2O storage capacity of hydrous wadsleyite and hydrous ringwoodite (Fig. 3; Smyth 1987; Inoue et al. 1995; Kawamoto et al. 1996; Kohlstedt et al. 1996; Kudoh et al. 1996; Smyth and Kawamoto 1997; Smyth et al. 1997; Demouchy et al. 2005). Therefore, the transition zone could play a significant role as a large H2O -reservoir formed by crystallization of hydrous wadsleyite and ringwoodite from a hydrous magma ocean. Kawamoto and Holloway (1997) measured the partition coefficient of H2O between hydrous wadsleyite/ringwoodite and hydrous partial melts of peridotite, and suggested the possible existence of a hydrous transition zone in the early history of the Earth. Upwelling from such a hydrous reservoir could generate partial melting at 410 km and produce komatiitic magmas. Through partial melting of a hydrous transition zone, in this hypothesis, the transition zone has been getting drier during the geological time, because the choke point prevents H2O from subducting into the transition zone. Therefore the present transition zone has much less ability to produce komatiite magmas. This hypothesis thus explains why komatiites were produced mainly in the Archean period.

STABILITY OF HYDROUS PHASES IN DOWNGOING PERIDOTITE There are two potentially-hydrated peridotite layers in subduction zones. One is the harzburgite/lherzolite of the subducting lithospheric mantle, which is overlain by oceanic basaltic crust and sediments. The other is downdragged mantle at the base of the mantle wedge. To what extent the peridotite layers are hydrated remains uncertain. Along transform faults, serpentine minerals (antigorite, lizardite, chrysotile) can be formed by seawater alteration. However, the rest of the subducting lithospheric mantle may not be hydrated. The downdragged mantle peridotite at the base of the mantle wedge should be hydrated by aqueous fluids liberated by dehydration reactions of hydrous minerals in downgoing sediment and basalt layers. Nicholls and Ringwood (1973) suggested that subducting basalt will be almost dry beneath the fore-arc region. Sakuyama and Nesbitt (1986), therefore, suggested that downdragged peridotite in the mantle wedge will be hydrated through H2O released by dehydration of the hydrous minerals in the basaltic layer and may carry H2O beneath the volcanic arc. Iwamori (2004) compiled the stability of hydrous phases in the MgO-SiO2-H2O, the MgO-Al2O3-SiO2-H2O, and KLB-1 peridotite systems, and presented the distribution of maximum H2O contents bound in mantle peridotite (Fig. 3). Komabayashi et al. (2004) also presented a similar stability diagram of hydrous phases based on Schreinemakers’ net analysis. They noticed two main differences of hydrous phase stability between the peridotite system and simple systems: (1) the addition of Al2O3 expands the stability field of phase E to the lower pressures and (2) the addition of TiO2 enhances the stability field of clinohumite and chondrodite (Fig. 2). The addition of fluorine is also found to expand the stability of clinohumite into a lower pressure range (Stalder and Ulmer 2001). According to Fumagalli et al. (2001), the10 Å phase (Table 1) is reported to be stable in the peridotite system at 5.2 GPa and 680 °C. Fumagalli and Poli (2005) found that the 10 Å phase has high Al2O3 contents (about 10 wt%) and suggested that this phase is a mixed layer of chlorite and pure 10 Å phase formed in the MgO-SiO2-H2O system. The stability field of this Al-rich 10 Å phase is close to the stability of Mg-sursassite (Bromiley and Pawley 2002). These phases cover some regions of the choke point (Fig. 3), though the H2O content contributed by Mg-sursassite and Al-rich 10 Å phase to peridotite is limited to 0.7 (Iwamori 2004; Fig. 3) and 1 wt% (Fumagalli and Poli 2005), respectively.

Figure 3. Phase diagram showing maximum H2O contents bound in hydrous minerals/phases in the peridotite system. The phase assemblages of fields numbered are shown on the right hand side. Abbreviations are in Table 1. This gray figure was made by the courtesy of Hikaru Iwamori. The original diagram is in full color with better resolution (Iwamori 2004).

278 Kawamoto

Hydrous Phases & Water Transport in the Subducting Slab

279

Amphibole was historically thought to be the most important phase to deliver H2O beneath the volcanic arc (fields 2, 3, 4, 6, and 7 in Fig. 3; Tatsumi 1986; Schmidt and Poli 1998; Niida and Green 1999). According to the compilation by Schmidt and Poli (1998), pargasite can be stable up to between 2.2 GPa and 3.0 GPa depending on the bulk rock chemistry of the system. Although Schmidt and Poli (1998) adopted the lowest pressure (2.2 GPa) for pargasite in harzburgite, it is important to realize that pargasite can be stable up to 3 GPa in more enriched peridotite such as enriched pyrolite (Niida and Green 1999). The 3.5 GPa for the high-pressure stability limit of pargasite adopted by Tatsumi (1986) seems overestimated as Schmidt and Poli (1998) suggested. In Figure 2, 2.8 GPa was adopted as a pressure limit for pargasite according to the recent experimental study by Fumagalli and Poli (2005). The stability of antigorite (line between fields 5 and 8 in Fig. 3) also depends on bulk composition and the effect of Al was evaluated by Bromiley and Pawley (2003). The stability of antigorite in Figure 2 is drawn with the data reported by Ulmer and Trommsdorff (1995).

Stability of hydrous phases in downgoing basalt and sediment There are many hydrous minerals observed in metamorphic basalt and sediments. Several experimental studies have explored their high PT stabilities. Concerning the stability of amphibole in the basalt system (Fig. 4), there is a discrepancy between Schmidt and Poli (1998) and Forneris and Holloway (2003). According to Schmidt and Poli (1998), in subducting basalt, amphibole and zoisite dehydrate first, then along a colder path, zoisite and chloritoid dehydrate, and finally lawsonite with or without chloritoid can retain H2O to Kawamoto the deep mantle (Fig. 4B). Along a warmer path, instead of lawsonite, zoisite becomes the Figure only hydrous4phase to possess H2O after amphibole dehydration, and then zoisite dehydrates liberating H2O (Fig. 4B). In contrast, according to Forneris and Holloway (2003), amphibole and zoisite at higher temperatures and amphibole with lawsonite at lower temperatures are stable up to 2.5-3.2 GPa (Fig. 4A). Then amphibole and zoisite dehydrate and lawsonite

Temperature (ºC)

500 550 600 2.0

650 700

Amp + Zo Zo

Pressure (GPa)

2.5 3.0

Lws + Amp

Zo

3.5

4.0

Temperature (ºC)

750 800 500 550 600 650 700 750 800

Lws

Amp

Cld + Zo

HT

Amp + Zo Zo

Lws + Cld

Cpx + Grt + Fluid

Lws

HT Cpx + Grt + Fluid

4.5 5.0

A, Forneris & Holloway (2003)

B,

Schmidt & Poli (1998)

Figure 4. Phase diagrams showing stability of hydrous minerals in the MORB system. (A) Forneris and Holloway (2003), Amp represents barroisite at high temperatures and glaucophane at low temperatures, (B) Schmidt and Poli (1998). This figure is after Forneris and Holloway (2003). According to Forneris and Holloway (2003), the chloritoid in B is likely to be formed by metastable crystallization, see text. HT represent a PT path of high-temperature subducting slab surface, and a PT path of low-temperature one (LT in Figure 2) is outside of this PT diagram (Peacock and Wang 1999).

280

Kawamoto

becomes the only hydrous phase (Fig. 4A). Forneris and Holloway (2003) suggested that a possible explanation for the discrepancy was the crystallization of metastable chloritoid (Fig. 4B) during short experimental run durations in the former experiments, perhaps due to the chemical difference between their starting materials: bulk compositions studied by Forneris and Holloway (2003) contained more MgO and Al2O3 than the starting materials of Schmidt and Poli (1998). The appearance of metastable chloritoid depresses the stability of amphibole. According to Schmidt and Poli (1998) and Forneris and Holloway (2003), basalt can possess 0.5-0.8 and 0.3 wt% H2O at 650 °C and 3 GPa. Lawsonite is the most important hydrous phase in subducting basalt because it is stable at relatively high temperature (Pawley and Holloway 1993; Pawley 1994; Poli and Schmidt 1995; Schmidt and Poli 1998; Ono 1998; Forneris and Holloway 2003; Schmidt et al. 2004). In particular, its stability covers the choke points in the stability of hydrous minerals of the peridotite system from 3 to 9 GPa (Fig. 2), and therefore lawsonite could re-hydrate the downdragged peridotite layer under those pressures when it dehydrates. At a temperature region higher than the lawsonite stability field, Schmidt and Poli (1998) observed phengite in basalt. The modal proportion of phengite is, however, limited in basalt because MORB has a low concentration of K and also if K is available in the system, K is partitioned preferentially into fluid. In the system relevant to sediments, Domanik and Holloway (1996) and Ono (1998) reported the stability of phengite, Mg-sursassite, topaz-OH, and phase Egg. These hydrous phases are characterized by their higher temperature stability than hydrous phases in peridotite and basalt systems as seen in Figure 2. Phengite has dehydration conditions similar to that of phlogopite (Fig. 2). The reaction boundary between topaz-OH and phase Egg is identical to the olivine - wadsleyite boundary (Fig. 2). Ono (1998) demonstrated that subducting sediment can bring 2 wt% H2O in phengite to 7 GPa, 0.7 wt% H2O in topaz-OH to 9 GPa, and 0.4 wt% H2O in phase Egg up to 15 GPa, and that subducting basalt can bring about 1 wt% H2O in lawsonite to 6 GPa and 800 °C. Phase Egg could be stable at least up to the transition zone, while lawsonite could dehydrate at around 10 GPa. This means that phase Egg could be formed in the sediment layer by H2O coming from dehydration of lawsonite in the basaltic layer. In addition to these phases, the δ-AlOOH phase, a high-pressure polymorph of diaspore, was proposed to be an important H2O host in sediment or basalt systems (Suzuki et al. 2000). However, it is still uncertain whether this phase is stable in sedimentary or basaltic systems (Litasov and Ohtani 2005).

PRESSURE - TEMPERATURE CONDITIONS AND DEHYDRATION REACTIONS IN THE SUBDUCTING SLAB Obviously the PT conditions of the downgoing slab are critical to determine the dehydration processes of hydrous phases in the slab. Furukawa (1993), Peacock (1993) and Peacock and Wang (1999) suggested several PT paths for subducting slabs (Fig. 2). These calculations have large uncertainties of 100-200 °C in the temperature at 90 km (3 GPa), because steep temperature gradients exist near to slab surfaces. Iwamori (2004) suggested that a kinematic critical parameter comprising the product of subduction angle, potential temperature, slab velocity and slab age, must be exceeded for PT paths to pass below the choke point at 6.2 GPa and 620 °C. When the downgoing hydrous peridotite follows relatively warm PT paths, antigorite breaks down, followed by talc and chlorite (HT path in Fig. 2). Beyond the chlorite-out reaction the subducting peridotite will be almost free of H2O bound in crystals except for a small amount in phlogopite at around 2.5 – 6.5 GPa. This means that when downgoing hydrous peridotite goes on paths like this, the hydrous minerals should encounter a “choke point” at 2.5 GPa (Fig. 2). If there is enough K2O to stabilize phlogopite in the mantle, the downgoing hydrous peridotite will carry a small amount of H into the deeper mantle. At 6.5-11 GPa, the

Hydrous Phases & Water Transport in the Subducting Slab

281

phlogopite breaks down into K-richterite, which has an equal H/K atomic ratio and is stable at least up to 13 GPa (Sudo and Tatsumi 1990) and then dehydrates again into another hydrous phase containing lower H/K (Trønnes et al. 1988; Inoue et al. 1998). The crystal structure of this phase remains to be investigated. When the downgoing hydrous peridotite follows relatively cold PT paths, antigorite breaks down at 6 GPa (LT path in Fig. 2) and beyond which small amounts of H2O may retain in phlogopite. Lawsonite in subducting basaltic crust contains ~11 wt% H2O and is stable beyond the choke point (Figs. 2, 4). Therefore, since the lower pressure stability of DHMS overlaps with the high-pressure stability of lawsonite, DHMS such as chondrodite, clinohumite, phase A and phase E in the downdragged base of the mantle wedge could absorb H2O from decomposing lawsonite in the basaltic layer and become H2O carriers in a cold subduction zone (LT path in Fig. 2) to the deeper mantle beyond the choke point. Phengite, topaz-OH and phase Egg in downgoing sediment could also pass H2O into DHMS because of their high-temperature stability (Fig. 2).

COMPOSITION AND DIHEDRAL ANGLES OF AQUEOUS FLUIDS IN MANTLE PERIDOTITE Since the pioneering work by Nakamura and Kushiro (1974), the chemical compositions of silicates dissolved in aqueous fluids have been assumed to be characterized by an SiO2rich component at relatively shallow depths corresponding to pressures between 1 and 3 GPa (Ryabchikov et al. 1982; Zhang and Frantz 2000). In contrast, recent experimental data above 3 GPa suggest that aqueous fluids coexisting with enstatite (MgSiO3) and forsterite (Mg2SiO4) exhibit higher Mg/Si ratios as the pressure increases from 3 GPa up to 10 GPa (Fig. 1; Stalder et al. 2001; Mibe et al. 2002). When the dihedral angles between crystals and fluids are smaller than 60°, permeable flow is allowed even if the porosity is small. The dihedral angles at triple junctions between forsterite crystals and aqueous fluid change from >60° to 90060 °1000

4 800

(c)

Mibe et al Ono et al

gt - fl 6

8

10

12

Pressure 900 1000 (GPa) 1100 1200 Temperature (°C)

3

4

5

Pressure (GPa)

6

(b)

55

700

800

50

Gt-fluid

45 1000

900

40 2

4

(c)

Mibe et al Ono et al

6

8

10

Pressure (GPa)

12

14

> 60 °

olin- fl Figure 5. (A, B) Dihedral angle (A) olivineolivine-fluid (Watson et al. 1990; Mibe et al. 1998, 4 1999) and (B) garnet-garnet-fluid (Ono et al. 2002; Mibe et al. 2003) versus pressure. < 60 ° Schematic contours at constant temperatures 6 (numerals) are drawn. The contours at 700 and 800 °C of gt-fluid > 60 ° are assumed to be parallel to the contour at 900 °C. gt (C)- The fl 60° isopleth of 8 the dihedral angle in olivine-olivine-fluid (olfl) and garnet-garnet-fluid (gt-fl). The plotted data in C are from A and B. 2

ol - fl

2

1000

60

0

> 60 °

1

65

35

6 14

700

6 45

0

Mibe et al Ono et al1000

1200

2 6

0

70 Ol- fluid

700

55 55

Pressureangle (GPa) Dihedral (degree)

Kawamoto35

5

Pressure (GPa)

(a) (b)

70 65 65 60 60

3

1200

40

Dihedral angle (degree)

35

1000

45

Pressure (GPa)

Dihedr

282

50

14 1300

10 700

800

900

1000

1100

Temperature (°C)

1200

1300

> 60 °

ol -this fl mechanism delivers H2O laterally to the partial melting H2O transport and suggested that zone. Such a process was also considered in numerical calculations by Iwamori (1998). 4 mineral is carried deeper by induced mantle flow until the stability limit of the A hydrous hydrous mineral is reached. < 60 °H2O is then liberated and fluids are trapped as structural H2O in the hydrous minerals and as immobile fluids due to their dihedral angle >60°. These would get 6 dragged down in the induced flow, and this process would be repeated until the fluid reaches > 60 ° the zone of partial melting. gt - fl Pressure (GPa)

2

8

10 SECOND CRITICAL ENDPOINT BETWEEN MAGMAS AND AQUEOUS 700 800 900 1000 1100 1200 1300 FLUID:Temperature IMPLICATIONS FOR SLAB-DERIVED COMPONENT (°C)

Simple silicate melts and aqueous fluids can mix completely under certain PT conditions (Fig. 6A,B). At pressure conditions equivalent to the Earth’s upper mantle, silicate melts and aqueous fluids cannot be distinguished from each other at the temperature-pressure conditions beyond a second critical endpoint, where a critical temperature meets its wet solidus (Kennedy et al. 1962; Paillat et al. 1992; Shen and Keppler 1997; Bureau and Keppler 1999). Following the visual demonstration of the complete mixing between albite melt and H2O (Shen and Keppler 1997), Bureau and Keppler (1999) reported complete miscibility between aqueous fluids and K2O-bearing nepheline melt, pure jadeite melt, haplogranitic melt, Ca-bearing haplogranitic melt and dacite in the SiO2-Al2O3-Na2O-K2O-CaO-MgO system. Sowerby and Keppler (2002) demonstrated complete miscibility between B2O3 - F enriched albite melt or pegmatite and H2O

Hydrous Phases & Water Transport in the Subducting Slab increasing pressure

Temperature

Tc

(B) supercritical fluid +A

A + H

H A

A + H

supercritical fluid +A

(C)

H2O

supercritical fluidobservations (B) (C) of (2004b) reported similar +A mixing relationships between aqueous fluids

su +A

and natural andesitic/dacitic melt (Fig. 7). The critical PT conditions observed insupercritica the supercritical fluid andesite/daciteH2O system are similar to those observed in the other simple silicateNo wet solidus Tc melt(Fig.melt but H2O systems 8; Shen and Keppler A +H fluid practical solid 1997; Bureau and Keppler fluid+melt +1999). H

fluid Experiments to determine H2O-satu- fluid +H temperatures often identify +H rated solidus H2Othem from abrupt changes H2OcomH inAchemical position of the minerals and/or the appearance of dendritic textures with increasing (C) supercritical fluidpressure (Inoue 1994; temperature at a given +A Kawamoto and Holloway 1997; Irifune et al 1998; Stalder et al. 2001; Mibe et al. 2002). supercritical fluid have distinguished two types Some workers of dendritic texture, one quenched from No wet solidus partial melt and the other from aqueous flubut ids (Irifune et al. 1998; Litasov and Ohtani A practical solidus 2002). However, they + mentioned that it is H difficult to distinguish between these types fluid of +H texture at pressures greater than 10-13 between H2O GPa. As the critical Htemperature A aqueous fluids and silicate melts decreases with increasing pressure (Paillat et al. 1992; Shen and Keppler 1997; Bureau and Keppler 1999), it should meet an H2O-saturated solidus temperature with increasing pressure (Fig. 6C).

A + H

supercritical fluid +A

fluid +H

No wet solidus but practical solidus

supercritical fluid

No wet solidus but practical solidus

H2O

supercritical fluid

increasing pressure

dry solidus

supercritical fluid

melt +H

melt +H

(C)

fluid +H

fluid +H H A

melt

fluid +H

Tc melt fluid fluid+melt

H A

A + H

Tc melt fluid fluid+melt

H2O

critical fluid

melt +H

supercritical fluid

H2O

H A

melt +A fluidabsent solidus

H A

melt +A fluidabsent solidus melt +H A + H

H A

supercritical fluid +A

(B)

dry solidus

t

melt +H A + H

fluid +H

H2O

melt +A fluidabsent solidus melt +H A + H

melt fluid +melt

fluid +H

wet solidus

fluid +melt

fluid

dry solidus

(A)

Tc

wet solidus

asing pressure

(Sowerby and Keppler 2002). Kawamoto

Temperature fluid

H2O

(A)

283

H A

A + H

H A

Figure 6. Schematic phase diagrams in the system of mineral A and H2O (Kawamoto et al. 2004). H is a hydrous mineral. (A, B) As pressure increases, a critical temperature (Tc) between H2O-bearing silicate melt and silicate-bearing H2O fluid decreases. (C) The Tc meets the H2O-saturated solidus temperature in the system at a second critical endpoint. At pressures beyond that of the second critical endpoint, there is no difference between melts and fluids. In this case there is no H2O-saturated solidus temperature. The practical solidus represents a temperature above which a detectable amount (more than a few percent) of silicate melt is formed (Iwamori 1998).

It is difficult to melt basaltic compositions in a Bassett-type diamond anvil cell due to its temperature limitation of 1100 °C (Bureau and Keppler 1999). Therefore, a supercritical behavior between aqueous fluids and mafic magmas equilibrated with mantle peridotite had remained to be investigated for years. Recently Mibe and his coworkers experimentally determined the PT conditions of a second critical endpoint between peridotite/basalt melts and aqueous fluids by the use of a Kawai-type large volume press and synchrotron X-ray radiography (Kanzaki et al. 1987; Mibe et al. 2004a). They reported that a second critical endpoint between peridotite/basalt melts and aqueous fluids may be located at 3.8 and 3 GPa, respectively (Mibe et al. 2004b, 2005, 2006). This pressure range is lower than that estimated by Kessel et al. (2005). Although

Kawamoto

284

50 °C

1000 °C

920 °C

910 °C

890 °C

950 °C

1025 °C

980 °C

970 °C

Figure 7. Microphotographs showing supercritical behavior between Fuji 1707 andesite and H2O using Bassett-type externally heated diamond anvil cell (Kawamoto 2004b). (A) Chips of the andesitic glass and water are in the rhenium gasket (gasket hole is 0.5 mm) with a small bubble (right) at 50 °C. (B) At 1000 °C and about 1 GPa, a homogeneous fluid, with several grains of unidentified crystals. (C) On cooling to 920 °C, a milky appearance due to tiny droplets of andesite melt in aqueous fluid is seen. (D, E) At 910-890 °C, melt globules are growing in the aqueous fluid. (F) Then during re-heating to 950 °C, the boundary disappears the fluid homogenizes. The crystals are also melting. (G) After heating at 1025 °C, there are no crystals left, and (H, I) during the subsequent cooling, the sample turns milky and separates into andesite globules and aqueous fluid. The difference among the critical temperatures on the first cooling (920 °C, in C, D), the heating (950 °C, in F), and the second cooling (980 °C, in H) could be due to a pressure decrease during the experiment.

Kessel et al. (2005) suggested that there is still a melt-fluid solvus at 4 and 5 GPa, they did not show the coexistence of two phases at 4 or 5 GPa. Therefore, the data shown in Figure 5 of Kessel et al. (2005) can be interpreted as evidence that the fluid compositions observed at 4 and 5 GPa vary continuously with temperature as in Figure 6C and these pressures are already beyond the second critical endpoint. In contrast, Mibe et al. (2004b, 2005, 2006) observed melts and fluids up to 3.8 and 3 GPa in peridotite- H2O and basalt- H2O systems, respectively, and found no coexisting two phases at higher pressures. Although X-ray radiography method is not able to detect a small difference between fluids and melts under certain conditions, Mibe et al. (2004a, 2006) tightly constrain the second critical endpoint between peridotite melt and aqueous fluids at 3.8 GPa, 1000 °C and with 55 wt% H2O. The pressure of 3.8 GPa is equivalent to the depth of the Wadati-Benioff zone beneath the volcanic front. If supercritical fluids are common at the base of the mantle wedge beneath volcanic arcs, the traditional H2O-saturated solidus temperature may represent a temperature where the concentration of silicate components dissolved into aqueous fluids increases drastically and should therefore be described as a practical solidus (Fig. 6C; Iwamori 1998). If supercritical

awamoto igure 8

Hydrous Phases & Water Transport in the Subducting Slab

0

Critical temperature (°C) 500 600 700 800 900 1000 1100 Andesite

Dacite

0.5 1 Pressure (GPa)

285

Ab

1.5

Jd

Ne 2

Ab Hgr

Ca-Hgr

2.5 3

Figure 8. Critical temperatures observed between aqueous fluids and albite (Ab, Shen and Keppler 1997), nepheline (Ne), jadeite (Jd), and haplogranitic (Hgr) melts, Ca bearing haplogranitic melts (Ca-Hgr) and dacite (Bureau and Keppler 1999), and natural calc-alkaline andesite/dacite (Kawamoto 2004b). The estimated second critical endpoints between albite (Stalder et al. 2000), basalt (Mibe et al. 2005), and peridotite (Mibe et al. 2004a, 2006) and aqueous fluids are also plotted.

Basalt

3.5 Peridotite 4

fluids commonly exist in the mantle wedge in subduction zones, such a supercritical fluid could separate into a silicate melt and an aqueous fluid when PT conditions become below the second critical endpoint along its migration to the surface (Fig. 8; Bureau and Keppler 1999). In this case, partitioning of elements between aqueous fluids and silicate melts should occur (Bureau et al. 2004). Such elemental fractionation may affect the chemical characteristics of the volcanic rocks. The existence of a second critical endpoint underneath the volcanic arcs suggests that dense supercritical fluids can come from the slab and separate into the aqueous fluid and melt in the mantle wedge. Otherwise, the slab component would be an aqueous fluid in cold subduction zones or a partial melt in warm subduction zones. Detailed studies of the critical curvatures in the peridotite, basalt, sediment systems will shed light on establishing a quantitative model for the magma generation and H2O transport in subduction zones (Manning 2004).

CONCLUDING REMARKS Our knowledge of the stability of hydrous phases in the downgoing slab has increased dramatically in the last decade (Figs. 2, 3, 4). Recently we have also learned much about the chemical features of aqueous fluids under upper mantle conditions. First, the chemical compositions of silicate components dissolved into aqueous fluids coexisting with mantle peridotite change from silica-rich at pressures lower than 3 GPa to magnesium-rich at pressures greater than 3 GPa (Fig. 1; Stalder et al. 2001; Mibe et al. 2002; Kawamoto et al. 2004). This means that the aqueous fluids in the mantle have peridotitic compositions beneath volcanic arcs. Second, dihedral angles formed between olivine and aqueous fluids change from >60° to 0 (m2s−1), relates the flux of one component J (mol m−2s−1) to its one-dimensional gradient of the concentration dC/dx (mol·m−4) following:  ∂C  J = − D   ∂x  t

Under general, non steady state conditions, the flux in each point, varies with time. In order to satisfy mass balance within the crystal the flux must obey the continuity equation: 1529-6466/06/0062-0013$05.00

DOI: 10.2138/rmg.2006.62.13

292

Ingrin & Blanchard ∂J ∂C =− ∂x ∂t

Combining both equations leads to Fick’s second law:  ∂   ∂C     ∂C   ∂t  =  ∂x  D  ∂x     t   t   x   

(1)

In this general law, the diffusion coefficient, D, is a function of concentration, C. If the diffusion coefficient is independent of concentration then Equation (1) can be simplified to  ∂ 2C   ∂C   ∂t  = D  2   x  ∂x  t

(2)

This assumption is usually considered valid when the concentration of the diffusion species is very small, for instance for hydrogen in nominally anhydrous minerals. For hydrogen diffusion, authors have argued only in few cases, for a dependence of D on concentration when the quality of the fits by Equation (2) was not satisfactory (see for instance Wang et al. 1996). In this work we consider only volume diffusion through the crystal lattice, which is the dominant diffusion process at high temperature. The isotope diffusion coefficients for hydrogen in nominally anhydrous minerals are regarded as impurity tracer diffusion coefficients, which refer to diffusion of species at infinitely small dilution, while isotope diffusion coefficients for hydrous minerals refer to self-diffusion of intrinsic components of the mineral. Most of these coefficients can be assimilated to the impurity-diffusion or selfdiffusion of hydrogen. On the other hand, the effective diffusion coefficients determined from experiments in which hydrogen is extracted or incorporated correspond to the interdiffusion of species with different diffusivities. The relationship between the effective diffusivity Deff and the individual diffusivities of the species is directly linked to the involved reaction (see for instance, Kohlstedt and Mackwell 1998). If a single mechanism of diffusion is involved, the temperature dependence of D can be described by an Arrhenius equation of the form:  − ∆H  D = D0 exp    RT 

(3)

where D0 (m2·s−1) is the pre-exponential factor, ΔH (J·mol−1) is the activation enthalpy, R is the gas constant (8.314 J·K−1mol−1) and T is the temperature (K).

EXPERIMENTAL METHODS All diffusion experiments aim to control the environment around the mineral (i.e., temperature; pressure; fugacities in oxygen, hydrogen, water; metal activities) in order to measure the kinetics of the mechanism of interest. Depending on the kind of minerals and reactions studied, the mineral sample can be placed directly into a furnace, which may be open or flushed by a gas (N2, H2, Ar + H2O, Ar + D2O...), or the sample can be sealed in a closed container with water and buffers, for instance, a silica ampoule placed in a furnace or a metal capsule which is then loaded in a piston cylinder or an internally-heated pressure vessel. Whatever the equipment used, two kinds of experiments can be distinguished: the bulk powder-fluid exchange experiments or experiments performed on single-crystals. In bulk powder-fluid (liquid or gas) exchange experiments, only the average extent of the exchange as a function of time is known from the analysis. The diffusion coefficients are then de-

Diffusion of Hydrogen in Minerals

293

termined from Fick’s second law (Eqn. 1 and 2) solved for the geometry and the boundary conditions imposed by the experimental settings. Some examples of these solutions are given in the next paragraph. As the material is crushed, an assumption must be made about the size and shape of grains. It is important then to have the most homogeneous grain population possible since depending on the grain geometry chosen, the uncertainties of the diffusivities can be significant. This technique also assumes that diffusion is isotropic which is not true for many minerals. Most of these experiments were originally performed by geochemists in order to measure the isotopic fractionation of hydrogen in hydrous minerals; determination of the diffusion coefficients was not the main objective of these studies. However, even though these results are based on relatively rough assumptions, they provide useful results on diffusion rates of hydrogen. Of the experiments performed on single-crystals, we distinguish the mass-loss experiments from those measuring diffusion profiles directly. In the first case, the sample undergoes successive periods of heating under the same conditions. After each heating step, the average hydrogen (or deuterium) concentration lost or remaining in the sample is measured, for instance, with a mass spectrometer, an infrared spectrometer or a thermobalance. In the second case, in profiling experiments, there is only one heating event, which lasts the time necessary for the formation of a full diffusion profile. This requires an assessment a priori of the diffusivity to choose the experiment duration. The diffusion profile is then measured directly by ion beam depth profiling analysis (nuclear reaction analysis or proton-proton scattering), through surface analysis following successive sectioning (scintillation counting) or the sample is cut in slices in order to measure the diffusion profile in the initial sample thickness by infrared absorption spectroscopy. In both mass-loss and diffusion profile measurements, the sample geometry is of great importance. First it simplifies the analytical treatment of the measurements. For mass-loss experiments, a sample with a plate shape with a small ratio of thickness over lateral sizes allows one to assume that the diffusion is unidirectional (i.e., normal to the plate surface) whereas for profiling experiments, one could assume that the diffusion profiles along the three perpendicular directions are independent by cutting the sample with appropriate length ratios. Second, once the sample is oriented with respect to crystallographic directions, it is possible to determine the anisotropy of diffusion. This point represents a great advantage of experiments using single-crystals over powder experiments. The equations used to fit the experimental measurements are solutions to Fick’s second law (Eqn. 1) for different geometries and boundary conditions (Carslaw and Jaeger 1959; Crank 1975). We consider here only the simplest situation where the diffusion coefficient is not a function of the concentration. All the following solutions correspond to diffusional transport-in experiments for some common sample geometries. For a semi-infinite solid with a homogeneous initial concentration, C0, in contact with an infinite, well-mixed outside reservoir of concentration, C1, the diffusion profile along the direction x perpendicular to the surface can be fitted by the following solution. C ( x, t ) − C 0 C1 − C 0

 x  = erfc    2 Dt 

( 4)

In experiments where the sample geometry can be compared to a solid with homogeneous initial concentration, C0, bounded by two infinite parallel planes (thickness, 2L) and in contact with an infinite reservoir, the diffusion coefficient can be determined from the diffusion profile along x, by using the following solution C ( x, t ) − C 0 C1 − C 0

=1−

4 π

 − D ( 2 n + 1)2 π2  exp ∑  4 L2 n =0 2n + 1  ∞

( −1)

n

 ( 2 n + 1) π x  t  cos     2L   

(5)

Ingrin & Blanchard

294

where the origin is located at the mid-point of the slab. If the measurements provide the average concentration rather than the concentration profile like in mass-loss experiments on single-crystals or powder, then Equation (5) has to be integrated over the sample thickness. The solution is then C ( t ) − C0 C1 − C 0

=1−

∞

8 π2

∑

n =0

1

( 2n + 1)

2

 − D ( 2 n + 1)2 π 2 exp   4 L2 

t   

(66)

For the same kind of geometry, if the sample is not thin enough to assume unidirectional diffusion, Equation (5) becomes C ( x, y, z, t ) − C0 C1 − C0

=1−

64 π3

( −1) ∑ ∑ ∑ ( 2l + 1) ( 2m + 1) ( 2n + 1) l =0 m =0 n =0 ∞

∞

l+m+n

∞

( 7)

 − π 2 t  D ( 2l + 1)2 D ( 2 m + 1)2 D ( 2 n + 1)2   y  x  exp  + + z  4   b2 L2 a2     ( 2l + 1) π x   ( 2 m + 1) π y   ( 2 n + 1) π z  cos   cos    cos  2L 2a 2b      

This corresponds to the case of a 2a × 2b × 2L parallelepiped where the diffusion coefficients along the three crystallographic directions are different. As before, the expression of the average concentration is obtained by integrating this equation over the volume analyzed. By integrating over the whole volume (2a × 2b × 2L), the average concentration is expressed as follows: C ( t ) − C0 C1 − C 0

=1−

512 π6

∞

∞

∞

∑∑∑

l =0 m =0 n =0

 − π2 exp   4 

1

( 2l + 1) ( 2m + 1) ( 2n + 1) 2

2

2

(8)

2 2 2 Dy ( 2 m + 1) Dz ( 2 n + 1)   t  Dx ( 2l + 1)  + +   b2 L2 a2  

The corresponding solutions for a spherical geometry (radius, R) are, respectively, C ( r, t ) − C0 C1 − C0

=1+

2R πr

C ( t ) − C0 C1 − C 0

∞

∑

( −1) n

n =1

=1−

n

6 π2

∞

 − D n 2 π2 t  nπr  exp   sin   2 R  R    1

 − D n 2 π2 R2 

∑ n 2 exp  n =1

t  

( 9)

(10)

MEASUREMENT TECHNIQUES In this section, we describe briefly the techniques of analysis that have been reported in the literature for measuring hydrogen diffusion (infrared spectroscopy, mass spectrometry, nuclear reaction analysis, thermogravimetry, scintillation counting) as well as some other new techniques that have not been used for diffusion measurements but are very promising like proton-proton scattering. More details on analytical methods used for measuring water in minerals can be found in this volume (Rossman 2006). We end this section with a short review of the theoretical simulations that contribute to our understanding of the diffusion mechanisms.

Diffusion of Hydrogen in Minerals

295

Infrared spectroscopy Infrared spectroscopy is the most frequently used method to measure hydrogen diffusion. The vibrational modes of the OH dipole within the sample interact with the infrared beam and give rise to absorption bands. The concentration of OH is directly related to the intensity of the bands, and the concentration can be determined if the spectra are measured accurately; IR spectra of anisotropic minerals should be measured in polarized mode (cf. Libowitzky and Rossman 1996a) and the relation between absorption and concentration must be calibrated against an independent hydrogen analysis method. The position in wavenumber of the absorption band depends on the strength of the hydrogen bond, bond geometry and neighbors. Therefore polarized spectra also provide information about the structure of OH (Libowitzky and Beran 2006). Advantages of the IR technique are very high sensitivity (> DMeCMe, DOCO, DSiCSi),

(

)

∇η h  = ∇µ h  + F ∇Φ ≈ 0

( A8a )

− F ∇Φ ≈ ∇µ h 

( A8b)

that is,

Recalling that the flux of silicon ions in Equation (9b) is given by jSi 4+ = −

(

DSiCSi D C ∇ηSi 4+ = − Si Si ∇µ Si 4+ + 4 F∇Φ RT RT

)

( A9a )

then using the relation Si4+ = Si + 4h• yields

(

)

DSiCSi D C ∇µ Si 4+ − 4∇µ h  = − Si Si ∇µ Si RT RT

jSi 4+ = −

( A9b)

Assuming that local thermodynamic equilibrium is established (consistent with the premises of irreversible thermodynamics), then from the reaction Si + O2 = SiO2, (local) equilibrium requires that μSi + μO = μSiO2. Thus, jSi 4+ = −

(

DSiCSi ∇µ SiO2 − ∇µ O 2 RT

)

(A9c)

Likewise j Me2+ = −

DMeC Me D C  1  ∇µ Me = − Me Me  ∇µ MeO − ∇µ O 2  RT RT  2 

( A10)

With DMeCMe >> DOCO, DSiCSi, 1 ∇µ MeO ≈ ∇µ O 2 2

( A11)

Therefore, from the above equations, jSi 4+ = −

(

DSiCSi ∇µ SiO2 − 2∇µ MeO RT

)

( A9d )

Kohlstedt

396

which is identical to Equation (15) and leads directly to Equation (17), jSi 4+ = −

CSi DSi DO ∇µ Me2 SiO 4 RT DO + 4 DSi

( A9e)

17

Reviews in Mineralogy & Geochemistry Vol. 62, pp. 397-420, 2006 Copyright © Mineralogical Society of America

The Effect of Water on Mantle Phase Transitions Eiji Ohtani and K. D. Litasov Institute of Mineralogy, Petrology and Economic Geology Tohoku University Sendai, Miyagi-ken 980-8578, Japan e-mail: [email protected]

INTRODUCTION Water has played an important role in the Earth’s evolution. The incorporation of water as hydroxyl into solid mineral phases or as coexisting hydrous fluids and melts affects the chemical and physical properties of crust and mantle constituents, i.e., it weakens rocks and minerals, reduces viscosity and strength of the materials, and depresses dramatically the melting temperature of silicate minerals (e.g., Karato 1990; Inoue 1994; Hirth and Kohlstedt 1996; Chen et al. 1998; Kubo et al. 1998; Mei and Kohlstedt 2000). Many recent studies have suggested the possible existence of water in the Earth’s mantle especially in the transition zone (e.g., Smyth and Frost 2002; Ohtani et al. 2004; Litasov et al. 2005a; Hae et al. 2006), where wadsleyite and ringwoodite can accommodate up to 3 wt% of H2O in their structures (e.g., Kohlstedt et al. 1996). Low-velocity zones observed seismologically at the top of the 410 km discontinuity may indicate the existence of trapped high-density melt (e.g., Revenaugh and Sipkin 1994; Song et al. 2004; Matsukage et al. 2005; Sakamaki et al. 2006), which is likely to be hydrous as it is not possible to melt the base of the upper mantle without water at these conditions. Electrical conductivity anomalies in the upper mantle and transition zone that are related to subduction zones have also been interpreted as an effect of water in the mantle (e.g., Fukao et al. 2004; Tarits at el. 2004; Hae et al. 2006; Koyama et al. 2006). Studies of the kinetics of the hydrous olivine-wadsleyite transformation (Ohtani et al. 2004; Hosoya et al. 2005) seem to be consistent with seismological observations (e.g., Koper et al. 1998) that indicate the absence of a metastable olivine wedge in subducting slabs. Such a metastable wedge would be expected as a result of the sluggish olivine-wadsleyite transformation under anhydrous conditions (Rubie and Ross 1994). Studies of the elasticity of hydrous wadsleyite and ringwoodite indicate that P- and S- wave velocities of the transition zone are consistent with the existence of hydrated wadsleyite and ringwoodite (Inoue et al. 2004; Jacobsen et al. 2004). These data indicate that a significant amount of water may be stored in the mantle especially in the transition zone. Seismic discontinuities at 410 and 660 km depths are well established on a global scale to the point where they occur in reference velocity models such as PREM (Dziewonski and Anderson 1981). These discontinuities are usually attributed to the phase transformations of olivine in mantle peridotite. Olivine α-(Mg,Fe)2SiO4 transforms to wadsleyite β-(Mg,Fe)2SiO4 at a depth of approximately 410 km, and ringwoodite γ-(Mg,Fe)2SiO4 decomposes to perovskite (Mg,Fe)SiO3 and magnesiowustite (Mg,Fe)O at approximately 660 km depth. The latter transformation is frequently termed the post-spinel transformation. The topography and sharpness of these discontinuities depend on mantle temperatures, chemical compositions and mineral proportions (e.g., Agee 1998; Weidner and Wang 2000; Frost 2003). According to most seismological studies the 410 and 660 km discontinuities are sharp, and the change of density and velocity occurs over a small depth interval, 4-35 km for 1529-6466/06/0062-0017$05.00

DOI: 10.2138/rmg.2006.62.17

398

Ohtani & Litasov

the 410 km discontinuity and