Methods and Instrumentations: Results and Recent Developments

To Be Published in This Series

Volume 1 Composition, Structure, and Properties of Mineral Matter: Concepts, Results, and Problems

Volume 2 Methods and Instrumentations: Results and Recent Developments

Volume 3 Mineral Matter in Space, Mantle, Ocean Floor, Biosphere, Environmental Management, Jewelry

Volume 4 Processes of Mineral Formation: Frontiers Experiment and in Evolution in Geological History

Volume 5 Minerals as a Source of Metals, Energy and Materials

A.S. Marfunin (Ed.)

Advanced Mineralogy

Volume 2 Methods and Instrumentations: Results and Recent Developments

With 120 Figures and 18 Tables

Springer-Verlag Berlin Heidelberg New York London Paris Tokyo Hong Kong Barcelona Budapest

Prof. Dr. A.S. MARFUNIN

Geological Faculty

University of Moscow

119899 Moscow

Russia, CIS

ISBN 978-3-642-78528-3 001 10.1007/978-3-642-78526-9

ISBN 978-3-642-78526-9 (eBook)

Library of Congress Cataloging-in-Publication Data Advanced mineralogy /Marfunin. Arnold S. (ed.). p. em. Includes bibliographical references and index. Contents: - v. 2. Methods and Instrumentations ISBN-13: 978-3-642-78528-3 1. Mineralogy. I. Marfunin. Arnol'd Sergeevich. QE363.2.A35 1994 549-dc20 94-13315

This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer-Verlag. Violations are liable for prosecution under the German Copyright Law.

© Springer-Verlag Berlin Heidelberg 1995 Softcover reprint of the hardcover 1 st edition 1995

The· use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use.

Typesetting: Macmillan India Ltd., Bangalore 25

SPIN: 10128541 32/3145/SPS-5 4 3 2 1 0 - Printed on acid-free paper

Preface

All existing introductory reviews of mineralogy are written according to the same algorithm, sometimes called the "Dana System of Mineralogy". Even modern advanced handbooks, which are certainly necessary, include basic data on minerals and are essentially descriptive. When basic information on the chemistry, structure, optical and physical properties, distinguished features and paragenesis of 200-400 minerals is presented, then there is practically no further space available to include new ideas and concepts based on recent mineral studies.

A possible solution to this dilemma would be to present a book beginning where introductory textbooks end for those already familiar with the elementary concepts. Such a volume would be tailored to specialists in all fields of science and industry, interested in the most recent results in mineralogy.

This approach may be called Advanced Mineralogy. Here, an attempt has been made to survey the current possibilities and aims in mineral mater investigations, including the main characteristics of all the methods, the most important problems and topics of mineralogy, and related studies.

The individual volumes are composed of short, condensed chapters. Each chapter presents in a complete, albeit condensed, form specific problems, methods, theories, and directions of investigations, and estimates their importance and strategic position in science and industry.

The following fields will be covered in the individual volumes:

Vol. 1 Composition, Structure, and Properties of Mineral Matter: Concepts, Results, and Problems

Vol. 2 Methods and Instrumentations: Results and Recent Developments

Vol. 3 Mineral Matter in Space, Mantle, Ocean Floor, Biosphere, Environmental Management, Jewelry

Vol. 4 Processes of Mineral Formation: Frontiers in Experiment and Evolution in Geological History

Vol. 5 Minerals as a Source of Metals, Energy and Materials

VI Preface

The book thus attempts to present a universal (or perhaps a wholistic) approach to the nature and role of mineral matter, by presenting frontier facts and hypotheses in as many fields of the mineral science as possible.

A complex set of volumes like this could never have been written by just one author. I am therefore happy that top specialists from all over the world and from different disciplines agreed to contribute. I have had the privilege of discussing the topics through extensive communication with the authors, orally and in writing, and I wish to thank them for their support and collaboration.

I am also grateful for the discussions of the different aspects of the book with R.J. Kirkpatrick (Urbana), G. Rossman (Pasadena), Chr. Amstutz (Heidelberg), W. Baur (Frankfurt), G. Amthauer (Salzburg), A. Beran (Vienna), S. Hafner (Marburg), Ch. Prewitt (Washington), Xie Xiande (Guangzhou), Y. Dusausoy (Nancy), W. Engel (Heidelberg, Springer-Verlag), L.V. Bershov and N.F. Chelishchev (Moscow).

Moscow. October 1994 A.S. Marfunin

Contents

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . XIX

Chapter 1 Systematics of the Methods of Investigation of Minerals: Logic of Development A.S. MARFUNIN . . . . . . . . . . . . . . . . .

Chapter 2 Diffraction Methods and Crystal Structure Analysis 15

2.1

2.2

Crystal Structure Analysis and X-Ray Diffraction Instrumentation A. GUINIER, TH. HAHN, and V.1. SIMONOV. X-Ray Diffraction Techniques for the Characterization of Minerals S. ALTANER and I.E. KAMENTSEV ...

2.3 Neutron Scattering, Neutron Diffraction: Hydrogen Location, Cation Distribution, Magnetic Structures

16

32

H. FUEss. . . . . . . . . . . . . 38 2.4 Electron Diffraction Analysis

B.B. ZVYAGIN ......... . 50

Chapter 3 Solid State Spectroscopy . . . . . . . . . . . . . . .. 65

3.1 3.1.1

3.1.2

3.1.3

3.1.4

3.2

Nuclear Gamma Resonance (M6ssbauer) Spectroscopy Summary of Theory and Important Results F.e. HAWTHORNE . . . . . . . . . . . . . . . . . . . . . Experimental Techniques and Spectrum Fitting F.e. HAWTHORNE, A.V. BYKOV, N.N. DELYAGIN, and V.1. NIKOLAEV ..................... . Iron-Containing Minerals, Ores and Glasses G. AMTHAUER, F.e. HAWTHORNE, and E. POLSHIN . M6ssbauer Spectroscopy of Sn, Sb, Eu, Au F.e. HAWTHORNE . . . . . . . . . . . . . . . . . . . . X-Ray and Photoelectron Spectroscopy of Minerals.

66

66

69

74

83 87

VIII

3.2.1 Parameters in Different Types of X-Ray Spectra

3.2.2 D.S. URCH .................... . Mineralogical and Geochemical Information from X-Ray Absorption Spectroscopy

Contents

87

A. MANCEAU and G. WAYCHUNAS . . . . . . . 91 3.3 Optical Absorption Spectroscopy

K. LANGER, A.N. PLATO NOV, and G.R. ROSSMAN. 109 3.4 3.4.1

Luminescence of Minerals . . . . . . . . . . . . . 124 Interpretation of Luminescence Spectra in Terms of Band Theory and Crystal Field Theory. Sensitization and Quenching. Photoluminescence, Radioluminescence, and Cathodoluminescence A.N. TARASHCHAN and G. WAYCHUNAS . . . ., 124

3.4.2 Selective Laser Excitation of Rare-Earth Luminescence Spectra M. IUEV and M. SENDOVA-V ASSIUEVA . . . ., 136

3.4.3 Origins of Luminescence in Minerals: A Summary of Fundamental Studies and Applications B.S. GOROBETS and G. WALKER. . . . . . . . . . . . . 138

3.5 Thermoluminescence and Exoelectron Spectroscopy of Minerals . . . . . . . . . . . . . . . . . . . . . . . 147

3.5.1 Mechanisms and Parameters; Factors Governing Thermoluminescence S.W.S. McKEEVER . . . . . . . . . . . . . . . . . .. . 147

3.5.2 Thermoluminescence Applications S.W.S. McKEEVER, V.K. VLASOV, O.A. KuuKov, and K.S.V. NAMBI . . . . . . . . . . . . . . . . . . . . . . .. 157

3.5.3 Exoelectron Spectroscopy of Minerals V.S. KORTOV. . . . . . . . . . . . . . . .

3.6 Infrared Spectroscopy ................... . 3.6.1 Band Assignments in Infrared and Raman Spectroscopy

A.N. LAZAREV, P.F. McMILLAN, and S.W. KIEFFER 3.6.2 Polarized Infrared Spectra

A. BERAN ........................ . 3.6.3 Applications of Infrared Spectroscopy

3.7

3.8 3.8.1

3.8.2

to Structure and Bonding in Minerals and Glasses and to Speciation of Hydrous Components W.B. WHITE and A.M. HOFMEISTER . . . . . Raman Spectroscopy in Earth Sciences J. DUBESSY, R.Y. ORLOV, and P. McMILLAN Electron Paramagnetic Resonance (EPR) . Principles, Technique, Applications in Mineralogy J.A. WElL, Y. DUSAUSOY, and S.L. VOTYAKOV .... Electron Nuclear Double and Multiple Resonance J.R. NIKLAS, A.B. BRICK, and I.-M. SPAETH ..... .

166 174

174

180

183

189 197

197

209

Contents IX

3.8.3 EPR: Improvement of Experimental Technique Y A.S. LEBEDEV. . . . . . . . . . . . . . . . . . . . . 211

3.9 Nuclear Magnetic Resonance (NMR) Spectroscopy R.J. KIRKPATRICK. . . . . . . . . . . . . . 213

3.10 Nuclear Quadrupole Resonance (NQR) LN. PENKOV and D. BRINKMANN. . . . . 224

3.11 Muon Resonance. Application to the Study of the Hydrogen Atom Position in Quartz J.A. WElL . . . . . . . . . . . . . . . . . . . . . . . . . .. 227

Chapter 4 Remote Sensing Methods: Visible, Infrared, and Microwave B. CERVELLE . . . . . . . . . . . . . . . . . . . . . .. 229

Chapter 5 Microprobe Analysis . . . . . . . . . . . . . . . . .. 239

5.1 Electron Probe Microanalysis S.J.B. REED and I.M. ROMANENKO . ...... 240

5.2 Trace Element Microanalysis by Proton-Induced X-Ray Emission (PIXE): The Proton Microprobe D.S. WOOLUM .................. . 246

5.3 Nuclear Microprobe and Microscopic Analysis P. TROCELLIER. . . . . . . . . . . . . . . . . . . . . . . .. 254

Chapter 6 Electron, Acoustic, and Tunneling Microscopy of Minerals . . . . . . . . . . . . . . . . . . . . . .. 263

6.1 Electron Microscopy of Minerals H.-R. WENK, A.C. McLAREN, G.M. PENNOCK, and V.A. DRITS . . . . . . . . . . . . ..

6.1.1 Fundamentals of TEM and HRTEM 264

A.C. McLAREN . . . . . . . . . . . . . . 264 6.1.2 Scanning Electron Microscopy and Image Formation

G.M. PENNOCK . . . . . . . . . . . . . . . . . . . . . . 273 6.1.3 Applications of Transmission Electron Microscopy

H.R. WENK . . . . . . . . . . . . . . . . . . . . . 279 6.1.4 Applications of Scanning Electron Microscopy

G.M. PENNOCK . . . . . . . . . . . . . . 296 6.2 High Resolution Acoustic Microscopy

U. BELLER . . . . . . . . . . . . . . . . . 298 6.3 Scanning Tunneling and Atomic Force Microscopy

A.V. ERMAKOV and S.V. TITKOV . . . . . . . . . . . . 300

x Contents

Chapter 7 Recent Developments in Analytical Methods in Mineralogy ..... 303

7.1 General Overview of the Methods of Analysis of Minerals, Rocks, Ores, and Materials PJ. POTTS ........... .

7.2 Classical and Rapid Methods P.J. POTTS ........... .

7.3 Atomic Absorption Spectrometry P.J. POTTS ............ .

7.4 Inductively Coupled Plasma -Atomic Emission Spectroscopy

304

308

311

J.G. CROCK and P. H. BRIGGS 315 7.5 X-Ray Fluorescence Analysis

V.P. AFONIN. . . . . . . . . . . 319 7.6 Neutron Activation Analysis

CHR. KOEBERL. . . . . . . . .. ........... 322 7.7 Nuclear Techniques for Uranium and Thorium Analysis

S.J. PARRY. . . . . . 329 7.8 Mass Spectrometry

P.F. McDERMOTT. . 332 7.9 Inductively Coupled Plasma Mass Spectrometry

K.E. JARVIS . . . . . . . . . 337 7.10 Ion Exchange Techniques

P.J. POTTS ......... . 340

Chapter 8 Isotopic Mineralogy. . . . . . . . . . . . . . . . . .. 345

8.1 Radioactive Isotopes in Mineralogy and Geochemistry Yu.A. SHUKOLYUKOV and K. WETZEL. . . . . 346

8.2 Isotopic Systems in Geochronology. . . . . . 357 8.2.1 The K-Ar Isotope System in Geochronology

Yu.A. SHUKOLYUKOV and H.J. LIPPOLT 357 8.2.2 40 Ar r 9 Ar and its Laser Variant

M.1. KARPENKO, and J.F. SUTTER. . . . . . . . . . 362 8.2.3 The Rb-Sr Method of Isotopic Dating

D.J. DEPAOLO, T.F. ANDERSON, and V.I. VINOGRADOV 364 8.2.4 The Sm-Nd Method of Isotope Dating

V.I. VINOGRADOV, DJ. DEPAOLO, and T.F. ANDERSON 366 8.2.5 The U-Pb System and Zircon

as Mineral Geochronometer E.V. BIBIKOVA and J.N. ALEINIKOFF . . . . . . . . . .. 368

8.3 Noble Gas Isotopes in Planetary and Earth Minerals Yu.A. SHUKOLYUKOV and M. OZIMA. . . . . . . . . .. 374

Contents

8.4 Radiogenic Isotopes as Indicators of Sources of Mineral Matter . . . . . . . . . .

8.4. t Pb Isotopy; The Lead Sources LV. CHERNYSHEV and B.L. GULSON

8.5 Light Stable Isotope Ratios as Indicators for Conditions of Mineral Formation ...

8.5.1 Theoretical Aspects of Isotopic Fractionation 1.R. O'NEIL and E.M. GALIMOV .....

8.5.2 Natural Variations in Stable Isotopes 1. HOEFS and V.I. VINOGRADOV .....

8.5.3 Oxygen and Hydrogen Isotopes in Mineralogy

XI

385

385

388

388

396

B.G. POKROVSKY and T.F. ANDERSON . . . . . . . 398 8.5.4 Carbon Isotopes in Mineralogy and Geochemistry

E.M. GALIMOV and D. RUMBLE. . . . . 401 8.5.5 Sulfur Isotopes in Mineralogy

V.L VINOGRADOV and T.F. ANDERSON. 410 8.5.6 Nitrogen Isotopes in Mineralogy

D. HAENDEL and B.G. POKROVSKY . . . . . . . . . . . .. 414 8.6.1 Geochemical Significance of 87Sr/86Sr Isotopic Ratios

T.F. ANDERSON, D.l. DEPAOLO, and V.L VINOGRADOV 416 8.6.2 Geochemical Significance of 143Nd/144Nd

Isotopic Ratios D.l. DEPAOLO, T.F. ANDERSON, and V.1. VINOGRADOV . 418

Chapter 9 Computer Databases in Mineralogy D.G.W. SMITH.

Subject Index . . . . . . . .

421

437

List of Contributors

AFONIN, V.P., Institute of Geochemistry, Favorky Str. la, Irkutsk 664033, Russia

ALEINIKOFF, J., US Geological Survey, Denver Federal Center, PO Box 25046, Denver, CO 80225, USA

ALTANER, S., Department of Geology, University of Illinois, Urbana, IL 61801, USA

AMTHAUER, G., Institut fiir Mineralogie, Universitat Salzburg, Hellbrunnerstrasse 34, A-5020 Salzburg, Austria

ANDERSON, T.F., Department of Geology, University of Illinois, Urbana, IL 61801, USA

BERAN, A., Institut fUr Mineralogie and Kristallographie der Universitat Wien, Dr-Karl-Liiger-Ring 1, A-10lO Wien, Austria

BIBIKOVA, E.V., Vernadsky Institute of Geochemistry, Kosygina 19, Moscow 117975, Russia

BRICK, A., Institute of Geochemistry and Physics of Minerals, Palladina 34, 252680 Kiev 142, Ukraina

BRINKMANN, D., Physik-Institute der Universitat Ziirich, Schonberggasse 9, CH-8001 Ziirich, Switzerland

BRIGGS, P.H., US Geological Survey, Denver Federal Center, Denver, CO 80225, USA

BRYZGALOV, J.A., Geological Faculty, Moscow University, Moscow 119899, Russia

BUSECK, P.R., Department of Geology, Arizona State University, Tempe, AZ 85287, USA

BYKov, A.V., Physical Faculty, Moscow University, Moscow 119899, Russia

CERVELLE, B., Laboratoire de Mineralogie-Cristallographie, Universite Pierre et Marie Curie Paris VI-VII, Tour 16, 4 Place Jussieu, F-75252 Paris Cedex 05, France

CHERNYSHEV, J.V., IGEM Academy of Sciences, Staromonetny 35, Moscow lO9017, Russia

CROCK, J.G., US Geological Survey, Branch of Geochemistry, Mail Stop 973, Denver Federal Center, Denver, CO 80225, USA

XIV List of Contributors

DELYAGIN, N.N., Physical Faculty, Moscow University, Moscow 119899, Russia

DE PAOLO, D.I., Department of Earth and Space Sciences, UCLA, Los Angeles, CA 90024, USA

DRITS, V.A., Geological Institute, Academy of Sciences, Pyzhevsky 7, Moscow 109017, Russia

DRURY, M.R., Research School of Earth Sciences, The Australian National University, Canberra, ACT 2601, Australia

DUBESSY, J., CREGU, 3, rue de Champelle, F-54501 Vandoeuvre les Nancy Cedex, France

DUSAUSOY Y, Laboratoire de Mineralogie-Cristallographie, Universite de Nancy, F-54037 Nancy Cedex, France

ERMAKOV, A.V., Institute of Physics, Leningrad University, Ulyanovskaya 1, St Petersburg-Petrodvorets 198904, Russia

FUEss, H., Technische Hochschule, Karolinenplatz 5, D-64289 Darmstadt, Germany

GALIMOV, E.M., Vernadsky Institute of Geochemistry, Kosygina 19, Moscow 117975, Russia

GORBATOV, G.A., Institute of Mineral Resources (VIMS), Staromonetny 33, Moscow 109017, Russia

GOROBETS, B.S., Institute of Mineral Resources (VIMS), Staromonetny 33, Moscow 109017, Russia

GORSHKOV, A.I., IGEM, Academy of Sciences, Staromonetny 35, Moscow 109017, Russia

GUINIER, A., Laboratoire de Physique des Solides, Universite Paris-Sud, F-91405 Orsay Cedex, France

GULSON, B.L., Division of Mineral Physics and Mineralogy, CSIRO, Institute of Energy and Earth Resources, North Ryde, N.S.W. 2113, Australia

HAENDEL, D., Umweltforschungszentrum, Leipzig/Halle GmbH, Sektion Hydrogeologie, Hallesche Str. 44, 06246 Bad Lauchstiidt, Germany

HAGGERTY, St.E., Morrill Science Center, Geological Department, University of Massachusetts, Amherst, MA 01003, USA

HAHN, T., Institut fUr Kristallographie, RWTH, Templergraben 55, D-52062 Aachen, Germany

HAWTHORNE, F.e., Department of Geological Sciences, University of Manitoba, Winnipeg, Manitoba R3T 2N2, Canada

HOEFS, I., Geochemisches Institut, Universitiit G6ttingen, Goldschmidtstrasse 1, D-37077 G6ttingen, Germany

HOFMEISTER, A.M., Department of Geology, 175 Physics/Geology University of California, Davis, CA 95616, USA

ILIEV, M., Faculty of Physics, University of Sofia, 1126 Sofia, Bulgaria

List of Contributors xv

JARVIS, K.E., Centre for Analytical Research in the Environment, Silwood Park, Buckhurst Road, Ascot, Berks SL5 7TE, England

KAMENTSEV, I.E., Chair of Crystallography, St Petersburg University, St Petersburg 199034, Russia

KARPENKO, S.F., IGEM, Academy of Sciences, Staromonetny 35, Moscow 109017, Russia

KIEFFER, S.W., Geology Department, Arizona State University, Tempe, AZ 85287-1404, USA

KIRKPATRICK, RJ., Department of Geology, University of Illinois, Urbana, IL 61801, USA

KOEBERL, CHR., Institute of Geochemistry, University of Vienna, A-10lO Vienna, Austria

KORTOV, V.S., Department of Experimental Physics, Ural Poly technical Institute, Ekaterinburg 62002, Russia

KULIKOV, O.A., Geographical Faculty, Moscow University, Moscow 119899, Russia

LANGER, K., Institut fUr Mineralogie und Kristallographie, Technische Universitiit Berlin, Ernst-Reuter-Platz 1, D-I0587 Berlin, Germany

LAZAREV, A.N., Institute of Silicate Chemistry, Makarov Quay 2, St Petersburg 199034, Russia

LEBEDEV, Y A.S., Institute of Chemical Physics, Kosygina 4, Moscow 117977, Russia

LIEBAU, F., Mineralogisch-Petrographisches Institut, Universitiit Kiel, Olshausenstrasse 40-60, D-24118 Kiel, Germany

LIPPOLT, HJ., Laboratorium fur Geochronologie, Universitiit Heidelberg, D-69121 Heidelberg, Germany

MANCEAU, A., Environmental Geochemistry Group, LGITIRIGM, University of Grenoble, BP53, 38041 Grenoble Cedex 9, France

MARFUNIN, A.S., Department of Mineralogy, Geological Faculty, Moscow University, Moscow 119899, Russia

McDERMOTT, P.F., Department of Earth Sciences, The Open University, Walton Hall, Milton Keynes MK7 6AA, England

McKEEVER, S.W.S., Department of Physics, State University of Oklachoma, Stillwater, OK 74078, USA

McKEOWN, D.A., National Institute of Standards and Technology, Gaithersburg, MD 20899, USA

McLAREN, A.c., Research School of Earth Sciences, The Australian National University, Canberra ACT 2601, Australia

McMILLAN, P.F., Department of Chemistry, Arizona State University, Tempe, AZ 85287, USA

XVI List of Contributors

NAMBI, K.S.V., Environmental Assessment Section, Bhabha Atomic Research Center, Trombay, Bombay 85, India

NIKLAS, LR., Experimentalphysik, UniversiHit-Gesamthochschule Paderborn, Warburgerstrasse tOOA, D-33098 Paderborn, Germany

NIKOLAEV, V.L, Physical Faculty, Moscow University, Moscow, 119899, Russia

O'NEIL, I.R., Department of Geological Sciences, The University of Michigan, 1006 c.c. Little Building, Ann Arbor, MI 48109-1063, USA

ORLOV, R.lu., Geological Faculty, Moscow University, Moscow 119899, Russia

OZIMA, M., Geophysical Institute, University of Tokyo, Tokyo 113, Japan

PARRY, S.J., Imperial College Reactor Centre, Silwood Park, Buckhurst Road, Ascot, Berks. SL5 7TE, England

PENKOV, LN., Geological Faculty, Kazan University, Kazan 420111, Russia

PENNOCK, G., Research School of Earth Sciences, The Australian National University, Canberra, ACT 2601, Australia

PLATONOV, A.N., Institute of Geochemistry and Physics of Minerals, Palladina 34, Kiev 252680, Ukraina

POKROVSKIY, B.G., Geological Institute, Academy of Sciences, Pyzhevsky 7, Moscow 109017, Russia

POLSHIN, E.M., Institute of Geochemistry and Physics of Minerals, Palladina 34, Kiev 252680, Ukraina

POTTS, PJ., Department of Earth Sciences, The Open University, Walton Hall, Milton Keynes MK 7 6AA, England

PREWITT, CH., Geophysical Laboratory, 5251 Broad Branch Road, Washington, DC 20015, USA

REED, S.J.B., Department of Earth Sciences, University of Cambridge, Cambridge CB2 3EQ, England

ROMANENKO, I.M., Institute of Experimental Mineralogy, Academy of Sciences, Chernogolovka 124432, Moscow Region, Russia

ROSSMAN, G.R., Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA

RUMBLE, D., Geophysical Laboratory, 5241 Broad Branch Road, Washington, DC 20015, USA

RYABEVA, E.G., Institute of Mineral Resources (VIMS), Staromonetny 33, Moscow 109017, Russia

SENDOVA-VASSILIEVA, M. Faculty of Physics, Sofia University, 1126 Sofia, Bulgaria

SHUKOLYUKOV, Ju.A., Vernadsky Institute of Geochemistry, Kosygina 19, Moscow 117975, Russia

List of Contributors XVII

SIMONOV, V.I., Institute of Crystallography, Leninsky Prospect 59, Moscow 117333, Russia

SMITH, D.G.W., Department of Geology, University of Alberta, Edmonton, Alta, T6G 2E3, Canada

SPAETH, I.M., Experimentalphysik, Universitiit-Gesamthochschule Paderborn, Warburgerstrasse 100A, D-33098 Paderborn, Germany

SUTTER, l.F., US Geological Survey, Reston, VA 22092, USA TARASHCHAN, A.N., Institute of Geochemistry and Physics

of Minerals, Palladina 34, Kiev 252142, Ukraina TITKOV, S.V., IGEM, Academy of Sciences, Staromonetny 35,

Moscow 109017, Russia TROCELLIER, P., Laboratoire Pierre Sue, Direction des Sciences de

la Matiere, Departement de Physique General, CEA/CEN Saclay, F-91191 Gif sur Yvette, France

URCH, D.S., Department of Chemistry, Queen Mary College, Mile End Road, London E1 4NS, England

VINOGRADOV, V.I., Geological Institute, Academy of Sciences, Pyzhevsky 7, Moscow 109017, Russia

VOTYAKOV, S.L., institute of Geology and Geochemistry, Pochtovy 7, Ekatezinburg 620644, Russia

WALKER, G., Department of Pure and Applied Physics, University of Manchester, Institute of Science and Technology, Manchester M60 lQD, England

W A YCHUNAS, G.A., Center for Materials Research, 351 McCullough Bldg, Stanford University, Stanford CA 94305, USA

WElL, l.A., Department of Chemistry, University of Saskatchewan, Saskatoon, Saskatchewan S7N OWO, Canada

WENK, H.R., Department of Geology and Geophysics, University of-California, Berkley, CA 94720, USA

WETZEL, K., Wintergarten Str. 2(226), D-04103 Leipzig, Germany WHITE, W., Material Research Laboratory, Pennsylvania State

University, University Park, PA 16802, USA WOOLUM, D.S., Physics Department, California State University,

Fullerton, CA 92634, USA ZOTOV, N.S., Institute of Applied Mineralogy, Bulgarian Academy

of Sciences, Rakovski Street 92, Sofia 1000, Bulgaria ZVYAGIN, B.B., IGEM, Academy of Sciences, Staromonetny 35,

Moscow 109017, Russia

Introduction

Changes in the methods of investigation have had a crucial impact on progress in mineralogy. Many important methods have emerged with their own parameters, interpretations, problems, new albeit restricted possibilities and scopes, and with their own instrumentation and measurement tools. Even the aims and contents of the investigations have changed.

Two processes can be discerned:

1. The logical completion of developing and elaborating the methods, i.e. determining all physical principles based on a single, general, multifaceted phenomenon: the interaction of radiation (particles) with matter.

These interactions and hence these methods can be presented schematically by two coordinates. One is energy (or wavelength of radiation) and parts of the electromagnetic spectrum (nuclear, X-ray and electron, UV, visible, IR, microwave, SHF and RF). The other coordinate is represented by the modes of interaction (spectroscopy-absorption, emISSIOn, scattering-diffractometry, microscopy). The various intersection points in this system of coordinates have already been determined and elaborated in detail according to theoretical and instrumentational aspects.

2. The natural, ultimate possibilities of the methods were then realized with the new generation of instruments. Their development from 1960-1990 puts into effect the principles of interaction of radiation with matter and brings them to the limits of detection, accuracy, and precise determination. Furthermore, it allows the complex measurement of small and diverse objects.

Completely new spectroscopic methods, developed in the last decades, resulted in an enormous volume of new data on mineral matter. These include: nuclear gamma resonance (Mossbauer) spectroscopy, all variants of X-ray and photo-electron spectroscopy, optical absorption spectroscopy, all types of luminescence, infrared and Raman spectroscopy, electron paramagnetic resonance, and nuclear magnetic and nuclear quadrupole resonance. In a short time, some decisive changes regarding these new methods have taken place, opening even greater possibilities: e.g. synchroton radiation in X-ray spectroscopy, development of EXAFS, XANES, ESCA, Auger methods, laser excited IR, Raman and luminescence spectra, MASSNMR, spin-echo in NQR, and electron-hole centers in EPR.

However, these developments have a more general significance. All methods together compose, in principle, a single, new field: solid state spectroscopy. This

xx Introduction

not only represents a pure analytical or diagnostic method, but also allows the direct observation of atomic properties of chemical bonding in crystals and molecules.

For more than a century microscopy was the basic tool in optical mineralogy and petrography as well as in ore research. Moving beyond the visible regions, today the whole spectrum from nuclear to radar radiation offers the possibility of direct imaging of atoms and point defects, visualization of crystal structures, modulated, hybrid and dissolution structures, and non-stoichiometric, nanometer scale surface studies, in combination with microprobe analysis and electron microdiffraction.

Methods now include nuclear microscopy, X-ray topography, scanning tunneling, and atomic force microscopy, high-resolution, analytical and highvoltage electron microscopy, scanning Auger microscopy, and high-resolution acoustic microscopy.

The major steps forward in the development of diffraction methods are listed below:

- major innovations in the determination of crystal structures: introduction of automated, direct methods for small-structure crystallography, improved Patterson, the Rietveld method for X-ray powder diffraction data;

- progress in electronics, computers and X-ray techniques: X-ray detection by position-sensitive, linear, two- and three-dimensional counters; electronically controlled, four-circle diffractometers; the impact of computers on calculation and measurement procedures; powerful, rotating-anode X-ray tubes; tunable synchroton radiation sources;

- development of neutron scattering methods and diverse electron microdiffraction methods combined effectively with electron microscopy and microprobes;

- construction of X-ray diffraction instrumentation for high-pressure and high-temperature crystal structure determinations;

- precise electron density measurements, calculations and interpretations; - computer-based X-ray scattering methods for non-crystalline materials; - extremely versatile powder diffraction methods for phase analysis and for

the determination of the most complicated characteristics of the ore and rockforming minerals, in microsamples, as poorly crystallized materials.

Technical and electronic discoveries include new sources of radiation (synchroton, laser, klystron, among others) which have led to new generations of spectrometers and advancements in EXAFS and XANES methods. Fine focusing for radiation of all wavelengths has led to the beginning of an epoch of direct microprobe analysis and identification of ore and rock-forming minerals.

Progress in analytical and isotopic methods and instruments has opened possibilities for direct measurements at ultrahigh pressures and temperatures (e.g. measurements of spectra of silicate melts). The integration of isotopic mineralogy in studies on diamonds and carbonates, sulfides and sulfates, zircons and galena may help to contribute to the understanding of such important

Introduction XXI

problems as mantle sources of ore matter, planetary evolution, and cosmochemistry.

Computerized databases (DB) on mainframes or personal computers are beginning to transform the processes of search, storage and treatment of mineralogical information. New software is available: e.g. crystal structure DB, computer-readable powder diffraction file, thermodynamic DB, some spectroscopic DB, and mineral collection catalogues. The accumulation of the enormous mass of analytical data for minerals, rocks, and ores is the consequence of applying rapid instrumental methods in automatized, robotized, and computerized chemical laboratories, thus producing hundred thousands of analytical data. The retrieval systems for the storage and manipulation of these data are indispensable.

CHAPTER 1

Systematics of the Methods of Investigation of Minerals: Logic of Development

2 Chapter 1. Systematics of the Methods of Investigation of Minerals

Systematics of the Methods of Investigation of Minerals: Logic of Development

A.S. MARFUNIN

An increasing number of the methods and techniques, diversity, and complexity of the instrumentations involved in analysis and investigation of mineral matter are becoming nearly boundless.

At the same time, the new developments in mineral science are more and more crucially dependent on the use of advanced methods. The extreme variety of classes and groups of minerals (from diamond to clays, from perovskite to zeoli ties); the complexity of the composition due to impurities, unmixing, ordering, microinclusions, trace elements, rare isotopes; the necessity of nondestructive determinations of different kinds of centers in smallest quantities and with limits of sensitivity; glasses and melts; surface states and adsorbed molecules; measurements at ultra-high pressures and temperatures; remote sensing; high-resolution imaging; time-resolved spectra and diffraction patterns; analysis and structure determinations in micron grains; many thousands of analyses of ores, rocks, and minerals ~ all these and similar purposes and conditions determine the need for applying the most diverse methods and instrumentations.

This concise chapter attempts:

1. to arrange them in an order and to organize them in a system presenting the hierarchy of the principal branches, methods, technical variants, and modifications;

2. to understand the most general essential features: ~ in the basic principles of the methods, ~ in the development of instrumentations.

Systematics Based on the Principles of the Methods

A first division of methods can be made into two groups:

~ methods based on the interactions of matter with radiations and particles; these are analytical and structural methods;

~ methods based on the effects of mechanical, optical, electrical, magnetic, thermal actions on matter, which determine the corresponding properties. The development of modern material science has led to extremely complicated manifestations and the use of the properties of solids and in particular minerals; their description and methods of determination have been considered in preceding chapters.

The second division: all the physical methods based on the interactions of matter with radiation can be classified depending on the two characteristics composing these interaction:

Systematics Based on the Principles of the Methods 3

- kind of radiation (energy of radiation or region of electromagnetic spectrum): nuclear, X-ray, ultraviolet, visible, infrared, microwave, super high frequency, radio frequency;

- mode of interaction: spectroscopy, microscopy, diffractometry (and microprobes combining two or three of these modes of interaction).

All types of radiation and all their intersections lead to the methods of investigation (Table 1). Hence it follows logically that all the methods in their principles are now established.

Table 1. Principles of the methods of investigation: crosswise radiation-interaction systematics

Interaction Radiation

Nuclear X-ray Optical Infrared Superhigh Radio electron Microwave frequency frequency

Spectroscopy + + + + + + Microscopy + + + + Diffractometry +

Third division (for the modes of interaction of radiation with matter): For spectroscopy: the interactions of radiation with matter giving rise to

spectra are described for all the regions of spectroscopy (Table 2) by the energy levels schemes (or energy bands): nuclear levels (ground state, excited states, nuclear spin, nuclear quadrupole levels), electron levels (inner electron, valent electron, nonoccupied excited levels, electron spin levels), and vibrational levels.

These energy levels represent a fundamental characteristic of condensed matter. Splitting and shifting of levels are related to the pecularities of chemical composition and crystal structure of the material or mineral.

All kinds of spectra are described as the result of transitions between corresponding energy levels. Within any region of spectroscopy the interpretations for all the techniques of the spectroscopy are based on the same energy levels scheme, but on the transitions between different levels of the energy levels scheme. Emission, absorption, excitation, scattering, luminescence spectra (and in X-ray spec~roscopy Auger, EXAFS, XANES spectra) can be obtained and all are explained by the same energy levels scheme. Registration of electrons instead of photons leads to electron spectra (interpreted by the same energy levels scheme).

There are two kinds of spectroscopy, depending on the state of an investigated specimen: purely analytical spectroscopy, when matter is dispersed into atoms (atomic energy levels schemes), and solid state spectroscopy (energy levels of atoms in a crystal).


Analytical spectroscopy is further divided, depending on the region of spectroscopy (nuclear, X-ray, UV and visible), and on the mode of excitation of the spectrum (flame, cathode lamp, plasma).

Solid state spectroscopy is divided into Nuclear Gamma Resonance (Mossbauer) spectroscopy, X-ray, and Electron spectroscopy, Optical Absorption and Luminescence spectroscopy, Infrared and Raman spectroscopy, Electron Paramagnetic Resonance, Nuclear Magnetic and Nuclear Quadrupole resonance spectroscopy. They are based on the transitions between corresponding energy levels, and belong to the corresponding parts of spectrum.

For microscopy: the types of radiation (nuclear, X-ray, UV, visible, IR, acoustic) or particles and the modes of the contrasting interaction transformed into a visible image or into a photo- or instrumentally registered pattern determine the basis for the division of microscopy. The modifications are related to the different uses of the corresponding ray diagrams. A natural combination

1 nm 1 IJ. mm 0

10 1 10 2 10 3 10 4 10 5 10 6 10 7 10 8 10 9 A 0 0

A A 108 EYE 10 8

107 SEM 10 7

CD 10 6 HREM ,-, LM 10 6 r-

as ,--r 0 til 105 10 5 r- I I HM as 0 10 4 I I I 10 4 .,.. .j.J

I EMP I I I 1 L CD 10 3 I 10 3 > ~ II I

10 2 L -L=: --tli-- J I 10 2

o • 00 0 0 0 0 II I o PIXE 0

10 1 0 o 0 II I I 10 1

o STM AFM &~...:::..PCM~J-J .... .. -0 . . . . 00

0

0 10 1 10 2 10 3 10 4 10 5 10 6 10 7 10 8 10 9 A

Lateral scale

Fig. 1. Detection limits of microscopes. LM low power optical microscopy; H M high power optical microscopy; DIM differential interference microscopy; PCM phase contrast microscopy; SEM scanning electron microscopy; HREM high resolution electron microscopy; STM scanning tunneling microscopy; AFM atomic force microscopy; EPM electron microprobe; PIXE particle-induced X-ray emission (and other types of nuclear microscopy). Parts of the figure concerning optical and electron microscopy are presented after Cosslet and H. Komatzu, see Crystal Growth and Characterization edited by Veda and Mullin, 1975, NorthHolland, p. 333)

The Development of Instrumentation 5

of microscopy with analytical determinations is possible due to the detection of corresponding effects of the interactions. The limits of resolution are determined by the wavelengths of the radiations or particles. (Fig. 1).

For diffractometry: The different nature of radiation or particles (X-ray, neutron, electron) of the same region of the spectrum corresponding to wavelengths comparable with interatomic distances determines the types of methods for crystal and glass structure investigations. The modifications of the methods depend primarily on the geometry of the diffraction pattern.

While considering the most complicated and advanced techniques (Table 2), it is useful to keep in mind these simple divisions of the principles of methods and their crosswise radiation-interaction systematics to understand their essential contents and the distinguishing features of the modifications of the methods.

These discoveries of the principles of the interactions of radiation with matter, i.e., the principles of the methods, compose the first stage in the logics of the development of the methods. Realizing these principles and reaching their ultimate possibilities lead to next stage: the development of instrumentation between 1970 and 1990.

The Development of Instrumentation

Modern scientific instrumentation accumulates and concentrates series of the most sophisticated inventions and extraordinary achievements of material science, electronics, and computer techniques. It is a sublimate of ultimate capabilities for structural and analytical investigation of materials and minerals.

The thread of the development of all kinds of instrumentation (indispensable for the technical realization of the scientific principles for all the methods which are based on the interactions of radiation with matter i.e., spectroscopical, microscopical, structural) includes four elements: source of radiation, specimen, detector, and whether the radiation is operating before or after the interactions.

The revolutionary changes in each of these elements lead to the contemporary generation of the instrumentation.

A systematics of the most important technical development in the methods of investigations can include the following directions.

Advanced Radiation Sources. Radiation in different regions of the continuous electromagnetic spectrum is generated by the discontinuous series of the differing radiation sources which are based on completely different phenomena: nuclear reactors, radioisotopes, ~~cderators, synchrotrons, X-ray tubes, electron guns, lasers, bolometers, klystrons, etc. The specimen itself is in many techniques a source of different kinds of secondary radiations excited by accelerated ions, neutrons, electrons, protons, X-ray, lasers, plasma, flame, etc.

Important improvements have been made in all these kinds of radiation sources. Several of them have led to radical changes and broadening the


horizons in the methods: synchrotron radiation, positive ion micro beams, laser radiation.

Synchrotrons are sources of X-ray radiation of high intensity, broad spectral range, natural collimation, polarized, pulsed time structure.

In the 1980s, 12 storage rings used as sources of synchrotron radiation were available in the world: in the USA (Stoughton, Standford, Ithaca, Chicago, Brookhaven), France (Orsay), the USSR (Novosibirsk), Germany (Hamburg), Italy (Frascati), Great Britain (Daresbury), and Japan (Tsukuba).

Their use is restricted to special cases, but these cases concern the utmost limits of all X-ray methods and of the Earth, planetary, soil, environmental, material sciences (as well as biology and chemistry). Challenging programs have been presented by the Consortium for Advanced Radiation Sources, CARS, (The University of Chicago and Illinois University) and by the University of Paris VI and VII.l

Extraordinary possibilities in the use of synchrotron radiation in mineralogy can be estimated from its following applications:

- ultra-high pressure research of micron-sized specimens in the diamond anvil cell up to 360 GPa (the pressure at the center of the Earth) and in largevolume press; in situ high resolution X-ray diffraction and X-ray spectroscopy of the phase transitions at simultaneously high pressure and high temperature;

- studies of extremely small-sized specimens, low concentrations, at very short time intervals; atmospheric particulates (atmospheric dust), micrometeorites (less than 1 ng), aerosol depositions, deep-sea metal-rich precipitates;

- refinements of crystal structures from 2-50 m crystal; - high-resolution, high-energy, time-dependent, energy-dispersive X-ray dif-

fraction; - non-Bragg scattering from non-crystalline materials, indentification of poorly

crystallized minerals, differential and anomalous, small angle scattering from gels, glasses, melts;

- EXAFS of glasses, solutions, silicate gels, clay minerals; - synchrotron radiation-induced X-ray fluorescence microprobe (very low

concentrations, less than 1 ppm, in very small volumes, less than 1000 nm 3);

- analysis of trace elements, platinum group metals, rare earths, transuranic elements;

- surface diffraction, time-resolved X-ray absorption spectroscopy of metalorganic surface reactions;

- X-ray microradiography; X-ray white-beam and monochromatic topography; X-ray fluorescence micro tomography.

1 Thanks to J.V. Smith, Executive Director of CARS, for providing a copy of the Proposal to the Advanced Photon Source, March 15, 1991, and to G. Calas for discussions of the synchrotron applications in mineralogy [see also Phys. Chern. Min. (1984) 11: 17-36].

The Development of Instrumentation 7

Positive ion microbeams sources (protons, deutrons, helium-3 or -4, ions heavier than 15N) from electrostatic Van de Graaf accelerators lead to the series of nuclear microscopy and microprobe methods (see Chap. 5.3)

The use of laser radiation sources transformed Raman spectroscopy and presented new possibilities for luminescence spectroscopy (see Chaps. 3.7 and 3.4)

Another of the most important improvements in instrumentation is related to obtaining focused radiation down to the micrometer level. This led to the construction of microprobes for different regions of the spectrum.

There are two generations of microprobes: (1) electron microprobe, (2) particle-induced X-ray emission microprobe (PIXE), nuclear microprobe, ion microprobe, synchrotron radiation-induced X-ray microprobe, as well as laser microprobe, Auger microprobe, acoustic microprobe. Different kinds of optics are used for focusing radiation (pinhole collimator, multilayer mirror, bent crystals, magnetic quadrupole, electrostatic lenses, etc.).

Radical changes in instrumentation are connected also with the use of solid state detectors: Si(Li) and high-purity Ge, with fast Fourier-transform techniques (infrared, NMR spectroscopy, microscopy etc.), energy-dispersive registration, pulse methods.

Contemporary generations of spectrometers, microscopes, diffractometers, and microprobes represent extremely sophisticated versatile systems containing as indispensable components powerful computers and series of associated microprocessors, with hardware on-line with the instrumentation and the operation software providing hundreds of substantial functions: collection of signals, their conversion into spectra or into digital form, scan control, image enhancement and analysis, versatile data handling and output routines, data display and storage on magnetic media, automatic or interactive operation modes, programming, etc.

Only by using these techniques it is now possible to obtain the ultimate capabilities for each of the characteristics of the methods: ultra high sensitivity, extra resolving power, extremely high accuracy, and rapid multielement analysis.

00

Tab

le 2

, S

yste

mat

ics

of t

he m

etho

ds o

f in

vest

igat

ion,

Acr

onym

s of

the

met

hods

(j ::r

po

Nuc

lear

X

-ray

E

lect

ron

Opt

ical

M

W,I

R

RA

S

HF

,RF

"0

c:;

..,

Spe

ctro

scop

y :--

' C/

J '<

N

GR

N

IS

XA

S A

ES

AA

S P

LS

M

RS

R

S E

PR

=E

SR

~

MS

UC

NIS

P

ES

A

ES

XL

S (1

)

XE

S P

MR

R

aS

EN

DO

R

3 E

VM

S

BN

S X

FS

X

PS

IC

P-A

ES

C

LS

L

ER

R

ME

NR

po

c,

CE

MS

S

EN

S

EX

AF

S

UP

S

OA

S =

EA

S P

EL

S

IR

SRS

FM

R

..., on

GS

IN

AA

X

AN

ES

E

SCA

D

RS

T

L=

TL

S

NIR

A

SRS

MR

0 -,

GR

ES

F

NA

A

EIX

E=

IP

ES

=

SRS

TrL

F

IR

RR

S :;- (1

)

FSS

EN

AA

E

MP

A

BIS

IR

S N

SL

S

FT

IR

HR

S

NM

R

~

RB

S R

NA

A

CH

EX

E

EE

LS

PA

S S

LE

LS

S

WIR

E

RS

PM

R

(1)

ER

DA

P

GN

AA

P

IXE

A

RP

ES

S

HG

E

EE

S

MW

IR

CA

RS

BL

NM

R

:;- 0 N

RA

A

S S

RIX

E

SIP

S

NL

O

EM

S

TIR

M

AS

NM

R

0-

on

PIG

E

MS

NQ

R

0 -,

SE

NQ

R

S <

Mic

rosc

opy

(1) ~

FT

AR

S

EM

X

RM

S

EM

S

TM

L

M

IRO

ciQ

' a A

AR

C

CM

X

RT

T

EM

A

FM

U

VM

0

' N

AIA

AR

H

EIM

BM

S

TE

M

IRM

::s 0

BA

R

ST

IM

HR

EM

M

BI

-,

NA

IBA

R

IMB

T

AE

M

DIM

~

5'

RIM

H

CP

M

HV

EM

P

CM

(1

) .., SA

M

HR

AM

eo.

on

Dif

frac

tom

etry

....,

::

r N

S

XR

D

ED

M

BE

D

(l>

0 N

D

SC

XR

D

HE

ED

C

BE

D

(l>

<

PD

L

EE

D

HC

BE

D

(l>

0-S

XR

D

RE

D

LA

CB

ED

"0

SAX

S T

ED

5M

BE

D

3 (l>

SA

ED

R

MB

ED

a

HR

ED

O

TE

D

0 ....,

AL

CH

EM

I 5'

~

Mic

ropr

obes

.... c 3

NM

P

PIX

E

LM

P

RM

P

(l>

a IM

P

EP

MA

P

AL

P

~

SA

HM

P

o· ::l

'Cl


Nuclear Spectroscopy NGR MS EVMS CEMC GS GRES FSS RBS ERDA NRA PIGE NIS UCNS BNS SENS INAA FNAA ENAA RNA A PGNAA AS MS TIMS GSMS SSMS ICP-MS

Nuclear Gamma Resonance Mossbauer Spectroscopy Emission Variant of Mossbauer Spectroscopy Conversion Electron Mossbauer Spectroscopy Gama Spectroscopy Gamma-Ray Emission Spectroscopy Forward Scattered Particles Spectroscopy Rutherford Backscattering Spectroscopy Elastic Recoil Detection Analysis Nuclear Reaction Analysis Proton Induced Gamma-ray Emission Neutron Inelastic Scattering Ultra-Cold Neutron Inelastic Scattering Backscattering Neutron Spectroscopy Spin-Echo Neutron Spectroscopy Instrumental Neutron Activation Analysis Fast Neutron Activation Analysis Epithermal Neutron Activation Analysis Radiochemical Neutron Activation Analysis Prompt Gamma Neutron Activation Analysis Aplha Spectrometry Mass Spectrometry Thermal Ionisation Mass Spectrometry Gas Source Mass Spectrometry Spark Source Mass Spectrometry Inductively Coupled Plasma Mass Spectrometry

X-ray Spectroscopy XAS X-ray Absorption Spectroscopy XES X-ray Emission Spectroscopy XFS X-ray Flourescence Spectroscopy WDXFS Wavelength Dispersive X-ray Flouorescence Spectroscopy EDXFS Energy Dispersive X-ray Fluorescence Spectroscopy EXAFS Extended X-ray Absorption Fine Structure XANES X-ray Absorption Near Edge Structure EIXE Electron Induced Emission Spectroscopy EMPA Electron MicroProbe Analysis CHEXE Channeling-Enhanced X-ray Emission PIXE Proton Induced X-ray Emission SRIXE Synchrotron Induced X-ray Emission

Electron Spectroscopy AES Auger Electron Spectroscopy PES PhotoElectron Spectroscopy XPS X-ray Photoelectron Spectroscopy UPS Ultraviolet Photoelectron Spectroscopy ESCA Electron Spectroscopy for Chemical Analysis IPES Inverse Photoelectron Spectroscopy BIS Bremstrahlung Isochromat Spectroscopy EELS Electron Energy Loss Spectroscopy ARPES Angle-Resolved Photoelectron Spectroscopy SIPS Sputter Induced Photoelectron Spectroscopy

Optical Spectroscopy AAS Atomic Absorption Spectroscopy AES Atomic Emission Spectroscopy ICP-AES Inductively Coupled Plasma-Atomic Emission Spectroscopy


OAS EAS DRS SRS IRS PAS SHG NLO

Luminescence PLS XLS CLS PELS TLTLS TrL NSLS SLELS EEES EMS

Optical Absorption Spectroscopy Electron Absorption Spectroscopy Diffuse Reflectance Spectroscopy Specular Reflectance Spectroscopy Internal Reflectance Spectroscopy PhotoAcoustic Spectroscopy Second Harmonic Generation Non-Linear Optics

Photo Luminescence Spectroscopy X-ray Luminescence Spectroscopy CathodoLuminescence Spectroscopy Proton Excited Luminescence Spectroscopy Thermally Stimulated Luminescence TriboLuminescence Nano-Second Luminescence Spectroscopy Selective Laser Excited Luminescence ExoElectron Emission Spectroscopy Ellipsometry and Modulated Spectroscopy

Microwave Spectroscopy MRS Microwave (Radar) Remote Sensing PMR Passive Microwave Radiometry

Infrared and Raman Spectroscopy IR InfraRed Spectroscopy NIR Near InfraRed Spectroscopy FIR Far InfraRed Spectroscopy FTIR Fourier Transform InfraRed Spectroscopy SWIR Short Wave InfraRed (in Remote Sensing) MWIR Middle Wave InfraRed TIR Thermal InfraRed RS Raman Spectroscopy RaS Rayleigh Spectroscopy LERS Laser Excited Raman Spectroscopy SRS Spontaneous (Stokes) Raman Scattering ASRS Anti-Stokes Raman Scattering RRS Resonance Raman Spectroscopy HRS Hyper Raman Spectroscopy ERS Electronic Raman Spectroscopy CARS Coherent Anti-Stokes Raman Scattering

SuperHigh Frequency Spectroscopy EPR' ESR Electron Paramagnetic (Spin) Resonance ENDOR Electron-Nuclear Double Resonance MENR Multiple Electron Nuclear Resonance FMR Ferro-Magnetic Resonance MR Muon Resonance FTEPR Fourier Transform EPR PEPR Pulse EPR TDEPR Two-Dimensional EPR VHFEPR Very-High Frequency EPR PENDOR Pulse ENDOR EPRI EPR Imaging

Radio Frequency Spectroscopy NMR Nuclear Magnetic Resonance PMR Proton Magnetic Resonance BL NMR Broad Line Nuclear Magnetic Resonance

11

12

MASS NMR NQR SE NQR

Chapter 1. Systematics of the Methods of Investigation of Minerals

Magic Angle Spinning Sample NMR Nuclear Quadrupole Resonance Spin-Echo NQR

Nuclear Microscopy FTAR Fission Track AutoRadiography AAR Alpha AutoRadiography NAIAAR Neutron Activation Induced AAR BAR Beta Autoradiography NAIBAR Neutron Activation Induced BAR RIM Radioactive Ion Microscopy SEM Secondary Electrons Microscopy CCM Channeling Contrast HEIMBM High Energy Ion Micro Beams Microscopy STIM Scanning Transmitted Ions Microscopy IMBT Ion MicroBeam Tomography HCPM Heavy Charged Particle Microscopy

X-Ray Microscopy XRM X-Ray Microscopy XRT X-Ray Topography

Electron Microscopy SEM Scanning Electron Microscopy TEM Transmission Electron Microscopy STEM Scanning Transmission Electron Microscopy HREM High Resolution Electron Microscopy AES Analytical Electron Microscopy HVEM High-Voltage Electron Microscopy SAM Scanning Auger Microscopy STM Scanning Tunneling Microscopy AFM Automic Force Microscopy

Optical Microscopy LM Light Microscopy UVM Ultra-Violet Microscopy IRM InfraRed Microscopy

Acoustic Microscopy HRAM High Resolution Acoustic Microscopy

Microwave Imaging IRO Imaging Radar Observation

Diffractometry ND NS XRD SCXRD PD SXRD ED HEED LEED RED TED HRED SAED MBED CBED

Neutron Diffraction Neutron Scattering X-ray Diffraction Single-Crystal XRD Powder (X-ray) Diffraction Synchrotron X-ray Scattering Electron Diffraction High-Energy Electron Diffraction Low-Energy Electron Diffration Reflection ED Transmission Ed High-Resolution RD Selected Area ED MicroBeam ED Convergent Beam ED


HCBED LACBED 5MBED RMBED OTED ALCHEMI

Microprobes PIXE NMP IMP EPMA SAHRMP LMP PALP

Hollow Cone Beam ED Large Angle Convergent Beam ED Scanning MicroBeam ED Rocking MicroBeam ED Oblique Texture ED Atom Location by Channeling Electron Microscopy

Particle Induced X-ray Emission Nuclear MicroProbe Ion MicroProbe Electron Probe Microanalysis Scanning Auger High-Resolution MicroProbe Laser MicroProbe Pulse Acoustic Laser Probe

13

CHAPTER 2

Diffraction Methods and Crystal Structure Analysis

16 Chapter 2. Diffraction Methods and Crystal Structure Analysis

2.1 Crystal Structure Analysis and X-ray Diffraction Instrumentation

A. GUINIER, TH. HAHN, and V.1. SIMONOV

Development of Crystal Structure Analysis

A crystal with its strictly periodic atomic structure represents a natural, very symmetrical three-dimensional diffraction grating for wavelengths of the order of the interatomic distances. Crystal structure analysis is thus based on the theories of symmetry (space groups) on the one hand and of interaction of radiation with solids (diffraction) on the other. The diffraction techniques, with the help of extensive computer calculations, lead to the atomic arrangements of crystalline materials. Depending on the problem, X-rays, electrons, or neutJ:ons are used, which provide the electron-density, electrostatic potential density and nuclear density (also magnetic spin density) distribution, respectively, in a crystal.

All possible arrangements of atoms in crystals are governed by the 230 types of space groups, derived in 1890 by the crystallographer E.S. Fedorov and the mathematician A. Schoenflies. X-ray diffraction of crystals was discovered by the physicists M. Von Laue, W. Friedrich, and P. Knipping in 1912. In 1913-1914, W.H. Bragg and W.L. Bragg first applied X-ray diffraction to the experimental confirmation of the structures of single crystals of NaCl, Cu, diamond, etc. previously predicted by W. Barlow. The application of X-ray diffraction was extended to polycrystalline materials in 1916 by P. Debye and P. Sherrer. Details of the history of crystallography can be found in a recent Historical Atlas by 1. Lima-de-Faria.

The power of a structure analysis of a synthetic crystal or a mineral and the reliability and accuracy of the results depend on many factors: sample quality, radiation source, apparatus and techniques available, especially for the measurement of the diffracted intensities. The first problem of a structural study is the determination of the symmetry and the lattice parameters of the crystal. The next step is the derivation of an atomic model. The structure is then refined from the diffraction data, taking into account the finer effects of the interaction between sample and radiation. The final stage is the crystallochemical analysis which comprises interpretation of the geometrical model, calculation of interatomic distances, valency angles, sizes, and orientations of thermal-motion ellipsoids and, finally, generation of visible structure models on a display or a plotter.

Further advances of the theory and the methods of structure analysis provide information not only on the geometry but also on more subtle features of the mineral structure, such as isomorphous replacement, disorder, variable atomic occupation of crystallographic sites, sizes and misorientation of mosaic blocks in single crystals, as well as thermal motion of the atoms including its anisotropy

2.1 Crystal Structure Analysis and X-Ray Diffraction Instrumentation 17

and deviations from the harmonic approximation. The anharmonic components of the thermal motion in crystals are related to phase transitions which involve changes of many crystal properties. Investigations of the spatial distribution of valency electrons in crystals are highlights of modern structural analysis: the socalled deformation electron density maps obtained from very accurate X-ray diffraction data yield insights into the nature of chemical bonds in crystals.

Modern structural databases playa major role in unraveling the relations of the crystal structure studied with other structures. Such databases contain literature references, chemical composition, symmetry, lattice parameters, structural models (atomic coordinates), and other crystallographic information. The best known are the following four databases: metals and alloys (Canada), inorganic materials (Germany), organic compounds (England), and protein macromolecules (USA). Such databases with automated search and retrieval procedures are essential for profound studies in the field of structural crystallography. For the phase analysis of poly crystalline samples and the interpretation of powder diagrams, the JCPDS powder data file (USA) is indispensable.

Scope of Diffraction Methods

Modern diffraction methods comprise a large variety of applications in different fields of science and technology, in addition to the "classical" and well-known determination of crystal structures by means of single crystals. Some of these applications are briefly characterized as follows:

1. In practice, by far the most widely used diffraction technique is the X-ray powder method for the identification of polycrystalline phases, and the powder diffractometer is the most common diffraction apparatus in industry, research laboratories, and teaching. In mineralogy, materials science, and materials engineering, the powder pattern is used as a "finger-print" of a solid. The widespread application of the powder method for qualitative and quantitative phase analysis is mainly due to the existence and steady growth of the JCPDS powder data file (PDF), which started out with individual file cards but now exists also in the form of a computer databank, CD-ROM's, etc., with very sophisticated search strategies. 2. High-resolution powder diffractometers for X-rays, neutrons, and X-ray synchrotron radiation provide increasing capabilities of detecting modulated phases, "weak"phase transitions, and pseudo symmetric phases, and of tracing rapid solid-state reactions in real time (time-resolved diffractometry).

The Rietveld method permits the refinement of crystal structures from highquality powder data. This and new approaches to ab-initio structure determination have opened new paths to the structural characterization of polycrystalline phases for which no single crystals are available. The recent remarkable success of powder-diffraction studies of high-Tc superconductors bears witness to the power of this method.


3. The most prominent application of single-crystal diffraction is the determination of crystal structures, mentioned in the previous section. At present, more than 100000 well-determined crystal structures are known with an increase of about 8% per year. This flood of data can only be handled by computer databanks and electronic retrieval systems. 4. Beyond the "static" crystal structure, inelastic neutron scattering (and recently also X-ray scattering) provides information about the dynamics of a crystal: phonon dispersion, soft modes and phase transitions, dynamic disorder, excitations in incommensurate phases (phasons and amplitudons), magnons in magnetically ordered structures, but also time-dependent processes during solidstate reactions. The method requires fairly large single crystals (several mm3) and involves a high degree of theoretical interpretation (lattice dynamics). 5. Less known as a diffraction method is X-ray and neutron topography, again requiring large single crystals. This method yields information about the perfection and the growth history of crystals and makes directly visible lattice deformations, dislocations, growth striations, growth-sector boundaries, twinand domain boundaries, and other kinds of defects.

Radiation Sources

Standard equipment for X-ray diffraction experiments in the laboratory are high-voltage generators and sealed X-ray tubes, operating at a rating of 800 to 2000 W. Usually employed characteristic wavelengths (Kcx) range from silver (0.55 A) to iron (1.94 A); p-filters or crystal monochromators are used to reduce or eliminate the white-radiation background.

The intensity of the radiation source is important since it reduces data collection time and enables investigations of small-sized crystals. The characteristric linear size of single cystals are tenths of a millimeter for a sealed X-ray tube. In this case the experimental data will be collected within several days. For X-ray tubes with rotating anodes (rating from 6 to 20 kW) the radiation intensity exceeds that of a sealed X-ray tube by an order of magnitude.

The application of X-ray synchrotron radiation provides qualitatively new possibilities for structural studies of crystalline materials. Storage rings of various types and sizes were developed for investigations in nuclear and fundamental particle physics. Synchrotrons, i.e., cyclic resonance electron accelerators, initially were not intended for application as X-ray sources. On the contrary, this radiation was considered as an unwanted but inevitable effect, causing severe energy losses due to electron acceleration. It is only in recent times that the synchrotron sources in use have been modified and new ones with powerful X-ray beams have been developed. Synchrotron sources yield polarized X-ray radiation whose intensity exceeds that of the K-lines of a sealed X-ray tube by several orders of magnitude. Thus, a synchrotron source permits a full structural study of a single crystal of only microns in linear size.


The intensity is an essential but not the only specific feature of a synchrotron X-ray source. In contrast to the characteristic radiation of X-ray tubes, a synchrotron source provides a continuous spectrum of radiation. With special monochromator systems the desired wavelength can be selected. If a mineral contains impurities, only synchrotron radiation enables the investigator to find and locate them correctly in the crystal structure, using anomalous scattering and a suitable wavelength for each element. The efficiency of this approach was brilliantly demonstrated in protein crystallography when one iron atom was located in a molecule which contained several thousand atoms of other elements.

Diffractometers and Detectors

An important part of any diffraction apparatus is the detector. For X-rays, either an inexpensive but reliable two-dimensional detector, i.e., an X-ray film, or various types of proportional, scintillation and semiconducting counters are used. A film is indispensable for the preliminary examination of a full diffraction pattern from a new crystal, but its sensitivity and accuracy of intensity measurements is much below that of any counter tube. The efficiency of data collection becomes much higher when one- and especially two-dimensional position-sensitive detectors are used. Two-dimensional detectors are most suitable when not only Bragg reflections but also satellite reflections or diffuse scattering are to be detected: complicated twinning laws in minerals, the presence of modulated structures, studies under various pressures and temperatures, often accompanied by phase transitions, can be carried out with higher efficiency using two-dimensional detectors.

A modern automated single-crystal diffractometer operates under control of a computer. Such a system not only performs routine analysis but also provides for immediate feedback during the measurement of each reflection. With position-sensitive detectors, this feedback is not possible because many reflections are measured simultaneously. For this reason area detectors are normally not used for precision structure determinations.

Present computer-controlled diffractometers provide extremely accurate intensity data from single crystals. The statistical error in the intensities can be as low as 1-2%. However, two circumstances should be taken into account: decreasing the statistical error by an order of magnitude increases the datacollection time by two orders of magnitude, other conditions being equal. Sometimes, the real accuracy limit is due to the quality of the sample and not due to the instruments and methods employed.

Intensities of Diffraction Reflections and Structure Amplitudes

Integrated intensities of diffraction reflections are collected during an X-ray experiment. The measured intensities are influenced not only by the atomic


structure ofthe sample, but also by many other factors due to the data collection scheme, character of radiation, shape, size and perfection of the sample, its temperature, and some other effects of the interaction between radiation and sample. All these factors should be taken into account when the integrated intensities Ihkl are converted to moduli of structure amplitudes Flhkll that bear information about the atomic structure of the single crystal studied.

The Lorentz and polarization factors allow for the geometry of the diffraction pattern obtained and the character of polarization of the primary X-ray beam. Of special importance is a correct account of radiation absorption by the sample. Absorption correction is essentially simplified if the sample can be machined into a sphere. However, mechanical properties of many crystals as well as small sizes of minerals often prevent this. There are several algorithms and computer programs that ensure absorption correction in samples of a random shape. If the crystal is a polyhedron, its shape can be described by crystallographic indices of the faces and the distances between these faces and a fixed point inside the crystal. These characteristics and the matrix of crystal orientation in the diffractometer are sufficient for the calculation of the pathways of the primary and diffracted beams for all reflections.

In most structural studies the effects of simultaneous reflections are ignored. When measuring the intensity of a particular reflection, one or more further crystallographic planes may simultaneously be in reflection position. This gives rise to a mutual energy transfer which is proportional to the intensity of the reflections involved and depends on the phase relationships between them. The result usually is a loss of intensity by the stronger and gain by the weaker reflections. The detailed account of the influence of diffraction beams on the intensities of simultaneous reflections is rather complicated because it depends on the spectrum of the radiation, primary beam divergence, mosaic spread of the sample and other factors. Simultaneous reflections are avoided experimentally by "l/I-scans", i.e., by rotating the sample about the normal of the reflecting plane, the diffraction condition being retained.

Usually, the kinematic theory of interaction between radiation and sample is used in structural studies. If high-accuracy data are desired, dynamical effects of the interaction should be taken into account, especially for highly perfect crystals. Several algorithms of corrections for extinction have been developed and implemented. All these methods require the knowledge of the crystal structure. For this reason, corrections for extinction are introduced not during the initial data-reduction process but at the final stages of the least-square refinement of the structure.

In the course of precision structural studies, in order to obtain accurate values of structure factors, it is necessary to estimate and subtract from the intensities measured not only the usual background but also the so-called diffuse scattering. If the components of the elasticity tensor or an array of rates of ultrasound propagation in the single crystal in anisotropic approximation are known, these data allow one to calculate the contribution of thermal diffuse scattering to each Bragg reflection. In recent times, algorithms for empirical estimates of thermal diffuse scattering from the profile analysis of diffraction


reflections and the behavior of background scattering near each Bragg reflection are being developed.

The best way to avoid the influence of thermal diffuse scattering is to make low-temperature measurements. The use of liquid nitrogen temperature lowers significantly thermal diffuse scattering, while measurements at liquid helium temperatures finally solve the problem. Low temperatures, decreasing the amplitudes of atomic oscillations, reduce the influence of the thermal motion on the intensities of the Bragg reflections, too. In this case much better data arrays are obtained including reflections from crystallogniphic planes with small interplanar distances. This is most important in detailed investigations of the atomic thermal motion parameters as such.

At present, very convenient and economic, closed helium systems have come into use. The systems with one vaporation cycle provide temperature to 20 K, and with two cycles to 10 K. This is quite sufficient for solving most precision problems. The use of low temperatures sometimes makes the use of a shorter wavelength efficient. This makes the data array even better. The only difficulty encountered in low-temperature measurements is a phase transition in the sample. If the investigator is concerned with the atomic structure of a hightemperature phase, he has to work in the temperature range where this phase exists.

The Phase Problem of Crystallography

From the hundreds or even thousands of measured intensities Ihkl of a crystal, a set of observed structure amplitudes IFhk11 is derived, which forms the experimental basis for the determination of the crystal structure. Only the modulus IF hkll of the complex structure factor F hkl = IF hkll exp 2niochkl can be measured, the phase OChkl being inaccessible to experiment. This is the famous "phase problem of crystallography".

In recent years a good deal of attention has been focused on methods for the experimental determination of the phases. The effect of simultaneous reflections in a crystal, described above, makes it possible, in special, favorable cases, to determine experimentally the phases of some structure factors. However, the measurements are very complicated because the sample and the instruments have to meet high requirements. The methods are not yet developed enough to serve as an effective and routine tool for the structure determination of minerals.

The structure factors F hkl are the coefficients of a Fourier series representing the scattering density in a unit cell of the crystal. Hence, knowledge of the phases OChkl would reveal the structure by means of a routine Fourier summation.

The Patterson Method

At present there exist two main approaches to the solution of crystal structures by means of diffraction data. Historically, the first method is the "Patterson


function", suggested in 1934 by A.L. Patterson. The Patterson function is calculated by a Fourier series with the squares of the structure amplitudes, IF hkJl2, as coefficients; thus no knowledge of the phases is required. The Patterson function displays all the interatomic vectors of the structure, originating from a common origin. It is the self-convolution (pair-correlation function) of the crystal structure.

The most simple technique of the application of the Patterson function refers to crystals whose structure contains one or few much heavier atoms. In this case, which is ideal for the heavy atom technique, the Patterson function displays peaks corresponding to interatomic vectors between the heavy atoms. Then come maxima at the ends of the vectors relating the heavy atoms with the light ones. A great number of maxima due to the distances between the light atoms constitute the background of the Patterson function and will not prevent its interpretation.

Superposition techniques are the most effective tool for the interpretation of the Patterson function for a wide range of crystal structures. The foundations of these techniques were laid by Dorothy Wrinch in 1938 and later developed by M.J. Buerger. They have also been automated and implemented on computers. They consist of the superposition with a parallel shift of two or more copies ofthe Patterson functions by specific vectors. The peaks that overlap on all copies correspond to atoms of the structure.

A further step in the development of modern superposition methods was made at the Institute of Crystallography, Russian Academy of Sciences. Instead ofthe interpretation of Buerger superposition functions that derive the structure from the Patterson distribution, it was suggested to use the Fourier transformation of these functions so as to calculate phases of structure amplitudes. Phases thus obtained are used together with the observed moduli of structure amplitudes to obtain the next approximation to the electron density distribution in the crystal. There is a still more effective and less time-consuming technique of locating peaks of the function of deriving the structure from the Patterson distribution or from any other approximate electron density distribution and the use of the coordinates and heights of these peaks for the calculation of the values and phases of the structure amplitudes from the conventional formulae used for atomic models of crystals. The calculated phases and observed moduli of structure amplitudes yield the next approximation to the electron density distribution in a crystal. The modification of such a distribution on the basis of the a priori knowledge of the fact that the crystal structure is built up of atoms which diffract X-rays according to fz(sin ()/J..) permits one to refine an approximate electron density distribution on a computer. This cyclic procedure, in fact, provides the refinement of the phases of structure amplitudes. It is effective when applying not only Patterson methods but also direct methods of the determination of phases of structure amplitudes which will be reported below. Modification of the approximate electron-density distribution is made at each iteration; first of all, it consists of removal of background and, particularly, negative


regions that obviously have no physical significance and are due to errors in the phases.

The Patterson method is most effective when atoms of different weight occur in a crystal, whereas structures containing atoms of nearly equal weight are too complex for the application of the method. Superposition techniques allow the investigator to rely on his experience, knowledge of crystal chemistry, and intuition. This is very important in studies of crystals with rather complex structures due to pseudo symmetry or twinning. In the latter case the Patterson method often is the only effective technique for solving the structure. It is very difficult to interpret the diffraction patterns from crystals twinned by merohedry, where the reflections from twin domains overlap exactly. Very often problems arise in studies of modulated structures and incommensurate phases, where the derivation of the first model requires specific techniques. If a model requires specific techniques. If a model is known, it can be refined under the control of the investigator, but not automatically. The interatomic Patterson function permits also the checking and refinement of the space group of a crystal when other techniques fail. When the statistical analysis of the reflection intensities does not resolve the space group unambiguously, a detailed inspection of the Patterson distribution may help: rotation axes, mirror planes, and inversion centers can be determined from characteristic features of the peak distribution in the Patterson function (Harker sections).

The analysis of Laue symmetry and systematic absences permits only the determination of one of the 122 diffraction symbols. In order to establish the space group of the crystal, one must prove the presence or absence of an inversion center and determine the absolute configuration if the sample exhibits chirality. Both these problems can be solved by anomalous X-ray scattering of a suitable atom in the crystal.

Direct Methods

Whereas the Patterson method involves crystal space and atomic models, the "direct methods" for the solution of the phase problem, developed since 1950, operate in reciprocal space. Direct methods prove to be the most effective tool for structure determinations, especially the versions developed by J. Karle and H. Hauptman (Nobel prize winners in chemistry in 1985), as well as M. Woolfson. The probability relations for the direct determination of phases are based on several fundamental properties of crystals:

1. Non-negativity of the electron density. This is true for X-ray and electron scattering. For neutrons, the nuclear density for some isotopes of some atoms can be negative, i.e., neutron scattering may involve a 180 phase reversal.

2. The crystal consists of atoms which correspond to maxima of the electron, potential or nuclear density. For neutron diffraction, positive and negative extrema are possible.


3. The number of symmetry-independent reflections of a crystal exceeds by at It:ast one order of magnitude the number of parameters defining the atomic model of the structure. This implies that the symmetry-independent structure amplitudes and their phases are related.

Statistical relations are the basis of modern direct methods. The simplest probability relation refers to a triplet of amplitudes whose indices are related as follows: hi + h2 + h3 = kl + k2 + k3 = 11 + 12 + 13 = O. For such a triplet it is most probable that the sum of the three phases is equal to zero. The probability for this phase relation is the higher the larger the product of the normalized moduli of the structure amplitudes of these reflections.

At present, much more complex statistical relationships and more rigorous estimates of their probabilities are used. These calculations are rather lengthy, the algorithms are complicated and can be effectively realized only on powerful computers. There exist five to seven program systems, differing in the techniques employed, and dozens of versions of such programs that are used worldwide. Most structures containing up to 100 atoms in the asymmetric unit can be solved by these methods. For a complex structure the model derived from these relations usually contains some peaks which approximately correspond to correct locations of atoms, some false peaks (that do not correspond to atoms) and, as a rule, some atoms of the structure are missing.

In recent times, several new approaches towards direct methods have been made. The first one is the application of more complex probability relationships that are based on correlations not only between three structure amplitudes, but also involve amplitudes of the type: (hi - h2)(kl - k2)(11 -12);(hl - h3) (kl - k3)(11 -13);(h2 - h3)(k2 - k3)(12 -13), or even ofa larger neighborhood of the initial triplet of amplitudes. The algorithms are more complicated and the computations take more time, but the phases can be determined more effectively. Another direction involves a complicated analysis of the higher-rank determinants of the relations between the structure amplitudes. This approach is inte:resting for itself, but the practical results are not yet convincing. Sometimes phase restrictions due to general crystal-chemical regularities are employed, for instance, when the structure contains rigid atomic groups such as Si04

tetrahedra in silicates or other stable atomic fragments. Even simpler information can be used, e.g., on the lower limit of acceptable interatomic distances, va1cmce angles, planar groups of atoms, etc.

A different approach to the solution of the phase problem involves the idea of maximum entropy. This method has yielded good results and convincing determinations of particular structures. A further direction in the development of direct methods is connected with the analysis of the atomic structure of protein crystals.

Coming back to structural studies of minerals, the most efficient way is the combination of direct and Patterson methods to the solution of structures. The advantages of such a combination result from both, the high degree of formalization of the direct methods, and the more visual nature of the Patterson (superposition) approach. The a-priori structural information can be used more


easily and naturally during the Patterson stage. Such an approach results in a quite efficient and fully automated phase-refinement technique. This procedure was successfully applied to studies of crystals containing 200 or even more atoms in the unit cell.

The new method requires only a very approximate electron density synthesis, where only 30-40% of the strongest maxima correspond to atomic locations. The other maxima may be false. This limitation, however, will not prevent automated location of all atoms of the structure, except for hydrogen atoms.

A modern structural laboratory should have and use several of the available program systems for the determination of crystal structures. Crystallographic program systems are so complex that even the same technique implemented by different authors can have different efficiency. This is due to the inevitable use of different approximations in the formulation of the algorithms.

Refinement of the Structures

When the atomic model of the crystal under investigation has been derived using Patterson or direct methods, the same observed moduli of structure amplitudes are employed to refine the coordinates and parameters of atomic thermal motion. The refinement is performed by the least-squares technique, by way of minimization of the functional

N

<P(P} = L Wi (IF oli - IF clit , i= 1

where usually n = 2, m = 1 or 2, N is the number of independent IF ol,w i are weighting multipliers making the w;lF oli values equally accurate. When the data array and specific features of the crystal exclude a strong correlation between the refined parameters, this procedure can be implemented on a computer. If the computer is powerful, the full-matrix LSM variant is recommended.

A strong correlation between the refined parameters can hinder the refinement of the atomic model of the structure. Very often such a correlation occurs between the reduction coefficient of the data array IF 01 to the absolute scale, the mean parameter of the atomic thermal motion and isotropic extinction parameter. Special attention should be paid also to a correlation between occupancy coefficients and appropriate parameters of atomic thermal motion. Finally, between coordinates of the main atoms related by pseudo symmetry and showing disorder.

Whenever there is a strong correlation between the parameters, the LSM refinement of the structure can yield physically meaningless results. In such a case the method of step scan should be applied. When the parameters strongly correlate a number of values differing by a certain step A are assigned to one of them. All the other parameters of the structure are refined for each such value. A curve showing the dependence of minimum R-factors-on the scanned parameter is plotted from the results of the refinement. The plateau of the minimum values


of the R -factor indicates the confidence interval of the values, including the best value of the scanned parameter and other related parameters of the structure. The choice of solution inside the confidence interval cannot be made only from diffraction data. Other independent information is needed.

Depending on the problem to be solved, two different approaches to the refinement stage should be considered. Usually, the ultimate aim of a structural study is to obtain an atomic model with the accuracy of atomic location within 0.01-0.02 A. Here rather reliable estimates of atomic thermal parameters can be made. The situation changes when a precision structural study should be carried out and fine specific structural features including deformation electron density distribution should be found. The sample should be of high quality, intensities should be measured in the entire reciprocal space. Then, after an account for absorption has been made, upon the averaging of symmetry equivalent reflections the averaging R-factors is calculated. Its value illustrates the sample and experiment quality. Besides, a number of fine effects of the interaction between radiation and sample should be taken into account. Such effects comprise the above-mentioned diffuse scattering and simultaneous reflections.

In structure refinement of crystals characterized by a large number of refined parameters, there can be problems with the use of the most efficient full-matrix variant of the least squares technique. In this case the block-diagonal version should be applied. When such blocks are composed, one such block should contain all the parameters that can correlate with one another noticeably. Upon the transition according to the block-diagonal scheme the number of iterations required increases as compared to the full-matrix refinement.

As for chiral structures, the least-squares technique applied to enantiomorphous modifications of the atomic model allows one to establish the absolute configuration of the sample structure. In this case the anomalous scattering components of the X-ray radiation the structure atoms should be applied. The reciprocal space including the Bijvoet reflection pairs should be measured. These pairs should not be averaged in the primary interpretation ofthe data array, thus allowing one to establish which of the models - the right or the left one -corresponds to the sample studied.

At the stage of a least squares structure refinement it is possible to make allowance for extinction correction in the values of measured integral reflection intensities. It is difficult to predict which of the approaches suggested by different authors most exactly expresses the specific features of radiation scattered by the given sample. It is convenient to try all the available algorithms of correction for extinction and to choose the best one from the refinement results.

The important problem of the determination of the presence of isomorphous dopants and their distribution over crystallographic sites can be solved during the total refinement of structure parameters. The diffraction data contain information about the effectiveness of radiation scattered by atoms occupying a particular crystallographic site. This problem is difficult because the scattering intensity is affected not only by the composition of the isomorphic mixture of


atoms occupying the site but also by the character of thermal atomic motion and a possibility of atomic static displacements from the mean equilibrium sites within the structure. All these parameters strongly correlate with one another. The appropriate confidence intervals should be calculated to determine the parameters reliably, a control over the results should be made by the construction of the difference electron density distribution near such sites. The difference syntheses can point to a case when the account of thermal atomic motion in the harmonic approximation is insufficient.

The attempts to determine the anharmonic components of the thermal atomic motion from diffraction data can be made only when there are highly accurate experimental data and with a proper correction for absorption, anomalous scattering, extinction, and other factor of the interaction between radiation and sample. The atomic thermal motion can be described in three ways: by means of a temperature factor, the probability density function of atom location at a certain point of space in the process of thermal motion, and, finally, potential distribution near the equilibrium atomic position. Each of these forms of describing atomic thermal motion ensures allowance for anharmonicity. From the practical point of view, the Gram-Charlier expansion into probability density distribution over quasimoments is most convenient to use. Usually the programs used for such aims are confined to sixth-rank tensors. Only some of these tensors are meaningful. At the first refinement stages such meaningful components are established and later they are refined. Separate components, as such, cannot be interpreted. The physical sense belongs to the combined characteristics of atomic thermal motion that should be analyzed.

The highest achievement of modern precision structure studies are the works devoted to the establishment of spatial distribution of valence electrons in crystals. Unfortunately, not all the publications on this subject are reliable. Sometimes the authors misinterpret the data. The experimental accuracy and a, proper account of all the effects of the interaction between X-ray radiation and sample in such studies should be even higher than during the analysis of anharmonic components of atomic thermal motion.

In works of finding valence electron distributions it is always difficult to separate the influence of anisotropy of atomic thermal motion from asymmetry in the distribution of valence electrons on the total electron density. These effects strongly correlate with one other. The asymmetry in the distribution of valence electrons depends on the character of the chemical bonds in a crystal, while the direction and bond strength account for the total pattern of the atomic thermal motion. There are two methods for an objective separation of these effects.

The first one consists of the use of both X-ray and neutron diffraction data obtained from the sample. Neutron diffraction data contain information on the location and thermal motion of the nuclei of the atoms within the crystal. Proceeding from the observed neutron diffraction structure amplitudes, the investigator determines and refines the coordinates of atoms (nuclei) and parameters of their thermal motion with allowance for anisotropy and, if


necessary, anharmonicity of these oscillations. Then, X-ray diffraction data are used to calculate the valence electron distribution within the crystal. In this c~se it is assumed that the thermal oscillations of the nucleus of any atom coincide with those of its electron shells. Such a supposition is not always justified. The second difficulty encountered in a simultaneous use of X-ray and neutron diffraction data consists ofthe fact that it is practically impossible to obtain these data from the same single crystal. The point is that the sample volumes that are optimum for measuring the diffraction pattern with either radiation differ by four to five orders of magnitudes. This difference can be reduced by two to three orders of magnitude provided special high-flux reactors are used, but it cannot be completely avoided so far. A negative effect of the use of two different samples consists in the fact that their real structures can differ; hence, correction coefficients for absorption, extinction, and other effects differ greatly in the course of the primary processing of the same reflections. All these circumstances lead to a lower accuracy of the ultimate results.

The second way to solve this problem consists of an independent use of highangle reflections and a full X-ray data array. The important factor is that X-ray scattering from external valence electrons diminishes with an increase in the diffraction angle much faster than the scattering from more compact internal electron shells. The internal shells do not contain valence electrons, and electron density distribution in them is similar for free atoms as well as for the atoms involved in chemical bonds within the crystal. It is possible to estimate the scattering angle for each particular sample of a known chemical composition and certain X-ray radiation. The contribution of valence electrons with a required accuracy to the reflection intensities at higher values of the scattering angle is insignificant. This technique can be used provided the sample yields enough intense high-angle reflections that would be sufficient for the refinement of all the structure parameters of the crystal, including coordinate and thermal ones. If this condition is fulfilled, then, using the least squares technique, one can refine the location of all the atoms within the structure and their thermal motion. In this case all the complications due to a correlation between the refined parameters are not removed. New problems arise, too. A high-angle data array does not allow one to refine a correction for extinction and scale factor, which strongly correlates with it. Several procedures of successive refinement of structure parameters from a full and high-angle array with some parameters fixed and other parameters refined have been developed. The lower limit of the high angle array and the approach to the application of least squares technique should be set for any object of the studies. These problems can be solved only if the investigator is highly qualified and experienced. Just a routine use of the programs without taking into account fine specific features of the sample and the problem to be solved can lead to errors that distort not only the quantitative but also the qualitative pattern of valence electron distribution in crystals.

A vivid qualitative pattern of chemical bonds within a crystal can be obtained upon the construction of the threedimensional deformation electron density distribution in a crystal. Such density can be obtained if the electron density


distribution in free (noninteracting) atoms contained in the structure is extracted from the real pattern of spatial electron distribution. In this case, free spherically symmetrical atoms should be extracted exactly from the sites occupied by atoms within the crystal. The second important condition consists of the fact that the extracted atoms should be assigned the parameters of the thermal motion of the appropriate atoms of the structure including anisotropy and anharmonism of these oscillations. A correct determination of scale factors for the observed and calculated data arrays is most important for obtaining the difference electron density distribution. The regions of positive and negative electron density in the deformation synthesis indicate redistribution of valence electrons in a crystal taking place when free atoms are involved in a chemical bond, thus forming a crystal. The electron density in this case moves from the negative to the positive regions of the deformation synthesis, and covalent bonding is most vividly seen on such syntheses. Positive electron density peaks concentrate between the atoms involved in such a bond. Crystals with ionic bonding exhibit a different electron density distribution. Usually one can locate reliably the lone electron pairs of the appropriate atoms.

Structural Studies Under External Actions

Investigations of mineral behavior at great depths in the earths crust require the creation of the experimental conditions involving high pressures and various temperature regimes. At present the apparatus are available that enable the realization of such conditions. One example is diamond anvils. Being light and small, they ensure pressures up to a million bars that can be retained sufficiently long. If the investigator needs pressures of several dozens of kilo bars, the corresponding diamond anvil is so tiny that is can be mounted on a goniometer head of a standard X-ray diffractometer for studying single crystals. Well known is the law ofluminescence line shift of a laser ruby under the influence of pressure. A grain of a single crystalline ruby placed together with the sample inside the anvils is an internal manometer for measuring pressures in anvils. The high transparency of diamond for X-rays and neutrons makes the anvils an ideal instrument for structural studies under high pressures. However, the use of diamond anvils does not ensure the high accuracy that can be achieved without them. Rather large diamonds should be used for high pressures; therefore, correction for absorption and extinction in them is required.

The metallic frame of the anvils prevents the investigator from obtaining the whole array of diffraction reflections from the sample. Two diamonds of the anvils are single crystals themselves and yield their own reflections, that can hinder reflection measurements from the sample. Special control programs have been developed from automated diffractometers with diamond anvils mounted on them. These programs take into account specific features of the design and geometry of the anvils. Optimum operating conditions are chosen by such a


program for each reflection. An account of shadows of metallic anvil parts, absorption of the incident and diffracted beams in diamonds is made; besides, reflections from the diamonds themselves are not superimposed or the investigator is informed that such superposition is inevitable. At pressures exceeding several dozens of kilo bars, the distance between the diamonds is of the order of a 100 J.L therefore the sample should be prepared as a disc. Its thickness should be less than the distance between the diamonds, while the square should be large enough to provide strong diffraction reflections, capable of passing through the diamond and retaining the intensity measured.

Some special problems require structural studies of crystals placed into electric fields or into magnetic fields in the case of neutron diffraction studies. The appropriate devices developed for diffractometers enable such studies. Fiber optics devices have been created that allow one to obtain an X-ray diffraction pattern from a laser irradiated crystal. However, samples are often studied in a wide temperature interval. The importance of low-temperature studies has already been mentioned. They remove intensive thermal atomic oscillations and provide a more informative diffraction pattern. Sample cooling down to liquid nitrogen temperatures is not difficult. The helium temperature technique is more complicated, but it is also becoming frequently used in structural studies. The devices intended for high-temperature studies at temperatures up to 500° Care rather simple. The principle of the operation is based on sample blowing by a torrent of heated and dry nitrogen. The nitrogen torrent temperature can be monitored within the required range. Devices for still higher temperatures are not that simple. Certain parts of the device should be protected from the heat, a special sample holder is required as well as special glue for fixing the sample.

One of the most interesting problems that can be solved by a structural study at various temperatures and pressures is the atomic mechanism of phase transitions in crystals. Such investigations are important for the crystals that are widely used in various devices. Semiconductors, superionic conductors, ferrolectrics, materials with nonlinear optical characteristics, solid state lasers, and many other crystalline materials are employed in power plants, communication systems, computers, and electronic units. All these applications require the knowledge about changes in the atomic structure and physical properties of these crystals depending on temperatures, pressures, and other external actions.

Structural studies are of major importance in mineralogy, as they serve as the basis for a rigorous classification, and genetic relations between natural crystalline objects can be established. Reliable data on the occurrence of rare and scattered elements, and details of isomorphous replacements in minerals also can be obtained from precision structure studies. The foundation of structural mineralogy, laid by W.L. Bragg, was developed by scientists in many countries. An original contribution to the crystal chemistry of silicates with large cations has been made by N.V. Belov. It is in mineralogy that diffraction techniques first came to be used for structural studies. Later, methods of structural crystallography were also employed in chemistry and solid state physics. Establishment of correlations between chemical composition, atomic structure, and physical


properties of crystalline materials open up prospects for a search of new crystals, and modifying known materials at the atomic level.

References

Albinati A, Willis BTM (1982) The Rietveld Method in neutron and x-ray powder diffraction. J Appl Cryst 15: 361-374

Andrianova ME, Kheiker DM, Popov AN et al (1982) A coordinate X-ray diffractometer based on a two-dimensional proportional chamber and a two-circle goniometer. J Appl Cryst 15: 626-631

Belov NV (1976) Reports on Structural mineralogy. Nedra, Moscow Bish DL, Post JE (eds) (1989) Modern powder diffraction. Reviews in mineralogy, Vol. 20.

Ribbe, P.H. Ser ed, Mineralogical Society of America, Washington, D.C. Blessing RH (1987) Data reduction and error analysis for accurate single crystal difraction

intensities. Crystal Rev 1: 3-58 Bragg WL (1937) Atomic structures of minerals. Cornel University Press, Ithaka Buerger MJ (1959) Vector space and its application in crystal-structure investigation. Wiley,

New York Catlow CRA (ed) (1986) High resolution powder diffraction. Mater Sci Forum 9: 1-l34 Chang S-L (1984) Multiple diffraction of X-rays in crystals. Springer, Berlin Heidelberg New

York Coppens P (1982) Concepts of charge density analysis: the experimental approach. In: Coppens

P, Hall MB (eds) Electron distributions and the chemical bond. Plenum Press, New York pp 61-92

Dachs H (ed) (1978) Neutron physics. Topics in current physics, vol 6, Springer, Berlin Heidelberg New York

Glusker JP, Patterson BNK, Rossi M (1987) Patterson and Pattersons. Fifty years of the Patterson function. Oxford University Press, Oxford

JCPDS - International Centre for Diffraction Data: methods and practices in X-ray powder diffraction (a regularly updated loose-leaf binder) 1601 Park Lane, Swarthmore, Pennsylvania USA

International Union of Crystallography (ed) (1987) Crystallographic databases. Information content, software systems, scientific applications. International Union of Crystallography, Chester, UK

Klug HP, Alexander LE (1974) X-ray diffraction procedures for polycrystalline and amorphous materials. Wiley, New York

Ladd MFC, Palmer RA (eds) (1980) Theory and practice of direct methods in crystallography. Plenum Press, New York

Giacovazzo C (1980) Direct methods in crystallography. Academic Press, London Lima-de-Faria J (ed) (1990) Historical atlas of crystallography. Kluwer Academic Publ,

Dordrecht, Holland Muradyan LA, Radaev SF, Simonov VI (1989) Precisional structural studies of crystals and

correlations of refined parameters In:Vainshtein BK (ed) Methods of structural analysis. Nauka, Moscow pp 5-6

Ramaseshan S Abrahams SC (ed) (1975) Anomalous scattering Sabine TM (1988) A reconciliation of extinction theories. Acta Cryst A44: 368-373 Schwarzenbach D Abrahams SC, Flack HD et al (1989) Statistical descriptors in crystallogra-

phy. Report of the International Union of Crystallography subcommittee on statistical descriptions. Acta Cryst A45: 63-75

Simonov VI (1982) Automatic Patterson superposition and minimum function methods. In Sayre (ed) Computational Crystallography. Clarendon Press, Oxford pp 150-158

Simonov VI (1987) Phase refinement by density modifications. In Schenk H,Wilson H (ed) Direct methods, macromolecular crystallography and crystallographic statistics. World scientific India, pp 349-359


Tanner BK (1976) X-ray diffraction topography. Pergamon Press, Oxford Tanner BK, Bowen DK (eds) (1980) Characterization of crystal growth defects by X-ray

methods. Plenum Press, New York Vainshtein BK (1981) Modern crystallography vol 1: Symmetry of crystals. Methods of

structural crystallography. Springer, Berlin Heidelberg New York Willis BTM, Pryor A W (1975) Thermal vibrations in crystallography, Cambridge Univ Press,

Cambridge

2.2 X-Ray Diffraction Techniques for the Characterization of Minerals

S. ALTANER and I. E. KAMENTSEV

Mineral investigations commonly use modern X-ray diffraction techniques to determine crystal symmetry (point and space group), unit cell parameters, atomic coordinates, solid solutions, distribution of atoms in crystallographic sites, domain structure, mixed-layer structures, and polytypes. The data provide important information on the crystal structure, crystal chemistry, and origin of minerals. In this chapter we will discuss methods of X-ray diffraction characterization of minerals and results for some of the important minerals and mineral groups.

X-ray single crystal analysis is used to obtain information on the structure of minerals; however, it is a very tedious process and is appropriate only for defectfree crystals. X-ray powder diffraction (using a Debye-Scherrer camera, Guinier camera, or powder diffractometer) is widely used for analyzing the phase composition of a sample and its crystal-chemical features. The method allows fine-grained and poorly crystallized samples to be studied. A Gandolfi camera is used to obtain a powder pattern from a small single crystal (down to about 30 fl), which is rotated simultaneously about two axes. A resultant powder diffraction pattern (a set of peaks related to interatomic distances or d-spacings and peak intensities) is compared with reference data for minerals. The most complete source of diffraction data is the Powder Diffraction File (PDF), which contains data for more than 50000 substances, including minerals. The database includes experimentally obtained patterns and patterns calculated on the basis of a known crystal structure.

Phase identification involves comparison of a set of experimental d-spacings and intensity values with those in a database. Comparison is often done using a computer program. New minerals are typically characterized at a minimum by a combination of optical data, crystal symmetry, unit cell parameters, calculated and observed densities, chemical composition, the number of atoms in the unit cell, and the X-ray powder diffraction pattern (including values of d-spacing, intensity, and Miller index).

Quantitative phase analysis is based on the correlation between the intensity of diffraction peaks for a phase and the abundance of the phase in a sample.

2.2 X-Ray Diffraction Techniques for the Characterization of Minerals 33

Investigations are carried out by a variety of procedures including the internal standard method (i.e., the reference intensity ratio or Chung method), the absorption-diffraction method, the method of standard additions, and the Rietveld method (Snyder and Bish 1989). The above methods have been used to investigate a wide variety of problems including the abundance of quartz in industry dust, rocks, and ceramics; iron oxides in iron-bearing ores; titanium oxides in rutile-bearing samples; gibbsite, boehmite, and diaspore in bauxite deposits; chalcopyrite in sulfide ores, and many others (Snyder and Bush 1989).

Clay minerals, which are usually extremely fine-grained and platey, are often analyzed after preferential orientation of the sample on the (00 1) cleavage face. Thus, the diagnostic 001 diffraction peaks are studied. Clays are also analyzed as randomly oriented powders in order to analyze hkl diffraction peaks (Moore and Reynolds 1989).

In recent years, methods of profile fitting of powder diffraction patterns have been developed for crystal structure determinations (Howard and Preston 1989). These methods are capable of considerable accuracy and do not require perfect crystals. Profile fitting is also used for the following determinations: quantitative phase analysis, accurate determination of lattice parameters, texture studies, particle-size analysis, characterization of microstrains, and others. Such studies require precise diffraction data, using monochromatic radiation that is stepscanned for long counting times. The peak profile can be characterized lily a variety of functions including Gaussian, Lorentzian, Pearson, and Voigt. Experimental data are initially corrected for experimental scanning conditions and conditions of sample preparation. Profile fitting of a powder diffraction pattern can yield a refinement of the atomic coordinates, occupancy of crystallographic sites, and thermal vibration coefficients. Observed profiles are compared to those calculated on the basis of a refined structural model.

A whole-profile fitting routine, termed Rietveld refinement, allows the refinement of crystal structures, and determination of highly accurate quantitative analyses and unit cell parameters (Post and Bish 1989).

X-ray diffraction analysis gives important information on solid solutions (the number and kind of atoms in crystal structures of minerals). Substitution of ions of different sizes causes differences in d-spacings and unit cell dimensions, which typically have a linear dependence with respect to ionic substitution as described by Vegard's Law. This law is valid for microimpurities as well. Thus, X-ray diffraction analysis is used to estimate the chemical composition of a mineral. Relations between composition and unit-cell parameters are documented for nearly all major minerals.

To study the abundance of microimpurities, highly accurate (± 0.01-0.001 %) determinations of unit cell parameters are needed. To achieve this kind of accuracy, the experimentalist must collect diffraction data over a large angular range, use well-aligned equipment, add a standard material (e.g., quartz, silicon, germanium, or wolfram) to correct for systematic shifts, and process the experimental data (e.g., using the least-squares method of calculation). Such methods can allow the detection of down to 0.01 wt% impurity. In addition, the


content of vacancies and interstitial atoms may be determined by comparing calculated and measured densities.

Various methods are used for studying the phenomenon of order-disorder, which is characteristic for minerals. For example, variations in AI-O and Si-O bond lengths in alkali feldspar allow determination of the distribution of Al and Si in the possible tetrahedral sites. Variations in atom distribution in crystallographic sites for minerals of similar chemical composition produce differences in the size and symmetry of the unit cell. For example, the transformation from a disordered state to an ordered state can cause a reduction in the symmetry of a mineral (monoclinic to triclinic for alkali feldspars and cubic to tetragonal for some sulfides). In minerals such as plagioclase and sulfides, particular atom distributions can cause superstructures, resulting in the appearance of additional reflections. The position and intensity of these reflections depend on the domain structure. X-ray structural analysis has been used to study cation distributions in pyroxene, amphibole, garnet, spinel, and other minerals.

Single crystal X-ray methods (those of Schulz, Berg-Barrett, Fujiwara, Borrmann, and Lang) allow the characterization of the presence of perfect and defective areas, boundary features and rotation of blocks, dislocation distributions, and other features of the real structure of crystals. This gives import,ant information regarding the mechanism of crystal growth, conditions of the origin of the mineral, influence of defects on crystal properties, and possibilities for the practical use of defects. The methods are widely used for the investigation of crystals of quartz, beryl, corundum, topaz, and other minerals.

Crystal imperfection is also estimated by analysis of the diffraction line profile. The measured peak width at half-maximum intensity at various Bragg angles is compared to that of a standard material in order to determine particlesize and microstrains. An example of a geological application for this kind of analysis is the relation between the degree of crystal perfection and conditions of shock metamorphism. Increasing particle size of minerals generally correlates with increasing shock pressure and increasing crystal perfection in quartz correlates with increasing temperature of origin (Snyder and Bish 1989).

In technical mineralogy studies, single crystal X-ray methods are used for the determination of crystal orientation and the degree of crystal perfection from the standpoint of their practical use.

X-ray methods are used to study the behavior of earth and lunar minerals in subsolidus cooling by examination of exsolution microstructures from the main rock-forming minerals (pyroxene, amphibole, alkali feldspar, plagioclase, and ore minerals). Single crystal investigations (Laue's method or the method of precise investigations) permit the determination of the interrelationship of ex solved lamellae, the size and form of separations, the features of intergrowing boundaries (coherent and noncoherent), structural features of exsolution lamellae, the mechanism of exsolution, and the rate of subsolidus cooling (Putnis and McConnell 1980).

High precision X-ray analysis on single crystals, quantitative and qualitative phase analysis, determination of lattice parameters, and techniques for deter-


mining crystal defects are widely used as analytical methods for characterizing crystal chemical properties of the mineral groups. A discussion of the results of X-ray investigations of various important minerals and mineral groups follows.

Pyroxenes and Amphiboles

From single crystal data, the cation distribution (AI, Fe, Mn, Mg, etc.) in octahedral positions has been established for both pyroxene and amphibole. It has been demonstrated that the a and b unit cell dimensions can be used to characterize the Fe/Mg occupancy of the Ml and M2 sites in orthopyroxenes. Al/Si distribution in T 1 and T 2 sites is usually determined from AljSi-O interatomic distances. The variations in lattice parameters of pyroxenes and amphiboles as a function of chemical composition are well studied: MgSi0 3-FeSi03, CaMgSi20 6-CaFeSi20 6, NaFeSi20 6-CaMgSi20 6, CaMgSi20 6- NaAlSi20 6, NaAlSi20 6- NaCrSi20 6, CaMgSi20 6-CaAI2Si06, CaMgSi20 6-CaFe + 3 AlSi06, CaMgSi206-Mg2Si206' Mg7SisOdOH}zFe~ 2SisOdOH)2' and Na2CaMgsSisOdOH)2-Na2CaFet 2SisOdOH}z. (Prewitt 1980; Veblen 1981).

Layer Silicates

Single crystal diffraction studies of coarse-grained layer silicates (e.g., muscovite) laid the foundation for powder diffraction studies of fine-grained layer silicates (i.e., clay minerals) (Bailey 1980). Samples are typically prepared as sizefractionated oriented aggregates or as randomly oriented powders. Oriented samples are X-rayed after a variety of treatments including air-drying, exposure to ethylene glycol, and heating. Analysis of the d(OO 1) spacing from oriented samples permits determination of the mineral group [d(OO 1) of ,..., 7 A = the kaolin and serpentine groups, d(OO 1) of ,..., 10 A = mica and illite, d(OO 1)of ,..., 14 A = chlorite and vermiculite, and d(OO 1) of ,..., 12.5-15 A = smectite].

The hkl reflections obtained from randomly oriented samples provide information on polytype.

Solid solutions can be estimated by determination of unit cell dimension. For example, the b-axis dimension is commonly used to estimate the relative amount of Mg and Fe substitution in trioctahedral micas and chlorites. In addition, the amount of tetrahedral Al in chlorite can be estimated by the d(O 0 1) value (Brown and Brindley 1980). A widely used relationship is the one between the cation occupancy of octahedral sheet (dioctahedral vs. trioctahedral) and changes in the b-axis dimension [d(O 6 0) spacing].

Mixed-layer clay minerals (those with more than one structure along the crystallographic c*-axis) have complex diffraction patterns with nonrational 001 peaks (Reynolds 1980). Characterization of these kinds of minerals involves determination of the kinds and proportions of layers involved and the precise


arrangement of ordering of layers. Identification of mixed-layer minerals can be done by a direct method of Fourier analysis or by comparing to computercalculated diffraction patterns. Recently, Eberl et al. (1990) characterized mixedlayer illite/smectite using an alternative method of analysis of X-ray diffraction data. To better understand the mechanism of formation of mixed-layer illite/ smectite, they determined the proportion and thickness of sequences of illite layers in illite/smectite (termed illite fundamental particles) using a profile-fitting method.

Minerals of Iron-Manganese Modules

X-ray analysis of fine-grained and poorly crystallized minerals such as Mn hydroxides consists of heating the sample to observe changes in the diffraction pattern. Mn-hydroxides are common in marine environments where Fe/Mn nodules contain the minerals buserite, disordered mixed-layer asbolan-buzerite, birnessite, vernadite, and others. It is possible to determine the composition of mixed-layer asbolan-buzerite from the d(OO 1) spacing of the heated sample. Methods to characterize the relative disorder in vernadite structure have been proposed.

Alkali Feldspars

The AljSi distribution in tetrahedral sites of alkali feldspars is based on changes in the AljSi-O bond lengths which can be established by means of an X-ray crystal structure analysis. The AljSi distribution (occupancy of Al in the T 1(0), T 1(m), T z{o), and T 2(m) sites) is based on the determination of the band c unit cell dimension as well as the angles, IX and y. This method uses the d-spacing of the 204, 060,131, and 131 peaks. X-ray diffraction methods (using the 010, 001, and 110 d-spacings) have been proposed for analysis of oriented single crystals which permit more precise calculations of the AljSi distribution, allow investigation of inhomogeneous samples, and yield estimates of the degree of perfection of individual phases on the basis of a peak width at half-maximum.

The composition of alkali feldspars (K vs. Na content) can be determined by the measurement of a single peak [e.g., d(20 1)] or unit cell dimensions. Determination of monoclinic and triclinic potassic feldspar [using d(l 3 1) and d(l 31) peaks] have also been reported.

Plagioclase

Lattice parameters in plagioclase vary as a result of solid solutions (coupled Na + Si for Ca + Al substitution) and AljSi ordering, which can result in the

presence of complex superstructure peaks. AljSi distribution in the four alumi-


num/silicon tetrahedra of the average plagioclase structure can be calculated from the y angle, and the d(O 1 0) and d(11 0), d(131), and d(131) peak positions. Single crystal diffraction studies (based on measurement of the d(O 1 0) and d(110) peaks) have been proposed for the determination of AljSi ordering, composition, and degree of perfection of exsolution structures in sodic plagioclases.

Quartz

Precise determinations of the c-axis value of quartz permit calculation of the amount of Al for Si substitution. In addition, the abundance of monovalent and bivalent ions in open cavities of quartz can be determined on the basis of the aaxis value. For these determinations, the powder diffraction technique (which can yield an uncertainty of 0.0001 A in the unit cell parameter) or the Bond method for single crystals (which can yield an uncertainty of 0.00001 A in the unit cell parameter) can be used. Perfection of quartz crystals (including the effects of both strain and particle size) can be estimated from the measurement of the width at half-maximum of diffraction peaks. The Lang method permits characterization of the zonal and sectorial structures in single crystals. Methods of quantitative phase analysis have been developed for the determination of the silica polymorphs (quartz, cristobalite, and tridymite) in a sample.

Garnet

Unit cell parameters are widely used to characterize the chemical composition of garnet. The a-axis value varies in the isomorphous series of andradite-grossular, andradite-uvarovite, almandine-grossular, and pyrope-grossular. The a-axis value varies as a function of the substitution of + 1, + 2, + 3, and + 4 cations. For garnets of complex composition, the unit cell parameters are studied in combination with the index of refraction and density.

References

Bailey SW (1980) Structures of layer silicates. In: Brindley GW, Brown G (eds) Crystal structures of clay minerals and their X-ray identification. Mineralogical Society, London, pp 1-124

Bond WI (1976) Crystal technology. Wiley, New York Brown G, Brindley GW (1980) X-ray diffraction procedures for clay mineral identification. In:

Brindley GW, Brown G (eds) Crystal structure of clay minerals and their X-ray identification. Mineralogical Society, London, pp 305-360

Chichagov AV, Sipavina LV (1982) X-ray parameters for solid solutions. Nauka, Moscow (in Russian)

Cullity BB (1978) Elements of X-ray diffraction. Addison-Wesley, Reading, 555 pp


Eberl DD, Srodon J, Kralik M, Taylor BE, Peterman ZE (1990) Ostwald ripening of clays and metamorphic minerals. Science 248: 474-477

Frank-Kamenetsky VA (ed) (1983) X-ray analysis of most important types of rock-forming minerals (layer and framework silicates). Nedra, Leningrad (in Russian)

Howard SA, Preston KD (1989) Profile fitting of powder diffraction patterns. In: Bish DL, Post JE (eds) Modern powder diffraction. Reviews in mineralogy, vol 20, Mineralogical Society of America, pp 217-176

Hutchinson EK (ed) (1989) The encyclopedia of mineralogy. Ross Publ Stroudsburg, Pennsylvania

Moore DM, Reynolds RC, (1989) X-ray diffraction and the identification and analysis of clay minerals. Oxford Univ Press, 333 pp

Newkirk JB, Wernick JH (eds) (1962) Direct observation of imperfections in crystals. Interscience Publ New York

Post JE, Bish DL (1989) Rietveld refinement of crystal structures using powder X-ray diffraction data. In: Bish DL, Post JE (eds) Modern powder diffraction. Reviews in mineralogy, vol 20, Mineralogical Society of America, pp 227-308

Prewitt CT (ed) (1980) Pyroxenes. Reviews in Mineralogy, vol 7, Mineralogical Society of America, 525p

Putnis A, McConnell JDC (1980) Principles of mineral behaviour. Blackwell, London Reynolds RC (1980) Interstratified clay minerals: In: Brindley OW, Brown 0 (eds) Crystal

structures of clay minerals and their X-ray identification. Mineralogical Society, London, pp 249-304

Ribbe PH (ed) (1983) Feldspar mineralogy. Reviews in mineralogy, vol 2, Mineralogical Society of America, 362 pp

Snyder RL, Bish DL (1989) Quantitative analysis: In: Bish DL, Post JE (eds) Modern powder diffraction. Reviews in mineralogy, vol 20, Mineralogical Society of America, pp 101-144

Veblen DR (ed) (1981) Amphiboles and other hydrous pyriboles. Reviews in mineralogy, vol 94, Mineralogical Society of America, 372p

2.3 Neutron Scattering, Neutron Diffraction: Hydrogen Location, Cation Distribution, Magnetic Structures

H. FUESS

Properties of the Neutron

Neutron scattering techniques are complementary tools to X-ray diffraction and spectroscopy for the study of structure and dynamics of the solid state. Neutrons are conventionally produced by fission of uranium-235 with thermal neutrons. About 2.5 neutrons are evicted by a fission process. These neutrons are high energy, fast neutrons. They are slowed down, moderated, by collision with light elements, e.g., D20 or graphite. Research reactors are optimized to produce large quantities of neutrons, characterized by the flux (about 1015

neutrons/cm2/s in the best available sources). Most of those neutrons never reach the sample, a maximum flux at the sample position is about 108 n/cm2/s. Moderation at room temperature produces so-called thermal neutrons (wavelength 0.5 A < A. < 2.0 A) with a distribution following a Maxwellian law.

2.3 Neutron Scattering, Neutron Diffraction 39

Suitable monochromatic neutrons are selected by a single crystal monochromator [Cu(200); Ge(111); Si(111)] or a chopper.

In contrast to X-ray diffraction, the neutron is scattered by the nucleus, hence the scattering powers of isotopes of the same element are different. This scattering is called nuclear scattering and gives rise to Bragg reflections. The power is described by a quantity called scattering length given in units of 10 - 12 cm in Fig. 2. Here the huge differences between isotopes, e.g., H 1 and D2 or Ni58 and Ni62 are seen. Differences are also pronounced, for some elements adjacent in the Periodic Table (e.g., Fe and Mn), whereas hydrogen has a scattering power which is almost identical to all other elements. Therefore the classical applications of neutron diffraction are (1) determination of hydrogen positions and (2) cation distribution, both important for mineralogy.

In addition to the interaction of the neutron with nuclei, a magnetic contribution arises from the scattering of neutrons by the magnetic moments of unpaired electrons of transition elements. In a paramagnetic material the magnetic scattering contributes to the background. If a material is magnetically ordered, the magnetic scattering gives rise to Bragg peaks. These Bragg peaks are superstructure peaks for an antiferromagnetic solid, whereas the scattered magnetic intensity has the same periodicity, hence identical scattering angles for ferro- and ferrimagnetic material. In any case are nuclear and magnetic intensities of unpolarized neutrons strictly additive. Whereas the nuclear scattering length does not depend on the scattering angle, the magnetic

Ni!;8 -12 1.4

10 cm 1.2

1.0

III 0.8 .c 0 -en 0.6 c: .! en 0.4 c: ·c 0.2 Q.I .... .-0

0 u Ni

en 20 40 80 100 -0.2 Li7 ret. atomic mass -0.4 H Ti Mn

-0.6

-0.8

Fig. 2. Scattering lengths of elements and isotopes as a function of atomic mass


scattering is governed by a form factor comparable to the form factor of X-ray scattering. Some properties of X-ray and neutron diffraction (= elastic scattering) are compared in Fig. 3. Paramagnetic scattering only contributes to the background in a neutron diffraction pattern.

Neutrons may not only be characterized by a wavelength (.Ie) but at the same time by an energy (E) or velocity (v) .

.Ie = 2n/k = h/mov = hlp = h/sqr(moE) [A.l The wave number k is most convenient for the description of the inelastic scattering process, mo is the mass of the neutron, and h is Planck's constant.

The energy of thermal neutrons is of the same order of magnitude as the vibrational energy of phonons. Therefore loss or gain of energy by neutrons are an indicator of dynamical processes in the sample. Inelastic scattering may arise from a collective motion of particles in a solid, hence to coherent scattering due to phonons. Inelastic incoherent scattering is mainly due to hydrogen and reveals slow librational or vibrational motion in solids (e.g., water diffusion in zeolites). The complete process in a neutron scattering experiment is therefore composed of an elastic part (nuclear and magnetic) and an inelastic part (almost only nuclear; inelastic magnetic scattering as rarely observed). In addition the scattering may be coherent (elastic coherent scattering = Bragg scattering) or incoherent. Figure 4 gives a schematic view of the range of interactions of

scattering centre

wave lengths

scattering power

f

scattering f power of elements

x - Rays

electron

o 0 0.71 A, 1.54 A IPtc

sin ell.

z

Neutrons

nucleus and .. magnetic" electrons

o 0.5 - 20A

fmo,~

1.0 .....

b

62 Ni

: sin ell.

A

Fig. 3. Comparison of some fundamental properties of X-ray and neutron diffraction


Neutron Atom Result

Absorption 0 - • Nucleus ~ Diffraction 0 - • •• _n

Inelastic 0 - • •• ___ n Scattering

,--Magnetic

/0;: . ..0<\ 6f .... ,,\0\ Scattering o -- L,~!.IJ}

\'0--/0 '/ (elastic and ..... :-=:. Y

inelastic I 3d, 4f, 4d, 5f-

Elements

Fig.4. Schematic representation of the possible interaction of neutrons with atoms

neutrons with matter. A more complete account of neutron scattering has been published (Fuess 1979, 1988).

Neutron Diffraction

Neutron diffraction experiments were first performed on powder samples in the 1950s, then mainly on single crystals, but the advent of the Rietveld techniques (Rietveld, 1969) brought powder diffraction back into competition. High resolution powder pattern from steady state reactors (Hewat, 1986) or from pulsed sources produce structural results of a precision almost equal to single crystal experiments.

Examples for the determination of hydrogens are the location of water molecules in natural zeolites or the OH-dipole in amphiboles (Joswig et al. 1989). Stuckenschmidt et al. (1990) have shown (Fig. 5 ) that even at 10 K a certain amount of disorder is still present in harmotome, a barium zeolite, whereas well-ordered water molecules were found in scolecite (Joswig et al. 1984). Systematical studies on water in zeolites were reported by Smith, Kvick, and coworkers (e.g. Kvick et al. 1986).

The location of organic molecules in the catalytic active zeolite Y was achieved by Czjzek et al. (1989) by powder neutron diffraction. Based on a powder pattern (Fig. 6), the cations (Na and Vb) were first located by Fourier difference syntheses and included in the refinement. In an additional step, the organic molecules were placed by a subsequent Fourier synthesis (Fig. 7). On the same system inelastic and quasielastic scattering revealed that the xylene molecules are well localized up to temperatures of about 100 K. The methyl groups start rotational motion at about 70 K, the molecule as a whole at about 120K.

42

22

Chapter 2. Diffraction Methods and Crystal Structure Analysis

H1

010 Ba

H1

Fig. 5. The barium water complex in harmotome at 15 K

700 r----------------------------- ------------------------------~

600

500

400

300

200

100

o I I I ll iI !UIIIIIU IIII III IIIIII I!lI l1l l1 l1' l lI illl'l iI 'IlfI Il !l IlIl Il !lIl'l!li lllllllll11 111 1111 111 11 111111 111 11 11 !I !I 111111111 11 111111 11 11

vrvY"+~,'v,»~!i~~IJr'N~)v"Y.riW1(}Y'''ly,ryJ...$''''''\l'..Ni'!IIV'·\/ll .• ~~, 20 40 60 80 100 120

28 (degrees)

Fig.6. Refined Neutron powder pattern of Yb- Y- Zeolite with 9.4 molecules para-Xylene


Fig.7. The location of deuterated P-Xylene molecules in zeolite Y

Both single crystal and powder neutron diffraction studies have contributed to our knowledge of cation distribution and magnetic order in minerals. Figure 8 shows the pure nuclear scattering of Fe2Si04, Mn2Si04, and MnFeSi04. The pronounced difference in scattering power between Fe and Mn (see Fig. 2) is clearly reflected. Figure 9 represents the powder pattern of the same samples below the antiferromagnetic ordering temperature. Coherent magnetic Bragg scattering is observed in addition to the nuclear scattering. The intensity differences in the powder patterns indicate an alignment of the magnetic moments parallel to different crystallographic directions. The powder pattern for three intermediate compositions of the solid solution (FexMnl-x)2Si04 with x = 0.29,0.50, and 0.76 is displayed in Fig. lOa, the resulting cation distribution in

Fig. lOb.

44

o 10

o 10

111 201

", ", ",

" " ,,' II'

"

" , '

,-,

"

,

20

20

301 121

221 401

I I "I" I ., , , "

: II

, "

"

, , , , , ,

30

30

Chapter 2. Diffraction Methods and Crystal Structure Analysis

FeMnSi04

40

40

A. = l.22l

FeMnSiQ,

A. = 2.46 A

Fig. 8. Powder diffraction pattern of three olivines (nuclear scattering)

Fig. 9. Powder diffraction of three olivines (nuclear and magnetic scattering)

o a

2.3 Neutron Scattering, Neutron Diffraction

x =0.29 h300K

10 20 30 40 50 60 70 28 [.)--........

0.8

0.6 nl(Fe)

0.4

0.2 I •

I .1 I·

I

45

......... -.-..... . / •• / I.

o 0 0.2 0.4 0.6 0.8 1.0

b nz(Fe)

Fig. 10. Magnetic and nuclear scattering at 4.2 K from (Fe.Mn1 _.hSi04 ; wavelength), = 2.46 (Ballet et al. 1987)

The change in the magnetic structure with composition and temperature is clearly indicated. A detailed study on the magnetic properties on the olivine structure has been carried out by the present author and his coworkers (Fuess et al. 1988).

Electron distribution in solids has been determined by a combination of xray and neutron scattering (X-N-techniques). This technique consists of several steps: (1) the complete electron distribution is measured by an X-ray diffraction experiment, (2) the position and the thermal motion of the nucleus is derived from neutron diffraction. (3) The electron density of a spherical atom is calculated by taking form factors calculated for a spherical atom into account and finally (4) the difference between the density of step (3) and step (1) is calculated. The result should present the deformation of electrons by chemical bonding. The density thus obtained (Fig. 11) is therefore termed deformation density and depicts electrons in the chemical bond as well as lone pair electrons not involved in bonding. It has been demonstrated that the deformation is affected not only by covalent bonding but also by the influence of hydrogen bonds or cations surrounding a water molecule (Bats and Fuess 1986).

These experimental densities are to be compared with calculated densities based on molecular orbital theories. Another method to elucidate electron distributions in solids is the determination of magnetic density distributions. Careful measurements of magnetic structure factors by polarized neutrons (neutrons with only one spin state) are the base of a calculation of a distribution of magnetic moments. Studies of the synthetic garnet structure not only indicated different magnetic moments and densities on the two different iron ions in that structure but also a density on the oxygen atoms. This is interpreted as a transfer of unsaturated spins from iron to the oxygen atoms. The density on


, /~

i/ ,il

~® { /'"

.... _. __ ....... ..

Fig.H. Static deformation density of the water molecules in MgS 20 3 *6H20. At left: the plane defined by the H-O-H group. At right: the plane bisecting the OH bonds

oxygen is thus a direct visualization of the superexchange pathway Fe(oct)-O-Fe(tetr) and an indication of partly covalent bonding in the garnet structure. Figure 12 shows the situation for erbium iron garnet (Hock 1990).

Inelastic Neutron Scattering

Hydrogen has a large incoherent scattering cross-section which is due to the two possible spin states of the hydrogen nucleus. The inelastic spectra of water


....... .. ' Fe-r "·,

.......

(~"j\: .. '

Er __ .-- ...

(.' .. ~ : " ' " .'

Fig. 12. Magnetization density in erbium iron garnet. Left: Plane of ER, Fe(tetr.), Fe(oct.). Contour interval 0.15IlB. Right: Magnetization density on two nonequivalent oxygen atoms, contour lines at 0.001 IlB

Glh",) (\., T = 295 K /~.; '\

./ t .. f ~\

!-v. ,/ I '\ ( \ .... /~..,. ~\,-'\ C~_AE!AZITE t : \ \ .' ! ("\ -!I:' I V\ STiLBITE l' i........ ,.........,

! \ V ~ .... ", /,' v,

i \ /"l \,

.!!! ,I v \ HEULANDITE '2 j \/-....... ~ I ' >. '-c '--1i '-c

{\

1\ ; , , i i , i , ;

: ! PHILLIPSITE

". " I ~\ ! I \0.''\ ./ J \....../j

o 20 LO 60 80 100 120 lLO

ne.> [meV)-

o 30 60 90 1 0 w[meV)

Fig. 13. Generalized phonon density of states G(hw) of zeolites. Left: Spectra for several zeolites. Right: Harmotome in different dehydration states


motion in zeolites are therefore essentially due to the dynamics of the water molecules and therefore a direct indication of the motion of water in zeolites. Figure 13 displays inelastic spectra for several zeolites. Most structure is observed for samples which contain water closely bound to cations and framework as in zeolite natrolite. Only for natrolite is a sharp separation between translational and librational motion seen. Here the energy observed corresponds to distinct vibrations of water molecules against the cation or the framework, whereas the huge unstructured spectra reflect, however, diffusional motion of water.

In Fig. 13 phonon density of spectra for harmotome in different dehydration states reveals more structure for low water content which corresponds to water molecules in close contact with the framework.

The inelastic coherent interaction of neutrons with matter is the basis of the determination of lattice vibrations by the measurement of phonon dispersion curves. Figure 14 displays the phonon dispersion curves for the mineral anhydrite CaS04 .

Both the acoustic and the optical phonon modes show reasonable agreement between the experiment and the model calculation (Schweiss et al. 1987). The calculations were carried out based on a breathing shell model for (1) a purely ionic model with charges on all atoms and (2) a model which treats the SOIgroup as a whole as a rigid body neglecting internal vibrations. The calculated phonon frequencies differ only little (about 1- 2 meV) for the two models. The fit of model parameters to the observed phonon dispersion curves was achieved by

( 0~01 (0)

[OO~I (~O~ I 10~0)

40 40

30

20

10

(b) (SOC) (OO~)

" -A, 8 , 8,

Fig. 14. Phonon dispersion curves ror CaS04 (a) experimental, (b) calculated


treating the charges on the atoms (or ions) as variables. This treatment indicates values of charges which are considerably lower as those produced by a pure ionic model. In fact the charges refined are approximately Ca (1.3), S (0.54), 01 ( - 0.37), 02 ( - 0.53).

It is not claimed that these values are to be taken as absolute values, but they agree with the tendency indicated by X-N deformation studies in minerals.

Conclusion

Neutron scattering offers a wide range of applications in studying minerals. The classical applications of hydrogen location and cation distribution were completed by the investigation of deformation and magnetization densities. Inelastic scattering both coherent and incoherent is a complementary method to the traditional spectroscopic techniques applied in solid state research.

References

Ballet 0, Fuess H, Fritsche T (1987) Magnetic structure and cation distribution in (Fe, MnhSi04 (olivine) by neutron diffraction. Phys Chern Minerals 15: 54-58

Bats JW, Fuess H (1986) Charge density distribution in thiosulfates: NazSz0 3 and MgSz0 3 * 6H zO. Acta Cryst B42: 26-32

Czjzek M, Vogt T, Fuess H (1989) Lokalisierung von para-Xylol in Yb-Faujasit (Zeolith Y) durch Neutronenbengung. Angew Chern 101: 786-787

Fuess H (1979) Application of neutron diffraction to chemistry. In: Fluck E, Goldanskii, (eds) Modern physics in chemistry, Vol 2. Academic Press, pp 1-193

Fuess H (1988) Pulsed and continuous sources in comparison. In: Carrondo MA, Jeffrey GA (eds) Chemical crystallography with pulsed neutrons and synchrotron X-rays. NATO Adv Study Inst. Reidel, Dordrecht 77-115

Fuess H, Stuckenschmidt E, Schweiss BP (1986) Inelastic neutron scattering studies of water in natural zeolites. Ber Bunsen-Ges Phys Chern 90: 417-421

Fuess H, Ballet 0, Lottermoser W (1988) Magnetic phase transitions in olivines. In: Ghose S, Coey JMD, Salje E (eds) Structural and magnetic phase transitions in minerals. pp Springer, Berlin Heidelberg New York, pp 185-207

Hewat A W (1986) D2B, a new high resolution neutron powder diffractometer at ILL, Grenoble. Mater Sci Forum 9: 69-79

Hock R (1990) Bestimmung der magnetischen Strukturen der Seltenen Erd-Eisen Granate des Erbiums und Terbiums durch elastische Neutronenstreuung. Thesis, Universitat Frankfurt, Germany

Joswig W, Bartl H, Fuess H (1984) Structure refinement of scolezite by neutron diffraction. Z Kristallogr 166: 219-223

Joswig W, Fuess H, Mason SA (1989) Neutron diffraction study of a one-layer monoclinic chlorite. Clay Clay Minerals 37: 511-514

Kvick A, Artioli G, Smith JV (1986) Neutron diffraction study of the zeolite yugawaralite at 13 K. Z KristalJogr 174: 265-281

Rietveld HM (1969) J Appl Cryst 2: 65-71 Schweiss BP, Dyck W, Fuess H (1987) A theoretical lattice dynamics model and phonon

dispersion measurements for CaS04 (anhydrite). J Phys C Sol State Phys 20: 651-670 Stuckenschmidt E, Fuess H, Kvick A (1990) Redetermination of the structure of harmotome

by X-ray (293 K, 100 K) and neutron diffraction (15 K). Eur Mineral 2: 861-874


2.4 Electron Diffraction Analysis

B.B. ZVY AGIN

The phenomenon of electron diffraction (ED) was discovered by Davisson and Germer in1927, confirming the idea of de Broglie, proposed in 1923-24 on the wave nature of moving elementary particles. This laid the foundation for an independent method of structure analysis similar to X-ray and neutron diffraction (XRD and ND). The specific features of ED distinguishing it from XRD and ND, are defined by the higher (105-106 times) interaction of electrons with the scattering matter. This in turn defines the kinds of suitable study objects (thin films, surfaces, fine-grained sediments, gas molecules), and it also allows investigation of extremely small substance volumes, giving ED patterns at exposures measured in seconds. Another prominent feature of ED is the possibility to regulate the electron trajectories using electromagnetic fields arranged in various lens systems. This makes it possible to deliver peculiar diffraction patterns (DP) and, hence, the solution of special problems.

Techniques of ED

In connection with the well-known relation between the wave length and the velocity of electrons A. = h/mv there are two distinct application ranges of ED using either high- or low-energy electrons: HEED (several tens KeV to some MeV) and LEED (tens to hundreds eV). Depending on experimental conditions, two forms of ED are distinguished: transmission (TED) and reflection (RED). TED can examine thin crystals (films, textures, and polycrystals) and RED examines surfaces of any materials.

Different kinds ofDPs are obtained depending on the lens systems used. The simplest but perhaps one of the most efficient optical schemes is to use a condensor between the electron source and the specimen, which focuses primary and diffracted beam electrons on the fluorescent screen, photo sheet, or scintilating detector (Fig. 15). This scheme formed the experimental base of ED structure analysis (EDSA). Because there are no lenses after the specimen, the DPs have no geometric distortions and contain mostly sharp reflections. Therefore, this technique is called high-resolution ED (HRED) or electronography (in Russia). It may be applied both in TED and RED modes, with the diffracting area being mm's in cross-section. Special apparatuses were constructed for HRED.

The progressive development of electron microscopy (EM) has been accompanied by the development of a variety of ED modes. The most widely used is selected area ED (SA ED). It is based on the action of an intermediate lens which transfers to the object plane of the projective lens either the image in the conjugate plane of the objective lense or the DP in its focal plane (Figs. 16, 17a).

2.4 Electron Diffraction Analysis 51

2

3

4 ..... ---+--L-

Fig. 15. Ray paths at HRED. 1 Electron source; 2 condensor lense; 3 specimen; 4 screen (detector plane)

Fig. 16. Ray paths at SAED-EM. a formation of the EM image. b Formation of the SAED pattern. 1 Electron source; 2 condensor aperture; 3 condensor; 4 specimen; 5 objective; 6 objective aperture; 7 selector aperture; 8 intermediate lense; 9, 12-magnified ED pattern; 10, 13.magnified EM image of the object; 11 projective


a b c d e

Fig. 17a-f. ED schemes realized by EM. dashed lines indicate lenses in a switched-out state. a SAED. b MBED. c HRED. d HDED. e CBED. f 5MBED

An aperture in the conjugate plane selects the diffracting area, an aperture in the focal plane selects a diffraction beam which forms the corresponding dark-field image. Thus DPs from areas 1-2 ~m in cross-section may be obtained, presenting a unique means for the study of micro-single crystals.

The modern EM presents a possibility to use a microbeam (MB) having a cross-section down to several nm (Fig. 17b) for observing both the image and DP. MBED not only results in smaller scale selectivity but also yields a close correspondence between the imaged and diffracting selected area. Such a correspondence is distorted by the spherical aberration in the case of SA ED.

The electron microscope may also be used for HRED. The specimen is then placed between the projective lens and the screen (Fig. 17c). In the case of highdispersion ED (HDED) the scale of the DP is increased by means of electronic lenses (Fig. 17). The close spaced reflections are then better resolved, which is especially important near the central spot.

A distinguishing feature of convergent beam ED (CBED) is that the DP is formed by a continuous set of primary beams included in a cone (Fig. 17e). They resemble the Kikuchi patterns (KP), obtained already in 1928 for sufficiently thick crystals. As a result of multiple scattering, electron incidence cones satisfying the Bragg condition arise for any set of lattice planes. The intersection of the plane of detection with the diffracted beam cones results in traces as Kikuchi lines. In the case of CBED, similar ED effects are obtained for thin crystals because the necessary set of primary beams is formed beforehand in the incident cone. The CBED patterns contain unique information on the real symmetry (presence or absence of symmetry centers and glide components for


planes and axes of symmetry), handedness, structure amplitude values, and lattice parameters.

The use of micro beams presents two additional forms of EO: EO using a scanning or rocking MB (SMBEO and RMBEO, Fig. 17f). Resultant OPs are similar to those with a static parallel or convergent incident beam respectively, but are applicable to the use of scanning electron microscopes and may display additional possibilities if the scanning or rocking is performed according to special programs.

These techniques form an efficient system of investigation methods. It may be applied in combination with other methods often realized in one and the same instrument (e.g., in an analytical scanning transmission electron microscope). Thus, diffraction information is combined with EM images (in particular, with direct structural images produced with high-resolution EM, HREM), and spectroscopic data on composition and electronic structure of the objects under study. The system of EO methods, their relations, and combinations with other methods are presented in the Table 3.

Functions and Applications of ED Methods

HREO is the main method of EOSA. It operates with a variety of EOPs obtained for single crystals, textures, and polycrystals (Figs. 18, 19). They are imaging nearly planar cross-sections of the respective reciprocal lattices passing through the origin normal to the primary beam. This facilitates their indexing and determination oflattice constants. The reflection intensities being estimated and analyzed a number of problems may be solved ranging from substance identification to complete structure determination.

The single crystal patterns (SCP) carry information on the structural projections on planes normal to the primary beam. If the structure has axes a and b and an interaxial angle y between them in ab plane the net of point reflexions has a cell with dimensions proportional to (asin y)-l and (bsin y)-l and angle y' = 1t - Y independent of the direction of the third period c. The distribution of intensities corresponds to the normal projection of the structure on the EO P plane.

When the specimens contain quantities of crystals having a preferential orientation with a common plane or a direct line but with random orientations around the normal to this plane or line, they represent lamellar or axial (fiber-) textures, respectively. The reciprocal lattice (RL) of a texture is a corresponding rotation body of a single crystal RL, and the EOPs image their planar crosssections.

In a texture-RL, the RL-rows parallel to the rotation axis (RA) describe cylinders and the lamellar OTEOPs contain reflections distributed along ellipses as oblique planar sections of cylinders. Such patterns present separately features related to the orientating plane, texture basis (TB), and to the direction normal to it, the texture axis (T A). If the TB coincides with the plane ab, the T A


Table 3. Relation between ED techniques

ED PR DR

(SEED) CCDR (MCR)

XPED I STM. AFM I + I FESEM I + I AES I

Abbreviations and comments ED - electron diffraction. SE - selected energy (realized by filtration). PR, DR, CCDR (MCR) - design three levels of ED techniques defined by photo-, direct (detector) - and charge coupled devices (multi-channel) registration. AS-, CS-, G- and L- define the areas of application of ED having particular instrumental, experimental, and theoretical features for amorphous- and crystal substances, gases, and liquids. T - transmission and R - reflection ED. XPED - X-ray photon electron diffraction (the diffracting electrons are formed in the substance by irradiating it with X-rays). HE - and LE - high- and low-energy ED. The latter gives striking results in combination with scanning tunnel and atomic force microscopy (STM and AFM), field emission scanning electron microscopy (FESEM) and Auger electron spectroscopy (AES). High resolution ED as all the ED modes realized by means of electron microscopy (HRED and EM-ED) increase their importance when combined with analytical abilities of the X-ray energy dispersive and electron energy loss spectroscopy (XEDS and EELS). HRED is not only a sepearate ED technique but is also included in EM-ED. High-dispersion ED may be realized both in HRED and EM-ED. Selected area (SA), convergent beam (CB), and micro-beam (MB) are other techniques of EM-ED, as well as electron chaneling (EC). The latter is a result of anomalous penetration of electrons near the Bragg incidence angle. It characterizes thicker areas (e.g., at the surface) and is well combined with CBED. KP -Kikuchi patterns; PC - polycrystal -; TP - texture; SC - single crystal Ps. LA - large angle CBED realizing separate CBED discs in a wider angular range containing an increased diffractional information. Hollow cone - (HC) ED is a particular case of CBED when the incident electrons are passing along the cone surface, and represents special areas of the reciprocal lattice. As mentioned, S - scanning and R - rocking are particular cases of MBED. ALCHEMI - atomic location by channeling enhanced microanalysis, a new tool for revealing fine details of the atomic distribution. L - lamellar and F - fiber involve particular cases of TPED with the texture axis along the normal to the lamella crystal or elongation of fibre crystals. RP - rotation patterns obtained for crystals rotated during the exposure around either·a crystal axis which is perpendicular to the beam or around a normal to a crystal plane in an oblique position to the incident beam. The DPs resemble TPs but give more detailed and definite information. Z - zero; H - higher; F - first; S - second order (0) Laue zones (LZ) presenting respective parts of the reciprocal lattice.


• • • • • • • • •

• • • • • •

• • • • •

• • • •

•

• • Fig. 18. Single cry tal ED pal\ern of the serpentine mineral, lizardite

Fig. 19. Oblique-texture ED pattern of a polytype combination IT + 2H I of Lizardite


is then the reciprocal axis c*. The ellipses defined by indices hk have small axes proportional to

Bhk = (h2/a2 + k2/b2 - 2hkcosy/ab)1/2/siny.

The reflections hkl have positions along their ellipses hk defined by distances from the small axis line

Dhkl = (ha*cosfJ*/c* + kb*coslX*/c* + l)c*

= ( - hXn - kYn + 1)/doo1 ,

where Xn = (c/a)(cosfJ - coslXcosy)/sin2Y,Yn = (c/b)(coslX - cosfJcosy)/sin2y are the components of the normal projection Cn of the axis c on the plane ab measured in the units a and b, doo1 is the projection of c on the normal to abo These formulae serve for reflection indexing and lattice constant determination.

The features of SCPs and TPs are combined in patterns obtained from single crystals rotated during the exposure around axes perpendicular to or in an oblique position to the incident beam (Fig. 20). Like TPs, they contain threedimensional reflection sets. However, they express structural features related to a definite crystal, not averaged over a great number of different crystals. By changing the rotation axes and rotation angular range, it is possible to obtain intensities of particular reflections which coincide in TPs. In particular, rotation around the normal to a lamella tilted relative to the electron beam results in patterns simulating OTEDPs. They have the same advantages but are more sensitive to weak diffractional effects, revealing slightly defined features of modulations, superperiodicity, and order-disorder. The objects are studied in a state free from any effects of texture specimen preparation, e.g., structure disordering or even phase transitions caused by grinding of the samples.

In polycrystal-like specimens, the particles have any orientations and the RL nodes of a single crystal describe concentric spheres. The DPs consist of concentric circles, the radii of which, r, are in a direct relation with the lattice interplanar distances d = LA/r, where L is the distance between the specimen and screen.

As in the cases of X-ray and neutron structure analyses, the interpretation of the pattern's geometry results in the unit cell determination and indexing of hkl reflections. The space symmetry is considered in correspondence with regular reflection absences for some hkl combinations. The next stage is the analysis of reflection intensities for the determination of atomic coordinates and interatomic distances or for testing some structural models. It is performed according to similar means (trial and error, Patterson and Fourier syntheses, least squares refinement, etc.). One has to take into account, however, the relations between intensities and structure amplitudes F(hkl) (the scattering power of the unit cell contents for hkl reflections) specific for ED.

Under the kinematic approximation, the diffracted beams are supposed to be formed by scattering of the much more intense primary beam only, their own scattering by the lattice being neglected. The intensities Ihkl are then propor-


Fig. 20. Rotation ED pattern of a layer sulfide Galn 2S4

tional to the primary intensity 10 and to 1F12, with some coefficients depending on the kinds of specimen. Thus it has been shown that for a single crystal

Ihkl/ 1o = ,FIF/QI2A2,

where Q is the unit cell volume and A is the crystal thickness. The structure amplitude F depends on the distribution of the electrostatic potential cp(r) in the unit cell as

F = f cp(r)exp2ni(hx + ky + lz)dvr = ~)jexp2ni(hxj + kYj + IZj) , Q j

where fj is the atomic scattering amplitude of the j-th atom in position Xj' yj' Zj'


It depends on its potential field ¢ as

00

fj = (8n 2me/h 2) J ¢ j(r) r2 (sin sr/sr) dVr where s = 4nsin.9/ A . o

The relation AIF/ QIA ~ 1 defines the limits of applicability of the kinematic approximation. It is more justified for higher accelerating voltages (shorter A), light atoms (smaller f- and F-values), greater Q (more complicated structures, containing many atoms in the unit cell) and thinner crystals.

For the I-F relations of OTEDPs and PCPs, there are additional factors: pdhkodhk' and pd~k' respectively, where p is the multiplicity factor for a combination hkl (number of lattice planes contributing in the same reflection).

The peculiarities of the structural data obtained by EDSA are based on the fact that the F-values present the structure as a three-dimensional distribution of its potential field q>(xyz) = IF hkl exp - 2ni(hx + ky + lz) characterizing not only the atomic positions by the maxima of q>(xyz) but also the kind of atoms and occupancy of atomic sites (by the maxima values), as well as the interatomic bonds (q>(xyz)-values in the intermediate points).

SA ED forms a happy combination with EM. Presenting local diffractional characteristics and respective EM images (Fig. 21), it suits very well to the

Fig.21. SAED pattern of antigorite di playing it uperperiodicity along the a-axi and the image of the diffracting elected microcry tal


homogeneity-inhomogeneity problem. It makes it possible to choose microareas and microcrystals giving D Ps best suited for the purposes of identification and even structure analysis, which is especially important for poorly ordered substances. In the particular case of layered structures, the reflection series obtained from bent edges of particles are a powerful means of distinguishing particles with different interlayer distances, intergrowths, and inter-stratifications. SAED may also serve as an effective control of phase transitions and transformations occurring under both natural and experimental conditions. It is also a means for the interpretation of the EM contrast variations caused by different kinds of defects (dislocations), influencing the properties of materials. In favorable cases it gives diffractional data sufficient for structural determinations.

MBED has essentially the same purpose as SAED but it studies much smaller areas and presents an exact correspondence between image and DP. The DPs may contain few or no sharp reflections, or lack them completely if the diffracting area is less than one unit cell. This complicates the interpretation of DPs including the diffuse background but the information obtained justifies the efforts required.

HDED is useful in the study of superlattices, modulations, spinodal decompositions, epitaxial intergrowths, secondary diffraction effects, short distance order, peculiarities of quasicrystals, etc.

CBED sets a new stage in EDA. The DPs consist of diffraction discs, each of which is a map of intensity variation (Fig. 22). As a result of their interpretation, one can distinguish polar and nonpolar crystals, allowing the unique identification of all point symmetry groups. They make it possible to reveal glide components of symmetry elements displayed by special extinction lines (G-M lines) and to identify most of the space groups. It is also possible to analyze the symmetry of incommensurate structures and quasicrystals in the four-, five- and six-dimensional space. The direct evaluation of the structure amplitudes F including their phases (signs) suggests the development of a crystal structure analysis method completely based on CBED. The fine micro-beam probe at

Fig. 22. CBED pattern of a silicon crystal in the (100) orientation


CBED permits to obtain information on the structure and its defects concerning nano-areas, e.g., at different kinds of boundaries (twin, grain, interphase, magnetic, domain, etc.). Some modifications of CBED: large angle (LACBED), hollow cone beam (HCB), displaying Laue zones of different order-, zero- and higher (first-, second-, etc.), i.e., ZOLZ, HOLZ, FOLZ, SOLZ have advantages in obtaining more detailed and definite information on the structure and symmetry. RMBED may also be applied in the same respect, being also useful in the realization of electron channeling (EC) effects.

All the ED techniques, modes, and schemes applied separately or in combination with other methods compose the arsenal of electron diffraction analysis (EDA). EDA is many-functional and may be divided into parts according to the kind of problems which it is to solve. The application areas of EDA and effective ED-modes are presented in Table 4.

Some Results Obtained by EDA

Many thousands of reports have demonstrated the efficiency of EDA in the study of substances ranging from elements to biological compounds. The examples given below concern mainly mineral substances.

Table 4. Functions and modes of EDA. (main functions marked *, others +)

ED modes

HRED * * * * + + + + + SAED + + + * + * + * + + CBED * + + + * + MBED + * + * + + + + + RED + + + + * + LEED + + * + + ALCHEMI + + * * AS ED +

XPED + * Abbreviations. EDA-electron diffraction analysis. SD - structure determination. PA - polytype analysis. SI - substance identification. H-IH - homogeneity - inhomogeneity. OD -order-disorder. DA - defect analysis. SA - surface analysis. SRO - short range order. CV -compositional variations. PT - phase transitions. ES - electronic structure.


The use of OTEDPs presents a unique possibility for the determination of the atomic structure of fine-grained minerals because only powder specimens may be prepared for XRD study. Thus, detailed data have been obtained on atomic positions, lattice and structure distortions, and distribution of atoms over available positions for phyllosilicates celadonite, kaolinite, nacrite, dioctahedral K- and Na- mica polytypes 1M, 2M!, 2M2, and 3T. The structural study resulted in the discovery of new crystallochemical kinds of minerals, as in the cases of chapmanite and bismuthoferrite. The improvement of intensity measurements (transition from photo- to electronometric registration) increased the accuracy and enabled to reveal the positions and measure interatomic distances for hydrogen atoms.

The spot patterns obtained for single crystals were also used for structure determinations although they contain only two-dimensional reflection sets. The main source for such patterns is SAED. It has been used in the structure determination of a wide range of hybrid minerals, ribbon-layer silicate sepiolite, and the three-chain silicate, jimthompsonite. By means of HRED, a set of SC patterns corresponding to different tilting angles has been obtained and used for the structure analysis of talc.

In the case of poorly crystallized substances, ED may present patterns which although not suitable for a complete structure determination are sufficient for the construction of structural models consistent with the main peculiarities of the intensity distribution. This is especially important in the discovery of new minerals. Thus, it was established that halloysite is an original mineral of the kaolin group, chrysocolla is a peculiar copper silicate in which Cu-octahedral sheets are connected by sheets of inverted Si-tetrahedra, and that the iron oxidehydroxide group includes the previously unknown phases ferrihydrite and feroxyhite. A great variety of new phases have been found and identified among manganese oxides-hydroxides.

Qualitatively estimated intensity relations and positions of reflections along ellipses permitted the discovery of a peculiar dioctahedral mica polymorph. The T-sheets of the TOT or 2:1 layers are related by two-fold rotation axes 2 only, the symmetry centers T characteristic for usual 2:1 layers missing. In other words, their vacant octahedra are in one of the two cis-positions and not in the trans-position as usual for micas. Similar layers were detected in some of the montmorillonite species after subjecting the specimens to a procedure of structural ordering as a result of a potassium saturation with subsequent cycles of wetting and drying. In such a way, previously hidden features of single layers were revealed, different forms of cation vacancies were found, and a new research area in clay mineralogy has been developed.

The EDPs, especially the OTEDPs, RPs of single crystals, and SAEDPs of crystal aggregates, form a basis for polytype analysis. It implies a combination of experimental diffraction studies with a theoretical consideration of the whole polytype family to which the studied object belongs. By means of the latter, the possible polytypes are derived and their distinctive diffraction features are


predicted. In its turn, the solution of particular problems promotes the development of the polytypism theory.

Thus, the identification ofthe exact stacking of 1:1 OT layers in the structure of nacrite has required the derivation of the whole kaolin polytype family, consisting of 36 homogeneous polytypes. The peculiar diffractional characteristics of halloysite were a reason to consider it as a separate kaolin polytype 2M different from kaolinite (lTC), dickite (2M 1), and nacrite(2M 2). It became clear that the realization of a "kaolin" 1 M is possible for the Fe-analogs containing Fe instead of Al in octahedra and Sb or Bi in the interlayers as in the structures of chapmanite and bismuthoferrite.

For 1:1 trioctahedral phyllosilicates, 12 polytypes were derived. According to the OTED data, the main component of Zn clays is zinalsite, a serpentine-like mineral having several polytypes. The OTEDPs have revealed the existence of a serpentine 2H 1, even if it is an admixture to IT.

The derivation of the polytype families of pyrophyllite and talc, molybdenite and astrophyllite, the identification of both their natural and synthetic polytypes, are the result ofEDA. Diagnostic intensity relations were recognized even in cases of imperfect structures and diffuse reflections of the DPs. This was the case with the discovery of ferripyrophyllite composed of 2: 1 TOT layers stacked as in the poly type 2M of the usual pyrophyllite and with Fe3 + -cations replacing AI. The efficiency of SAED in polytype analysis is especially valuable when applied to intimate mineral mixtures. SAED of aggregated or tubular crystallites has revealed the original kaolin polytype realized for halloysite and the diversity of polytypes combined in single chrysotile particles.

Not only chrysotile but serpentine-like minerals in general cannot be studied satisfactorily without SA ED. Fine variations of lizardite microcrystals and superperiodicities of antygorites ranging from 20 to 100 A may be revealed only by SAED.

The combination of SAED with modern EM (high resolution, analytical attachments, ion thinning of specimens, etc.) has given abundant information. The variety of problems may be illustrated by the following few examples considered in some recent publications: superstructures in mullite; symmetry relations between single domains and twins in microcline; radiation-induced lattice defects in natural zircon; antiphase domains in CaAI2Ge20s-feldspar; structural defects in microcrystalline silica (chalcedony, moganite, opal, etc.); melting of diamond and graphite at 50 to 300 kbar; structural and orientational characteristics of spherules formed, and structural variations of the carbon phases composing them; influence of grinding on the dissolution kinetics of calcite; structural details including modulations of the sulfosalts, cylindrite and franckeite; display of periodic and aperiodic saphirine polytypes relative to planar and linear defects; reconstruction of a dynamic model for the pI-II phase transition in anorthite; orientational relationships between high-pressure phases synthesized from natural olivine, and many others.

The newest techniques of ED, such as nano-diffraction, allow reconstruction of crystal structures with resolution better than 1 A. Striking example is the use


of CBED in combination with ALCHEMI in the study of local symmetry and AI, Si ordering in K-feldspars. The cooperation of mineralogists and physicists made it possible to study the atomic structure and nanometer scale morphology of hematite and galena crystal surfaces by means of LEED in combination with STM and FESEM. More results and successes are to be expected from the combination of ED with HREM. HREM is presenting the structural information which is included in the limited number of reflexions selected by the objective apperture. It reveals the arrangement of atoms together with structural defects and local variations but is restricted in resolution. The Fourier syntheses constructed as a result of the analysis of intensities have better resolution because they are based on all the reflections of the diffraction field, but the structural information is averaged over all the unit cells. The above-mentioned study of cylindrite and franckeite, as well as of the saphirine poly types are impressive illustrations of such a combination.

References

Christy AG, Putnis A (1988) Planar and line defects in the saphirine polytypes. Phys. Chern Mineral 15: 548-558

Cowley JM (1967) Crystal structure determination by electron diffraction. Pergamon Press, Oxford

Cowley JM (ed) (1992) Electron diffraction techniques. Oxford Univ Press, Oxford Drits V A (1987) Electron diffraction and high-resolution electron microscopy of mineral

structures. Springer, Berlin Heidelberg New York Goodman P (ed) (1981) Fifty years of electron diffraction. DReidel, Dordrecht Hochella MF Jr, Eggleton CM, E1ings VB, Parks GA, Brown GE Jr, Chao Ming Wu, Kjoller

K (1989) Mineralogy in two dimensions: scanning tunneling microscopy of semiconducting minerals with implications for geochemcial reactivity. Am Mineral 74: 1233-1246

JEOL news (1977) 15E. Analytical TEMSCAN. JEOL Ltd, Tokyo McLaren AC, Fitz Gerald JD (1987) CBED and ALCHEMI investigation of local symmetry

and AI, Si ordering in K-feldspars. Phys Chern Mineral 14: 281-292 Pinsker ZG (1953) Electron diffraction. Butterworth, London Steeds JW (1986) Convergent beam electron diffraction. Edizione Enea, Roma Tanaka M, Terauchi M (1985) Convergent beam electron diffraction I. JEOL Ltd, Tokyo Tanaka M, Terauchi M, Kaneyama T (1988) Convergent beam electron diffraction II. JEOL

Ltd, Tokyo Vainshtein BK (1964) Structure analysis by electron diffraction. Pergamon Press, Oxford Williams TB, Hyde BG (1988) Electron microscopy of cylindrite and franckeite. Phys Chern

Mineral 15: 521-544 Zvyagin BB (1967) Electron diffraction analysis of clay mineral structures. Plenum Press, New

York Zvyagin BB, Vrublevskaya ZV, Zhukhlistov AP, Sidorenko OV, Soboleva SV, Fedotov AF

(1979) High-voltage electron diffraction in the study oflayered minerals. Nauka, Moscow (in Russian)

CHAPTER 3

Solid State Spectroscopy

66 Chapter 3. Solid State Spectroscopy

3.1 Nuclear Gamma Resonance (Mossbauer) Spectroscopy

3.1.1 Summary of Theory and Important Results

F.e. HAwTHORNE

Introduction

The Mossbauer effect is the recoil-free emission and resonant absorption of y-rays by specific atomic nuclei in solids. The y-rays can be used as a probe of nuclear energy levels which are sensitive to the local electron configuration and the electric and magnetic fields of the solid. Thus Mossbauer spectroscopy can differentiate between oxidation states of atoms, electron spin states, and structural environments. Of particular mineralogical and petrological interest are the abilities to derive oxidation ratios and site-occupancies of elements (isotopes) sensitive to the technique.

)I-ray Emission

Radioactive isotopes are unstable and spontaneously decay by emission of radiation; y-ray emission is one of these decay processes and is of central importance in Mossbauer spectroscopy. When a nucleus emits a y-ray, the nucleus must recoil such that the conservation of momentum principle is satisfied. When the nucleus is part of an atom of a solid, the recoil energy is too small to break any chemical bonds, and the recoil energy transfers to the phonon spectrum of the solid. As phonons are quantized, the momentum is transferred in integral amounts, and there is a finite probability that in some cases there will be no momentum transfer. The energy of the emitted y-ray in this zero-phonon event is equal to the energy of the transition as the emission process involves no recoil energy. If the emitted y-ray encounters another nucleus of the same type, it can be absorbed by a zero-phonon process, raising the nucleus into an excited state. The probability of such zero-phonon events is designated the recoil-free fraction.

Resonant Mossbauer Absorption

In a solid, the energy levels of a nucleus are a function of its local environment; they differ from one material to another, and also from one crystallographically distinct site to another in the same material. Thus (zero-phonon) y-rays emitted by one material will not necessarily be absorbed by another material with the same type of isotope present. However, we can modulate the energy of an

3.1.1 Summary of Theory and Important Results 67

emitted y-ray by vibrating the source material, thus applying a continuously varying Doppler shift to the y-ray energy. This Doppler shift can bring the y-ray energy into coincidence with the transition (absorption) energy of the same type of isotope in a different material; when this occurs, there is resonant absorption. Thus if we monitor the energy spectrum of the modulated y-ray(s), there is absorption at an energy characteristic of the nuclear state ofthe active isotope in the absorbing material. The modulated y-ray is thus a probe of the nuclear energy levels of a specific isotope in a solid.

For several reasons (e.g., unsuitable half-life, low cross-section for absorption, unsuitable decay scheme), most isotopes are not suitable for Mossbauer spectroscopy. From a mineralogical perspective, 57Fe is by far the most important isotope, but 119Sn, 121Sb, 153Eu, and 197Au are also of interest.

Mossbauer Parameters

Mossbauer spectroscopy involves interactions between the nucleus and extranuclear electric and magnetic fields. These are called hyper fine interactions; of principal importance are monopole and quadrupole interactions and magnetic dipole interactions.

The electric monopole interaction arises from the interaction between the positive nuclear charge and the electric field of the surrounding electrons. This interaction shifts the nuclear energy levels (Fig. 23a) according to the details of the local electronic structure. In Mossbauer resonant absorption, one compares the relative energy levels between nuclei in a y-ray source (emitter) and an absorber (sample). This energy difference is called the chemical (or isomer) shift, c5, and is commonly measured relative to some standard material (often IX-Fe for 57Fe).

The isomer shift is sensitive to any factor that affects the number and/or distribution of valence-shell electrons, and is thus a probe of oxidation state, spin state, coordination, and covalency. Only s-electrons have a finite probability of overlapping with the nuclear charge density and directly affecting the isomer shift. Thus Mossbauer-sensitive isotopes with different oxidation states involving variations in the number of valence s-electrons (e.g., 119Sn, 121Sb) show large differences in isomer shift (Fig. 24a). On the other hand, valency differences involving p- or d-electrons only affect the isomer shift indirectly via shielding effects, and so such isotopes (e.g., 57Fe) show much smaller variations in isomer shift with valence (Fig. 24a,b). Changes in coordination (coordination number, type of ligand) affect the details of the electron arrangement of the Mossbauer-sensitive isotope, and thus affect the isomer shift. Thus isomer shift can also be a probe for these parameters; in particular, the isomer shift for 57Fe is sensitive to coordination number (Fig. 24b), and this can be of particular use in the characterization of poorly crystalline and amorphous phases.

The electric quadrupole interaction involves interaction between the nuclear quadrupole moment and the ambient electric field, which leads to a splitting of

68

__ --JI Isomer Shift

<t1 I I

Quadrupole Splitting

Q.S.

m, +J

I

+' _I

±~

C.S. (from source) a

-I Velocity (mm s-')

0.94

0.92

Chapter 3. Solid State Spectroscopy

ISOMER SHIFT

MAGNETIC DIPOLE

SPLITTING

0.9~'-B--_6'--.--J_4'--~_2:-~0-~2--4~:-6~----iB b

Vrloclty tmm/stc)

Fig. 23a,b. Nuclear energy level diagram and M6ssbauer spectrum showing the combined effects of electric monopole and electric quadrupole interactions after Bancroff (1974); b Nuclear energy level diagram and M6ssbauer spectrum showing the combined effects of electric monopole and magnetic dipole interactions, after Wertheim (1964); in both diagrams, allowed transitions are shown by arrows

the nuclear energy levels (Fig. 23a). This splitting is called the quadrupole splitting (QS) and is a function of the electric field gradient (EFG) in the vicinity of the nucleus. Transitions between energy levels are controlled by a set of selection rules. Each allowed transition will give rise to a single absorption in the M6ssbauer spectrum. For 57Fe [I :s;; (3/2)], two transitions are allowed and a (quadrupole-split) doublet results (Fig. 23a); for 121Sb (with ground and excited states of I = (5/2) and (7/2) respectively, eight transitions are allowed and an octuplet results. The quadrupole splitting is sensitive to details of atom coordination, particularly the amount of deviation from regularity.

The interaction of the magnetic dipole moment of the nucleus with the magnetic field at the nucleus further splits the nuclear energy levels, giving rise to magnetic (or nuclear Zeeman) splitting. The magnetic field removes the spin degeneracy to form 21 + 1 energy levels (Fig. 23b). The resulting spectrum (Fig. 23b) is considerably more complex than when a magnetic field is not present. The magnetic field at the nucleus can be imposed through an externally

3.1.2 Experimental Techniques and Spectrum Fitting 69

Fe [6]

Sn4' ~ i [5] ,

? Sn2• ':.-..... Sb5•

Fe3'[4jt

Fe 2'[B] Sb3'

~ [7] Eu3' ....... Eu 2' . [6]

>----<

?

Au 3' [5]

Au l •

i [4]'

AuO [ 4]'P

I

-15 -10 -5 0 5 10 05 1.0 1.5 (a) ISOMER SHIFT(mm!s) ISOMER SHIFT (mm!s) (b)

Fig. 24a,b. Isomer shift ranges for selected isotopes in various valence states: 57Fe relative to Fe-foil, 119Sn relative to SnOz, 121 Sb relative to InSb, 151 Eu relative to Euz0 3' 197 Au relative to pure 197 Au. a All isotopes, note the wide range for species whose valence changes involve selectrons (with the exception of 197 Au). b Results for 57Fe, showing the effect of coordination number and valence. (After Hawthorne 1988)

applied magnetic field, or it can be intrinsic and due to unpaired orbital electrons. Thus M6ssbauer spectroscopy is an important tool in the study of the magnetic properties of minerals.

Determination of Site Occupancies

The most common application of M6ssbauer spectroscopy in mineralogy involves the determination of 57Fe site-occupancies in minerals. In a mineral, for each crystallographically unique site partly occupied by Fe in either valence state, there will be a quadrupole-split doublet. Assuming that the recoil-free fraction of Fe is the same at each site, the relative intensities of the doublets give the relative amounts of Fe (Fe2+ and Fe3+) at the different sites, provided the total Fe content of the mineral is known from a chemical analysis. Such results playa significant role in thermodynamic modeling of minerals.

3.1.2 Experimental Techniques and Spectrum Fitting

F.e. HAWTHORNE, A.V. BYKOV, N.N. DELYAGIN, and V.1. NIKOLAEV

The experimental set-up of a M6ssbauer spectrometer is fairly simple; a scheme is shown in Fig. 25a. A radioactive y-ray source is attached to a vibration mechanism (drive) that imparts a Doppler shift to the emitted y-ray energy. The modulated y-ray passes through the sample where that component with the

70

a

b

z o H fa.. ct: o (f)

aJ IT

fZ w U ct: w a..

o

4

..

~ :::: -../\../'-- .: II .., ra(liallon :~: ...

-2

Source. :::" moving ---

-1 o

Sample

1


---·11 Counter

2 3

VELOCITY. MM /SEC

Fig.25a,b. Scheme of an experimental arrangement for transmission Miissbauer spectroscopy. b An experimental Miissbauer spectrum; the data are represented by vertical dashes, the length of which represents ± (J based on counting statistics, the individual doublets are the fitted components of the spectrum, and the line through the data points is the envelope of the fitted spectrum. (After Hawthorne 1988)

appropriate energy is absorbed. The y-ray then passes into a detector and the resulting signals are accumulated (as a function of source velocity) in a multichannel analyzer (MeA); this is the raw spectrum.

Drive Mechanism

The drive transmits a constant acceleration (alternately positive and negative) to the source such that a range of velocities is scanned linearly and repeatedly. The energy of the resultant y-ray at any instant is related to the velocity of the source, and the spectrum is recorded in terms of y-ray intensity as a function of source velocity. The vibration of the source is driven by a sawtooth waveform of a symmetric (V\N\) or asymmetric (.....-vl) form. In the symmetric form (which is


more common), the source moves towards the sample with a constant acceleration while the MeA accumulates counts over half the channels, and then the source moves away from the sample while the MeA accumulates counts over the other half of the channels. This generates mirror image spectra in each half of the MeA; during processing, this double spectrum is folded back upon itself to give a single spectrum of the sample. In the asymmetric form, a single spectrum is accumulated over most ( > 95%) of the channels in the MeA; this waveform is more difficult to generate reproducibly.

Detectors

For 57Fe, good resolution can be obtained with a scintillation counter or a proportional counter; a proportional counter is more commonly used, usually positioned such that it senses the transmitted y-rays (Fig. 25a).

Sample Preparation

For most applications, powdered samples are used. The sample must be of uniform thickness and the grains of the powder must be randomly oriented. An inert matrix can help alleviate problems of preferred orientation. Ideally, it should be inert, granular, softer than the sample, iron-free and easy to remove (preferable soluble, e.g., sugar, salt); graphite is often used for high-temperature spectroscopy.

The theoretical arguments concerning derivation of site-occupancies from peak intensities are generally developed for infinitely thin absorbers. Excessive y-ray absorption leads to peak-shape degradation, a nonlinear baseline, peak broadening, and saturation effects. For many years, there has been a general rule that for Fe-bearing minerals, an Fe content of '" 5 mg/cm2 is a good compromise between a thin absorber and good counting statistics. However, thickness corrections are necessary for very accurate work; these are applied by running the spectra at various sample thicknesses and extrapolating the results to an infinitely thin sample.

Calibration

In the experiment, the peak positions are measured not in terms of their true energies but in relation to a zero energy point and an energy scale derived from a standard absorber spectrum. For absorbers of mineralogical interest, the following standards are commonly used: 57Fe: iron foil, stainless steel, sodium nitroprusside; 119Sn: Sn02; 121Sb: InSb. For 57Fe, the relative conversion factors are: stainless steel: 0.10 mm/s; sodium nitropruside: 0.257 mm/s.


Spectrum Quality

The variance of the counts in a specific channel of the MeA is equal to the number of counts. Hence, assuming only random error, the relative precision can be increased by counting for longer times, but this improvement tails off with increasing time. The optimum baseline (off-resonance) count is of the order of 1-5 x 106 counts/channel, usually taking 24-48 h for 57Fe in most minerals, and providing an acceptable balance between the need for precision and the desire for experimental efficiency.

The Mossbauer Spectrum

An experimental spectrum is shown in Fig. 25b. Each vertical dash represents the number of counts recorded in that specific channel (source-velocity interval) together with its associated standard deviation based on counting statistics. The counts at the edges of the spectrum (zero absorption) are the intensity of the ')'ray beam over those energy ranges where no absorption has occurred. Towards the center of the spectrum, the counts decrease due to resonant absorption of ')'rays by the sample. The ideal line shape is Lorentzian (see below), and an observed spectrum ideally consists of a series of Lorentzian lines, the number and characteristics of which are a function of the Mossbauer nucleus and the crystal structure of the sample. There is often complex overlap of individual lines, and the derivation of quantitative information (peak position, widths, and areas) from such spectra requires numerical spectrum fitting.

A Mathematical Description

The ideal shape of a peak is Lorentzian. However, there are a number of factors that can result in a Gaussian component in the peak shape. Thus the description of the peak shape can be represented as a combination of these two forms (other more complex functions are also possible). This being the case, the intensity of the ,),-radiation transmitted by the sample as a function of its energy, x, can be written as

- ~ exp { - 4ln2C ~iXiY}J where I is the number of lines, b is the background intensity, rx is the fractional Gaussian character of the line, Ai is the area (intensity) of the ith line, ri is the half-width of the ith line, Xi is the position of the ith line, y is the channel count, and x is the channel number.


Slight linear and sinusoidal deviations can occur in the background intensity due to source movement and instrumental drift; this can be incorporated by writing the background intensity, b, as a function of channel number:

b = bo + blx + b2sin(nx/w),

where bj are refinable parameters and w is the width of the spectrum in channel numbers. More complex background models are possible.

Spectrum Fitting

The equations for the spectrum given above are not linear in all variables, and they must be linearized via some expansion approximation. The variable parameters of the equation are then adjusted by least-squares refinement in order to achieve agreement between the calculated spectrum and the observed spectrum. As the initial equations of the spectral model are not linear, the process has to be iterated, gradually approaching the optimum values for which there is a good fit between the equation of the envelope and the observed data (Fig. 25b).

When there is significant peak overlap in the spectrum, as in Fig. 25b, the refinement procedure is more difficult as the variables can now interact with one another in the refinement procedure. In this case, the spectrum inherently contains less information than is the case when there is no peak overlap. As a consequence of this, the precision of the calculated parameters is less due to significant correlation in the fitting procedure; in these circumstances, it is of crucial importance that all variable be refined simultaneously in the final cycle of refinement, and that the full variance-covariance matrix be used in the calculation of the standard deviations. Linear constraints can be used in the fitting process to reduce correlation between variable parameters. However, it is important that such constraints be correct, or else the precision will be improved but the accuracy will be degraded.

Goodness-or-Fit Criteria

The least-squares method minimizes the weighted sum of the squares between the observed and "calculated" data:

where Yj is the observed count in the ith channel, O"j is the weight assigned to the ith observation, f(xj) is the calculated count in the ith channel, and N is the number of channels.

We may define the ideal residual xl as

n 1 xl = L 2{yj - f,(x j)]

j= 10"j


where fJ(x) is the true function. xl' follows the chi-squared distribution, and x~ is a value from this distribution if the parameters of the fitted function f(x) are a valid approximation to the true function fJ(x). To test whether this is the case, one uses the distribution to assess the probability that Ro < Rj at a certain percentage confidence limit. At the 1 % point on the chi-squared distribution, there is a 1 % probability that xl' will exceed this value. Thus if X~ > 1 % point, then f(x) is generally accepted as a good approximation to fb) and the "fit" is acceptable. Within the 1 % and 99% points of the chi-squared distribution, there is no statistical justification for preferring a fit with a lower X~ value as the probability that xl' will exceed X~ is quite high within these limits. There are other statistical indicators (reduced X2, MISFIT) that are also used as goodness-of-fit parameters.

Statistical acceptability is no guarantee that the derived model is correct; it merely indicates that the model adequately (but not exclusively) explains the observed data. One must judge whether or not the observed model is correct on the basis of whether or not it is physically and chemically reasonable.

3.1.3 Iron-Containing Minerals, Ores and Glasses

G. AMTHAUER, F.e. HAWTHORNE, and E. POLSHIN

There has been a large amount of work on iron-bearing minerals, focusing primarily on the derivation of site occupancies and oxidation ratios. We will first survey some site-occupancy results on important rock-forming minerals, and then go on to examine the use of Mossbauer spectroscopy in characterizing next-nearest-neighbor effects, structural phase transitions, and magnetic properties.

Garnets

There are three cation sites in the garnet structure: {X} (dodecahedral), [Y] (octahedral) and (Z) (tetrahedral); all of these can be occupied by Fe in one or more valence states, and Fig. 26 shows some resultant spectra. In the derivation of site-occupancies, it is usually assumed that the recoil-free fractions in the sample are the same at each non-equivalent site in the structure. The recoil-free fraction is related to the mean-squared vibrational displacement of the Mossbauer nucleus. Crystal structure refinements of garnets show that the meansquared vibrational displacement of atoms at the {X}, [V], and (Z) sites are distinctly different (being related to coordination number and mean bondvalence). Thus there should be significant differences in recoil-free fractions at the three sites in garnet. Mossbauer spectra of garnets recorded at different temperatures show that this is indeed the case. There is a strong differential


0.· W

0.07

0.0.4

0.06

0.10. a i 0..1]

Fe2+(dodec)

-] -I o ] mm/sec 4

0.06

01J9

0.1]

0.15 b 0.18

0.]1

Q]4

-4~~-3~~~~~-I--~--~--~]--m-m~~-.c~4

z Q Ii:

O~ r 80..0] • I

10.04 \ ~ !

i QD6 \,fJ \\ II &0.08 C

0.10. -..5 Fe3+(oct)

- Fe3+(tetr)

75

"" 5..0

§ d #.

lD.oW--f-:--'-'.l...--~:'--'..I...J.---'---:,-::-'---'--'-,,=--, -2 . .0 .0 2 . .0 '.0

mm/sec

Fig. 26a-il. Mossbauer spectra of 57Fe in garnets. a Fe2 + at {X} in pyrope b Fe3 + at [YJ in andradite; c Fe3 + at [YJ and (Z) in a synthetic Ti-andradite. d Fe2 + at {X}, Fe3 + at [YJ and (Z) in a melanite. Temperatures: a, b, 77 K; c 15 K; d 295 K. Spectra modified from Amthauer et al. 1976 (a,b,c) and Schwartz et al. 1980 (d)

temperature dependence of the recoil-free fractions at the three sites, suggesting that the use of room-temperature spectra to derive site-occupancies will introduce significant errors (at least 20%).

In some garnets, the doublet due to Fez + at the {X} site shows evidence of splitting into two closely overlapping doublets; this suggests the existence of two (or perhaps more) subsites, possibly correlating with the noncubic symmetry of some garnets.

Olivine

There are two crystallographically distinct octahedral sites, Ml and M2, in the olivine structure, both of which are occupied by Fe2 + in nearly equal amounts.


Because of the similarity of the two sites, the resultant two quadrupole-split doublets overlap very closely, and only a single doublet is apparent at room temperature. However, there are better methods than Mossbauer spectroscopy for deriving site-occupancies in ferromagnesian olivines. Ferromagnesian olivine can also incorporate significant amounts of Fe3 + as fine-scale intergrowths of laihunite, Fe2+Fe~+(Si04h; this is easily detected by Mossbauer spectroscopy.

The nonmagnesian olivine series liebenbergite (Ni2Si04)-fayalite and knebelite (Mn2Si04)-fayalite both show significant long-range ordering. The Niolivine spectra are well resolved and show Ni to strongly order at the M I site (i.e., Fe2 + strongly ordered at M2). The Mn-olivine spectra are less well resolved, but show Mn to order at M2 (i.e., Fe2+ orders at MI).

Pyroxenes

The MI and M2 sites are often occupied by Fe; details oftheir coordination vary from one series to another, and with differences in space group, these sites can each split into nonequivalent pairs of sites. In addition, the tetrahedral site(s) can also be occupied by Fe. The resulting spectra (Fig. 27) show considerable variability, depending on the structural and chemical complexity of the pyroxene.

Ferromagnesian pyroxenes, (Mg,Fe)Si03, can be orthorhombic (Pbca) or monoclinic (P2dc) but their spectra are similar. Most work has focused on the orthopyroxenes in which Fe2+ occupies Ml and M2, both of which are [6]coordinated. The resulting spectra (Fig. 27a) have two doublets which show greater resolution at low temperatures; the more intense inner doublet corresponds to Fe2+ at M2. The ordering of Fe2+ over Ml and M2 is very temperature-sensitive, and there have been extensive studies of both equilibrium and kinetic aspects of this ordering. With increasing substitution of Al and Fe3 +, the spectra become more complex, showing evidence of next-nearest-neighbor effects.

Clinopyroxenes close to the diopside-hedenbergite join can consist of either one or two (poorly resolved) doublets corresponding to Fe2+ at Ml (outer doublet) and M2 (inner doublet), respectively (Fig. 27b). In the hedenbergiteferrosilite series, the spectra cannot be satisfactorily fitted to two doublets, suggesting the occurrence of significant next-nearest-neighbour effects. Figure 27c shows the spectrum of a synthetic ferrian diopside with two doublets due to Fe3+; the isomer shift values (0.42, 0.18 mm/s) show Fe3+ to be octahedrally and tetrahedrally coordinated, respectively, and this behavior is observed in many volcanic pyroxenes.

The most complex pyroxene spectra are of the P2/n omphacites with four distinct M sites, all of which can be occupied by Fe, giving rise to quite complex spectra (e.g., Fig. 27d).


100

98

96

94

92

90

100

96

92

88 (a)

84 -4 -3

0

0.01

0.02

0.03

0.04 (c)

OD5

-2 -I 0 mm/sec

~:~~:':) , I L...-.-J

2 3 4

~

c .. V L .. 4.

... z w ~ w u.

77

(b)

2 2'

-2D o 2.0 40 Velocity(mmtsec)

2·0

40 (d)

-2·0 o VELOCITY (mm/sec)

Fig. 27a4l. M6ssbauer spectrum of 57Fe in pyroxenes. a Orthopyroxene (upper spectrum at 293 K,lower spectrum at 77 K). b Diopside; note the weak doublet assigned to FeZ + at M(2). c Ferrian diopside, with significant tetrahedrally co-ordinated Fe3 +. d Titanium ferro-om ph acite. b, C, d at room temperature. (After Virgo and Hafner 1970; Bancroft et al. 1971; Hafner 1971; Aldridge et al. 1978)

Amphiboles

The amphibole structure has four nonequivalent sites, M(l), M(2), M(3), and M(4), that may be occupied by Fe. Because of the structural and chemical complexity of these minerals, the fitting and site-assignment of the spectra are often not straightforward.

5

10

15


The spectra of the ferromagnesian amphiboles are the simplest to interpret. The spectra of anthophyllite and cummingtonite-grunerite (Fig. 28a) consist of two fairly well-resolved doublets corresponding to Fe2 + at the M(l, 2, 3) sites (outer doublet) and M(4) site (inner doublet), respectively. Solid solution towards gedrite leads to a decrease in resolution, and the spectrum of gedrite consists of a single unresolved doublet.

Calcic amphibole spectra show considerablt: variability, and details of spectral fitting and peak assignment by different authors are at variance. The spectrum of tremolite (Fig. 28b) has been resolved in two doublets due to Fe2 +

at M(4) and M(l, 2, 3), respectively. Conversely, spectra of actinolites show more fine structure, and have been resolved into four doublets due to Fe2 + at M(l), M(2) and M(3), and Fe3 + at M(2). Figure 28c shows the spectrum of a chemically more complex magnesio-hornblende with a peak assignment similar to that of actinolite; however, different peak assignments have been proposed for such amphiboles, and it is possible that such spectra cannot be satisfactorily resolved from the Mossbauer data alone.

Alkali amphibole spectra consist of three doublets assigned to Fe2+ at M(l) and M(3), and Fe3 + at M(2) in vacant A-site amphiboles. With significant A-site occupancy (e.g., eckermannite-arfvedsonite series), there should be Fe2 + occupancy of M(2) and a third Fe2 + doublet in the spectra.

The distribution of Fe2+ over the M(4) and M(l, 2, 3) sites seems to be strongly temperature-dependent both in ferromagnesian and calcic amphiboles, although kinetic effects must be taken into account when applying such results to natural assemblages.

GRUNERITE

a

HIH2.HJ

-2 o 2

·1 3

o~ • JI'" '\ ,.vr-... 1

\ f \ f '\ ,; ,II.; ~z . , "/ :; 1/' I " ~ •

i! i · i li

2 r b TREMOLITE

1.9 ,., F,O

2

3

VELOCITY (mm/sec) ~

MAGNESIOHORNBLENDE

c

-2 0 2

Fig.28a-c. Mossbauer spectra of 57Fe in amphiboles. a Grunerite (at 78 K, the inner doublet is due to Fe2+ at M(4), the outer doublet is due to Fe2+ at M(1, 2, 3). b Tremolite, the inner doublet is assigned to Fe2+ at M(4) and the outer doublet is assigned to Fe2+ at M(1, 2, 3). c Magnesio-hornblende, AA' = Fe2+ at M(1), BB' = Fe2+ at M(2), CC' = Fe2+ at M(3), DD' = Fe3+, band c are at room temperature. (After Hafner and Ghose 1971; Goldman and

Rossman 1977; Bancroft 1975)

3.1.3 Iron-Containing Minerals, Ores and Glasses 79

Micas

Ignoring subtle differences in symmetry, spectra have been interpreted in terms of two octahedral sites, M(l) and M(2), and one tetrahedral site, T. There has been confusion in the literature as to the specific structural assignment of the M(l) and M(2) sites, and this should always be checked [M(l) has a trans arrangement of OH anions, M(2) has a cis arrangement].

Dioctahedral mica spectra are dominated by a single doublet due to Fe3+ at M(2), together with two minor doublets due to Fe2+ at M(l) (inner) and M(2) (outer doublet), respectively (Fig. 29a). The large half-width ofthe Fe3+ doublet

"0 Q)

E E II) c: as F ~ 0

2 ~-+-__ -+-__ -+-__ ..J

97

c 96

Velocity (mm/sec)

o

MM/SEC

2 4

o 1 234

MM/SEC

Fig.29a-c. Mossbauer spectra of 57Fe in micas. a Muscovite containing significant Fe3+ and Fe2+. b Synthetic fluor-annite. c Annite with Fe2 + doublets from M(I) and M(2) [(1, 1) and (2,2) respectively], and Fe3+ doublets from M(I), M(2) and T sites [(4,4) shaded, (3,3), and (T, T) respectively]. (After Finch, et al. 1982; Dyar and Burns 1986, b c)


is probably due to NNN effects, as the half-width increases with increasing trioctahedral mica substitution.

The spectra of trioctahedral micas are much more complicated and show great variability in fitting and peak assignment. Figure 29b shows the spectrum of a synthetic fluor-annite with a doublet that can be resolved into two components; these have an area ratio of 2: 1, corresponding, to Fe2 + at M(2) and M(I), respectively. Natural biotites are generally more complicated than this (Fig. 29c), with additional doublets that can be assigned to Fe3 + in both octahedral and tetrahedral coordination. Extensive work has been done on the mechanism of oxidation/dehydroxylation and weathering in micas.

Glasses and Amorphous Materials

The Mossbauer effect is well suited to the study of systems lacking long-range order as it is sensitive to short-range effects, and IS and QS are good indicators of Fe valence state and coordination in such materials.

Of considerable interest is the valence and coordination of Fe in silicate glasses. Glassy materials are characterized by a broad range of site geometries, leading to broadened lines and greater difficulties in interpretation than is generally the case for crystalline materials. Nevertheless, systematic studies on series of synthetic glasses have now led to a fairly consistent spectral interpretation. Fe2 + is generally in octahedral coordination, whereas Fe3 + may be in either tetrahedral or octahedral coordination, depending on the degree of oxidation of the glass.

In the study of cryptocrystalline and amorphous materials, Mossbauer spectroscopy can be very useful in interpreting chemical reactions (alteration, oxidation, dehydroxylation) that cannot be easily examined by diffraction or microchemical techniques.

Magnetic Properties of Minerals

Magnetic hyperfine splitting arises from the interaction between the magnetic dipole moment of the nucleus and a magnetic field; for 57Fe, this results in a characteristic six-line pattern (Fig. 23b). Together with traditional methods such as neutron diffraction and magnetic susceptibility measurement, Mossbauer spectroscopy is an excellent tool for the study of magnetic properties and magnetic phase transitions. It is especially well suited to measure the internal magnetic fields at Mossbauer nuclei at different sites in a structure, as well as the onset of magnetic ordering. By applying an external field, the magnetic character of minerals that have an intrinsic magnetic moment in the absence of an applied field (ferromagnetic, antiferromagnetic, ferrimagnetic) can be derived. Singlecrystal studies are particularly useful in obtaining the orientation of electric and magnetic fields relative to the crystallographic axes.

3.1.3 Iron-Containing Minerals, Ores and Glasses 81

There has been much work done in the last few years on the magnetic properties of the rock-forming minerals. These are generally paramagnetic at ambient temperature (300 K), but show magnetic ordering at low temperatures. The garnets almandine, Fe~+ A12Si30 12, and andradite, Ca3Fe~+Si3012' both order antiferromagnetically at 15 K and 8 K respectively. However, almandine shows two resolved hyperfine patterns, suggesting some sort of canted spin structure. Fayalite becomes antiferromagnetic at 65 K, but undergoes another transition at 23 K to a canted antiferromagnetic state. In aegirine, NaFe3+Si20 6, and hedenbergite, CaFe2+Si20 6, there is antiferromagnetic ordering at 7 K and 42 K, respectively.

Magnetic hyperfine splitting in the amphiboles and the sheet silicates is somewhat more complicated. In cummingtonite, (Mg,Fe2+hSigOzz(OH}z, the magnetic species, Fe2 +, are separated by diamagnetic cations, and this is a spin glass; different local configurations lead to different magnetic interactions between neighboring magnetic ions. The resulting magnetic state has both a random character and spin frustration, indicating that adjacent magnetic moments cannot set at optimum orientations with regard to all their magnetic neighbors. This is not the case in grunerite, FeVSigOzz{OH}z, which orders antiferromagnetically at 47 K and then undergoes a spin-canting transition at 7 K. In the range 47-8 k, there is ferromagnetic coupling within the octahedral strip and antiferromagnetic coupling between adjacent octahedral strips. Similar behavior is observed in many of the Fe-rich sheet silicates. Within the octahedral sheet, there is ferromagnetic coupling and the magnetic moments tend to lie in the plane of the sheet; coupling between the sheets is antiferromagnetic.

If the size of magnetic particles is very small (::; 100 A), the rate of the thermal fluctuations of the magnetic moments of domains is approximately the same as the reverse mean life-time of the exited 57Fe nucleus, and the magnetic field at this nucleus is reduced, often effectively to zero. This phenomenon is called superparamagnetism, because the magnetic spectrum is transformed into a paramagnetic spectrum; this behavior has been observed in Fe-oxides and hydroxides in soils and sediments.

Next-Nearest-Neighbor Effects

Solid solution gives rise to different NNN arrangements around the Mossbauersensitive species in the structure. As Mossbauer spectroscopy is a probe of local structure, different NNN arrangements give rise to different absorptions in the Mossbauer spectrum. This can produce broadening of peaks or give rise to discretely resolvable absorptions in the spectrum, depending on the local details of the structure. Such effects are best observed in structures where the environment of the Mossbauer-sensitive species is very regular (often the case in highly symmetrical structures).

In wiistite, Fe2 + occupies octahedral sites with cubic point symmetry; in this case, there should be no quadrupole splitting and the Mossbauer spectrum


should show a single line. However, the Fe: 0 ratio in wiistite is less than 1, and the resultant clustering of defects produces very complex spectra with both singlets and quadrupole-split doublets.

In spinels and thiospinels, a similar situation exists. Ordered end-members show an Fe2+ singlet (e.g., for chromite, Fe2+Cr20 4, or hercynite, Fe2 + AI20 4). Intermediate solid solutions of the form Fe2+(Cr,Alh04 show complex spectra that have been resolved into a singlet and two doublets due to various NNN configurations. This effect is very prominent in the thiospinel series Fe2+(Cr,RhhS4' in which the very good resolution along the series has allowed a quantitative interpretation of the local structure.

In both wiistite, spinels and thiospinels, the situation is optimum for observing NNN effects as the end-member structure has cubic symmetry, giving rise to a singlet, and the quadrupole-splitting is very sensitive to slight deviations from cubic symmetry. When dealing with less symmetric (regular) environments, the spectrum is less sensitive to NNN effects, although they have been observed in synthetic alkali-calcic pyroxene solid solutions, the hedenbergite-ferrosilite series and some omphacites. Presumably they are also present in complex amphibole and mica solid solution, but are possibly obscured by the inherent complexity of the spectra.

In pyrrhotite, Fe1 - xS, NNN variations affect the magnetic spectrum. Cation vacancies affect the magnetic interaction between neighboring Fe2 + cations, and additional six-line magnetic patterns are discernable, with distinctly different magnetic fields and quadrupole splittings.

Intervalence Charge Transfer

In crystal structures with Fe2 + and Fe3 + at neighboring cation sites, chargetransfer may occur via electron-hopping from one cation to the other. Optically activated intervalence charge-transfer between Fe2 + and Fe3 + through common edges or faces of neighboring polyhedra is observed in the optical absorption spectra of many minerals but not in their Mossbauer spectra. Thermally activated electron delocalization between Fe2 + and Fe3 + occurs only if Fe2 + and Fe3 + occupy geometrically similar sites which share common elements, forming multiple or extended structural units such as the chain of octahedra in magnetite, Fe2+Fe2 3 +04, or the ribbon of octahedra in ilvaite, CaFe/+Fe3+Si20 70(OH). If these valence fluctuations are faster than the reverse mean life-time of the excited Fe-nucleus (i.e., 108 s -1), one pattern with averaged hyperfine interaction and averaged parameters is observed in the Mossbauer spectrum, and is assigned to "mixed-valent" iron ("Fe 2 .S +"). The probability of such an electron exchange depends on the energy barrier between the Fe2 + and Fe3 + ions at their different sites. If this energy barrier is in the order of kT, then electron exchange can be thermally activated. Thus in magnetite, mixed-valent iron is observed at T m ~ 120 K, and in ilvaite at T m ~ 350 K. In both minerals, the onset of extended electron delocalization is

3.1.4 Mossbauer Spectroscopy of Sn, Sb, Eu, Au 83

combined with a crystallographic phase transition. Similar mixed-valence states of iron have been observed by the Mossbauer effect in many other minerals: deerite, vonsenite, aegirine-augite, melonjosephite, and lipscombite.

3.1.4 Mossbauer Spectroscopy of So, Sb, Eu, Au

F.e. HAWTHORNE

Although 57Fe work has dominated Mossbauer spectroscopy of minerals and other geological materials, there has been a small amount of work on other Mossbauer-sensitive nuclei. Of particular interest in this regard are 119Sn, 121Sb, 151Eu, and 197Au. Work on these species has been of considerable importance as it has provided some of the first direct information as to the valence states of these species in minerals.

Tin may be divalent (5s25pO) or quadrivalent (5s05pO) in minerals. In some tin minerals, the valence state is fixed by stoichiometry (e.g., cassiterite, Sn4+02), whereas in other more complex minerals, the situation is less straightforward.

Sn has been detected in minor quantities in some garnets. The Mossbauer spectrum (Fig. 30a) looks like a singlet, but is a poorly resolved doublet. The IS value is 0.0 mmls relative to cassiterite, showing the Sn to be quadrivalent, Sn4+; in addition, the very small QS is consonant with very regular octahedral coordination, as expected if Sn4+ occupies the [V] site in the garnet structure.

Sn is a common constituent in complex sulfosalt minerals, and because of the multiplicity of cations with possible variable valence states, it is important that the cation valences be characterized as their roles in the structure(s) are very sensitive to differences in valence state. Franckeite, ideally Pb5FeSn3Sb2S14, shows considerable chemical variability, with solid solution towards incaite, essentially a Ag- and Sn-rich franckeite. The 119Sn spectrum of near endmember franckeite shows a strong singlet due to Sn4+ and a weak doublet due to Sn2 + (Fig. 30b). In incaite (Fig. 30c), the Sn2 + doublet has greatly increased in intensity relative to the Sn4+ peak, showing that the substitution Sn2+~ Pb2+ is an important aspect of this sulfosalt series.

Antimony may be trivalent (5s25pO) or pentavalent (5s05pO) in minerals, and its role in the structure is sensitive to its valence state.

.l!! 'iii .". 0 c:: ~ :b a.

84

0.04 (a)

-6

0 ci I

~ ci

(b)

en 0 ci -12.00

o ci I

-3

I SHoo oct

o 3 6

v (mmjsec)

II Sn'+

Sn4 +

-2.40 7.20

I I I Sn'+

Sn4 + 8l ci~----~--~----,-----~--~

-12.00 -2.40 7.20


100.0

98.0

97.0

96.0

100.0

99.0

98.0

~ 97.0

c:: (e) 0 'iii 96.0 U)

'e 100.0 U) c:: 11I ... I-

100.0

99.9

99.8

99.7

99.6 (g)

99.5

-8 -6 -4 -2 2 4 6

Velocity (mmjs)

Fig.30a-g. Mossbauer spectra of: a 119Sn in garnet; b 119Sn in (Pb-rich) franckeite; c 119Sn in incaite (Sn-rich franckeite); d 197 Au in native gold; e 197 Au in aurostibite (AuSb2); f 197 Au in nagyagite [PbsAu(Te,Sb)4Ss]; g 197 Au in sylvanite (AuAgTe4). (After Amthauer et al. 1979)

8

3.1.4 Mossbauer Spectroscopy of Sn, Sb, Eu, Au 85

Antimony is pentavalent in ordonezite, ZnSb/+06, as indicated by the electroneutrality principle. The Mossbauer spectrum shows a single line (actually a poorly resolved quadrupole-split octuplet) with an IS of 9.1 mm/s, diagnostic of Sbs +. In franckeite, the situation was not clearcut because of the possible variable valences of Sn, Sb and Fe. However, the spectrum showed a single line with an IS of - 13.7 mm/s, diagnostic of Sb3+.

151Eu

Europium may be divalent or trivalent in minerals; it is usually (assumed to be) trivalent, but the geochemical Eu anomaly common in many rocks has been explained as the substitution of Eu2 + for Ca in plagioclase feldspars. Synthesis of Eu-bearing anorthite showed the Eu was divalent in the feldspar, the first direct conformation of this important point.

Gold may be metallic, monovalent, or trivalent. However, the IS values do not reliably distinguish between these different valence states (Fig. 24a), and a consistent interpretation of the spectra of gold minerals has not yet been advanced.

The spectra of several gold minerals are shown in Fig. 30. Native gold gives a singlet with an IS of - 1.23 mm/s (relative to the Pt source); alloying with silver gradually increases the isomer shift which can reach values of 1.0 mm/s for dilute gold in a silver matrix. Aurostibite, AuSb2, is cubic and has the pyrite structure. It gives a singlet with a large positive IS. Nagyagite, PbsAu(Te,Sb)4Ss, gives a doublet, presumably the result of quadrupole-splitting, with a large positive IS. Sylvanite, AuAgTe4, is monoclinic and gives a significantly more complicated spectrum consisting of (at least) two quadrupole-split doublets with very different IS and QS values.

Mossbauer spectroscopy has been used effectively in the characterization of gold ores, and to follow the progress of gold extraction in the smelting and roasting processes used in commercial gold recovery.

References

Amthauer G, Rossman GR (1984) Mixed valence of iron in minerals with cation clusters. Phys Chern Miner 11: 37-51

Hawthorne FC (1988) Mossbauer spectroscopy. Rev Mineral 18: 255-340 Maddock AG (1985) Mossbauer spectroscopy in mineral chemistry. In: Berry FJ, Vaughan DJ

(eds) Chemical bonding and spectroscopy in mineral chemistry. Chapman and Hall, London pp 141-208


Seifert F (1988) Recent advances in the mineralogical applications of the 57Fe Mossbauer effect. Phys Prop Thermodyn Behav Miner, Proc NATO Adv Study Inst, Cambridge, July 27-Aug 8, 1987 Dordrecht, pp 687-703

Seifert F (1990) Phase transition in minerals studied by 57Fe Mossbauer spectroscopy. In: Mottana A, Burregato F (eds) Absorption spectroscopy in mineralogy. Amsterdam, Elsevier pp 145-170

References for Figures

Bancroft GM (1974) Mossbauer spectroscopy: an introduction for inorganic chemists and geochemists. McGraw-Hili Maidenhead

Wertheim GK (1964) The Mossbauer effect. Principles and applications. Academic Press, New York

Hawthorne FC (1988) Mossbauer spectroscopy. Rev Mineral 18: 255-340 Amthauer G, Annersten H, Hafner SS (1976) The Mossbauer spectrum of 57Fe in silicate

garnets. Z Kristallogr 143: 14-55 Schwartz KB, Nolet, DA, Burns RG (1980) Mossbauer spectroscopy and crystal chemistry of

natural Fe-Ti garnets. Am Mineral 65: 142-153 Virgo D, Hafner SS (1970) Fe2+, Mg order-disorder in natural orthopyroxenes. Am Mineral

55: 201-223 Bancroft GM, Williams PGL, Burns RG (1971) Mossbauer spectra of minerals along the

diopside-hedenbergite tie line. Am Mineral 56: 1617-1625 Hafner SS, Huckenholz HG (1971) Mossbauer spectrum of synthetic ferridiopside. Nature

233: 9-11 Aldridge LP, Bancroft GM, Fleet ME, Herzberg CT (1978) Omphacite studies, II, Mossbauer

spectra ofC2/c and P2/n omphacites. Am Mineral 63: 1107-1115 Hafner SS, Ghose S (1971) Iron and magnesium distribution in cummingtonites

(Fe,MghSis0 22(OHh. Z Kristallogr 133: 301-326 Goldman DS, Rossman GR (1977) The identification of Fe2+ in the M(4) site of calcic

amphiboles. Am Mineral 62: 205-216 Bancroft GM, Brown JR (1975) A Mossbauer study of coexisting hornblendes and biotites:

quantitative Fe3+ /Fe2+ ratios. Am Mineral 60: 265-272 (1975) Finch J, Gainsford AR, Tennant WC (1982) Polarized optical absorption and 57Fe Mossbauer

study of pegmatitic muscovite. Am Mineral 67: 59-68 Dyar MD, Burns RG (1986) Mossbauer spectral study offerruginous one-layer trioctahedral

micas. Am Mineral 71: 955-965 Amthauer G, McIver JR, Viljoen EA (1979) 57Fe and 119Sn Mossbauer studies of natural tin

bearing garnets. Phys Chem Mineral 4: 235-244 Amthauer G (1986) Crystal chemistry and valences of iron, antimony and tin in franckeites.

Neues Jahrb Mineral Abh 153: 272-278 Wagner FE, Marion Ph, Reynard J-R (1988) A 197 Au and 57Fe Mossbauer study of the

roasting of refractory gold ores. Hyperfine Interactions 46: 681-688

3.2.1 Parameters in Different Types of X-Ray Spectra 87

3.2 X-Ray and Photoelectron Spectroscopy of Minerals

3.2.1 Parameters in Different Types of X-Ray Spectra

D.S. UReH

In all types of X-ray spectra, X-ray photoelectron (XP), X-ray absorption (XA), and X-ray emission (XE), there is a common primary step, the ejection of an electron from a core orbital. This can be achieved by irradiation with photons, electrons, or high energy particles. The parameters that determine the aspects of these different types of spectra will be discussed below.

X-Ray Photoelectron Spectra

For the photoelectron effect to yield spectra of any importance it is essential to use monochromatic radiation (hv) so that the basic equation, hv = EK + Eb + () can be applied (EK = kinetic energy of ejected photo electron; Eb = binding energy of electron before ejection; () a correction term which depends upon sample work function, sample charging, etc.). The efficiency (and the peak intensity in the spectrum) with which an electron will be ejected will depend upon the spatial overlap of the core orbital from which it came and of the plane wave of the photoelectron. The critical factor in determining the magnitude of this integral is the "effective wavelength" of the orbital (i.e., the radial function) and of the photoelectron. This latter wavelength A.K will be related to EK by the De Broglie relationship A.K(A) = [150/EdeV)]1/2. For the core orbitals, whose radial functions give effective wavelengths of tens of picometers, photoelectrons ejected by Al or Mg Koc X-rays will have energies of hundreds of eV, corresponding to wavelengths of tens of picometers, but valence shell orbitals will correspond to longer wavelength functions, (hundreds of picometers), whilst the ejected photoelectron will have an energy greater than a 1000 V and so a wavelength of less than 40 picometers. Thus more intense peaks are seen in XP spectra from core orbitals than from valence shell orbitals. This argument can be extended to show that, within a given quantum shell, the most intense XP peaks are usually associated with the largest momentum quantum number.

Other factors than can affect XP spectrum ccncern additional processes during excitation and the interaction of unpaired electrons with core or valence band vacancies. The former can give rise to extra peaks in XP spectra, on the low kinetic energy side of the main photo peak. The energy brought to the impacted atom or ion can be used not only to eject a specific photoelectron, but also to excite, even to ionize, one or more of the other electrons. The energy required for such excitation (shake-up) or ionization (shake-off) is taken from


the kinetic energy of the photoelectron, hence peaks on the low energy side of the main peak. This is a simple description of an aspect of configurational interaction which can in some extreme cases lead to XP peaks being completely displaced from their anticipated positions. Of more direct interest in mineralogy is the "shake-up" structure which is present in transition metal ions in some valence states. The 2P1A, 2P3!~ peaks of Cu (II) are, for example, invariably associated with very intense ( '" 60% intensity of main peaks) shake-up/off peaks separated from the main peaks by 6 '" 7eV. Such satellites are not observed for Cu (I).

X-Ray Absorption

Whilst in the photo electron process attention is focused upon the ejection of a core electron from an atom, in X-ray absorption the excited electron is retained by the atom in a bound state. The interest in this type of spectroscopy lies in determining the energies of such excitation processes, and their relative probabilities. In the simplest experiments, absorption is measured as a function of incident photon energy. The intensity of absorption will be clearly determined by the spatial overlap of the core and virtual orbitals and by the electro magnetic transition operator. For long wavelength absorption (say A. > 3 A) this latter term will correspond to the electric dipole vector giving rise to the selection rule ,11' = ± 1 where I' is the angular momentum quantum number. Furthermore, as core orbitals have a very limited spatial extent the integrals that determine the intensity of absorption will themselves be determined by the local density of states of the virtual (molecular) orbitals. XA spectra thus probe particular aspects of empty orbitals in a molecule or in the band structure of a mineral: such spectra will be both atom-specific and orbital type-~pecific and will compliment VXE spectra (see below) in determining electronic structure. XA spectra of this type which are observed up to the absorption edge are known as NEXAFS (Near Edge X-ray Absorption Fine Structure) or XANES (X-ray Absorption Near Edge Structure). Similar fine structure can also be observed in Electron Energy Loss (EEL) spectra in which the discrete energy losses suffered by a beam of monoenergetic electrons passing through a sample are measured. In this case the electric dipole selection rule does not operate and so a comparison of EEL and XA spectra can indicate which absorption processes involve transitions other than ,11' = ± 1.

Rather surprisingly, variations in the intensity of X-ray absorption are observed on the high energy side of the absorption edge. This is caused by the back scattering of the ejected photoelectrons by neighboring ligand atoms. Clearly this will only happen when the wavelength of the photoelectron corresponds to a specific bond distance, but when it does the back scattered electron can be "re-absorbed" by the emitting atom. It is as though the photoelectron had never been ejected, the initial photon not absorbed. Thus absorption at that particular energy is reduced. This modulation of X-ray

3.2.1 Parameters in Different Types of X-Ray Spectra 89

absorption with frequency gives rise to structure that typically extends to a few hundered eV above the absorption edge: extended X-ray Absorption Fine Structure (EXAFS). A detailed study of EXAFS can thus provide information about the local environment of each atom in a solid. Whilst for simple crystals this duplicates structural data determined by X-ray crystallography, EXAFS can also provide information about amorphous, glassy and noncrystalline solids, it can also be used to study the local environment of impurity atoms or trace elements which are distributed at random throughout a lattice. EXAFS is therefore of great potential value to mineralogists.

X-Ray Emission Spectroscopy (XES)

Despite the similarity in name to X-ray absorption spectra, XE really has more in common with XP spectroscopy. XES studies one of the relaxation processes open to the ions made in XPS.

The parameters that determine the intensity of X-ray emission can be summarized as:

1. probability of creating a core vacancy, 2. probability of relaxation by X-ray emission, 3. overlap of initial and final state wave functions.

Core Vacancy. The basic factors that determine the creation of a core vacancy have been discussed above under XP spectroscopy. The situation is, however, more complicated in most XE experiments because bombardment with mono chromatomatic radiation (or mono energetic electrons) is not necessary, Thus, XE is usually initiated by radiation from an X-ray tube, a mixture of characteristic lines and bremsstrahlung, or by electron bombardment. In both cases energy is lost in penetrating the sample so that the probability of ejecting a specific type of electron from its orbital will vary widely with depth. Very roughly it can be estimated that the maximum efficiency for core hole creation is at about three times the ionization energy.

X-Ray Emission - Auger Electron Emission. The ions with core vacancies can relax in two possible ways, by X-ray emission or Auger electron emission. It is found that for relaxation energies in excess of 5-10 keV X-ray emission predominates, and that the X-ray emission increases with energy. For low energies, however, the converse is true; if the excitation energy is only a few hundred volts [e.g., C (ls-l)+], then more than 99% of the relaxation events lead to the ejection of an Auger electron.

Initial and Final State Wave/unction Overlap. The intensity of an XE peak will be determined by an integral of the type Jt/Ir P t/li and t/lr are the wavefunctions for the initial and final states respectively and P is the transition operator. This


operator will be dominated by the electric dipole term for energy differences between t/li and t/lc that correspond to X-ray emission with a wavelength greater than a few atom diameters (A. > 6 A, i.e., AE (t/li' t/lc) < 2000 eV). Under these conditions At = ± 1, but for shorter wavelengths "forbidden" quadrupole, and magnetic dipole lines, etc. can also be observed. Line width is determined by lifetime ofthe shortest lived excited state. As spontaneous relaxation probability is related to the cube of the frequency ofthe emitted radiation, it follows that the larger AE (t/li' t/lc), and the shorter the wavelength of the emitted radiation, the broader will be the X-ray peak. Conversely, for long wavelength XE very sharp peaks should be observed, [Ar x r - 10- 16 where r is peak width in eV and Ar the lifetime of the excited state in seconds]. The lifetime of those states where AE (t/lb t/lc) is small (say < 2000) eV is, however, mostly determined by Auger decay rather than by X-ray emission, even so soft X-ray emission spectra (e.g., N KtX) are characterized by natural line widths of the order of 1/10 eV.

Most XE spectra are generated by electronic relaxation processes between core orbitals. The parameters that determine relative intensity are those that control the magnitude of the overlap of t/I i and t/I c; i.e., their radial function. Thus in relaxation to a 1s- 1 state the intensity will be KtX (ls-I--+2p-l) > KP1,3 (is -1 --+ 3p -1) > KP2 (is -1 --+ 4p -1), etc. When relaxation is from a valence shell orbital, then valence X-ray (VXR) spectra result. In principle any valence shell atomic orbital (a-o) can participate in molecular orbital (m-o) formation. However, it can be shown that since the core hole to which relaxation takes place is well localized on a particular atom, VXR spectra reflect the local density of states. The fine structure observed in VXR spectra is a direct indication of the extent to which specific atomic orbitals on the emitting atom participate in molecular orbital formation and the relative intensities of the component peaks measure the extent of this participation in each m.o. If the ionization energy of the initial (core hole) state is known then the energy of each molecular orbital can be calculated, E(m,o) = E(core hole) - hv (X-ray line).

VXR spectra can thus be used as direct probes of electronic structure. An s vacancy on atom A will attract transitions from A, p orbitals and the VXR spectrum will indicate the participation of valence shell p orbitals from A in the molecular orbitals of the molecule or ion in which A is bound. Similarly, transitions to a p vacancy (on A) would generate VXR spectra that would, in turn, indicate the bonding roles of valence shell sand d orbitals from atom A.

Complications arise in open-shell structures such as minerals which contain transition metal ions. The spin state of the ion can interact with the initial state core vacancy and, more importantly, with the final state, causing peak shifts and peak splitting. The most thoroughly investigated, (but not necessarily best understood) spectra are the KP1.3 - KP' peaks from compounds of first transition metal ions. In some cases it has proved possible to relate valency to the relative intensity of the satellite peak (KP') to the main peak (KP1,3)' This works well for Cr and Mn, but not for Fe. In the case of manganese (II) it is also possible to relate the structure observed Mn, 3p XP spectrum with that of the Mn KP1,3,Kp'.

3.2.2 Information from X-Ray Absorption Spectroscopy 91

Other features that are observed in X-ray emission spectra involve relaxation processes in multiply ionized atoms. In K spectra there are usually high energy satellites but in L spectra high and low energy features have been identified. The simplest peaks of this type are the so-called K<X3.4 satellites generated by 1 s -1, 2s - 1 or 2p - 1 -+ 2s - 1 2p - 1 or 2p - 2. If the initial irradiation is by electron or photon bombardment, these satellites have an intensity only a few percent of the main peak. High energy ionised particle bombardment however greatly increases their relative intensity, so that sometimes the satellite peaks (K<X3.4) can be more intense than K<X 1.2'

Often a study of the fine structure of high energy satellites can yield chemically interesting information, e.g., the Al K<X3,4 spectrum from alumina is distinctly different from that of aluminum metal and can therefore be used analytically to identify oxide layers.

References

Agarwal BK (1979) X-ray spectroscopy. Springer, Berlin Heidelberg New York Azaroff LV, Pease DM (1974) X-ray absorption spectra. In: Azaroff LV (ed) X-ray spectro

scopy. McGraw-Hill, Montreal, Canada Fadley CS (1978) Basic concepts in X-ray photoelectron spectroscopy. In: Briindle CR, Baker

AD (eds) Electron spectroscopy - theory, techniques and applications. Academic Press, New York, 2:1

Price WC, Potts A W, Streets DG (1972) The dependence of photoionization cross-section on the photoelectron energy. In: Shirley DA (ed) Electron spectroscopy. North-Holland Publ Co, Amsterdam, Netherlands, p 187

Siegbahn K, Nordling C, Fahlman A, Nordberg R, Hamrin K, Hamrin J, Johansson G, Bergmark T, Karlsson S-E, Lindgren I, Lindberg B (1967) ESCA, Atomic molecular and solid state structure studied by means of electron spectroscopy

Nova Acta Reg Soc Sci Upsaliensis ser. IV 20 (1) Teo BK, Joy DC (1981) EXAFS spectroscopy, Plenum Press, New York Thompson M, Baker MD, Christie A, Tyson JF (1985) Auger electron spectroscopy, J Wiley,

London Urch DS (1988) PAX (photoelectron and X-ray spectroscopy): basic principles and chemical

effects. In: Gomes Ferreira J, Teresa Ramos M (eds) X-ray spectroscopy in atomic and solid state physics. Plenum Press, New York

3.2.2 Mineralogical and Geochemical Information from X-Ray Absorption Spectroscopy

A. MANCEAU and G. WAYCHUNAS

A great deal of chemical and structural information can be obtained from X-ray Absorption Spectroscopy (XAS). This includes: electronic structure, local order of ions in the bulk or at the surface of minerals, site locations, intersite and intracrystalline distributions of trace/minor elements, and local structure of


non- or poorly-crystalline materials. This diversity of information explains the large number of mineralogical and geochemical applications that have been made. These applications have been extensively reviewed in the last few years, and recent compilations ofresults can be found in Brown et al. (1988) and Calas et al. (1990). This chapter below synthesizes representative examples classified according to some of the main geochemical processes the mineralogist confronts: nucleation and crystal growth in both homogeneous and heterogeneous conditions, structural chemistry of non-, poorly-, and well-crystalline materials, sorption and speciation of chemical species at solid/water interfaces, and the nature of dissolution phenomena.

Nucleation Processes

Because of its sensitivity to short range order, XAS is a suitable method to follow structural changes during nucleation phenomena. Classically, two kinds of nucleating processes are distinguished: homogeneous and heterogeneous.

Homogeneous Nucleation. Little is known about the local structure of soluble metastable species derived from the partial hydrolysis of cations in aqueous media. However, it is generally considered that these minute phases are important components of natural solutions. Owing to their high surface reactivity, these species also playa key role in the geochemical cycle of many trace/minor elements. Because of the abundance of ferric iron in surficial aqueous systems, a large amount of work has been devoted to elucidate the local and longer-range structure of Fe polymers obtained through the hydrolysis of ferric salt solutions. This case study provides a good example of the possibilities offered by XAS spectroscopy, particularly when complemented by other techniques.

XAS studies have shown that Fe (III) remains six-fold coordinated up to the completion of hydrolysis. Low weight Fe polymers are well ordered at the local scale (5-10 A), and posses a local structure similar to that in aFeOOH/ pFeOOH (Fig. 31a). This arrangement is preserved until the completion of hydrolysis. Changes in the longer-range structure of Fe polycations during hydrolysis have been investigated by small angle X-ray scattering (Fig. 31b and c). Fe colloids are made up of aggregates of subunits, the diameter of which is 16 A. The structure of these colloids formerly depends on the hydrolysis ratio, i.e., NaOH/Fe ratio. Low weight polymers are formed from linear aggregates whose stability originates from long-range magnetic dipolar interactions. With increasing hydrolysis, the sticking of linear polycations leads to aggregates with a branched shape. As the fractal dimension of the aggregates increases, reaching two by the end of hydrolysis, a hydrous ferric oxide is formed (HFO). These two complementary structural approaches have confirmed the existence of a structural continuity all along the hydrolysis process between aqueous Fe species and ferric gels. This finding is at variance with the hydrolysis of aluminum where a transient AI13 species precedes the formation of gibbsite (Bottero et al. 1987).


A

:;:::::: ~

:. ~ u.:

2 3

B P(R)

15

10

5

20 40 60

c P(R)

500

400

300

200

100

o 300 600 900 1200 R (A)

Fig. 31. Homogeneous nucleation of hydrous ferric oxide from a ferric chloride solution. A XAS radial distribution function (uncorrected for phase shift) showing changes in the short-range order around iron during hydrolysis, and structural model of the local structure of Fe polymers. Band C Small-angle X-ray scattering-derived distance distribution function P(R) of Fe colloids at different hydrolysis ratio. B NaOH/Fe = 1.0; C NaOH/Fe = 2.6. (After Combes et a\. 1989,

Tchoubar et a\. 1991)

When the freshly precipitated HFO is aged in solution at 92 °C for 2-6 h, a new phase, not discerned by X-ray diffraction, has been detected by XAS (Combes et al. 1990). This transient product is characterized by the appearance of face-sharing octahedra, and possesses a <5FeOOH-like local order. It is apparently an intermediate step in the conversion of non-aged HFO into hematite, and thus catalyses the HFO -+ hematite transformation. At the macroscopic level, the existence of face-sharing octahedra in the aged gel could


substantiate the dissolution-reprecipitation mechanism often invoked to explain the formation of goethite from natural ferric gels (Schwertmann and Murad 1983) inasmuch as the local structure of this transient phase has been recognized in natural gels (Manceau and Combes 1988; Manceau et al. 1992a).

Heterogeneous Nucleation. Under the thermodynamic and kinetic conditions that prevail at the surface of the Earth, nonequilibrium processes are the rule rather than the exception. Heterogeneous nucleation is one of these, and for a complete understanding of its effect on the macroscopic properties of solids, the aspects of this phenomenon on the molecular-structure level must be known. This sort of information allows a better prediction of phase stability in natural systems because the stability fields of most low temperature materials have been determined from equilibrium studies (Tardy and Nahon 1985; Trolard and Tardy 1987, and references therein). Only one study of this type has been achieved to date by XAS, and concerns the heterogeneous nucleation of hydrous chromium oxide (HCO) onto HFO (Charlet and Manceau 1992). Chromium was chosen because of its high affinity for ferric oxides, which are used in water treatment plants and in analytical chemistry to remove chromium from dilute solutions. Cr-containing Fe oxides are also constituents of soils and weathering profiles (Stucki et al. 1988). The removal of Cr(III) by HFO is achieved either through the formation of surface complexes on already formed Fe gels, as in an adsorption/surface precipitation, or through the simultaneous hydrolysis of Fe(III) and Cr(lll), as in coprecipitation. These two mechanisms differ by the solubility of the Cr-rich end-products. At high Cr /Fe content of the solid phases, the coprecipitate solubility remains an order of magnitude lower than the solubility of a surface precipitate. Furthermore, the solubility of the latter is equal to that of HCO synthetized in the absence of HFO, i.e., homogeneous precipitation. These differences in macroscopic behavior have been related to distinct local structures, as derived by EXAFS spectroscopy. In coprecipitation, Fe and Cr form an IXFeOOH-IXCrOOH solid solution, whereas the surface precipitate possesses a yCrOOH-like local structure as does HCO precipitated in the absence of substrate, i.e., homogeneous precipitation (Fig. 32). Thus, a difference in structure (IX- vs y-CrOOH) accounts for the difference in solubility between the coprecipitated Cr-rich phase and the surface precipitate, and thus a difference in the removal efficiency of the two processes.

Structural Chemistry of Non- and Poorly Crystalline Materials

Structure of Fe and Mn Hydrous Oxides. Hydrous oxides, such as those of iron, aluminum, manganese, and silicon, are abundant in nature. They can sorb large concentrations of chemical species because of their tendency to be finely dispersed. They were long referred to as "amorphous" products, as their X-ray diffraction patterns display few and ill-defined reflections. In their most disordered state, HFO and hydrous manganese oxides (HMO) XRD patterns


LA TrICE SUBSTITUTION

Cr (llI) for Fe (III) isomorphic ~UU'~Ul'UU~'"

HETEROGENEOUS PRECIPITA nON

Multinuclear yCr 0 0 H-like surface complexes

Fig. 32. Fe(III)-Cr(III) coprecipitation vs. Cr(III) adsorption onto hydrous ferric oxide (After Charlet and Manceau 1992). In the former mechanism, Fe and Cr atoms form an IXFeOOHIXCrOOH solid solution, and in the latter, Cr build multinuclear yCrOOH-like surface complexes

consist of two broad hk bands at 2.5- 2.6 A and 1.5 A. The lack of long range order causes concern about the uniqueness of structural determinations by XRD (Drits et aI., 1993). However, the local structure of Fe and Mn gels has been successfully derived from XAS thanks to a polyhedral approach, which has provided a unified framework to rationalize data on poorly-crystallized Fe and Mn oxides (Manceau and Combes 1988; Manceau and Drits, 1993). The polyhedral approach to the structure of ferric and fourvalent Mn oxides is based on the recognition that interpolyhedral bond angles depend uniquely on the way in which octahedra are linked to each other. Hence, each structure can be readily identified by the knowledge of the two or three nearest metal-metal distances, and the coordination number of metal atoms at each of these distances (Fig. 33).

From XAS measurements HFO and HMO appear to have well-defined local structure, and their "amorphous" character as judged by XRD arises only from the small size of coherently scattering domains. The local structure around iron is dependent on formation conditions. In synthetic gels precipitated from Fe2 + -

containing solutions, Fe06 octahedra were found to be coupled as in y-FeOOH; however, in gels precipitated from Fe3 + - containing solutions, the local structure around iron is similar to that in (1.- or {J-FeOOH, depending on whether the counterion is, respectively, nitrate or chlorine (Combes et ai. 1986, 1989). However, the most interesting result is that natural Fe gels instead possess a bFeOOH-like local structure, as do thermally aged synthetic HFO (Manceau

crys

tal

stru

ctur

e po

lyhe

dral

app

roac

h

AK

AG

AN

EIT

E ~lt9

I ~ ~~".A

~ ~FeOOH

b 1

2 3

" {

'5 R

(A)

'C>

0

\

GO

B11

lfm

~

I. _

EI A

w;5 t

o:Pe

OO

H

~\/\ 7

'{ jl

7'\W.-~

'

-\.7\~

,/ ~

II

\

4'5R

{A.)

I ..

..

(l;>~

. ;::::

~-I ~

HE

MA

TIT

E

r"'~

t*-..:

>I !

-<1 "

..

('~..

..llii

~

~

(J

t ir

O:P

ep3

~\

r ~W

~ J----~~'---.--"--

R (A

) ~

01

23

45

~

w

5:

>C

CIl ~

<>

yFeO

OH

'(I~~~

~DTY

~EY.7

gll

\/\/\~

! o '" n o "0

'<


et al. 1992a, Manceau and Drits, 1993). The physical and chemical parameters which govern the formation of these hydroxide gels have thus a direct influence on the resulting mineralogy by control of the local structure. Another example of XAS results is the structure of the Mn gel vernadite (oMn02), which can be viewed as a three-dimensional anionic framework of cubic and hexagonal anionic packings distributed at random (absence of hkl reflections). In this structure each pair of close packed anionic layers possesses a very similar cation content (absence of basal reflection), and Mn06 octahedra are linked to each other through edges and corners (Chukhrov et al. 1988, Manceau et al. 1992a).

Structure of Silicate Glasses and Melts. The study of the local structural environment of cations in glasses and melts has been one of the major applications ofXAS, and several reviews have been published on this subject (see e.g. Calas et al. 1987; Greaves 1985). Most of the XAS studies that have been performed focused on the coordination geometry of poorly characterized transitional metal ions in silicate glasses, although actinides, alkali metals, and network-forming Ge4 + have also been investigated. As examples here we consider the XAS results for Ti4 + and Fe2+ in silicates. Ti4 + usually occurs in octahedrally coordinated sites in silicates and oxides, but can occupy tetrahedral sites in rate situations (Waychunas 1987). Its structural role in silicate glasses was problematical until the studies by Sandstrom et al. (1980) and Gregor et al. (1983) on Ti02-Si02 glasses. These workers showed that the Ti4 + occupied mainly tetrahedral sites at low concentrations, but quickly saturated these sites and became mainly octahedrally coordinated with significant Ti4 + content. In contrast, Yarker et al. (1986) showed that Ti4 + in K20-Ti02-2Si02 composition glasses occupied square pyramidal sites. Still other studies of Ti4 + in cordierite glasses (Dumas and Petiau 1986) and spodumene glass (Ramos et al. 1985) suggest four or five coordination. Hence the coordination environment of Ti4 + in silicate glasses tends to differ from most silicate crystals.

Another transition metal ion which has been difficult to characterize in glasses is Fe2 +. Mossbauer studies have long showed that presumably octahedral Fe2 + in alkali silicate glasses has an uncharacteristically low isomer shift, suggesting either very short Fe2 + -0 bonds, or a smaller coordination number. Recent XAS studies have shown that this low isomer shift is indeed correlated with short metal-oxygen distances and reduced average coordination numbers (Waychunas et al. 1988), suggesting that Fe2+ can actually be, at least in part, a network-former in some silicate melts.

Silicate melts have resisted direct structural investigation because of the obvious problems of sample containment. However it is crucial to determine the melt structure at high temperatures, because the corresponding glass actually

Fig. 33. Polyhedral approach of the structure of Fe oxides. The structure of 0(-, p-, y-FeOOH, IX-Fe203 can be differentiated on the local scale by EXAFS owing to their distinct octahedral linkages and/or Fe-Fe distances. (After Manceau and Combes 1988, Manceau and Drits, 1993)


only records the melt structure at the glass-transition temperature, which may be hundreds of degrees lower than the fusion temperature. The fist XAS studies of transition metals in silicate glasses were done by Waychunas et al. (1988) where Fe2 + in Na2 FeSi30 s and K2 FeSi30 s melts at temperatures up to 1200 K was probed. The Fe-O distances in these liquids was 1.94-1.96 A with an average coordination number of 4. Pre-edge Is -+ 3d features were intensified consistent with this non-centrosymmetric coordination. The quenched glasses had longer Fe-O bonds of 1.99-2.02 A, but similar coordination numbers. Since these studies, Jackson et al. (1991) have been able to examine the environment of Fe2 + in fayalite melt at near 1600 K. Their work also indicates four and possibly five coordinated Fe2 +. Because of these investigations, the nature of ironbearing silicate melts in the lower crust and upper mantle may need to be rethought.

Structural Chemistry of Minerals

Our knowledge of the crystal chemical behavior of trace and minor elements is extremely limited because very few structural techniques yield information on atom environments at high dilution. However, EXAFS spectroscopy has emerged as one of the most powerful of these methods as it is capable of assessing the local symmetry, site occupation, and intracrystalline distribution of very dilute species. The manganese oxide lithiophorite illustrates this point. This mineral possesses a mixed Mn02-AI(OHh layered structure, and is known to contain various 3 d elements (Wadsley 1952; Pauling and Kamb 1982). Owing to the high Z contrast between Al and Mn as X-ray backscatterers, EXAFS spectroscopy is able to distinguish between these two cations as second nearest neighbors of substituting 3 d elements. Thus Ni and Cu have been found to be located in the aluminous sheet in lithiophorite, whereas Co, which is trivalent and in a low-spin electronic configuration, either randomly substitutes for Mn or segregates in CoOOH domains possessing a phyllomanganate-like local order (Fig. 34). Other examples of site location determination concern REE ions in epidote (Cressey and Steel 1988), and Ge4 +, which has been found to substitute for Fe in goethite and hematite with appropriate local charge-coupled substitutions (Bernstein and Waychunas 1987).

More interesting however is that EXAFS studies have demonstrated that the actual structure of minerals at the 2-5 A scale can differ from the average structure derived from diffraction-based techniques. It is now known that most minerals, and more especially those formed at low temperature, possess heterogeneous structure on a very fine scale. Three types of structural situations have been identified. (1) A simple deviation from the random distribution of atoms in a solid-solution. Examples include: Fe3+ -doped MgO (Waychunas 1983); NiFe-Mg phyllosilicates (Decarreau et al. 1987); Ni-Mg and Fe-Mg phyllosilicates (Manceau 1990). (2) Existence of a discrete phase intimately mixed with the major one, e.g., Ni(OHh in asbolane (Chukhrov et al. 1980; Manceau et al.


eo 3+ 0

Mn02~OMn4+

CU 2+ Ni 2+ 0 H AI(OH)3~OHAI 3+

C0 3+ 0 Mn02~OMn4+

Ni 2+ CU 2+ OH

AI(OH)3~OHAI 3+

LITHIOPHORITE

Fig. 34. Location of Co, Ni, and Cu within the structure of lithiophorite. Co is trivalent and in a low-spin configuration. Whether Co is randomly substituted for fourvalent Mn or segregated in CoOOOH domains possessing a phyllomanganate-like local order is not yet firmly established. (After Manceau et al. 1987, Manceau et al. 1990a)

g oethite framework

o O,OH • Mn

Phyllomanganata cluster.

C Fe

Fig. 35. Structural model of natural Mn-containing goethite. Hexagonal anionic close-packed oxygen arrays contain Fe(III) and Mn(IV) domains with a goethite-like and a phyllomanganate-like structure, respectively. (After Manceau et aI., 1992a)

1987). (3) Intergrowth of two distinct phases, e.g. bFeOOH in bMn02 (Manceau and Combes 1988; Manceau et at. 1992a), phyllomanganate clusters in aFeOOH (Fig. 35), and aFe20 3-like clusters in aAIOOH (Fig. 36).

Even though many studies have assessed the usefulness of EXAFS spectroscopy for probing the short range order (SRO) of minerals, difficulties are often encountered in accurately interpreting the EXAFS spectra of complex structures. These complications must be emphasized as they have largely contributed

100

a

b


Hematite-like cluster in diaspore

o AI 2,59~

Formation mechanism of multinuclear Fe complexes

Fig. 36. a Structural model of natural Fe-containing diaspore as determined from P-EXAFS. Fe atoms are segregated within the diaspore framework, and build I1Fe203-like clusters in the direction of channels ([OIOJ). b Mechanistic interpretation of the formation of Fe multinuclear surface complexes at the time of diaspore growth. (After Hazemann et al. 1992)

to the restriction of XAS applications to relatively simple mineral structures. Among the major limitations are: (1) the difficulty in discriminating elements with similar atomic numbers; (2) complications created by many differing site occupations of the X-ray absorber; (3) the necessity of having well-known samples from which phase shifts, amplitude functions, and Debye-Waller parameters can be extracted; (4) texture effects; (5) the overlap between the EXAFS contributions of the first two or three first cationic shells. This last drawback is especially severe in silicates where sites of different nature and of different chemical composition are linked to each other, e.g., tetrahedra and octahedra filled by Si, AI, Mg, and transition elements. The non-uniqueness ofleast squares


fits to the EXAFS spectra explains why the application of EXAFS spectroscopy to complex primary silicates like pyroxenes is still largely limited to the analysis of the first coordination shell (Waychunas et al. 1986; Waychunas and Brown 1990). One way to overcome this difficulty is to take advantage of the plane polarization of the synchrotron beam and selectively collect information along particular directions in the sample (Manceau et al. 1988, 1990b; Waychunas and Brown 1990; Hazemann et al. 1992). The angular dependence of the K-shell EXAFS, X, can be written in the following way

NJ

X(k,O) = L L 3 cos2 (O\)· X lso (k), j i= 1

where the index j runs over all the shells around the absorbing atom, the index i runs over all the N j atoms ofthej shell, O~ is the angle between the electric vector e and the vector r~ that binds the absorbing atom to the ith atom of the jth shell, and Xiso holds for the powder. Unlike isotropic EXAFS, which yields a onedimensional map (modified correlation function) of the local structure about the target atom, a three-dimensional map can be constructed by measurements with an orientable single crystal in a polarized X-ray beam. Such measurements give access to directional structural and chemical information. This selectivity permits one the separation of individual atomic pair contributions in favorable situations, and offers the possibility of unraveling overlapping cationic contributions (Manceau 1990). Furthermore, when e, is parallel to a given absorberneighbor bond, the back scattering amplitude is three times enhanced. Thus polarized EXAFS measurements make feasible the detection of weaker contributions, than would be unperceived with powder spectra. Given the low symmetry and multielement composition of most minerals, polarized EXAFS allows inroads towards a more complete analysis of their local structure. This technique has been used together with proton and fluorine NMR to investigate the octahedral distribution of Fe in trioctahedral micas. In these minerals it has shown that F is exclusively associated with Mg, while OH can be bonded to Fe or Mg ions. In addition to this Fe-F avoidance, both F and Fe tend to segregate, resulting in a non-random distribution of Fe and Mg in the octahedral sheet (Sanz and Stone 1977, 1983a and 1983b, Manceau et al. 1990b).

When applied to chain frameworks, polarized EXAFS enables us to look at the site location and distribution of atoms within and in between octahedral chains. In addition, in the case of multiple octahedral chains, as exist in e<FeOOH and analogous structures, it is even possible to filter the contributions of the nearest cations along the single chain wherein the X-ray absorber resides, form those of cations located in the adjacent single chain. Because of the enhancement in the sensitivity and accuracy of structural data in a polarized experiment, such studies give access to an unequalled three-dimensional description of the local structure of chain-minerals. This technique is especially interesting in the case of trace elements, since such a description of their local structure may not be accessible by other methods. In a certain sense, these EXAFS experiments can be compared with ALCHEMI (atom location by


channelling enhanced microanalysis, Tafto and Buseck 1983). In this TEMbased method the site location of trace elements in minerals can be identified by taking advantage of the channelling effects of ineiastically scattered fast electrons. The only polarized EXAFS experiment of the type described has been conducted on a Fe-containing diaspore, and has revealed unexpected results (Hazemann et al. 1992). Instead of being randomly distributed, or segregated, within the ocAIOOH lattice, Fe atoms have been found in a mixed oc(AI,Fe)OOH and oc(AI,Feh03 microstructure. Part of the Fe(III) ions are segregated in aluminous octahedral chains, the others being located in the channels of the lattice (Fig. 36). The structure of these domains has been modeled assuming hematite-like clusters of three Fe octahedra to po tactically grown onto aluminous chains. These Fe clusters are viewed as ocFez0 3 nuclei formed by heterogeneous nucleation onto an ocAIOOH substrate. The formation process of these Fe clusters can be easily understood in the light of results on the heterogeneous nucleation and further crystal growth of yCrOOH or ocFeOOH (cf. "heterogeneous nucleation" section above). The Fe clusters are thought to represent ancient multinuclear Fe complexes that formed at the time of diaspore growth, and which are now sealed in its bulk structure. In addition to the well-known cases of atom segregations, ion vacancies, and intergrowths of discrete phases, this result provides a new example of a nonequilibrium crystallization processes.

Complexation Mechanisms at Mineral/Water Interfaces

The mobility, dispersal, and speciation of ions at the Earth's surface is largely controlled by their interactions with mineral surfaces. Indeed, as usually thought, it is not mineral solubility which controls most elements availability, but rather the adsorption and desorption equilibria in which they are involved. A detailed, molecular-level description of surface reactions is thus essential to an understanding of many geochemical processes. XAS spectroscopy is one of the very few methods capable of probing the environment of surface complexes at concentrations below the monolayer coverage. Furthermore, at the difference of many other spectroscopic or scattering methods, a large variety of systems can be studied in situ.

Given the recent emergence of this new field of research, very few studies have yet been performed (see, e.g., the review of Brown 1990). The first study of this kind appeared in the literature a few years ago, and concerned the sorption mechanism of selenium oxo-anions on ocFeOOH (Hayes et al. 1987). It was shown that selenite sorbs strongly as an inner-sphere complex, whereas selenate bonds weakly as an outer-sphere complex. This difference in structural behavior provided the first molecular-level explanation for the difference of affinity of selenite and selenate complexes as macroscopically measured with adsorption isotherms. This pioneering work has made inroads towards a comprehensive microscopic understanding of worldwide macroscopic studies of sorption phenomena.


Ferric gels have long been recognized as efficient scavengers of heavy metals. Among these, radionucleides have been extensively studied with respect to either exploration or environmental geochemistry. Manceau et al. (l992b) showed that U sorb as uranyl complexes. U bonds directly to surface functional groups as mononuclear bidentate complexes, i.e., by edge-sharing with Fe(O,OH)6 octahedra (Fig. 37). The overwhelming UO/ + affinity for HFO is explained at the molecular level by the similarity of edge lengths, (i.e., OH-OH distances), in Fe(O,OH)6 octahedra and of the uranyl equatorial coordination shell. This strong affinity thus arises from a quite perfect steric fit between the adsorbate and the adsorbent.

Dissolution

Many mineral dissolution processes are controlled by a chemical mechanism at the solid-water interface (Stumm and Wollast 1990). Although the application of solution coordination chemistry concepts has been successful in providing phenomenological models and pedagogic tools for interpreting solution chemistry experiments, very few macroscopic models have yet been validated at the molecular level. The only two present studies of this type concern the aqueous corrosion of a uranium-containing borosilicate glass (Greaves et al. 1989), and the dissolution of Mn02 by the coupled Cr(Ill) oxidation to Cr(VI) (Manceau and Charlet 1992).

fu.:

2

~:'0'\ : ~

......... OH

3 4 5 R (A)

Hydrous Ferric Oxide

~ ~. ~ .

• , • , .. I ..

...... :,0\/0 .. /0 .. .. . .. C . " • ' .'. ---" r . 0 '. 0 . 0 ': ..

l1li - - '. '. ' •••. ..-

Fig, 37. XAS radial distribution function (uncorrected for phase shift) of uranium sorbed on HFO (U LIII edge), and structural model of the sorption mechanism. Mononuclear uranyl species are adsorbed as inner-sphere complexes which are edge-linked with the Fe(O,OH)6 octahedra of the gel. (After Manceau et aI. , 1992b)


The leaching of uranium by water leads to the formation at the glass surface of hydrated uranyl units with incoherent U-O distances. The increasing splitting of the U-O distances has been logically interpreted as due to an ion exchange of sodium for protonated water. Prolonged corrosion results in the development of a surface structure resembling hydrated uranyl silicates. This in-situ XAS study confirms the qualitative scenario for glass corrosion obtained from analytical measurements. This commences with the release of alkali into solution, leads to the dissolution of the network with the formation of a gel layer, and culminates in surface mineralisation (Abrajano et al. 1990).

Kinetics studies of the Cr(lII) to Cr(VI) oxidation on Mn(IV) and Mn(III) oxides have identified Mn as the oxidizing agent (Eary and Rai 1987; Johnson 1991). These studies have also shown that the oxidation rate is independent of solution parameters such as pH, P02, and ionic strength, but is proportional to the solid suspension concentration. Whereas a surface controlled oxidation mechanism was demonstrated in this way, uncertainties remained as to the type of surface complex involved in the electron transfer reaction step. The independence of the oxidation rate from pH would argue in favor of an outer sphere complex (Johnson and Xyla 1991), whereas its independence from ionic strength indicates the formation of an inner sphere complex (Johnson and Xyla 1991; Hayes et al. 1988). EXAFS spectroscopy has been used to demonstrate that the second hypothesis was indeed the correct one, at least at the birnessite (Na4Mn14027· 9H 20) surface/ solution interface. The molecular mechanism of the oxidation process has been mechanistically sketched as follows (Fig. 38): (1) Cr(lll) aqua ions diffuse towards Mn(lV) vacancies in the lattice; (2) the coupled Cr(lII) oxidation/ Mn(IV) reduction takes place; (3) Cr(VI) and Mn(II) ions are released to the solution. This study has provided the first evidence of double solid-state diffusion, first towards, and then backwards from the sorbent. This mechanism is rendered feasible owing to a favorable steric fitting between

t = 0 s. t = 30 s . t = 2 min.

Cr QII)

Birnessite

Fig. 38. Molecular mechanism of the Cr(lll) oxidation to Cr(VI) by birnessite at different periods of solid-solution equilibrium times. From left to right. Step 1 Cr(III) ions are in the solution; step 2 Cr(III) ions have diffused into the octahedral vacancies of the birnessite; step 3 Cr(III) ions have been oxidized by Mn(IV) ions, and products of the reactions, i.e., Cr(VI) and Mn(II), have diffused into the solution. Birnessite layers are vartially dissolved. (After Manceau and Charlet 1992)


the active site of the sorbent and the sorbate, as both Mn(IV) and Cr(III) ions have similar ionic radii. In a certain sense, the vacant octahedral cages of birnessite act as a molecular sieve for Cr(III) free ions, because polynuclear Cr(III) species are too large to enter into reactive sites. Furthermore, a comparison of the oxidation rate on different polymorphs of Mn(IV) and Mn(III) oxides indicated that the reaction is speeded on two-dimensional as compared to three-dimensional substrates, and also when the concentration of active surface sites, i.e., Mn(IV) vacancies, increases.

Despite the limited number of studies as yet reported in the field of sorption reactions at mineral/water interfaces, XAS has largely proven its ability to decipher the mechanistics of elementary surface reactions. When coupled with aqueous chemical studies, these molecular-level investigations can provide a solid basis for discussing macroscopic measurements more confidently. More specifically, both uranium and chromium studies illustrate how the structure of the sorbent, through the nature of the reactive surface site, may control the dispersal and speciation of metals in the environment.

Towards the 21st Century

In the last few years great efforts have been made by Earth scientists to develop sample conditions which allow the direct study of natural systems. These improvements have made possible new research areas, three of which are noteworthy: bonding changes and phase transitions under high pressure, the structure of silicate melts at high temperature, and sorption reactions at mineral/water interfaces.

In situ high-temperature and/or high-pressure experiments are particularly necessary for the understanding of dynamic processes within the Earth's interior. More specifically, an understanding of the dependence of physical and chemical macroscopic properties of earth materials on molecular-level structure, chemical bonding, and composition derived from such experiments allows exacting models of earth-processes to be developed. Technical advances included the building for a controlled atmosphere furnace whose temperature can rise to 2000 K (Jackson et al. 1990), and the use of diamond anvil cells (DAC) capable of extremely high pressures. DAC measurements are performed either in a conventional transmission mode (Ingalls et al. 1978, Sueno et al. 1986) or by using a dispersive XAS spectrometer (ltie et al. 1989). The recent application of a DAC with a dispersive EXAFS system has been used to follow coordination changes at high pressures in Ge02 glass and crystal, which are models for understanding the behavior or magmatic liquids. A transition from four-fold to six-fold coordination has been observed near 70 Kbar in both solids. But the most interesting result is that this high pressure phase transition is reversible in the glass and not in the crystal. This result clearly identifies the limits of ex-situ experiments performed on quenched glasses, and outlines the necessity of performing in-situ experiments. Likewise, the study of sorption processes at


mineral surfaces has been possible owing to the availability of high count-rate fluorescence X-ray detectors and the installation of insertion devices like multipole wigglers, and undulators which greatly enhance X-ray flux.

Three third generation synchrotron storage rings are currently being built and will be put into use in the last 5 years of the 20th Century. In a chronologie order these are: ESRF, France; APS, USA, and Spring-8, Japan. These storage rings with their various types of synchrotron radiation sources such as bending magnets, wigglers, and undulators, will provide the scientific community with Xray beams of unprecedented properties. Gaining about 10 orders of magnitude in brilliance with respect to a rotating anode X-ray tube, and 3 to 4 orders with respect to the most brilliant existing synchrotron X-ray sources, seems to be a peine croyable (hardly believable) to the "experienced experimenter". All of the present-day XAS experiments will benefit from this tremendous increase in brilliance, but in addition new research opportunities will undoubtedly emerge. For instance, it will permit time-resolved XAS studies of transient phenomena, such as solid-solid phase transitions and the dynamics of many geochemical processes.

References

Abrajano TA, Bates JK, Woodland AB, Bradley JP, Bourcier WL (1990) Secondary phase formation during nuclear waste-glass dissolution. Clays and Clay Minerals 38: 537-548

Bernstein LR, Waychunas GA (1987) Germanium crystal chemistry in hematite and goethite from the Apex Mine, Utah, and some new data on germanium in aqueous solution and in stottite. Geochimica et Cosmochimica Acta 51: 623-630

Bottero JY, Axelos M, Tchoubar D, Cases JM, Fripiat JJ, Fiessinger F (1987) Mechanism of Formation of Aluminium Trihydroxide from Keggin Al 13 Polymers. Journal of Colloid and Interface Science 117: 47-57

Brown GE Jr (1990) Spectroscopic studies of chemisorption reaction mechanisms at oxidewater interfaces. In "Mineral-Water Interface Geochemistry", M.F. Hochella and A.F. White Ed., Reviews in Mineralogy 23, Mineralogical Society of America, Washington DC, pp 309-364

Brown GE Jr, Calas G. Waychunas GA, Petiau J (1988) X-ray absorption spectroscopy: applications in Mineralogy and Geochemistry. In "Spectroscopic Methods in Mineralogy and Geology", F.e. Hawthorne Ed., Reviews in Mineralogy 18, Mineralogical Society of America, Washington DC, pp 431-512

Calas G, Brown GE Jr, Waychunas GA, Petiau J (1987) X-ray absorption spectroscopic studies of silicate glasses and minerals. Physics and Chemistry of Minerals 15: 19-29

Calas G, Manceau A, Combes JM, Farges F (1990) Applications of EXAFS in Mineralogy. In: A. Mottana and F. Burragato (eds) Absorption spectroscopy in Mineralogy. Elsevier, New York 172-204

Charlet L, Manceau A (1992) X-ray absorption spectroscopic study of the sorption of Cr(I1I) at the oxide/water interface. II Adsorption, co precipitation and surface precipitation on ferric hydrous oxides. Journal of Colloid and Interface Science 148: 425-442

Chukhrov FV, Gorshkov AI, Vitovskaya IV, Drits VA, Sivtsov AV, Rudnitskaya YeS (1980) Crystallochemical nature of Co-Ni asbolan. An SSSR Izvestiya, Seriya Geologicheskaya, 6, 73-81. (Translation in Int Geol Rev 24: 5, 598-604)

Chukhrov FV, Manceau A, Sakharov BA, Combes JM, Gorshkov AI, Salyn AL, Drits VA (1988) Crystal chemistry of oceanic Fe-containing vernadites. Mineralogicheskii Journal 10: 78-92

3.2.2 Information from X-Ray Absorption Spectroscopy to7

Combes JM, Manceau A, Calas G (1986) Study of the local structure in poorly-ordered precursors of iron oxi-hydroxides. Journal de Physique C8: 697-701

Combes JM, Manceau A, Calas G, Bottero JY (1989) Formation of ferric oxides from aqueous solutions: a polyhedral approach by X-ray absorption spectroscopy. I. Hydrolysis and formation of ferric gels. Geochimica et Cosmochimica Acta 53: 583-594

Combes JM, Manceau A, Calas G (1990) Formation of ferric oxides from aqueous solutions: a polyhedral approach by X-ray absorption spectroscopy. II. Hematite formation from ferric gels. Geochimica et Cosmochimica Acta 54: 1083-1091

Cressey G, Steel AT (1988) An EXAFS study of Gd, Er and Lu site location in the epidote structure. Physics and Chemistry of Minerals 15: 304-312

Decarreau A, Colin F, Herbillon A, Manceau A, Nahon D, Paquet H, Trauth-Badaud D, Trescases 11 (1987) Domain segregation in Ni-Mg-Fe smectites. Clay and Clay Minerals 35: I-to

Drits VA, Sakharov BA, Salyn AL, Manceau A (1993) Structural model for ferrihydrite. Clay Minerals, 185-208.

Dumas T, Petiau J (1986) EXAFS study of titanium and zinc environments during nucleation in a cordierite glass. J. Non-crystalline Solids 81: 201-220

Eary LE, Rai D (1987) Kinetics of chromium (III) oxidation to chromium (VI) by reaction with manganese dioxide. Environmental Science and Technology 21: 1187-1193

Greaves GN (1985) EXAFS and the structure of glass. Journal of Non Crystalline solids 71: 203-217

Greaves GN, Barrett NT, Antonini GM, Thornley FR, Willis BTM, Steel A (1989) GlancingAngle X-ray Absorption Spectroscopy of Corroded Borosilicate Glass Surfaces Containing Uranium. Journal of the American Chemical Society 111: 4313-4324

Greegor RB, Sandstrom DR, Wong J, Schultz PC (1983) Investigation ofTi02-Si02 glasses by x-ray absorption spectroscopy. J. Non-crystalline Solids 55: 27~43

Hayes KF, Roe AL, Brown GE Jr, Hodgson KO, Leckie JO, Parks GA (1987) In situ X-ray absorption study of surface complexes: selenium oxyanions on IXFeOOH, Science 238:783-786

Hayes KF, Papelis C, Leckie JO (1988) Modeling ionic strength effects on anion adsorption at hydrous oxide/solution interfaces. Journal of Colloid and Interface Science 125: 717-728

Hazemann JL, Manceau A, Sainctavit Ph, Malgrange C (1992) Structure of the IXFexA11 - x-

OOH solid solution. I. Evidence by polarized exafs for an epitaxial growth of hematite-like clusters in diaspore. Physics and Chemistry of Minerals 19: 25-38

Ingalls R, Garcia GA, Stern EA (1978) X-ray absorption at high pressure. Physics Review Letters 40: 334-336

Hie JP, Polian A, Calas G, Petiau J, Fontaine A, Tolentino H (1989) Pressure- induced Coordination changes in crystalline and vitreous Ge02. Physics Review Letters 63: 398-401

Jackson WE, Brown GE Jr, Waychunas GA, Mustre J, Conradson SD, Combes JM (1991) In situ high-temperature x-ray absorption study of divalent iron in orthosilicates, crystals, and liquids. In: S.S. Hasnain (ed) XAFS VI, Sixth Internat. Conf. on X-ray Absorption Fine Structure. Ellis Horwood Ltd. Publishers (in press)

Johnson CA, Xyla AG (1991) The oxidation of chromium (III) to chromium (VI) on the surface of manganite (yMnOOH). Geochimica et Cosmochimica Acta 55: 2861-2866

Manceau A (1990) Distribution of cations among the octahedra of phyllosilicates: insight from EXAFS. Canadian Mineralogist 28: 321-328

Manceau A, Charlet L (1992) X-ray absorption spectroscopic study of the sorption of Cr(III) at the oxide/water interface. I Molecular mechanism of Cr(III) oxidation on Mn oxides. Journal of Colloid and Interface Science 148: 443-458

Manceau A, Combes JM (1988) Structure of Mn and Fe oxides and oxyhydroxides: a topological approach by EXAFS. Physics and Chemistry of Minerals 15: 283-295

Manceau A, Drits VA (1993) Local structure of ferrihydrite and feroxyhite by EXAFS spectroscopy. Clay Minerals 165-184

Manceau A, Llorca S, Calas G (1987) Crystal chemistry of cobalt and nickel in lithiophorite and asbolane from New Caledonia. Geochimica et Cosmochimica Acta 51: to5-113


Manceau A, Bonnin D, Stone WE, Sanz J, Kaiser P Fretigny C (1988) Polarized EXAFS of biotite and chlorite. Physics and Chemistry of Minerals 16: 180-185

Manceau A, Buseck PR, Miser D, Rask J, Nahon D (1990a) Characterization of Cu in lithiophorite from a banded Mn ore. American Mineralogist 75: 490-494

Manceau A, Bonnin D, Stone WE, Sanz J (1990b) Distribution of Fe in the octahedral sheet of trioctahedral micas by polarized EXAFS. Comparison with nmr results. Physics and Chemistry of Minerals 17: 363-370

Manceau A, Gorshkov AI, Drits VA (1992a) Structural Chemistry of Mn, Fe, Co, and Ni in Mn hydrous oxide. II. Information from EXAFS spectroscopy, electron and X-ray diffraction. American Mineralogist 77: 1144-1157

Manceau A, Charlet L, Boisset MC, Didier B, Spadini L (1992b) Sorption and speciation of heavy metals on Fe and Mn hydrous oxides. From microscopic to macroscopic. Applied Clay Science 7: 201-223

Pauling L, Kamb B (1982) The crystal structure of lithiophorite. American Mineralogist 67:817-821

Ramos A, Petiau J, Gandais M (1985) Crystalline nucleation process in (SiOr AI 20 3-Li20) glasses. J. de Physique C8: 491-494 Sandstrom DR, Lytle FW, Wei PSP, Greegor RB, Wong J, Schultz PC (1980) Coordination of Ti in TiOz-SiOz glass by x-ray absorption spectroscopy. J. Non-crystalline Solids 41:201-207

Sanz J, Stone W.E.E. (1977) NMR study of micas. I. Distribution of Fe2+ ions on the octahedral sites. Journal of Chemical Physics 67: 3739-3743

Sanz J, Stone W.E.E. (1983a) NMR study of minerals: III. The distribution of Mg2+ and Fe2+ around OH groups in micas. Journal of Physics C: Solid State Physics 16: 1271-1281

Sanz J, Stone W.E.E. (1983b) NMR applied to minerals: IV. Local order in the octahedral sheet of micas: Fe-F avoidance. Clay Minerals 18: 187-192

Schwertmann U, Murad E. (1983) The effect of pH on the formation of goethite and hematite from ferrihydrite. Clays and Clay Minerals 31: 277-284

Stucki JW, Goodman BA, Schwertmann U (1988) Iron in Soils and Clay Minerals. NATO Series ASI, vol. 217, Riedel Publishing Company.

Stumm W, Wollast R (1990) Coordination chemistry of weathering: kinetics of the surfacecontrolled dissolution of oxide minerals. Reviews of Geophysics 28: 53-69

Sueno S, Nakai I, Imafuku M, Morikawa H, Kimata M, Ohsumi K, Nomura M, Shimomura o (1986) EXAFS measurements under high presure conditions using a combination of a diamond anvil cell and synchrotron radiation. Chern. Letters 1663-1666

Tafto J, Buseck PR (1983) Quantitative study of AI-Si ordering in an orthoclase feldspar using an analytical transmission electron miscroscope. American Mineralogist 68: 944-950

Tardy Y, Nahon D (1985) Geochemistry of laterites, stability of AI-goethite, AI-hematite and Fe3 + -kaolinite in bauxites and ferricretes: an approach to the mechanism of concretion formation. American Journal of Science 285: 865-903

Tchoubar D, Bottero JY, Qienne P, Arnaud M (1991) Partial Hydrolysis of Ferric Chloride Salt. Structural Investigation by Photon-Correlation Spectroscopy and Small-Angle X-ray Scattering, Langmuir 7: 398-402

Trolard F, Tardy Y (1987) The stabilities of gibbsite, boehmite, aluminous goethites and aluminous hematites in bauxites, ferricretes and laterites as a function of water activity, temperature and particle size. Geochimica et Cosmochimica Acta 51: 945-957

Wadsley AD (1952) The structure of lithiophorite, (AI,Li) MnOz(OH)2' Acta Crystallographica 5: 676-680

Waychunas GA (1983) Mossbauer, EXAFS and X-ray diffraction study of Fe3+ clusters in MgO : Fe and magnesiowstite (Mg, Fe), _ xO - Evidence for specific cluster geometries. Journal of Material Science 18: 195-207

Waychunas GA (1987) Synchrotron radiation XANES spectroscopy of Ti in minerals: effects of Ti bonding distances, Ti valence and site geometry on absorption edge structure. American Mineralogist 72: 89-101

Waychunas GA, Brown GE (1990) Polarized X-ray absorption spectroscopy of Metal ions in Minerals: Applications to site Geometry and electronic structure determination. Physics and Chemistry of Minerals 17: 420-430

3.3 Optical Absorption Spectroscopy 109

Waychunas GA, Brown GE Jr, Apted MJ (1986) X-ray K-edge absorption spectra of Fe minerals and model compounds: II EXAFS. Physics and Chemistry of Minerals 13: 31~47

Waychunas GA, Brown GE Jr, Ponader CW, Jackson WE (1988) Evidence of networkforming Fe2+ in molten alkali silicates. Nature 332: 251 ~253

Yarker CA, Johnson PAY, Wright AC, Wong J, Greegor RB, Lytle FW, Sinclair RN (1986) Neutron diffraction and EXAFS evidence for TiOs units in vitreous K 20-Ti02-2Si02 . J. Non-crystalline Solids 79: 117~ 136

3.3 Optical Absorption Spectroscopy

K. LANGER, A.N. PLATONOV, and G.R. ROSSMAN

Optical absorption spectroscopy is an effective method to investigate the local atomic and electronic structure of minerals. Whereas diffraction methods average a volume of several thousands of A 3 in a mineral structure, optical spectroscopy may elucidate the structural and electronic properties of local polyhederal centers and thus forms a valuable addition to the diffraction methods. Optical absorption spectra of minerals result from the interaction with electromagnetic radiation in the energy range of 40000 cm ~ 1 to 4000 cm ~ 1

corresponding to the wavelength range of 250 nm to 2500 nm. Recent reviews of this field have been presented (Burns 1970; Platonov 1976; Marfunin 1979; Nassau 1983; Langer 1984, 1988; Berry and Vaughan, 1985; Rossman 1988; and Vaughan 1990).

Optical Spectroscopy

Figure 39 presents a spectrum of Cr3 + in garnets and compares various abscissa scales where the energy units are in wavenumbers, electron volts, and joules.

The energy of the bands is related to the type of electronic systems or species that undergo excitation, as outlined below. Band intensities are related to the number of species per volume unit. A third quantity characterizing an absorption band is the width of the band at half-height. Bands in optical spectra vary widely in width. Spin forbidden bands and vibrational overtones can be very narrow compared to dd and charge transfer bands.

The excitation processes, giving rise to bands in the range 40,000 to 4000 cm ~ 1 of the electronic spectra of oxygen based minerals are:

1. LM-CT: ligand to metal charge transfer or an interband transitions, which involve transfer of electron density from ligand (usually oxide ion) to the metal cation. Such transitions cause bands of very high intensity in the ultraviolet region.

2. MM-CT: metal to metal charge-transfer (also called intervalence charge transfer, IVCT) between transition metal ions differing in valence state and


A(nmJ-300 400 500 600 800 1000 2000 1000000

,-UV- .. ,. VIS -I' NIR -I' IR

t , 1 I I , I I I ,

log (1./1} , - 24500_22700 , 1\ I I

or \ \ I I \ 0(= \ \ I I

log !I./Il/d \ I \ I \ -280001 I \ -17500 -16200 I \ " ,-29000 II \ \ I I " I I \-14500 \ I '\.. / -20700 \ / \' -I -.J 1 -19200

" I I I

30000 25000 20000 15000 10000 5000 10000 --v (cm-1J

4 3 2 1 0.5 0 -E(eVJ

i i , 6 5 4 3 2 1 0

_El'10-19JI

Fig. 39. Composite absorption spectra of chromian-pyrope (solid line) and uvarovite (broken line) comparing wavelength and energy units. (Langer 1988)

accommodated in interconnected polyhedra of the mineral structure. Such transitions give rise to broad bands in the visible range and are strongly polarized along the MM vector.

3. dd- (or ft)-transitions: electronic transition between crystal field split spectroscopic states of the free transition metal ions (or rare earth ions). These bands are have much lower intensities than those of mechanisms (1) and (2).

4. Electron-hole centers: excitation of trapped electrons and holes, introduced into the mineral structures, in connection with other substitutional defects of the structure. Such centers are usually at low concentrations.

From the optical spectra of minerals it is immediately possible to:

1. interpret the origin of color and pleochroism of minerals. When the bands from a spectroscopically active ion are properly assigned, it is often further possible to:

2. identify the ion and its oxidation state; 3. determine the concentration of the ion; 4. determine the site symmetry of the ion;


5. determine the distribution of the ion in multisite minerals; 6. determine the distribution of the ion among co-existing phases; 7. examine the influence of the ion on phase transformations.

Evaluation of Electronic Transitions

Theoretically best understood are dd-transitons, where crystal field or ligand field theory applies. (Burns 1970; Marfunin, 1979; SchHifer and Gliemann 1980; Berry and Vaughan 1985). In a cation site in a crystalline mineral, the valence electrons of transition metal ions such as Cr3 + and Fe2 + experience the electrostatic influence of the electrons associated with all the anions (ligands). As a result of these interactions, a number of electronic states involving the valence electrons are produced whose number and energies depend upon the number of valence electrons of the cation and the geometric details of the coordination site. Electronic transitions of the 3d electrons between pairs of these states produce the dd transitions.

The theoretical treatment of crystal field theory has been well developed for some time and is treated in a number of texts. (e.g. Schliifer and Gliemann 1980). Two concepts from this theory are especially important in the field of mineral spectroscopy.

The first concerns the decrease in the energy of an ion when it is incorporated in a crystalline host. The decrease in the ground state in the crystal relative to the free ion is the crystal field stabilization energy (CFSE). This quantity plays a prominent role in the partitioning of transition metal ions in minerals. (Burns 1970; Langer 1988)

The second concept relates to selection rules which govern the intensity of electronic transitions. Transitions between states of equal parity, i.e. same sign of the respective wave functions, are parity forbidden. As this selection rule applies to all dd-transitions of transition-metal ions in centro symmetric sites, the intensity of dd-bands is generally much lower than that of charge-transfer bands. The parity selection rule is partly removed in non centrosymmetric sites. An additional selection rule applies to transitions involving changes of more than one quantum number, namely transitions between states with different spin multiplicities. Such transitions are spin-forbidden. In practice, they give rise to bands with very low intensity. In noncubic crystals, the absorption bands caused by dd-transitions are polarization-dependent. This effect is due to symmetryrelated selection rules, whereby the symmetry under which the electrical dipole moment vector transforms in the respective point group and the symmetry of the ground and excited state must be properly related for the transition to be allowed.

The intensity of dd-transitions may be enhanced by magnetic coupling effects between two ions in adjacent polyhedra. A variety of such effects, known as ion pair effects, can increase the intensity of absorption bands by more than an order of magnitude (Bakhtin and Vinokurov 1978, Rossman 1988). Ion pair


aspin !3spin

-7 Ti(eg ) =::=========

-9

Ti(hg ) :::=::::::::::::.;; Fe(eg )

--------... :11.

Fe(hg ) > -11 --~ >. OJ

Iii -13 Fe(eg ) c

16a1 orbital

IJ.J o Fe(12g ) -- =

-15 !!!!!!!!I!B

=-- ---

- -- O(2p) -17 ---- -- ,,' -- u

a -- b

Fig. 40. a Calculated molecular orbital diagram for the [FeTiO lOJ1 4 - cluster. Energy levels indicated by dashed lines are unoccupied. b Wavefunction contours for the 16a, orbital. (Sherman 1987b)

effects also influence the absorption band intensity of cations involved in intervalence change transfer. (Amthauer and Rossman 1984).

The interpretation ofligand-metal and metal-metal charge transfer processes is more complicated. Quantitative theoretical approaches applicable to minerals were started only about 20 years ago when K.H. Johnson and his group at MIT introduced the SCF-Xa and related methods to describe the energy states of clusters of atoms in mineral structures such as the octahedral cluster (Fe2+06)lO- (Sherman and Waite 1985 see also chap. 5.5). Recently, the method has been extended by Sherman (1987a, 1987b see also Chap. 5) on clusters of interconnected polyhedra, containing transition metal ions in different valence state. The molecular orbital diagram for a (FeTi010l)14- cluster of edge connected octahedra and the wave function contours of the 16a1 molecular orbital are shown in Fig. 40. The latter is involved in the Fe2 + -Ti4 + charge transfer.

Experimental Aspects

Optical spectra of minerals are now obtained by conventional double beam grating spectrometers and diode array spectrometers. Either type of instrument can be incorporated into a microscope spectrometer system. FTIR systems are used for near-infrared measurements, but have seen only limited application to visible spectroscopy. Conventional instruments normally work with samples larger than 100 /lm; FTIR systems without special focusing optics use samples


down to 50 Jlm diameter; microscope systems have succeed in measuring sample areas down to 5 x 5 Jlm (Langer and Frentrup, 1979) and have proven particularly useful for studies of small synthetic crystals from hydrothermal and piston cylinder syntheses.

Fine particle size and heterogeneous samples can be measured with diffuse reflectance spectroscopy (Kortum, 1969). This method measures the radiation diffusely scattered from a "thick layer" of powder relative to that of a white standard, usually MgO, LiF or a fluorocarbon. Often these data are processed with the Kubelka-Munk function to obtain a spectrum resembling those obtained in the transmission mode. As mineral grains are randomly oriented in the diffuse reflectance experiments, spectra with mixed polarization are obtained and, hence, the information obtained from polarized single crystal spectra is lost. In addition, the thickness of the mineral penetrated by radiation is unknown, making determinations of the concentration of the absorbing species difficult.

Under optimal conditions, a range of 40000 to ~5000 cm -1 is available now for crystals with a minimum size of about 30 Jlm. Crystal preparation for polarized microscope spectrometer measurements proceeds as described by Langer (1988). A disadvantage of microscope spectrometers is that measurements are done with convergent radiation. Thus a certain degree of mixing of the pure polarizations occurs which, in the case of strongly polarized bands, may result in an artificial decrease of band intensity. (Goldman and Rossman 1977) For quantitative evaluation of band intensities, the convergence angles must be kept as low as possible and measuring conditions have to be kept constant. These problems are less important in many cases where relative band intensities are sufficient for quantitative purposes.

Application of Electronic Spectroscopy in Geosciences

Color of Minerals

The determination of the origin of color in minerals is among the fundamental problems addressed with optical spectroscopy. Color is one of the diagnostic properties of a mineral, a prospecting indicator, and a major value criterion of gem minerals and pigments. The color of most minerals is determined by absorption phenomena in the 380-760 nm region. Table 5 presents typical examples of the classification of the origin of color of minerals. More detailed reviews on mineral color have been presented by Nassau (1983), Platonov et al. (1984), and Fritsch and Rossman (1988).

The color of many minerals is determined by a complicated interplay of the absorption bands of several chromophores. In the spectra of many of the common rock-forming ferro-magnesian silicates, absorption bands from LMCT, MM-CT as well as dd transitions are observed. LM-CT bands, in particular, usually have an important role in defining the limits of transmission in the blue end of the spectrum. The pleochroism of low-symmetry minerals is determined by the polarization properties of the individual bands.


Table 5. Typical examples of the causes of mineral color

1. Colors dominated by absorption involving dd transitions of metal ions y3 + _ vanadian grossular (green), variscite (green) Cr3 + - ruby (red), uvarovite (green), emerald (green), kammerererite (violet) Mn2+ - rhodonite (pink), rhodochrosite (pink), spessartite (orange) Mn3+ - muscovite (pink), piemontite (red), manganian andalusite (green) Fe2+ - olivine (green), almandine (red), staurolite (brown) Fe3+ - epidote (green), andradite (green), ferrian-orthoc1ase (yellow) Co2 + - erythrite (pink), spherocobaltite (pink) Ni2+ - annabergite (green) Cu2+ - malachite (green), azurite (blue), dioptase (green) U4+ - zircon (blue)

II. Colors dominated by absorption involving ligand-metal charge transfer 0 2 - -Fe3+ - goethite (yellow), haematite (red), lepidocrocite (brown) 0 2 - -Cr6 + - crocoite (orange)

III. Colors dominated by absorption involving metal-metal charge transfer Fe2+ -Fe3+ - glaucophane (blue), vivianite (blue) Fe2 + - Ti4 + - sapphire (blue), dravite (brown), kyanite (blue) Ti3+ - Ti4 + - hibonite (blue meteoritic)

IY. Colors dominated by electron-hole centers and inorganic free radicals AlO:- (Na) - quartz (smoky) S; - sodalite (blue) CO; - calcite (yellow) SO;3 - celestite (blue)

Concentration Determination

The amount of absorption is related to the number of species per unit volume which undergo excitation. In many cases, the absorption is a linear function of the concentration of absorbing species. Thus, if the spectral features to be evaluated are unambiguously assigned to the species of interest, if the thickness of the crystal plate is accurately known, if the sample is in the correct orientation, and if the band is independently calibrated against the concentration of the target species, quantitative analyses of metal ions in minerals are readily accomplished. The calibrations are often the most difficult aspect of such studies and are available for only few mineral systems. When calibrations are available, the spectrometric methods have several advantages over other analytical procedures: (1) Most importantly, they are valence-specific. Other methods commonly used are atomic number-specific. Thus, it is possible, in principle, and often in practice to determination the concentration ofV3 +, Mn3 +, Fe2 +, etc. by spectrometric methods. (2) The determination may be obtained on the small volumes available to microscope spectrometry.

An example is the Fe2 + determination in pyrope-almandine solid solutions based on the calibration data in Fig. 41. The left part shows the absorption spectrum of Fe2 + in the triangular dodecahedral garnet site with bands at about 8000, 6300, and 4600 cm - 1. The linear absorption coefficients of the maxima of


2.8 r---------------, e(e) = 1.11

2.4

WAVE NUMBER (em -') 2.0

20000 10000 7000 5000 4000 E 0.8 r'-r--t-r"",-,""r1i--r-.-:-r.,...,......-..,........-t-..-.-.....-r-t .s

~ 1.6 Q) PYROPE

UJ 0.6 u

u <: til .0 o 1.2

z « '" .0

~ 0.4 o

«

en al « 0.2

0.0

500 1000 1500 2000

WAVELENGTH (nm)

a

OL--~~-~---~----~ 10 20 30 40

WEIGHT PERCENT FoO

b

Fig. 41. a Pyrope absorption spectrum with Fe2 + bands. b. Relationships between absorption band intensities and Fe2+ concentration in pyrope. (Rossman 1988)

20 40 60 80

y,.cm-'

11000

10800

10600

10400

10200

10000

Fig. 42. Variation of orthopyroxene Fe2 + band positions with mol fraction Fe2 + and Al content. (Khomenko and Platonov 1987)

these bands are plotted as a function of the FeO weight fraction of the garnet in the right part of Fig. 41. The concentrations ofFe2+ and Fe3+ in chloritoid were determined based on the shift of the ultraviolet edge towards lower energies when Fe3+ contents increased (Halenius and Langer, 1980) and on the intensity


of an Fez + ~Fe3+ charge transfer band (Halenius et a1.1981). Traces of Fe3+ involved in fayalite point defect chemistry were likewise determined, based on shifts of the UV absorption edge with increasing concentration of point defects. (Cemic et al. 1986). This effect was also used to study the kinectics of point defect annihilation (Langer 1988).

Quantitative determinations of the chemical composition of minerals can also be obtained from the energy of absorption bands which change with composition. Fig 42 presents the dependence of dd bands of Fez + in orthopyroxenes for the vI band (Ella, - 10000cm- 1) and the v2 band (Ellb, - 5000 -1) on the atomic fraction, f, of Fe2+ and on tetrahedral AI-contents (Khomenko and Platonov 1987).

Determination of Effective Site Symmetry

Figure 43 shows an orthopyroxene spectrum and the various interpretations, proposed in the literature, of the strongly polarized, low-energy bands originating from Fe2+ in Ml and M2 sites of the structure, both with the site symmetry C 1 (Steffen et al. 1988). From the assignment of bands derived from Fez + in the M2 site, it was possible to determine the effective site symmetry

., u c o .0 ::; III .0 «

-- Wavelength Inml----

1.00 500 600 800 1000 1500

25000 20000

Fe2+-Fe3+CT

(S.L.SI

15000 10000 • Wavenumber [cm-'j--

5000

Fig. 43. Orthopyroxene spectra with proposed assignments of Fe2+ and Fe2+ _Fe3+ bands. (Steffen el al. 1988)


of M2 from the polarization scheme of the pair of bands at about 11 000 and 5000 cm -1 (Goldman and Rossman 1977). The procedure by which the effective site symmetry was evaluated is described by Rossman (1988). The effective symmetry (pseudo symmetry) of the orthopyroxene M2-site is C2v (C/'), a supergroup of the actual point group of the crystallographic site, C1. Such pseudo symmetries are usually obtained from optical spectra because minor deviations from higher symmetries which reflect the actual point group of the crystallographic site have little effect upon the optical spectra.

Recent examples for the application of this procedure have been concerned with the effective symmetries of Mn3 + in andalusite (Abs Wurmbach et al. 1981), piemontite (Kersten et al. 1988), clinopryroxene, and amphibole (Ghose et al. 1986), Cr3+ in clinopyroxene (Abs Wurmbach et al. 1985, Khomenko and Plato nov, 1985), and Ni2+ in olivine (Rager et al. 1988; Hu et al. 1990).

Intracrystalline Cation Distribution

Information on the intracrystalline cation distribution in minerals with multiple cation sites in their structure can be obtained from electronic spectra by evaluating bands originating from both dd-transitions and metal-metal chargetransfer (MM-CT) transitions.

The site occupancy of the transition metal ions in minerals has often been deduced from their dd bands. Information available from the spectra includes the energy of the bands, their polarizations, intensities, widths, and the splittings between the bands. An example is provided by solid solutions of Ni in the synthetic olivine (Mg1 - xNixhSi04 • Figure 44 shows the spectra of crystals with x = 0.02, 0.30, 0.75 and 1.00. The strong band (b) at about 23500 cm- 1 is assigned to transitions of Ni2+ in the M2 site. The intensity of this band was used to derive the site fraction ofNi in M2, XNi(M2)' and Ni in Ml, XNi(M1) (Hu et al. 1990). The determination is based on the fact that 1X23000 has its maximum value when X = 1.00. The validity of these determinations is confirmed by the agreement of the spectroscopic cation distribution with those derived from X-ray structural refinements.

The energy and polarization of MM-CT bands depend on the sites the cations occupy. The energies of the CT bands most commonly found in mineral spectra depend on the particular cations involved and decrease in the order:

E(Mn2+Ti4 +) > E(Fe2+Ti4 +) > E(Fe2+Fe3+).

The energy is further generally dependent on the distance between the coupled ions if systems with the same polyhedral units are compared. Although there are no systematic expressions which relate the energy of a MM-CT band to all of these parameters, the general trends can be used in assignments of CT bands especially when they are combined with structural information derived from the polarization of MM-CT bands. MM-CT bands are strongly polarized along the metal-metal vector of the interacting cations. Because the metal-metal

118

8 <.>

---'" '" <.>

'" 0

'" 0

0-~

0 ~

.0

~ ... '" -"

1500

1000

500

750

500

250

0

10] 50

30

12,5 -

300

\. "-

\

\

"-

"-

"-

\

"-

\ \

Chapter 3, Solid State Spectroscopy

400 600 1000 2500

b --'I, -----y ------- -- X

NOIiIOO

-.,

NOD075

/ ''-fj

o .0 r--r-I I I I I I I I I i I I I I

36000 28000 20000 12000 4000 Wavenumber [I/em]

Fig. 44. Spectra of synthetic (Mg_xNix)2Si04 olivines presented by Hu et al. (1990)


vector is not necessarily parallel to either the crystallographic axes (a,b,c) or the principal optical directions (X,Y,Z), the spectra will often show components in multiple directions which are proportional to cos2<jJ, where <jJ is the angle between the metal-metal vector and the respective principal optical direction or crystallographic axis.

A prominent example ofMM-CT in minerals is the Fe2+Ti4 + MM-CT band in blue kyanite near 16000 cm -1 which is strongly polarized along the c-axis. This polarization proves that the dominant substitution of two Al Sites by Fe2+Ti4 + pairs occurs in interconnected Ml and M2 octahedra, forming chains parallel to C

Crystal Field Stabilization Energy in Inter Crystalline 3dN-Distributions and Phase Transformations

Both ionic radius considerations and the crystal field stabilization energy (CFSE) determine the specific site that a transition metal ion will occupy in a multi site mineral. The higher the CFSE of an ion, the higher will be the site fraction for that ion. This general statement holds not only for intra-crystalline partitioning, but also for the inter-crystalline distribution of 3d ions between coexisting phases, and between melts and crystals.

The influence of CFSE was demonstrated in the early period of the application of crystal field theory and reviewed by Burns (1970). A recent example concerns the distribution of Cr3 + minerals of kyanite-bearing eclogites. Cr3 + concentrations in the grospydite xenolith follow the sequence of CFSE in the metamorphic eclogite with lower total chromium contents (Langer 1988). More research is needed relating composition and CFSE in co-existing minerals with regard to the concentration dependence, and with regard to both compression and thermal expansion as they relate to the conditions of formation.

The most elaborate example of the influence of CFSE on partitioning behavior relates Mn3 + partitioning between solid solutions of the aluminum silicate polymorphs and related minerals in the AI203-Mn203-Si02 system (AbsWurmbach et a!. 1981, 1983). Because the CFSE is exceptionally high in the Ml octahedra of andalusite compared to most other minerals. (Langer 1988), this ion fractionates most strongly into the andalusite phase creating wide, divariant fields of andalusitejkyanite and andalusite/sillimanite co-existence which replace the univariant phase transformations in the pure A1 2SiOs system. A direct result of the CFSE is an increase in the thermal stability of the Mn3 + -substituted andalusite phase at constant pressure and an increase in the pressure stability at constant temperature.

Colorimetry

The color measuring system of the International Commission of Illumination (ICI), based on the physiological perception of color, has found use to describe


0.9 -

O.~

520 525 51~.,.;e.;.K 530

-lsi ...... 535 510 '.5'0

0.7 ,,-~5

550

505 ~.555

0.6 ~,O

r~ 500 570

>- 0.5

0.4

".:,:5

~ X 585

t::t--_ ~O 595

_\ , - , 600 -- -- --. 605

0.3 \ -, ,

P AD .~15 'C"S.: 620 '90 a ~~ r _____ 650

700

0.2 '85\ -.......

I'--- /v \ V

0.1 '80 ~ v""""'" 05\ //

l. 70.

~ W)'~,oo/

~ '50 ...... :..;- , ~ I I I , o 0,1 0.2 0,3 0.4 0.5 0.6 0.7 0.8

x

Fig. 45. The ICI color chart

the color characteristics of minerals. The color of a mineral is assigned to a point in the x-y coordinates of the leI color chart (Fig. 45). Special computer programs are used to calculate the color parameters of a mineral directly from the measured optical absorption or diffuse reflectance spectrum. The color parameters are an objective measure of the mineral's color which expresses the spectral peculiarities such as the relative absorption band intensities and widths of the different chromophores.

This approach has proven useful in the comparative analysis of garnets from different mantle parageneses in kimberlite from Siberia and South Africa. It was possible to quantitatively characterize the evolution of the color of garnets which result from the complicated process of upper-mantle differentiation (Matsuk et al. 1985). This is of particular importance in the identification of garnets from the various parageneses, including those which bear diamonds (Fig. 46). In turn, it helps to solve some important genetic and practical problems regarding upper mantle evolution. Similar investigations proved fruitful for the genetic correlation of hornblende- and biotite-bearing Precambrian metamorphic rocks (Platonov et al. 1988).

3.3 Optical Absorption Spectroscopy

Pe r.u

0.9

0.8

0.7

480 3BO (567) 500' 700 600

01 () 2 Q3 $4 • 5

121

Fig. 46. Color diagram of garnets from different mantle rocks: 1 harzburgites; 2 dunites; 3 wehrlites; 4 chromite-garnet uItrabasites; 5 pyroxenites; 6 garnet nodules; 7 i1mentite peridotites; 8 equigranular garnet Iherzolites; 9 broken down prophiric Iherzolites; 10 websterites; 11 Mg-Fe eclotites; 12 kyanite eclogites; 13 inclusions in diamonds (Matsuk et al. 1985) Pe the purity of the color line is given in relative units. r.u.

Colorimetric methods are widely used for the evaluation of gemstone color, although the optical properties of faceted stones make instrumental color determination with absorption spectroscopy difficult (Platonov et al. 1984). Instrumental colorimetric methods are being actively introduced as a diagnostic procedure for ore mineral characterization in reflected light (Peceff et al. 1992). Such applications of color evaluation from mineral spectra come close to another important application of the color and electronic spectra of oxygenbased minerals, the attempt to gain knowledge on the mineralogy of the planets' surfaces by means of the methods of remote sensing (e.g. Adams 1975; Burns 1989 and additional reference in Vaughan 1990). These methods are based on the interpretation of diffuse reflectance spectra. Such applications are increasingly important for problems of remote sensing, both of the Earth from airborne and space-borne platforms, and of extraterrestrial objects through telescopes. Interpretation of these data is based on laboratory studies conducted on powdered minerals which record the same absorption bands seen in single crystal spectra but also include effects of grain size and specular reflectance of grain surfaces. Polarization properties, of course, are not recorded. Extensive collections of reflectance spectra have been assembled which serve as a database against which observational data can be compared. Some of the most extensive


are the low resolution data of Hunt et al. (1973) and the high resolution data of Clark et al. (1990)

Applications have included identification of pyroxenes on planetary surfaces, hydrous minerals in asteroids, estimation of iron content in ferromagnesian minerals, and airborne prospecting for precious metals in regions of hydrothermal alteration.

References

Abs-Wurmbach I, Langer K, Seifert F, Tillmanns E (1981) The crystal chemistry of (Mn3+Fe3+)-substituted andalusties (viridines and kanonaite), (Al1_x_yMn; + Fe; +}z(OjSi04): cystal structure refinements, Mossbauer, and polarized optical absorption spectra. Z Krist 155: 81-113

Abs-Wurmbach I, Langer K, Schreyer W (1983) The influence of Mn3+ on the stability relations of the Al 2SiOs polymorphs with special emphasis on manganian andalusites (viridines), Al 1_xMn; -h(OjSi04): an experimental investigation. J Petrol 24: 48-75

Abs-Wurmbach I, Langer K, Oberhiinsli R (1985) Polarized absorption spectra of single crystals of of the chromium bearing clinopyroxenes cosmoclore and Cr-aegirine-augite. Neues Jahrb Mineral Abh 152: 293-319

Adams JW (1975) Interpretation of visible and near-infrared diffuse reflectance spectra of pyroxenes and other rock-forming minerals. In: Karr C (ed) Infrared and Raman spectroscopy of lunar and terrestrial minerals. Academic Press, London

Amthauer G, Rossman GR (1984) Mixed valence of iron in minerals with cation clusters. Phys Chern Mineral 11: 37-51

Bakhtin AI, Vinokurov VM (1978) Exchange-coupled pairs of transition metal ions and their effect on the optical absorption spectra of rock-forming silicates. Geokhimiya 1: 87-95 (Translation in Geochem Int (1978): 53-60)

Berry FJ, Vaughan DJ (1985) Chemical bonding and spectroscopy in mineral chemistry. Chapman Hall, London.

Burns RG (1970) Mineralogical applications of crystal field theory. Cambridge Univ Press, Cambridge. During the preparation of the present book, the 2nd edn appeared: Tsurus RG (1993)

Burns RG (1989) Spectral mineralogy of terrestrial planets: scanning their surfaces remotely. Mineral Mag 53: 135-151

Cemic L, Grammenopoulou-Bilal S, Langer K (1986) A microscope-spectrometric method for determining small Fe3 + concentrations due to Fe3 + -bearing defects in fayalite. Ber BunsenGes Phys Chern 90: 654-661

Clark RN, King TVV, Klejwa M, Swayze G, Vergo N (1990) High spectral resolution reflectance spectroscopy of minerals. J Geophys Res 95: 12, 653-12, 680

Ghose S, Kersten M, Langer K, Rossi G, Ungaretti L (1986) Crystal field spectra and Jahn Teller effect of Mn3+ in clinopyroxenes and clinoamphiboles from India. Phys Chern Mineral 13: 291-305

Goldman DS, Rossman GR (1977) The spectra of iron in orthopyroxene revisited: the splitting of the ground state. Am Mineral 62: 151-157

HAienius U, Langer K (1980) Microscope photometric methods for nondestructive Fe2 + -Fe3 + determination in chloritoid, (Fe2 + ,Mn2 + ,Mgh(Al, Fe3 +)4Si2010(OHk Lithos 13: 291-294

Halenius U, Annersten H, Langer K (1981) Spectral studies on natural chloritoids. Phys Chern Mineral 7: 117-123

Hu X, Langer K, Bostrom D (1990) Polarized electronic absorption spectra and Ni-Mg partitioning in olivines (Mg1- xNi.}z[Si04]. Eur J Mineral 2: 29-41

Hunt GR, Salisbury JW, Lenhoff CJ (1973) Visible and near infrared spectra of minerals and rocks, VI, additional silicates. Mod Geol 4: 85-106


Kersten M, Langer K, Almen H, Tillmamms E (1987) the polarized single crystal spectra and structures of synthetic thulite and piemontites. Z Kristallogr 185: III

Khomenko VM, Platonov AN (1985) Electronic absorption spectra of Cr3+ ions in natural clinopyroxenes. Phys Chern Mineral 11: 261-265

Khomenko VM, Platonov AN (1987) The optical spectra, color and pleochroism of rock forming pyroxenes. Naukova Dumka, Kiev (in Russian)

Kortum G (1969) Reflexionspektroskopie. Springer, Berlin Heidelberg New York. Langer K (1984) Die Farbe von Mineralen und ihre Aussagefaehigkeit fuer die Kristallchemie.

Rheinisch-Westfael Akad Wiss N332: 7-60 Langer K (1988) UV to NIR spectra of silicate minerals obtained by niicroscope spectrometry

and their use in mineral thermodynamics and kinetics. In: Salje EKH (ed) Physical properties and thermodynamic behaviour of minerals. Reidel, pp 639-685

Langer K, Frentrup KR (1979) Automated microscope-absorption spectroscopy of rockforming minerals in the range 40000-5000 cm - 1 (250-2000 nm). J Microsc 116: 311-320

Marfunin AS (1979) Physics of minerals and inorganic materials Springer, Berlin Heidelberg New York

Matsuk SS, Plato nov AN, Khomenko VM (1985) The optical spectra and color of the mantle minerals in kimberlites. Naukova Dumka, Kiev (in Russian)

Nassau K (1983) The physics and chemistry of color: the fifteen causes of color. John Wiley and Sons, New York

Peckett A, Phillips R, Henry NFM (1992) The colours of opaque minerals. John Wiley, New York

Plato nov AN (1976) Tile nature of the color of minerals. Naukova Dumka, Kiev (in Russian) Plato nov AN, Taran MN, Balitsky VS (1984) The nature of color of gemstones. Naukova

Dumka, Kiev (in Russian) Platonov AN, Matsuk SS, Khomenko VM, Taran MN, Litvin MA (1988) Optical activity as

an indicator of the evolution of mineral matter. In: Theory of Mineralogy. Leningrad, 76-86. (in Russain)

Rager H, Hosoya S, Weiser G (1988) Electron paramagnetic resonance and polarized optical absorption spectra of Ni2+ in synthetic forsterite. Phys Chern Mineral 15: 383-389

Rossman GR (1988) Optical spectroscopy. In: Hawthrone FC (ed) Spectroscopic methods in mineralogy and geology. Rev Mineral 18: 207-254

Schlafer HL, Gliemann G (1980) Einfiihrung in die Ligandenfeldtheorie, 2nd edn. Akad Verlagsges. Frankfurt/M

Sherman OM (1987a) Molecular orbital (SCF-XIX-SW) theory of metal-metal charge transfer processes in mineral. I. Applications to Fe2 + - Fe3 + charge transfer and "electron delocalization" in mixed-valence iron oxides and silicates. Phys Chern Mineral 14: 355-363

Sherman OM (1987b) Molecular orbital (SCF-XIX-SW) theory of metal-metal charge transfer processes in mineral. ii. Applications to Fe2 + - Ti4 + charge transfer transitions in oxides and silicates. Phys Chern Mineral 14: 364-367

Sherman OM, Waite TO (1985) Electronic spectra of Fe3 + oxides and oxide hydroxides in the near IR to near UV. Am Mineral 70: 1262-1269

Steffen G, Langer K, Seifert F (1988) Polarized electronic absorption spectra of synthetic (Mg,Fe)-orthopyroxenes, ferrosilite, and Fe3+ -bearing ferrosilite. Phys Chern Mineral 16: 120-129

Vaughan OJ (1990) Some contributions of spectral studies of the visible (and near-visible) light region to mineralogy. In: Monttana A, Burragato F (Eds) Absorption spectroscopy in mineralogy. Elsevier, Amsterdam, pp 2-38


3.4 Luminescence of Minerals

3.4.1 Interpretation of Luminescence Spectra in Terms of Band Theory and Crystal Field Theory. Sensitization and Quenching. Photoluminescence, Radioluminescence, and Cathodoluminescence

A.N. TARAsHcHAN and G. WAYCHUNAS

Luminescence is nonequilibrium light emission, in excess of that produced by thermal black body radiation, which has been excited by any kind of incident radiation source. Luminescence is peculiar to an extremely wide circle of objects of inorganic and organic nature in various states of aggregation, and includes all processes taking place when radiation or particles interact with matter and cause emission at optical wavelengths. The mechanisms responsible for luminescence phenomena are extremely varied. However, it is possible to distinguish the following common stages: (1) absorption of excitation energy and stimulation of the system into an excited state; (2) transformation and transfer of the excitation energy; (3) emission of light and relaxation of the system into an unexcited condition. The first two stages are always dependent on the manner of excitation, while the last stage is conditioned in mineral systems by the types and structures of the luminescence centers that are present.

From a practical standpoint, the various types of luminescence mechanisms are distinguished by the kind of incident radiation or particles which excite emission, and by the kinetics of the overall emission process. Thus we speak of cathodoluminescence as a response to incident high energy electrons, photoluminescence as excited by optical (UV -Vis-IR) photons, radioluminescence as excited by X-ray photons, gamma rays and nuclear particles, chemiluminescence as excited by chemical radicals, and triboluminescence as excited by the breaking or disruption of bonds in crystals.

Further, luminescence is historically called fluorescence or phosphorescence, depending on whether the absorption-emission process is relatively rapid, or slow, respectively. However these names disclose nothing about the actual kinetics of the emission, or the energy transfer processes which cause the luminescent behavior, and any definition of what is fast or slow is arbitrary.

Luminescent emission may be from the bulk of the material itself, or from isolated impurities, clustered defects, exsolved phases, fluid inclusions, or grain boundary contaminents. In general, the emission process involves the transition of an electron from an excited state to one lower in energy in the luminescence center. Interpretation of luminescence in minerals begins with the characterization of the luminescence centers, including the identity of the ions involved, their locations in the crystal structure, their energetic interactions, and their modes of energy transfer with each other, with other ions in the structure, or with vibrational states.

3.4.1 Interpretation of Luminescence Spectra

Excitation Mechanisms. Photoluminescence, Radioluminescence, and Cathodoluminescence

125

Despite all the types of excitation sources for luminescence, it is possible to categorize luminescence mechanisms into two general classes. In the first the process of excitation is localized near an isolated center and occurs without ionization of any species in the crystal. This type of luminescence is called intracentric. Excitation and emission occur as energy transfers due to electronic transitions at the center. The second mechanism occurs when the excitation is light of higher energy than the band gap, or consists of very high energy radiation or particles. Ionization events then occur, and the emission is produced by recombination of electrons or holes at ionized centers. A variation on the second class is excitation involving exciton formation and energy transfer. There is no net charge transfer in exciton-produced luminescence since the electron-hole pair are never separated in the crystal.

Photoluminescence is distinguished from other types of luminescence as a method of direct excitation of centers of emission without their ionization. It is possible to selectively excite each energy level of the center. Most commonly, photoluminescence is excited in minerals with ultraviolet (UV) excitation. Most (UV)-excited luminescence is produced using low pressure mercury discharge lamps, in which the gas pressure is optimized to preferentially excite the UVemitting energy states of the mercury ions. The commonly used transitions emit short wave ultraviolet at 253.7 nm, and long wave ultraviolet at 365.0 nm. Other sources of UV are xenon and hydrogen discharge lamps. Laser excitation in the UV has also been utilized with the advantage of very selective excited state generation and high excitation energy density.

Radioluminescence is excited by high energy particles, gamma rays and also by X-ray photons from ordinary X-ray tubes. In the case of an X-ray tube source, its intensity depends on the voltage and current applied to the tube. Compared to UV excitation, most luminescent minerals show additional emission bands when exposed to X-ray stimulation. Both exciton energy transfer and electron-hole recombination processes are common in X-ray-induced luminescence processes.

Cathodoluminescence is created by stimulation from high energy electrons, and is commonly observed under the electron microprobe beam. Direct electron stimulation has several advantages over UV excitation. The energy density is much greater than for UV photons as the penetration depth of 15 keY electrons is on the order of microns, whereas UV photons penetrate from a fraction to several mm. Additionally, in the case of the microprobe beam, focusing laterally to a few microns is possible. Thus single mineral grains, inclusions or grain boundaries can be excited for spectroscopy. Hence microprobe-based cathodoluminescence is much brighter than UV -excited luminescence and is able to discern spatial variations on a relatively fine scale. Due to the great excitation energy density, many more minerals are observed to cathodoluminescence than luminescence under UV.


Physical Characteristics

The emission spectrum shows the distribution of energy emitted by an excited system in terms of the intensity of emitted optical photons as a function of wavelength or photon energy. Usually the spectra consist of a series of lines or bands corresponding to transitions between electronic states. Figure 47 shows the relationship between an emission spectrum (the curve labeled number 2) and such electronic states. In the example given, Mn2+ replacing Ca2+ in calcite, emission is from only one transition, hence the spectrum is deceptively simple. In general, emission spectra are simpler than excitation or optical absorption spectra for the same system or center because only particular electronic transitions may be favored for emission. The maximum energy band in the emission spectrum has longer wavelength than the minimum energy excitation band (Stokes' shift), thus UV-excited photoluminescence in minerals is usually in the visible and near infrared.

The excitation spectrum shows the dependence ofluminescence intensity in a particular emission band as a function of the excitation wavelength or photon energy. Excitation spectra are similar to optical absorption spectra, but can differ from the latter by variations in intensity or even absence of particular bands. The intensity of absorption bands is related only to the oscillator strength of the single responsible electronic transition, whereas the excitation band intensity is related to the oscillator strength of the absorption, the

J

200

'00

200

.... Jl>

... 31900 cm-'

rl1~~IIII~~gll~~I-~~} 7 ~- Absorption 24700 em- (ezcifiztiOn) 22500 18200 ~-I

300

I

" I I I I I

400

15600~-1 F;;..;..~Emi$Sion (630nm)

~stokes shift

I I

500 600 700 8001,nm

Fig. 47. The emission spectrum (2) and excitation spectrum (1) of Mn2+ in calcite. (After Medlin 1964). Note the Stokes' shift of about 100 nm toward longer wavelengths

3.4.1 Interpretation of Luminescence Spectra 127

efficiency of internal electronic energy transfer processes to the final state from which emission occurs, and to the efficiency of the emission transition. Because of their similarity to absorption spectra, excitation spectra can be used for the characterization of luminescence centers. In principle, the identity, crystallographic position, local symmetry, valence state, and electronic structure of the ion or ions responsible for luminescent activity can be determined if suitable absorption spectra are available for comparison. In Fig. 47 the curve labeled number 1 is the excitation spectrum for Mn2 + in calcite.

The kinetics of luminescence emission depend on the nature of the specific processes involved, their temperature dependence, the intensity of the incident excitation source, and other factors. The historical terms fluorescence and phosphorescence are used to distinguish between quickly decaying and slowly decaying emission, once excitation has ceased. Popular usage defines fluorescence as operating on a time frame of 10- 6 s or less, slow fluorescence as operating on the scale 10- 5_10- 2 s, quick phosphorescence, 10- 2 _10° s, and phosphorescence inclusive of anything with longer decay rates. The difficulty with this picture is that rather different processes can produce emission with about the same time dependence.

A more practical distinction is one that separates fluorescence and phosphorescence in terms of temperature dependence. By this description, fluorescence is due to electronic transitions within single ions, or energy transfer between closely associated ions in a structure. Its characteristic lifetime is relatively independent of temperature. In contrast, phosphorescence involves energy transfer associated with energy bands and the trapping of electrons and holes. The luminescent centers that donate the electrons or holes or serve as recombination centers may be widely separated in the structure. Energy transfer thus cannot occur, and result in an emission process, until thermal energy frees the holes or electrons from the traps. Thus phosphorescence decay rate is dependent on the amount of thermal energy available, and is highly temperature-dependent.

The duration of luminescence emission also depends on the specific electronic mechanism. The lifetimes of excited states depend on the probabilities of radiative (light emitting) and nonradiative (heat producing) energy transfers, which are specific to site symmetry, ion electronic structure, accessible phonon states, and other factors. Fluorescence decay is generally exponential, but luminescence decay can be hyperbolic or more complex in phosphorescent systems.

Interpretation of Luminescence Spectra

The luminescence emission spectrum of a mineral is determined mainly by the presence of particular luminescence centers. This term includes any kind of point defect or clustered defect in a crystal structure which can absorb and emit optical photons. For further discussion we can divide the centers depending on their chemical, electronic, and spectroscopic properties.


The main and most abundant group consists of activator centers. These are ions of various valence that substitute for cations of the host structure. The most common possibilities are:

1. Transition metal (d-d electronic transitions); especially Cr3+ (oxides, silicates), Mn2+ (found in all types of minerals), Fe3+ (silicates).

2. Rare earth~ with f-f electronic transitions (Pr3+, Nd3+, Sm3+, Eu3+, Gd3+, Tb3+, Dy3+, Ho3+, Er3+, Yb3+), and with d-f electronic transitions (Ce3+, Sm2+, Eu2+, Yb2+); both types found in fluorides, phosphates, sulfates, tung states, silicates, and oxides.

3. Actinides; mainly U6 + found in fluorite and carbonates. 4. Mercury-like metals; Hg +, TI +, Pb +, Pb2+ (silicates, sulfates).

The optimal concentration for the activators noted is about 0.01 -1.0%. However, in some cases the main species in a mineral may act as an activator, especially at low temperatures, such as Mn2 + in rhodochrosite and rhodonite, Pb2 + in cerussite and anglesite, and Hg + in calomel.

The spectrum of a luminescing mineral activated with transition metal or rare earth ions can be interpreted only if the effects of the crystal field on the energy levels of the ion, and the static and dynamic properties of the host matrix are taken into account. These factors are common to absorption and luminescence spectroscopy, and hence indicate why ion energy level schemes in various symmetries are the starting points for luminescence spectral analysis. Although the intensities of bands may vary widely between absorption and excitation spectra, still the mode of identifying activators, their environments, and bonding attributes is completely analogous.

Two types of ions, rare earths and transition metals, have qualitatively different luminescence spectra. The f-f transitions in rare earths give rise to sharp emission and excitation lines, while the d-f transitions in rare earths and the d-d transitions in transition metals give rise to broad bands of varying width.

The best examples of sharp f-f transition luminescence are in minerals activated by lanthanide rare earths, e.g., apatite, fluorite, and zircon. In these minerals, the luminescence is often characteristic of the type of rare earth. This is true because the electronic energy levels responsible for much of the absorption and emission oflight have highly shielded 4f or Sf orbital character, and thus are not involved with bonding to nearby atoms. The influence of the crystal field results in small splitting of the nf energy levels (about 100 em -1) and the removal of symmetry restrictions for interconfigurational f-f transitions. The number of split sublevels, and thus the number of lines in the luminescence spectrum, depends on the crystal field symmetry. This picture is complicated by activator substitutions that are heterovalent and require charge compensation. In this case the symmetry of the luminescent center may be dictated by the position of the neutralizing charge, or other structural adjustment. This yields polycentric systems and reduced local symmetry, and the fine structure of the spectrum can be much more complicated. Examples are the cubic, tetragonal,


trigonal and rhombic centers for Dy3 + in fluorite. Deciphering the spectrum can be accomplished with free ion energy level schemes, but it is possible that there will exist differing types of emitting levels for particular ions even in minerals of very similar composition and structure.

Electronic levels in a rare earth may also be of mixed character, such as 4fk -15d. Such levels will be much more sensitive to variations in site geometry and local crystal field than 4f levels, and lead to variations in the associated emission bands for the same RE ion in different structures. For example, the maximum emission band of Eu2 + is at 385 nm in anhydrite, at 400 nm in oligoclase, at 425 nm in fluorite and at 460 nm in apatite. 4fk- 15d-4fk transitions yield broad (about 1500 cm -1), intense luminescence emission bands with relatively large Stokes' shifts (102-103 cm- 1). The intensity of the bands is due to the partially allowed nature of the transitions, and the corresponding finite oscillator strengths.

Among ions of transition metals the most widely observed activator of minerals is Mn2 +. Emission bands due to Mn2+ have been observed in practically all classes of minerals. Less abundant are cases with octahedrally coordinated Cr3 + activation (ruby, spinel, kyanite, topaz, and spodumene, where Cr3+ replaces AI3+, and diopside, where Cr3+ replaces Mg2+), and tetrahedrally coordinated Fe3 + (feldspar, scapolite, and other silicates where Fe3+ replaces AI3+; and quartz, zircon, beryl, phenakite, topaz, tourmaline, and other silicates, where Fe3+ replaces Si4 +).

Transition metal-activated spectra are analyzed mainly from excitation spectra with the construction of energy level diagrams, and the semiempirical calculation of crystal field parameters Dq, B, and C. For example, the excitation spectrum of red (ca. 690 nm) luminescence from orthoclase is similar to the absorption spectrum of Fe3 + in this mineral. The crystal field parameters (Dq = 900 cm - 1, B = 557 cm - 1, C = 3446 cm - 1) are indicative of tetrahedral Fe 3+ , and the transition 4T 1 -+ 6 A1 is responsible for the emission band.

The Stokes' shift of the emission band to wavelengths longer than the corresponding absorption band is about 2000-2500 cm -1 in Fe3 + activated silicate minerals. The magnitude of this shift depends on the details of vibrational relaxation in the excited and ground states, which, in turn, depend on the interaction of the emitting ion with adjacent coordinating ions in the structure. The cause of Stokes' shift, the nature of nonradiative transitions and the character of temperature dependence of luminescence are more easily discussed with diagrams involving a configurational coordinate. Such diagrams show the energy of particular electronic levels in a luminescent system as a function of average distances between, e.g., metal ions and their coordinating anions. Individual states in the diagram thus form configurational curves with shape dictated by the central ion's interaction with the structure. Strong interaction, perhaps due to bonding orbital overlap, results in configurational curves with a potential well shape, while nonbonding noninteracting states may be essentially flat, i.e., independent of interatomic distances. In practice, diagrams of this kind can be calculated theoretically (KCI : TI +) or built in a semiquantitative manner


from excitation, absorption, and emission spectra, and from the known dependence of energy levels of varying symmetry on the strength of the crystal field, Dq.

An important example for mineralogical systems is the Mn2 + ion in different coordination in silicates, oxides, carbonates, phosphates, and other species, where this ion is responsible for green, yellow, and orange-red luminescence. For the case of octahedrally coordinated Mn2+, as in calcite, the position of the minima of the configurational curves occurs at differing Mn2 + -0 distances for states which are dependent on the crystal field strength, Dq, such as 4T I and 4T 2

(Fig. 48). For states whose energies are independent of Dq, the minimum occurs at the same interatomic distance as in the ground state, 6AI (states 4Eg and 4A1).

The consequence of this is that there is a progressive shift in the energy of the 6 Al -+ 4T I and 6 Al -+ 4T 2 transitions as Dq increases, producing band broadening and a shift toward longer wavelengths for both emission and absorption bands. Fig. 49 shows the configurational coordinate scheme in more detail, along with a range of interatomic distances for which absorption normally occurs (a to b). It is easy to see that the curvature of the excited states leads to the width of the bands in the excitation spectrum. States not involved in bonding give rise to sharp absorption lines because they are only weakly coupled to host structure ions, while strongly coupled ion states vary considerably in energy with interatomic distance and yield broad bands.

Changes in coordination, such as from octahedral to tetrahedral, produce considerable change in Dq. The tetrahedral Dq value is smaller and the energy of all of the transitions in Mn2+ increase. This creates a shift toward short

's 'A, ('s) Dq-

c f

d f

Fig.48a-d. The sensitivity of the electronic states of the Mn2 + ion in octahedral coordinaion to changes in Dq and representation in a configurational coordinate diagram. (After Marfunin 1979). a Schematic representation of the states as defined by group theory at zero crystal field intensity. b The effect of increasing crystal field strength. Some states are independent of Dq, others shift up or down in energy. c For a particular value of Dq (that appropriate for calcite) the energy states are redrawn for fixed Mn2+ -0 distance. d Configurational coordinate diagram showing changes in state energy with variation in Mn2+ -0 distance. The center of the ground state (6 At) is nominally the equilibrium distance. The shift in the 4T 1 state leads to emission at smaller interatomic distance, and due to state curvature this is at smaller energy (and longer wavelength) than the absorption (Stoke's shift)


6.0

5.0

/ -r .... 6 ___ ....:..I.u....uu... i1'9

-3.0 -2.0 -/.0 0 1.0 2.0 3.0 Configuration coordinate

Fig. 49. Enhanced configurational coordinate diagram as in Fig. 48d showing emission and excitation spectrum, Stokes' shift, and the effect of state curvature on spectral band width. (After Medlin 1968)

wavelengths. Octahedral Mn2+ coordination in calcite thus has red-orange emission at 630 nm, while the analogous transition for Mn2+ on a tetrahedral site in willemite yields yellow-green emission at 525 nm. Prominent changes in the color of the emission may also occur from a change in the emission state. The 4T 1 state is responsible for emission in calcite and willemite, but the 4T 2 state is responsible for the green emission of fluorite (500 nm) and anhydrite (505 nm) even though Mn2 + has larger coordination sites. In other species the relative site distortion and crystal field strength variation can cause significant changes in the energy of the 4T 1 ~ 6 A 1 transition, e.g., yellow emission at 580 nm in apatite and datolite, compared to red-orange emission from forsterite and wollastonite. Finally, the energy of emission is also influenced by the Mn2+ concentration. Calcite Mn2 + emission is at 630 nm, but this increases to 660 nm in manganocalcite, and to 680 nm in rhodochrosite. Thus the color of emission depends on coordination number, site distortion (which affects the energy of electronic states), crystal field strength, activator concentration, and the nature of the electronic state from which emission occurs.

The influence of the crystal field on the luminescence of Cr3 + in minerals results mainly from changes in the energy of the 4T 2 and 2E states. For strong crystal fields, where DqjB > 2.3, the 4T 2 multiplet is situated higher in energy than the 2E state and below 300 K the 4T 2 is unoccupied. The luminescence observed is thus connected with transitions ofthe type 2E ~ 4 A2, although these are forbidden by symmetry and spin selection rules. Nevertheless, this transition produces the red emission in ruby, spinel, alexandrite, beryl, and topaz. In the


case of diopside, as well as in MgO and many glasses, the crystal field is weaker with Dq/B < 2.0. In this situation below 300 K strong broad-band luminescence comes only from the 4T z --+ 4 Az transition, which is only symmetry forbidden. In kyanite the luminescence due to the R lines as well as the strong broad band emission is observed.

Donor-Acceptor Pairs. This type of luminescent system is formed from activator centers in minerals with semiconducting properties. The sulfide minerals with narrow band gaps are the best examples, i.e., zinc, cadmium, and mercury sulfides. The activator ions have energy levels in the band gap. The donors, mainly Ga3+, In3+, TI3+, Ti3+, and so forth, are situated just below the conduction band; while the acceptors, Cu +, Ag+, V Zm and so on, have levels just above the valence band. The donors, on ionization, produce electrons in the conduction band, and the acceptors accept electrons from the valence band by ionizing the host structure leaving hole carriers. Excitation is followed by ionization of host structure or donor ion, and the mechanism of emission involves recombination of free electrons and localized holes, or free holes and localized electrons. In sphalerite the following luminescent centers are well established: Ag+ -CI- (460 nm), Cu + -CI- (AI3+) (520 nm), VZn-Cu + (Ag+) (590 nm), Ag+ -ln3+ (620 nm), Cu + -ln3+ (Ga3+) (640-670 nm), Cu + - TI3+ (Ti3+) (830 nm).

Defects in Crystals and Larger Activator Complexes. These are divided into the following types:

1. Tetrahedral and Octahedral complexes of closed shell transition metalsWO~ - MoO~ - , TiO~ -, VOl-, CrO~ - (scheelite, powellite, wulfenite, benitoite, etc.). In contrast to transition metals absorbing energy and emitting via d-d transitions, these complexes utilize transitions involving chargetransfer bands and molecular orbital states not localized on the cation. Luminescence can be very strong.

2. Radiation-produced centers: F and F aggregate (M, N, R) centers, V k and V F

centers (fluorite, fluorides, some phosphates and silicates). 3. Autolocalized excitons. In the oxides and the salts of oxyacids (mainly

silicates) holes located on oxygen ions and captured free electrons can create autolocalized exciton states. Recombination of hole and electron occurs with photon emission. The luminescence spectra of this group of centers are characterized by very broad structureless bands with large Stokes' shifts and well-defined temperature dependence.

Molecular Ions. These are centers that have characteristic and unique properties: UO~ + (secondary uranium minerals, silicates, and carbonates), S2" and 02" (minerals of the sodalite group, scapolite, sulfates). The luminescence spectra of molecular ions have a unique periodic appearance at low temperature resulting from the modulation of the broad emission band by local intermolecular vibrational structure.


Adsorbed Molecular Complexes. These are usually organic substances or aqueous complexes that are localized in micro defect areas (gas-liquid inclusions, dislocations) mainly in hypergene minerals (carbonates, borates, sulphates, hydroxides, etc.). Luminescence is observed only by photon stimulation and is characterized by bluish-white fluorescence (singlet-singlet Sl ----+ So transitions) and by green phosphorescence (triplet-singlet T ----+ So transitions). The duration of the phosphorescence is dramatically increased upon cooling of the mineral.

Energy Transfer, Sensitization, and Quenching

A set of important issues appear due to the fact that activator centers cannot be treated as isolated ions or complexes. In addition to the properties created by simple multiplication of the actions of many activators, other nonadditive properties appear by virtue of the interaction of the activators. These properties include the strengthening of emission band intensity in the presence of a second kind of activator, the redistribution of the intensity of emission bands, and the quenching of luminescence. All of these phenomena are connected with processes of energy transfer between substituent ions or between substituent ion and the host structure.

Sensitization. The luminescence of ions which have been excited as a result of energy transfer to their absorption band from other ions is called sensitized luminescence. The absorption center in such a case is now called a sensitizer or co-activator, and both species must be present in sufficient proximity for the energy transfer to occur. There are several kinds of excitation energy transfer from sensitizer to activator.

Emission-Reabsorption. This occurs when the emission from the sensitizer is partly or wholly absorbed and then emitted by the activator. It is also called cascade luminescence. The duration and kinetics of the emission process in each center are unchanged. Examples are the emission from Dy3 + ions in scheelite that are excited from emission from the W04 complex, and the strengthening of Nd3+ emission in the presence of Sm3+, Tb3+ and Dy3+ in fluorite.

Inductive Resonance. This process involves energy transfer without emission, and occurs as a result of interaction of the multipole electric fields at emission centers. The coupling may be by dipole-dipole, dipole-quadrupole, or quadrupole-quadrupole interactions. The appearance of this mechanism requires close proximity of the affected ions or complexes. The resonance transfer probability is proportional to the degree of superposition of the emission spectrum of the sensitizer and the absorption spectrum of the activator, and decreases with increase of the separation between them as R - 6 (dipole-dipole coupling) and R -8 (dipole-quadrupole coupling). This sort of energy transfer process generally shortens the lifetime of the excited state in the activator, thus affecting the kinetics of emission.


Examples of this process include the luminescence of Mn2+ in calcite, which is sensitized by Pb2 + or Ce3 +, and the emission of Fe3 + in feldspars, which is sensitized by Pb2 +. Practically all trivalent rare earths activators may be sensitized by various energy sources including other rare earths, Tl +, Sn2+, Pb2+, Cr3 +, Mn2+ ions, and UO~+, W04 , and Mo04 complexes.

A special case of resonance energy transfer is that of cooperative sensitization, where two or more neighboring sensitizers, excited at the same time, transfer energy to a near neighbor luminescence center. This process can result in emission which is of shorter wavelength than the absorption, and in such case is called anti-Stokes luminescence. In fluorite containing both Tb3 + and Yb3 + in sufficient concentrations, luminescence from Tb3 + at 380-490 nm and at 490-680 nm can be observed following excitation of the Yb3 + ions in the infrared at 980 nm. Another example is BaF 2: Vb, Tm where excitation of Yb3 + at 960 nm leads to emission from Tm3+ centers at 470 nm.

Nonresonance Energy Transfer. This can occur between centers in a structure where one center's absorption is of higher energy than the other's emission. The difference in energy is then lost to the host structure in the form of lattice vibrations, or excites an electronic transition in a third ion. This kind of energy transfer requires an exchange interaction between the interacting ions and is most probable between nearest or next-nearest neighbors, and generally not beyond the range of 6-8 A.

Quenching of luminescence occurs by several processes. Essentially all excited states can have their energy dissipated to the lattice by various types of thermal quenching. This is the fundamental reason why luminescence is always limited at high temperatures. Thermal quenching refers to the increasing probability of activating new lattice vibrational modes as the temperature increases, some of which will interact with an excited state to produce a nonradiative (not lightemitting) transition. Among activator ions the strongest thermal quenching of luminescence is observed for some divalent rare earths. Sm2+ and Yb2+ emission in minerals is observed only at or below 77 K. In feldspars the red luminescence of Fe3 + is partially quenched at 300 K, and is entirely quenched at about 450 K.

Among the various kinds of so-called external quenching, connected with the transfer of excitation energy, one can distinguish the following main types:

Concentration (or Self) Quenching. This quenching is due to energy transfer between nearby ions of the same type, where the excitation of the recipient ion is into a state which can decay nonradiatively. It depends on the nature of the interaction between the ions (or centers), and includes the types noted above. Some ions, such as Mn2 + in rhodochrosite or WOi + in scheelite, display very weak concentration quenching. Other ions, such as the trivalent rare earths, where the self-quenching is due to resonance multipole energy transfer between ions of the same type or to exchange interactions, are quenched at moderate concentrations (several %). Another type of self-quenching involves self absorp-


tion of the emission. This occurs in rare earth activation, but can also occur in transition metal activation, as with Cr3 + in ruby.

Quenching as Another Aspect of Sensitization. The strengthening of the emission from the activator is accompanied with partial or complete quenching of any emission from the sensitizer center. This type of quenching is highly selective, i.e., it acts between a pair of ions whose electronic energy levels have nearly the same separation. Hence it is specific to particular activator-sensitizer couples.

Quenchingfrom Ions with Intense Charge-Transfer Bands. Strong charge-transfer absorption bands can quench emission from any activator or system of activators. The only requirement is overlap of the emission band with the charge-transfer band. This kind of quenching is strongest for Fe3 + ions in octahedral sites, where concentrations of as little as 0.1 % can completely quench Mn2+ emission, and also for Fe2+, Co2 +, and Ni2+.

References

Blasse G (1980) The luminescence of closed-shell transition-metal complexes, New developments. Struct Bonding 42: 1-42

Blasse G, Aguilar M (1984) Luminescence of natural calcite (CaC03). J LumiI). 29: 239-241 Deb SK, Gallivan JB (1972) Photoluminescence of O2- and S2-ions in synthetic sodalites. J

Lumin 5: 348-360 Gorobets BS (1981) Spectre luminescence of minerals. Moscow, 154 pp Krasilshchikova OA, Tarashchan AN, Platonov AN (1986) Colour and luminescence of

natural fluorite. Naukowa Dumka, Kiev, 224 pp Krasnobaev AA, Votyakow SL, Krochalev VJ (1988) Spectroscopy of zircons. Nauka,

Moscow, 150 pp Kuznetsov GV, Tarashchan AN (1988) Luminescence of minerals of granitic pegmatites.

Naukowa Dumka, Kiev, 178 pp Marfunin AS (1979) Spectroscopy, luminescence and radiation centers in minerals. Springer,

Berlin Heidelberg New York, 352 pp McKeever SWS (1985) Thermoluminescence of solids. Cambridge Univ Press, Cambridge,

367 pp Medlin WL (1964) Trapping centers in thermo luminescent calcite. Phys Rev 135: 1770-1779 Medlin WS (1968) The nature of traps and emission centers in thermoluminescent rock

materials. In: McDougall DJ (ed) Thermoluminescence of geological materials. Academic Press, New York, pp 193-223

Tarashchan AN (1978) Luminescence of minerals. Naukowa Dumka, Kiev, 296 pp Telfer DJ, Walker G (1978) Ligand field bands of Mn2+ and Fe3+ luminescence centers and

their site occupancy in plagioclase feldspars. Mod Geol 6: 199-210 Walker G (1985) Mineralogical aspects of Luminescence techniques. In: Berry FJ, Vaughan DJ

(eds) Chemical bonding and spectroscopy in mineral chemistry. Chapman and Hall, London, p 103

Waychunas GA (1988) Luminescence, X-ray emission and new spectroscopies. Rev mineral 18: 639-698

White WB (1975) Luminescent materials. Trans Am Crystallogr Assoc 11: 31-49 White WB, Masako M, Linnehan DG, Furukawa T, Chandrasekhar BK (1986) Absorption

and luminescence of Fe3+ in single-crystal orthoclase. Am Mineral 71: 1415-1419 Williams FE (1966) Theoretical basis for solid-state luminescence. In: Goldberg P (ed)

Luminescence of inorganic solids. Academic Press, New York, pp 1-52


3.4.2 Selective Laser Excitation of Rare-Earth Luminescence Spectra

M. ILiEvand M. SENDOVA-V ASSILEVA

The rare earth (RE) impurities are typical of a number of minerals such as fluorite, anhydride, calcite, etc. and are often responsible for their optical transmission and for their luminescence. As the orbitals of the inner f-electrons practically do not overlap with the electron orbitals of the other ions making up the crystal, the f-electron energy level scheme, e.g., of RE3 +, is similar to that of isolated ions. The absorption and emission bands due to the f-f electron transitions are very narrow and the effect of surrounding atoms manifests itself mainly in the crystal field splitting. The latter depends also on the way of excess charge compensation, i.e., on the type of the impurity site.

The narrow spectral bandwidth of laser radiation makes possible a direct selective excitation of both different RE3+ ions and different impurity sites of the same RE3+ ion via tuning the laser photon energy to coincide with the energy difference between the ground and one of the excited states. This is the so called site-selective spectroscopy developed by Tallant and Wright, 1975; Seelbinder and Wright, 1979, Hamers, 1982. Luminescence due to more than one type of site appears only when their absorption spectra overlap. There is no energy transfer between luminescence centers containing a single RE ion. On the other hand if a luminescence line characteristic of a given site is chosen and its intensity as a function of excitation wavelength is followed, the excitation spectrum of that site is obtained.

If the laser photon energy does not correspond to energy difference between the ground and one of the excited multiplet levels, indirect phonon-assisted excitation processes are also possible, their probability decreasing with the number of phonons required in a single act of excitation. Iliev et al. (1988) have demonstrated that indirect excitation with emission of one phonon will be quite effective for photon energies situated at several hundred wavenumbers above an excited crystal field level. Since the multiplet splitting itself is of the order of a few hundred wave numbers, one comes to the conclusion that the luminescence line of a given RE3+ could be excited in a relatively broad region and hence more than one type of RE3+ center could be excited by a single laser line. Knowing the maximum energy of the phonons in a crystal and the approximate situation of the energy levels of the RE3+ ions, one can estimate which level of which ions will be excited with a given line (e.g., Table 6). This mechanism is illustrated in Fig. 50 for the excitation of Er3 + with commonly used Ar+ and Kr+ laser lines. At low temperatures excitation is possible with the 476.2 nm, 514.5 nm, 530.9 nm, and 647.1 nm laser lines at the 4F7/2' 2Hll/2' 4S3/2, and 4F9/2 level, respectively, but not with the 413.1 nm and 488.0 nm lines. When the temperature rises, excitation, assisted by the absorption of one phonon, becomes also possible. In addition, the higher components of the ground multiplet become thermally populated and excitation can take place with a photon energy lower than the one of the level. The latter case is demonstrated by the excitation of the

3.4.2 Selective Laser Excitation of Rare-Earth Luminescence Spectra 137

25

23

22

488.0nm-----20

514Snm-----

530.9nm-- -!~

18

17

16

6471nm-----15 ... ,

E u

(Y"\

~

>-1.9 cr w Z UJ 0

2H ___ _ 9/2

4F ___ _ 4F312 __

5/2

4F ----712

~ 411512 _

A-site B-site

Fig.50. Energy level scheme of A- and B-sites of Er3 + in CaF 2 (Tallant and Wright 1975) and the position of the excitation energies with the phonon density of states in CaF 2 below them

Table 6. Expected laser excitation of RE3 + via the process of one-phonon emission. The efficiency ( + ) or inefficiency ( - ) of a particular laser line has been confirmed by the experimental data on CaF 2: RE3+. (Iliev et al. 1988)

Laser line Expected RE3 + excitation

Pr Nd Sm Eu Tb Dy Ho Er

413.lnm + + + 476.2 nm + + + + + + 488.0nm + + + 514.5 nm + + + 530.9 nm + + + + 647.1 nm + +

4F 7/2 level of Er3+ by the 488.0 nm line only at high temperatures (Send ovaVassileva et al. 1988, 1991).

The selective laser excitation of RE3 + luminescent centers in crystals can be used for qualitative and in some cases quantitative analysis of these centers in minerals. The quantitative measurements can be facilitated by the use of a phonon Raman line to exclude the effect of the sample shape and the apparatus on the intensity of the luminescence.


References

Hamers RJ, Wietfeld JR, Wright JC (1982) Defect chemistry in CaF 2: Eu3+. J Chern Phys 77: 683-692

Iliev M, Liarokapis E, Sendova M BI (1988) Laser excited luminescence of rare earth impurities in natural and synthetic CaF 2. Phys Chern Mineral 15: 597-600

Seelbinder MB, Wright JC (1979) Site-selective spectroscopy of CaF 2: Ho3+. Phys Rev H20: 4308-4320

Sendova-Vassileva M, Iliev M, Liarokapis E (1988) One-phonon-assisted laser excitation of Er3+ luminescence in CaF 2. Bulg J Phys 15: 367-373

Tallant DR, Wright JC (1975) Selective laser excitation of charge compensated sites in CaF2:Er3+. J Phys Chern 63: 2074-2085

Sendova-Vassileva M, Iliev M, Chadwick AV (1991) Laser-excited luminescence ofCaF 2: Ho. The role of phonons. J Phys Condensed Matter 3: 5407-5414

3.4.3 Origins of Luminescence in Minerals: A Summary of Fundamental Studies and Applications

B.S. GOROBETS and G. WALKER

Luminescence Centers and Their Characterization

Because most minerals, excluding metals, are large bandgap materials (i.e., insulators), luminescence usually originates from the electronic states of localized centers in the structure rather than de-localized states. The luminescence characteristics of most minerals are, in fact, determined by the presence of impurity metal cations in minor or trace amounts substituting for lattice cations of similar ionic size, although intrinsic complexes of uranium, tungsten, and molybdenum can themselves luminesce. These impurity ions are usually transition-metal ions of the d- and f-groups. The following ions are known to form luminescence centers in minerals: (1) 3d ions of the iron group, e.g., Mn2+, Fe3+, Cr3+, Ti3+, Ti4 +, possibly Ni2+ (at low temperatures) and Ag+ of the precious metal group; (2) 4f elements of the rare-earth group (TR): Ce3 + , Pr3 + , Nd3+, Sm2 +, Sm3+, Eu2 +, Eu3+, Gd3+, Tb3+, Dy3+, Er3+, Yb2 +, Yb3+; (3) 4d-, 5d elements, e.g., M06 +, W6 +; (4) 5f elements, e.g., U 6 + and (5) 6p element TI +. (see Fig. 51). The commonest and best understood of these centers are the 3d group ions, particularly Mn2 +, which is probably the most ubiquitous of all luminescence centers in minerals.

Apart from such ions, other causes of luminescence in minerals are defect centers which may be intrinsic (e.g., F-centers in halides, oxygen-related defects in quartz and other oxygen-dominated lattices), or impurity-related (e.g., compensated AP+ centers in quartz). Such defect centers, which may give rise to luminescence emission, are more difficult to identify. The emission bands from such centers are often very broad even at very low temperatures, indicating a strong interaction with the lattice.

3.4.3 Origins of Luminescence in Minerals 139

d Kn-=k o- n·~ks 5 5·5

TR3~ Ce Dy Tb SmEuPr Nd em. 1 ~ ~~t !~ e Crd :i: lf:if ~ ~ 4fh_4f" ~ -0 r- I .. t<'l

I~ ID ; ~ t d~~O I~ >',100 nm Lr' V") 0(' dj.

3 5 6 7 10 11

Fig. 51a-e. Photoluminescence spectra of centers in minerals formed by metal ions. a Halite. b Pollucite. c Apatite. d Datolite. e Strontianite (Ce, Gd) and scheelite (the rest TR 3+). f Apatite. g Eucryptite. h Baddeleyte. i Chalcedony (silica). All spectra are measured at room temperature except d which is at liquid nitrogen temperature. Wavelength is marked in 100 nm units

The luminescence of ex-quartz and other forms of silica have been extensively studied, and many different types of center proposed for the characteristic broad blue emission. It now seems likely that at least two or three types of defect center may be responsible for blue luminescence in quartz. The alkali (or hydrogen)compensated aluminum center has often been suggested as being responsible for a deep blue emission band with a maximum intensity at just below 400 nm, whereas the intrinsic defect center involving an oxygen-oxygen linkage with a nearby oxygen vacancy has been associated with the very broad emission which peaks around 460 nm and increases dramatically in intensity and decay time on cooling to temperatures below about 150 K.

Occasionally, organic impurities may also be observed in some minerals, mainly supergene, giving rise to a very broad "white" emission. Experimentally,


there are a number of important properties and parameters which help to characterize the luminescence centers responsible for emission. The most obvious is the luminescence emission spectrum, but also its variation with temperature as the sample is cooled. For example, sharp line spectra at room temperature are indicative of the presence of either Cr3 + or trivalent rare-earths and the emergence <'f sharp lines (phonon-structure) at low temperature may also give important information. Excitation may be by ultraviolet light (UV) or even visible light, if it is strongly absorbed by the mineral in question (photoluminescence), but often electron-beam and X-ray excitation (cathodoluminescence and X-ray luminescence) are more effective.

The time for luminescence to decay after removal of the excitation, i.e., the luminescence lifetime or decay time, is also a useful parameter, long decay times indicating "forbidden" transitions. In the case of transition-metal ion centers, particularly 3d5 ions (Mn2+ and Fe3+), luminescence excitation spectra, which in certain circumstances yield the absorption profile of the luminescence center, have proved to be unequivocably diagnostic. In general, minerals containing appreciable amounts of iron are non-luminescent owing to the intrinsic quenching effect of Fez +. In such minerals, the energy of excitation is efficiently transferred from any luminescence centers which might be present to the quenching centers such as Fez + which provide a pathway for rapid nonradiative de-activation. It is, nevertheless, possible for ions such as Fe3 + (e.g., in feldspars) to act as luminescence centers in the near infrared when present in low concentration. Other "quencher" ions such as Ni2+, Co2+, and even Fez+ may also give rise to luminescence in the infrared at low concentrations particularly at very low temperatures. Ni2+ in synthetic forsterite, for example, has been shown to give two emission bands in the infrared, one showing a strong zerophonon line at about 7250 cm -1 at temperatures below about 150 K (Fig. SId). A broad featureless emission is still detectable at room temperature, although very much weaker. So far little work has been done to explore possible luminescence emission in minerals in the infrared, although potential tunable IR laser materials such as chromium-doped forsterite have been discovered recently.

Survey of Luminescent Minerals

Luminescence spectra of over 250 minerals have already been observed; clearly only a representative selection of different types of luminescent minerals can be mentioned here.

Probably the most well-known luminescent minerals are those of zinc; willemite was one of the earliest known examples. Moreover, the wurzite and sphalerite forms of ZnS are perhaps the most studied of all luminescent materials. Mnz+ readily substitutes for Zn2+ in tetrahedral co-ordination in these minerals to give a green-yellow emission, but many other luminescence centers have also been studied in ZnS, particularly donor-acceptor pairs. Magnesium minerals are often luminescent owing to the substitution of Mnz +


for Mg2 +, usually giving rise to a strong yellow, orange or red emission. Forsterite and enstatite (with less than about 1 % iron) both exhibit broad red emission under electron excitation (cathodolumines~ence) and show phononstructure at low temperatures. Magnesite and dolomite also show red emission due to Mn2+. In minerals such as dolomite, diopside, and monticellite Mn2+, ions may occupy both Mg2+ and Ca2+ sites, resulting in a double band in the luminescence spectrum in many instances, particularly in low-iron diopsides, where the two bands are well separated. The technique of luminescence excitation spectroscopy has enabled the characteristic d5 absorption profile of Mn2+ ions in each of the two sites to be measured separately in such minerals (Fig. 52).

Another characteristic of Mn2 + luminescence is a long luminescence decay time, which ranges from a few milliseconds up to several tens of milliseconds; this is indicative of the fact that the electronic transition involved (4T 1-6 A) is spin-forbidden and in some cases (e.g., carbonates) Laporte-forbidden also. Other magnesium minerals showing Mn2 + emission include brucite, sellaite and neighborite, spinel, cordierite, and borates such as fluoborite, halurgite, and ascharite. Occasionally, augite may show a red emission if the iron content is low enough. The presence of iron reduces not only luminescence intensity but also the luminescence decay time. Moreover, long-lived luminescence such as that from Mn2 + centers suffers a larger quenching effect than shorter-lived emission.

Manganese minerals themselves such as rhodonite, rhodochrosite, triplite, tantalite, and serandite, and iron-less manganese-rich varieties of calcite (manganocalcite), pectolite, and kurchatovite show a dark red luminescence of

a

13 15

d

6

, , , ,', , , , \./ 2 \

/, 17K

, , , , , , , , , 17

\

'--

b Fig. 52. a Cathodoluminescence spectrum of Mn2+ in a dolomite (BM1933, 448) showing a double band, more evident at 77 K, due to Mn2 + in both (1) Mg and (2) Ca sites. (Walker et al. 1989). b Luminescence excitation spectrum at room temperature of band 1 in a, isolated using a filter, showing the characteristic d 5

absorption bands of Mn2 + in Mg sites. c Luminescence excitation spectrum of band 2 in a showing the absorption profile ofMn2+ in Ca sites (lower crystal field). d Cathodoluminescence infrared emission spectrum of Ni2+ in synthetic forsterite at 77 K showing a large zero-phonon peak (Z). All wavenumber scales are in units of 1000 em - 1


intrinsic Mn2 +, although this may only be apparent at low temperatures. Compared with impurity Mn2 + centers, this emission is distinguished by a larger Stokes shift, somewhat lower emission energy, a shorter decay time and a low quenching temperature. Rhodonite, for example, is virtually nonluminescent at room temperature; quenching temperature is about 90 K, below which an emission band with a maximum at about 690 nm is evident with a decay time of 30 MS. This effect is due to energy transfer by excitons between Mn2+ ions. The excitation energy is eventually trapped by either a quenching impurity ion such as Fe2 + or a perturbed Mn2 + ion adjacent to a defect or impurity. The latter is more effective at low temperatures and the emission is therefore characteristic of the perturbed Mn2 + ions, rather than the "normal" ones.

Minerals containing aluminum in octahedral coordination such as corundum (and of course ruby), spinels, and beryl show sharp-line emission due to Cr3 + - the so-called R-lines near 690 nm. This emission, like Mn2 + emission, is also spin-forbidden, resulting in a lifetime of several milliseconds. Grossularite also shows Cr3 + R-line emission at low temperatures but a broad band emission at room temperature. This behaviour is due to a lower crystal-field Al3 + site in which the field-independent 2E excited state from which the R-lines originate, is only slightly lower in energy than the field-dependent 3T 1 state. At room temperature the latter state becomes thermally populated, resulting in a broad emission band with a much shorter decay time. Jadeite shows a broad emission in the red-near infrared region at room temperature, which is also probably due to a low-field Cr3 + center. In the case of alexandrite (chrysoberyl), both R-lines and broad band emission are evident. Spodumene, polylithionite, muscovite, amblygonite, and elbaite often show emission due to Mn2 +, probably substituting for aluminum.

Many aluminosilicates, in which silicon has been partially replaced by aluminum, show a strong emission near the red-infrared boundary which is due to Fe3 + in tetrahedral aluminum sites. Such Fe3 + centers occur in almost all feldspars, as well as in zeolites, sodalite, muscovite, elbaite, and eucryptite. The emission is distinguished by a broad band with a maximum in the range 690-790 nm and a decay time of around 1 ms. Such Fe3 + centers may occur due to substitution for silicon in non-aluminous silicates and silicas. The luminescence excitation spectrum of tetrahedrally coordinated Fe3+ is very characteristic and noticeably different from the isoelectronic Mn2 + centers (Fig. 51g). The minerals of calcium represent the largest group ofluminescent minerals owing to the possibility of substitution of Ca2+ by Mn2+ and divalent or trivalent rare-earths (TR). Calcite, dolomite, aragonite, plagioclases, apatite, fluorite, anhydrite, meionite, datolite, borcarite, the beryllium silicates leucophanemeliphane, and milarite all show Mn 2+ emission, TR 3+ , and sometimes TR 2+ , emission. TR 2 + ions such as Eu2 + show broad band emission (in the blue region) although Ce3+, also gives rise to a broad band emission in the UV (320-400 nm). Emission from TR 3+ ions such as Sm3+, Gd3+, Dy3+, and Eu3+ give sharp lines throughout the visible spectrum (see Fig 51e) and have decay


times around 1 ms or more because the transitions are usually spin-forbidden. Moreover, because the coupling with the lattice is so weak, the spectral positions of the TR 3 + lines are almost independent of the mineral host. Ce3 + and Eu2 +

emissions are much more dependent on the crystal host and have much shorter decay times of less than a microsecond.

The substitution of the smaller Mn2+ ion for Ca2 + in sites ranging from octahedral to seven- or eightfold coordination results in lower crystal fields giving emission colors from orange to green. Calcium minerals such as wollastonite, svabite, calciborite, nordenskioldine, harkerite, bavenite, and hydrated borates such as frolovite exhibit Mn2 + emission but not apparently TR 3 +

emission whereas danburite, scheelite, pro sopite, creedite, howlite, and shortite appear to show only TR emission.

Strontium and barium minerals such as strontianite, celestite, barite, and witherite may show luminescence due to Eu2+, Ce3+, Sm3+, or Dy3+ when genetically related to hydrothermal veins, but not when related to sedimentary rocks.

Lanthanide and yttrium minerals, namely gagarinite, yttro-fluorite, still wellilte, and monazite, show cathodoluminescence, their spectra showing sharp TR 3 + lines due to Sm, Eu, Gd, Tb, Dy, and Er. Some minerals of potassium and cesium contain TI + centers which emit in the UV near 300 nm and have a short decay time of about 1 jlS. Such minerals include K-feldspars, muscovite, rhodizite and pollucite (Fig. SIb). Many titanium and zirconium minerals have similar luminescent properties showing broad blue or green bands associated with (TiOnr- complexes where usually n = 6, m = 8. In particular, members of the sphene group, such as ramsayite and fersmanite, and leucosphenite exhibit such emissions as well as baddeleyite and zircon (Fig. 51h). Some tin minerals such as malayite and sorensenite also show very broad blue-green bands due to (Ti06)8 -, whereas narrower green and yellow bands in malayite and nordenskioldine are assigned to Mn2 + in Ca2+ sites. Cassiterite, however, shows a characteristic yellow emission with a short decay time (0.3 jls), which is thought to be due to a transition in the Sn2+ ion itself. Molybdenum minerals such as wullfenite and powellite show yellow luminescence due to the (M004)2-complex (decay time of 100-200 jls), whereas the analagous tungsten minerals, stolzite and scheelite, show either intrinsic blue emission (Fig. 53) due to the (W04)2- complex (decay time 6-12 jls) or emission due to impurity (M004)2-. Uranium-containing minerals show a variety of luminescence colors ranging from blue-green in carbonates to red in hydroxides (on cooling) due to the uranyl complex (U02)2 +. Perhaps the most well known is the phosphate autunite, which shows a strong green-yellow luminescence. The shift in the emission towards the red is accompanied by a decrease in: (1) the vibrational frequency of the O-U-O from 830 cm -1 in schrockingerite to 680 cm -1 in schoepite; (2) the decay time from around 1 ms in carbonates to around 0.1 ms in hydroxides: (3) the quenching temperature, which may be below room temperature. These effects are explained by the extension of the uranyl complex


Fig.53. X-ray excited luminescence spectra at 298 K of the minerals in which the predominant bands serve as selective properties for radiometric ore enrichment: 1 fluorite; 2 datolite; 3 danburite; 4 plagioclase; 5 microcline; 6 scheelite; 7 quartz; 8 apatite; 9 spodumene; 10 calcite; 11 barite and celestite; 12 laser-induced luminescence spectrum of cassiterite. (l"O*-superposition of oxygen-related emission bands). Wavelength scales are marked in 100 nm intervals.

due to increased interaction with the ligands (from [C03]2- to OH-), the loss of water of crystallization shielding the uranyls and disordering of the crystal lattice - factors which result in increased energy transfer between uranyls.

Applications - Technological and Geological

Luminescence properties of minerals have found many applications. For example, synthetic analogs of luminescent minerals or mineral structures have proved to be important luminophore and laser materials, e.g., the very first laser material, ruby. Indeed, recent developments in solid-state lasers have been in tunable infrared devices such as the alexandrite and Ti: sapphire lasers. The chromium-forsterite system has also been shown to lase and there are likely to be further developments in this area.

In sedimentology, cathodoluminescence has become a standard analytical tool for revealing information on provenance, growth fabrics, diagenesis, and mineral zonation. Authigenic and detrital components are easily distinguished and the chronological order of cementation events can be ascertained. The


Fig. 54a-d. Photoluminescence spectra of apatite from the deposits genetically related to the Earth mantle (on the left) and the crust (on the right) 1 carbonatite of Kovdor (Kola peninsula); 2 nepheline syenites of the Khibine mountains; 3 rare metal granite pegmatites (river Menza, East Siberia); 4 hydrothermal quartz-hubnerite veins (Bom-Gorkhon, East Siberia)

cathodoluminescence of quartz and carbonates is particularly important in this respect. Authigenic and detrital quartz often have very different luminescence properties which are currently the subject of much debate and interest.

Luminescence can be used to detect and identify minerals in sands, rock pieces, ores, outcropping in quarries, etc. It can be used to establish genetic relationships between mineral samples and their source (mantle or earth's crust) and the type of rock (Fig. 54). The key minerals serving as luminescent tracers are apatite, calcite, plagioclases, fluorite, zircon, and scheelite containing Mn2 +

or rare-earth impurity centers. A new airborne method of exploration of ore deposits has been realised with the help of luminescence lidar. Lidar embodies high power ultraviolet excimer laser and a coaxial telescope detection system installed aboard a helicopter. The tunable laser yields pulses having peak power 0.4 mJ, width range 10 ns, repetition rate 30 Hz, wavelength varying from 275 up to 400 nm. The laser "footprint" is about 400 cm2 and the distance between the closest footprints about 1 m. Airborne survey is carried out from an altitude of 30-60 m, i.e., to the scale 1: 10 000 or 1: 25 000. Data are recorded on magnetic tape.

The favorable areas are mountain and arid regions covered by vegetable and soil background up to 80-90%. Sensitivity is 0.2-0.5% by exposed area for most of the minerals. The luminescence of organic materials has a lifetime '"C much shorter than 1 J1.s. A signal from mineral targets is recorded in several spectral and time channels. To date, lidar has been test-flown and has been able to detect some ore minerals of U (minerals of uranyl), W (scheelite), Mo (powellite), and Zn (hydrozincite).

A photoluminescent signal recorded from a mineral substance can be related to some types of ore deposits. A nonluminescent background, represented by


magmatic, metamorphic, and sedimentary rocks, is related either to the rockforming minerals, containing iron in the crystal lattice (pyroxenes, amphiboles, micas) or to the rock-forming minerals containing no more than average concentration of activator elements which is not sufficient to yield noticeable luminescence (feldspars, quartz, carbonates). Thus, luminescence anomalies can be related to the processes of differentiation of some ore-forming elements and their accumulation in local structural and geomorphological traps. The most favorable targets are carbonatites (calcite, apatite, barite), skarns (calcite, scheelite, cassiterite, datolite), hydro-thermal veins containing rare metals (calcite, fluorite, barite), zones of oxidation of polymetals, silver and gold ores, and placers containing zircon and cassiterite.

Another important application is in the processing ofliquid-Iess radiometric ore enrichment (Fig. 53). The majority (99.5%) of diamonds of more than 0.5 mm in size are extracted from kimberlite ores by X-ray-excited luminescence processing. Ore enrichment of sodium-lithium pegmatites enables a spodumene concentrate (2.5% LizO) to be obtained using the Mnz + emission from spodumene for separation. Enrichment of apatite ores from carbonatites and nepheline syenites is achieved by using the Ce3+ or Mn2+ luminescence of apatite enabling the PzOs concentration to be increased from 5-15% to 10-18%. Boron silicate ores can be extracted from skarns using similar techniques by looking for Ce3 + and oxygen-related defect emissions in danburite and the Ce3 +

and Euz + emissions in datolite. Boron concentrates of 14-17% Bz0 3 can be obtained in this manner.

References

Gorobets BS (1981) Spectra ofluminescence of minerals. All Union Institute of Raw Materials (VIMS). Moscow (in Russian)

Gorobets BS, Gaft ML, Laverova VL (1978) Photoluminescence of manganese minerals. J Appl Spectrosc 28(6): 750

Marfunin AS (1979) Spectroscopy, luminescence and radiation centers in minerals. Springer, Berlin Heidelberg New York

Seigel HO, Robbins JC (1985) Luminescence method - new method of air and ground exploration of ore deposits. ITC J 3: 162 - 168

Tarashchan AN (1978) Luminescence of minerals. Naukova Dumka (in Russian) Walker Q (1985) Mineralogical applications of luminescence techniques. In: Berry FJ,

Vaughan DJ (eds) Chemical bonding and spectroscopy in mineral chemistry. Chapman and Hall, London, pp 103-40

Walker G, Abumere OE, Kamaluddin B (1989) Luminescence spectroscopy of Mn2 + centres in rock-forming carbonates. Mineral Mag 53: 201 - 11

Walker G, Burley SD (1991) Luminescence petrography and spectroscopic studies in diagenetic minerals. In: Barker CE, Kopp OC (eds) Luminescence microscopy: qualitative and quantitative applications

Spec Pub Soc Econ Pal Mineral (SEPM short course no. 25): 83 - 96 Waychunas GA (1988) Luminescence, X-ray emission and new spectroscopies. In: Hawthorne

FC (ed) Spectroscopic methods in mineralogy and geology. Rev Mineral 18: 639-694

3.5.1 Mechanisms and Parameters

3.5 Thermoluminescence and Exoelectron Spectroscopy of Minerals

3.5.1 Mechanisms and Parameters; Factors Governing Thermoluminescence

S.W.S. McKEEVER

147

Thermoluminescence (TL) is a member of a family of techniques collectively known as Thermally Stimulated Processes (TSP). Included in the TSP family are Thermally Stimulated Conductivity (TSC), Thermally Stimulated Polarization and Depolarization Currents (TSPC and TSDC), Thermally Stimulated ExoElectron Emission (TSEE), Thermal Gravimetry (TG), Differential Scanning Calorimetry (DSC), Deep Level Transient Spectroscopy (DLTS), and several more. These techniques, although differing in detail, share two essential stages: (1) the stimulation of the sample into a metastable state, and (2) the thermally stimulated relaxation of the system to equilibrium.

In the case ofTL, the stimulation takes the form of the absorption of ionizing radiation. For wide band gap insulators the irradiation may be gamma rays, X-rays, charged particles, neutrons, or anyone of many other radiation types. For small band gap semiconductors the stimulation can take the form of ultraviolet or visible illumination.

In each of the above cases, a small fraction of the incident energy may be stored within the sample via electronic processes (bound exciton creation or free carrier generation followed by trapping at defect sites) and/or atomic processes (vacancy/interstitial creation). During subsequent thermal stimulation of the specimen a proportion of this energy is released as visible light due to the radiative recombination of electrons and holes or the annealing of the atomic defects. The light emitted is thermoluminescence.

It may be concluded from the above that TL is no more than thermally activated phosphorescence with the stability of the stored energy being dictated by the physical nature of the defects concerned. As with all thermally activated processes, there is a finite probability that the energy stored will be released slowly at any temperature, with the probability being described by an Arrhenius law. The rate of release is determined by the character and parameters of each particular defect species involved in the process.

It is the purpose of this chapter to discuss the relevant parameters associated with TL production and to relate these parameters to the intensity of the recorded TL signal. The reverse problem - namely that of extracting information about the parameters from the monitored TL - is much more difficult. Nevertheless, it will be seen that by adopting simplified, phenomenological models to describe the TL process, numerical values for some of these parameters can be obtained and, in addition, much useful information about the dynamics of the thermally stimulated processes can be learned.


Common Processes of TL Production; Simplified Models

Electron-Hole Recombination. Excitation of an insulating or semiconducting sample by ionizing radiation results in the creation of excitons and free electrons and holes. These entities are delocalized in the lattice and considerable migration can take place before they become localized, or "trapped", at specific defects sites within the material. Localization of electrons and holes occurs via the process of nonradiative charge trapping (Henry and Lang 1977). Each charge may undergo several trapping and thermal detrapping processes until it is localized at a center with a binding energy of about 25 kT or greater (where T is temperature and k is Boltzmann's constant). As a result there will exist a nonequilibrium, metastable concentration of charge trapped at localized energy levels within the material's forbidden gap. Lattice disorder (impurities, vacancies, interstitials, dislocations) is thus a prerequisite for charge localization in these materials.

Upon thermal stimulation, sufficient energy may be absorbed by the trapped charges, with the result that they are thermally freed from their traps and become able once again to take part in migration processes along with further trapping and recombination. Those charges that recombine radiatively with charge of the opposite sign are lost from the process and give rise to luminescence - specifically, thermoluminescence.

The intensity of the luminescence emitted from a sample at a given temperature is related to several factors. Among these are the nature of the host lattice, the concentrations of the various traps and recombination centers directly involved in the process, the concentration of other "competing" centers, the capture cross-sections of the various centers, and the type and dose of the absorbed radiation. In order to extract information about any or all of these parameters from the thermoluminescence signal, it is necessary to assume simple models for the possible processes.

The choice of a suitable model is not trivial since it places severe restrictions on the variety of TL phenomena that one can describe. The first to present a simplified description of the TL phenomenon were Randall and Wilkins (1945), who adopted a model wherein the probability of charge retrapping was negligible compared with the process of recombination. The model invoked one type of trap and one type of recombination center only. This lead to the so-called first order description of the TL process.

A modification of this model by Garlick and Gibson (1948) led to a description of the "second-order" TL process. Here the probabilities of retrapping and recombination are made comparable. Several early theories of TL dealt only with this simple, two-level model (Haering and Adams 1960; Halperin and Braner 1960) and produced sets of rate equations to describe the process.

In the present paper we shall describe the TL process using a necessarily simplified, but slightly more realistic model than the two-level system described above. The model is that used by Dussel and Bube (1967) and later by Saunders (1969), and consists of an active electron level (AT), a set of shallow electron

3.5.1 Mechanisms and Parameters

e e e

E C ST ==

E a _________________ AT

OET -.--.-_e_ ---E f ------------------------------------------------~DHT

E v

~

=s= --9--

149

Fig. 55. The model used to describe the process of TL emission. Ec and Ev represent the edges of the conduction and valence bands, respectively. E f and E. represent the Fermi level and the energy of the active electron trap, respectively. The trap depth E is thus Ec - E •. The other terms are defined in the text

traps (ST), a set of deep, thermally disconnected electron traps (DET), and a set of thermally disconnected hole traps (DHT). The model is sketched in Fig. 55.

Irradiation of the system at a temperature To results in the creation of free electrons and holes, many of which become trapped at the various electron and hole traps within the material. Excitation of the sample continues in this way until equilibrium is reached, at which point there exists a large concentration of trapped charge carriers. Thermal stimulation of the sample results in the release of trapped electrons from the shallow traps. Over the temperature region of interest (the "active" temperature range) electrons from AT are released and the freed electrons recombine with holes trapped at DHT, with the subsequent emission of thermoluminescence, of intensity I(T).

The probability per unit time p of electron release at temperature T is given by the Arrhenius equation

p = sexp{ - E/kT} , (1)

where s is the "attempt frequency" and is related to the lattice phonon vibrational frequency and the entropy change associated with the transition. Defining nc to be the concentration of free electrons, n the concentration of trapped electrons, and N the concentration of available active electron traps, we have the rate equations:

dn/dt = - nsexp{ - E/kT} + nc(N - n)A

and

dnc/dt = - dn/dt - nJr: ,

(2)

(3)

where A is the trapping transition probability and t is the recombination lifetime and is equal to (nhAr)-l. Here nh is the concentration of holes trapped at the recombination sites and Ar is the recombination transition probability. Approximate solutions to these equations normally proceed via the introduction of the quasi-equilibrium (QE) approximation which states that dnc/dt ~ nc/t, or dn/dt = nc/t. Since the intensity of the thermoluminescence emission is given by I(T)


= cjJnclr:, where cjJ is the luminescence efficiency, then:

I(T) = cjJnsexp{ - E/kT}/[r(1/r + (N - n)A)]. (4)

With the inclusion of the slow-retrapping or first-order assumption that r- 1 ~ (N - n)A, we have:

T

I(T) = cjJnosexp{ - E/kT} exp [ - (s/P) f exp{ - E/kB}dB] , (5) To

which is the Randall and Wilkins (1945) expression for thermoluminescence with no as the original population of filled electron traps. An example of a firstorder thermoluminescence peak (a "glow-curve") is shown in Fig. 56.

A characteristic of thermoluminescence curves exhibiting first-order kinetics is that the temperature of the peak is independent of the level of trap filling, no. This is, in fact, found to be the case for most minerals, and nonfirst-order kinetics are somewhat rare. A well-known example of an exception to this, however, is feldspar for which material it is observed that the peak position shifts to lower temperature as the level of trap filling is increased. Most feldspars have been shown to follow second-order kinetics for which the assumptions used in the above analysis do not apply. Specifically we now have r -1 ~ (N - n)A. Expressions for the thermoluminescence intensity only become available if one also assumes that no/N ~ 1. With no thermally disconnected traps this leads to

4~-----r------~----~r------'-------r----~ a b c

3

Cii 2 -'c

::I

.ri ... ~

..J 1 ~

OL.~~-L~~--~~ __ ~~~ __ ...A-~~-...A~~ __ ~

o 100 0 100 0 100 200 Temperature I"CI

Fig. 56a-c. A typical, first-order, thermoluminescence peak from gamma-irradiated natural quartz. The shape is asymmetrical, dropping steeply on the high temperature side compared to the low temperature side. Peak fitting reveals an activation energy of 0.8 eV. (After McKeever et al. 1985)

3.5.1 Mechanisms and Parameters 151

the Garlick and Gibson (1948) expression:

T

I(T) = (n5s/ N)exp{ -E/kT}/[1 + (n5s/NfJ)fexp{ -E/ke}de]2. (6) To

Vacancy-Interstitial Recombination. In addition to the creation of electron-hole pairs, ionizing radiation also creates excitons. Nonradiative relaxation of the excitons results in energy transfer to the lattice with the formation of lattice vacancies and their associated interstitial atoms or ions. The details of the necessary mechanisms have been worked out for several material types (e.g., oxides (Hughes and Henderson 1972; Griscom 1978), halides (Hoh 1982; Hoh and Tanimura 1986) and fluorites (Hayes and Stoneham 1974)). In many materials the vacancies and interstitials so-formed are unstable and recombine quickly. However, in materials with a large degree oflattice disorder (impurities, dislocations, etc.), stabilization of the defects is possible. Thus, stable vacancies and interstitials perturbed by nearby impurities are prevalent in mineralogical materials following irradiation.

Heating after irradiation causes the vacancies and the interstitials to recombine. The energy emitted as a result of this recombination may be lost nonradiatively, or it can be released via the emission of a photon, i.e., thermoluminescence. The exact mechanisms by which the photon is emitted may be complex and vary from material to material. The emission wavelength will depend upon the nature of the nearby perturbing impurity. Thus, intrinsic triplet-state emission perturbed by the impurity may be observed, or, via the process of energy transfer, the emission may be characteristic of the impurity itself. For example, CaF 2 doped with rare earth impurities produces perturbed, intrinsic thermoluminescence emission at 280 nm and emission characteristic of the particular rare earth or other impurity ion (Jassemnejad and McKeever 1987).

The mobile species during these processes are interstitial atoms (i.e., interstitial ions with trapped electrons, so-called Hz-centers) and the recombination sites are electrons trapped at cation vacancies (F z-centers). The equations describing the process of recombination are of exactly the same form as those given above, with appropriate changes in the definitions of the terms. H is important to realize that since the vacancy and the interstitial are highly spatially correlated, then the thermoluminescence peak which results from these processes is first-order.

Analysis of Thermoluminescence Curves

Model Analysis. Before one can proceed to analyze an experimental thermoluminescence curve, it is necessary to verify whether or not the varied assumptions used in the derivation of the simple equations are, in fact, justified. The most important of these is the QE approximation (dnc/dt ~ nc/r). To test for the


validity of this assumption it is useful to define a function q(T) thus:

q(T)nclr = dnc/dt (7)

or

Q(T)nc/r = - dn/dt, (8)

where Q(T) = q(T) + 1. From the definition of q one can see that if one can measure simultaneously the thermally stimulated conductivity (T(T) = nceJ.l. and the thermally stimulated luminescence I(T) = nc/r, then the differential of the

12

10

8

.I!! 'c ::J

..ci 6

~ ...J I-

4

2

0

E o

100 120 140

Temperature (K)

160

Fig.57. An example of a Q(T) function for a set of theoretical first-order thermoluminescence and peaks, as a function of heating rate. (Mter Lewandowski and McKeever 1991)


conductivity curve divided by the thermoluminescence curve will give directly the shape of the q (T) function assuming no other recombination processes are active. Furthermore, since q(T) = 0 at the maximum of the conductivity peak and -+ - 1 as T -+ 00, then one can estimate, from the q (T) function thus obtained, how far the system is from QE and over what temperature ranges. A theoretical shape for a Q(T) function is shown in Fig. 57 for the model of Fig. 55, assuming slow retrapping.

The second assumption of importance is that of kinetic order. As noted above, first-order thermoluminescence peaks are characterized by an invariance of the position of the peak as a function of initial trap population. Thus, this property can be tested by following the variation in the peak position as a function of irradiation time (i.e, absorbed dose). Any shifts in the peak with absorbed dose should lead one to suspect nonfirst-order kinetics. An empirical estimate of the kinetic order was devised by Chen (1969) who used a geometrical factor JJ.g = fJ/w, where w is the full width of the peak at its half height (in QC) and fJ is the high temperature half-width (in QC). Theoretical first-order peaks were shown to have JJ.g = 0.42; second-order peaks gave JJ.g = 0.52. Intermediate values of JJ.g correspond to intermediate kinetics for which, unfortunately, physical meaning is lost.

The simple models for charge trapping and de-trapping can also be examined experimentally. In the model described in Fig. 55, it was explicitly assumed that only the electrons are thermally unstable over the "active" temperature range. The thermal release of holes from the recombination centers was not allowed for. If this, in fact, were not the case, then the thermoluminescence intensity would no longer be given by I(T) = nJr: = dnb/dt. Several situations may cause this, including hole release (giving rise to thermal quenching; Fillard et al. 1978), temperature-dependent capture-cross-sections (de Murcia et al. 1980) and refilling of recombination centers (Fillard et al. 1978). These difficulties can also be tested for from a simultaneous measurement of the conductivity, since the ratio of thermally stimulated conductivity to thermoluminescence is given by R = constant x (l/nb) and thus dnb/dt = constant d(I/R)/dt. Thus a plot of I(T) against d(I/R)/dt should reveal if I(T) = dnb/dt (Gasiot and Fillard 1977).

One final piece of information to be extracted from simultaneous conductivity and luminescence measurements relates to the recombination lifetime r. Since I(T) = nc/r, then dI/dt = (l/r)dnc/dt - (nc/r2)dr/dt. Thus if r is a constant during the luminescence production the thermally stimulated conductivity and luminescence peaks will have their maxima at the same temperature. However, if r is nonconstant (i.e., decreasing with increasing T) then the conductivity peak will appear at a lower temperature than the luminescence peak.

Glow-Curve Cleaning. Before analysis of a thermoluminescence glow-curve can be attempted, one practical difficulty has to be overcome, namely that real glowcurves rarely consist of isolated glow-peaks. This, in turn, gives rise to problems


of two kinds. First, the presence of more than one glow peak points immediately to the possibility that the traps will interact with each other. That is, charge detrapped from one level will either recombine at the recombination site, or be retrapped in deeper traps. The models applied to analyze these cases, however, rarely take into account interactions of this type and treat the glow peaks as isolated entities. Levy (1984) and Bull et al. (1986) have analyzed theoretically the effects of interactive kinetics of this type and discussed the potential errors that can accrue from using the simple models.

The second difficulty is a more practical one. The observed glow-peaks are often overlapping to the extent that clean peaks, necessary for analysis, are seldom obtained. The usual course is the adopt of a "thermal cleaning" program in which the sample is partially heated to a temperature which removes the low temperature, interfering signals. Reheating the sample will then reveal a clean, higher temperature peak. Several variants of these cleaning processes exist. Each are reasonably useful in arriving at clean, rising portions of the TL signals, but rarely can a complete, clean peak be isolated (making application of peak shape methods of analysis difficult).

Methods of Analysis. Several methods of analysis have been devised to extract parameter values from real glow-curves. Different classes of analysis exist, based on peak shape methods, peak fitting methods and heating rate methods. There are too many to review adequately in this chapter, so only the most popular techniques will be mentioned. Complete reviews of the available methods can be found in McKeever (1985) and, especially, Chen and Kirsh (1981).

All of the methods described here have been developed from the equations which describe thermoluminescence emission assuming the simple model described in the previous section. The Initial Rise technique, devised by Garlick and Gibson (1948), relies on the fact that at low temperature, just as the trap starts to empty (i.e., when n is approximately constant) all of the thermoluminescence equations can be described by:

I(T) = constant exp { - E/kT} . (9)

Thus a plot of In(I) versus 1/kT for this initial rise part of the curve will reveal a straight line of slope - E/k, from which E can be calculated. The frequency factor, s, can then be calculated from the position of the peak maximum, T m

using

PE/kT! = s exp { - E/kT m} , (10)

where P is the rate of heating the sample during the recording of the thermoluminescence signal.

Since Eq. (9) applies in all cases, this method of calculation is independent of kinetic order. Recently, it has also been shown to be independent of the QE approximation (Lewandowski and McKeever 1991), making it a valuable analytical method. Its accurate use, however, relies on a clean rising part of the thermoluminescence peak - fortunately something that can often be achieved using the thermal cleaning techniques described above.


The second family of techniques is known as the heating rate methods. Here the most popular is that of Hoogenstraaten (1958), who made use of Eq. (10) to plot In(T~p) against I/Tm , from which a slope of Elk and an intercept of In (E/sk) is obtained. This method is useful because it relies only upon an accurate assessment of the shift in the peak as a function of heating rate p and does not require clean peaks. Too much overlap with neighboring peaks, however, can give rise to false peak positions. Lewandowski and McKeever (1991) examine the influence of the QE approximation on this method.

The final popular experimental method to extract parameter values is that of peak fitting. Here values for E, s, kinetic order parameter b (b = 1 for first-order and b = 2 for second-order), and capture-cross-section for the trap involved can be extracted (using the fact that s = NcvS, where Nc is the density of states in the delocalized bands, v is the carrier thermal velocity and S is the capture-crosssection.) This is a powerful method and one that gives substantial information about the traps under study. Example applications of this method to mineral structures are shown in Figs. 56 and 58. Figure 56 gives an example of fitting a single, first-order peak, whereas fig. 58 illustrates the deconvolution of the complex glow-curve of an ordinary chondrite meteorite using peak fitting to the expression for second-order kinetics. The major luminescent mineral in these materials is feldspar.

Several of the above methods rely on the QE approximation and the simple, noninteracting model described earlier. These and other difficulties can mean that physical interpretation of the evaluated parameters may be uncertain. However, for mineralogical studies, it is not always necessary to interpret on a microscopic scale the actual meaning of, say, the parameter E. Instead E may be simply be interpreted macroscopically as the "activation energy" for the thermoluminescence process. In this case an important question than becomes "can we

60

~ 40 c: ~

.ci ~ ..J 20 I-

0 4

200 300 400 500 T(Oe) -+

Fig. 58a, b. Computerized glow-curve fitting of the complex thermoluminescence signal from two ordinary chondrite meteorites. a Soko Banja. b Estacado. The individual peaks revealed by the deconvolution routine are shown by the dotted lines and each peak is characterized by second-order kinetics. (After McKeever 1980)


use this parameter to predict the thermal stability of the luminescence signal under a variety of conditions?" In the case of the meteorite analysis described above this was shown to be the case. The parameters obtained from the analysis were used to predict accurately the thermal stability properties of this material at a variety of temperatures. This information is now being used to estimate the terrestrial age and orbital temperature of this class of meteorite.

Summary

This brief review of the parameters affecting thermoluminescence glow-curves is intended to show the reader how practical information about the thermoluminescence process can be extracted using simple models. In favorable cases, clear interpretations and understanding of the calculated parameters may be possible; in other cases, all that may be required are empirical parameters which can be used to predict the thermal stability properties of the system. The review has overlooked situations involving trap level distributions, tunneling transitions, and energy transfer processes, and has only mentioned, but barely discussed, charge transfer interactions, luminescence quenching, and simultaneous hole and electron release. All of these processes may be encountered in practice and lead to complications which may not be easy to unravel. Nevertheless, analysis of thermoluminescence glow-curves based on the simple models described herein has proved to be of value in a wide variety of circumstances.

References

Bull RK, McKeever SWS, Chen R, Mathur VK, Rhodes JF, Brown MD (1986) Thermoluminescence kinetics for multi peak glow curves produced by the release of electrons and holes. J Phys D Appl Phys 19: 1321-34

Chen R (1969) Glow curves with general order kinetics. J Electrochem Soc 116: 1254-7 Chen R, Kirsh Y (1981) Analysis of thermally stimulated processes. Pergamon Press, Oxford de Murcia M, Braunlich P, Egge M, Mary G (1980) Thermally stimulated relaxation of

photoconverted CdF 2: Sm3+. Sol State Commun 33: 737-41 Dussel GA, Bube RH (1967) Theory of thermally stimulated conductivity in a previously

photoexcited crystal. Phys Rev 155: 764-79 Fillard JP, Gasiot J, Jimenez J, Sanz FL, de Sala JA (1977) Evidence of refilling of

recombination centers during thermal stimulation. J Electrostat 3: 133-8 Fillard JP, Gasiot J, Manifacier JC (1978) New approach to thermally stimulated transients:

experimental evidence for ZnSe:Al crystals. Phys Rev B18: 4497-508 Garlick GFJ, Gibson AF (1948) The electron trap mechanism of luminescence in sulphides

and silicate phosphors. Proc Phys Soc Lond A60: 574-90 Gasiot JP, Fillard JP (1977) Correlation in simultaneous TSC and TSL measurements. J Appl

Phys 48: 3171-2 Griscom DL (1978) Defects and impurities in alpha-quartz and fused silica. In: Pantelides ST

(ed) Physics of Si02 and its interfaces. Pergamon Press, N.Y., pp 232-52 Haering RR, Adams EN (1960) Theory and application of thermally stimulated currents in

photoconductors. Phys Rev 117: 451-4 Halperin A, Braner AA (1960) Evaluation of thermal activation energies from glow curves.

Phys Rev 117: 408-15

3.5.2 Thermoluminescence Applications 157

Hasan F, Haq M, Sears DWG (1987) Natural thermoluminescence of Antarctic meteorites: a study of thermal/radiation history and pairing. Proc 17th Lunar and Planet Sci Conf, Part 2, J Geophys Res 92: E703-6

Hayes W, Stoneham AM (1974) Color centers. In: Hayes W (ed) Crystals with the fluorite structure. Clarendon Press, Oxford, pp 185-280

Henry CH, Lang DV (1977) Nonradiative capture and recombination by multiphonon emission in GaAs and GaP. Phys Rev B5: 989-1016

Hoogenstraaten W (1958) Electron traps in zinc sulphide phosphors. Philips Res Rep 13: 515-693

Hughes AE, Henderson B (1972) Color centers in simple oxides. In: Crawford JR, Slifkin LM (eds) Point defects in solids. Plenum Press, N.Y., pp 381-490

Itoh N (1982) Mechanisms of electron-excitation-induced defect creation in alkali halides. Radiat Eff 64: 161-9

Itoh N, Tanimura K (1986) Radiation effects in ionic solids. In: Wilson IH, Webb RP (eds) Proc 3rd Int Conf Radiation Effects in Insulators, Surrey, 1985. Gordon and Breach, London, pp 435-53

Jassemnejad B, McKeever SWS (1987) Photoreversible charge transfer processes and thermoluminescence in CaF2 :Ce. J Phys D Appl Phys 20: 323-8

Levy PL (1984) Thermoluminescence kinetics in systems more general than the usual 1st and 2nd order kinetics. J Lumin 31/32: 133-5

Lewandowski AC, McKeever SWS (1991) Generalized description of thermally stimulated processes without the quasi-equilibrium approximation. Phys Rev B 43: 8163-8178

McKeever SWS (1980) On the analysis of complex thermoluminescence glow-curves: resolution into individual peaks. Phys Stat Sol a 62: 331-40

McKeever SWS (1982) Dating of meteorite falls using thermoluminescence: application to Antarctic meteorites. Earth Planet Sci Lett 58: 419-29

McKeever SWS (1985) Thermoluminescence of solids. Cambridge Univ Press, Cambridge McKeever SWS, Chen CY, Halliburton LE (1985) Point defects and the pre-dose effect in

natural quartz. Nuc1 Tracks 10: 489-95 Randall JJ, Wilkins MHF (1945) Phosphorescence and electron traps I. The study of trap

distributions. Proc R Soc Lond A184: 366-89 Saunders J (1969) Thermally stimulated luminescence and conductivity of insulators. J Phys C

Sol State Phys 2: 2181-98 Sears DWG, DeHart JM, Hasad FA, Lofgren GE (1989) Induced thermoluminescence and

cathodoluminescence studies of meteorites. In: Coyne LM, McKeever SWS, Blake DF (eds) Spectroscopic characterization of minerals and their surfaces. American Chemical Society, ACS Symp Ser 145: 190-222

3.5.2 Thermoluminescence Applications

s.w.s. McKEEVER, V.K. VLASOV, O.A. KULIKov, and K.S.V. NAMBI

Introduction

The technique of thermoluminescence (TL) has found application in a wide variety of areas (McKeever 1985), especially relating to the study or use of minerals. Studies of mineralogical classification, radiation dosimetry, age determination, defect structure, and many others have all benefited from the use of TL. In this chapter we briefly review some of the mainstream applications, highlighting the successes and the limitations.


Age Determination

Principles. All matter on Earth is subject to ionizing radiation. The source of this continuous irradiation is from external gamma and cosmic rays, and alpha, beta, and gamma irradiation from the material's own internal radioactivity. Although the net irradiation rates experienced by matter in nature will be quite small, the integrated radiation doses absorbed by them over archeological or geological periods can be significant, and the induced microscopic effects can be measurable. Such possibilities have led directly to the application of TL for dating archeological and geological artifacts. Techniques of TL dating have been reviewed in several publications, most recently by Aitken (1985).

Confining our considerations to electronic effects in minerals, irradiation produces a population of trapped electrons and holes, as described in Chapter 3.5.1. Depending upon the binding energy of the trapped charge, the residence time (at a lattice defect) can vary as widely as a fraction of a second to several millions of years. However, a thermal stimulus (typically heating to 500 DC) can cause the release of the trapped charge and the subsequent emission of luminescence. The proportional dependence of the intensity of the luminescence emitted to the cumulative radiation dose received since the last heating paves the way for dating. The rate of irradiation from environmental radioactivity practically remains constant, thereby causing a steady increase of trapped charge concentration with time. This increase in the stored (or latent) luminescence continues until excavation when the thermoluminescence intensity from the freshly excavated sample can be related to its age via the age equation:

TL age (years) = TL acquired/TL acquisition rate. (1)

Dating Methodology. The basic methodology of TL dating is illustrated in Fig. 59. Here the latent luminescence accumulated over geological times may be removed by a suitable "zeroing" event. This may be a thermal event such as

TL Accumulated in Antiquity

NTL

Sample Retrieval

Time _

Zeroing Event

ED

Fig. 59. Schematic representation of the thermoluminescence dating procedure

Dose -(Laboratory Irradiation)


metamorphism, volcanism, or, in the case of archeological pottery, kiln firing. Other possibilities include zeroing via the optical release of trapped charge, as may occur during the deposition of terrestrial or marine sediments. During antiquity, there then follows a period of trapped charge accumulation, via the absorption of radiation from the environment, until the sample is retrieved for analysis.

In a simple experiment, the natural, accumulated dose (or "equivalent dose", ED) can be determined by measuring the growth of the thermoluminescence signal above the natural level (NTL) after the addition of known doses in the laboratory delivered to equivalent aliquots.

If additional gamma doses are administered during the above procedure, the ED estimated is strictly the gamma equivalent of the natural dose. However, it has been realized for some time that the TL production efficiency is not the same for all radiation types. Specifically alpha particles are known to be less efficient (by approximately a factor of 10) than gamma or beta particles at inducing a TL signal. Thus a relative efficiency factor for alpha particles (k) has to be determined experimentally (Zimmerman 1972).

Gamma (and cosmic ray) dosimetry will be a straightforward problem if the environment around the object is homogeneous to at least a radius of about 30 cm. The net dose could then be conveniently found using environmental thermoluminescence dosimeters (TLDs) or sensitive environmental radiation survey instruments. The dose rate evaluation problem could be further simpli. fied by eliminating the alpha component after surface etching of large mineral grains if they are present in the sample. It will be evident then that the dose absorbed by fine grains ( < 10 J.lm) will be greater than that of the larger grains ( '" 100 J.lm) whose outer surface has been removed by etching; the difference is due to the alpha component. The age equations can then be written:

ED Age t = Cg

, kDa + Db + Dg + Dc bDb + Dg + Dc' ED j

(2)

where D is absorbed dose and the subscripts a, b, g, and c refer to the contributions due to alpha, beta, gamma, and cosmic ray irradiation, respectively; b is the beta attenuation factor appropriate for the large inclusions being used in the measurements. EDcg and ED j are the equivalent doses in the fine grain and large inclusions, respectively. Using the above equations, one can eliminate the need for determining the external radiation component by the technique of "subtraction dating", thus:

A EDcg - ED j ge t = :-::-----=----,---, kDa + (1 - b)Db

(3)

The assessment of the internal alpha and beta doses then proceeds using conventional alpha-counting and K-analysis.

Fading. It is very important that the thermoluminescence measurements and calibrations should pertain to the nonfading part of the TL signal. Typical glow

160

o w

.0 ..... !1l

...J I-


500 600

Temperature,oC

Fig. 60. Typical glow curves from a crushed pottery fragment and the delineation of the plateau region. a Thermal background from the thermoluminescence oven; b natural thermoluminescence NTL signal; c thermoluminescence following irradiation in the laboratory to a dose of p rad on top of the natural dose; d a plateau test in which ED is evaluated as a function of glow-curve temperature

curves from two aliquots of a crushed pottery sample prior to and after artificial irradiation (with a dose of P rad) are shown in Fig. 60. An important feature of the glow curve is the presence of peaks at low temperature when the sample is irradiated in the laboratory compared to that resulting from irradiation in antiquity. Only the deep, thermally stable traps which have not undergone any charge loss ("fading") during antiquity contribute to the NTL signal. When the ED is evaluated at different temperatures in the glow curve, an ED or age "plateau" is observed over that region of the glow curV6 for which no fading of the signal has taken place during antiquity.

The fading observed can be both thermally induced or nonthermal (i.e., "anomalous"). Thermal fading follows the straightforward Arrhenius law [Eq. (1) in chapter 3.5.1]. Anomalous fading is more difficult to deal with since its cause may be related to several different phenomena (quantum mechanical tunneling, localized transitions, defect clustering, and others).

Disequilibrium and the Effect of Moisture. There are at least two important problems that need to be considered when assessing the internal dose rate. First, there has to be a guarantee that the long chain of radioactive daughter products of 238U and 232Th were in equilibrium throughout antiquity so that the dose rates remain constant. Equilibrium can easily be disturbed by the escape of gaseous radon and thoron, and these can lead to errors of 50% in the dose rate. Second, the dose rate factors need a wetness correction depending upon the


moisture content of the site. If the site was well preserved during antiquity, the best way to account for this is to bury a TLD at the site and leave it exposed to environmental radiation for about 1 year. The effect of the moisture can then be taken into account.

Zeroing Event. The event to be dated should be well defined, guaranteeing a zeroing of the TL signal. Generally, heating to a sufficiently high temperature (e.g., during the manufacture of pottery or during volcanic activity) will ensure that all prior radiation history of the sample has been erased. In the case of sediment dating (MejdahlI976), the zeroing event is not heating, but rather the exposure of the sample to sunlight, thereby causing the optically stimulated release of the trapped charges. With this zeroing mechanism, however, a residual signal is usually seen, and procedures have been adopted to account for it (see discussions by McKeever 1985; Aitken 1985; Mejdahll986; Vlasov and Kulikov 1987). The in-situ crystallization of minerals (e.g., calcites, gypsum) is also a TL zeroing event and thus crystallization ages can be evaluated.

Radiation Dosimetry

General. The correlation between the intensity of the thermoluminescence produced from a material following irradiation and the dose of radiation absorbed has led to the use of this technique in radiation dosimetry. The sensitivity and wide dynamic range of TL in this application is unrivaled. It is routinely used to measure, on the one hand, doses as small as those delivered by the natural environment over a period of days, and, on the other, doses as large as those delivered inside a nuclear reactor (a dose difference of a factor of 1011). The high sensitivity enables TLDs to be used in medical dosimetry, whereby the physician is able to insert small TLDs into the body, expose the patient to therapeutic radiation, and retrieve the sample for assessment ofthe dose actually delivered to an internal organ. No other form of dosimetry is able to do this.

Required Properties of TLD Materials. An obvious, desirable property of a TLD material is that it exhibits a linear response to absorbed doses. Unfortunately, most common TLD materials display the property of supralinearity - that is they display a TL ex Dn law, where n ~ 1. At high doses, sublinearity is observed wherein the signal saturates (or even decreases for high enough doses). The cause of supralinearity is not easy to pinpoint, but in the most common TLD material in use today (LiF doped with Mg and Ti), it appears to be caused by competition effects during the charge trapping and de trapping processes (Mische and McKeever, 1990). Sublinearity is related to the filling of the available traps by the radiation dose, while the decrease in the response at high doses is caused by unknown mechanisms of radiation damage.

At sufficiently low doses, however, most TLD materials display an approximately linear response to the absorbed radiation. The slope of a TL versus D curve defines the sensitivity of the material and the most sensitive materials


presently known (LiF:Mg, Cu, P, and A120 3 ) are able to record background dose levels over a period of a few days.

Energy Response. Since the intensity of the emitted thermoluminescence is related to the energy absorbed, the variation in the material's absorption coefficient as a function of incident energy is an important material characteristic. The photon energy response is defined as the response of the system at a given energy divided by the response for 1.25 Me V gamma photons. The higher the effective atomic number of the material and the lower the energy of the incident photons, the larger the value of the photon energy response (Horowitz 1984). Thus, CaF 2 has an "over-response" of approximately 3 at an energy of 10- 1 MeV. LiF, on the other hand, has a very weak over-response and is said to be "tissue equivalent" - hence its popularity in medical dosimetry applications.

LET Response. TLDs are also sensitive to irradiation by energetic particles (alpha particles, protons, fission fragments). In these applications the important parameter is the mass stopping power of the material, or the Linear Energy Transfer (LET) of the particle. Particles with a high LET deposit their energy in tight, localized tracks of high ionization density, whereas low LET irradiations (beta, gamma photons) produce an isotropic "sea" of charged particles. As a result, the energy deposition pattern is entirely different and each particle has a different efficiency in producing thermoluminescence. This was already noted in the above section on dating, when the efficiency of alpha particles was discussed relative to the other radiation types. Once the LET of the particles increases beyond approximately 2 ke V p.m - 1, the TL production efficiency begins to drop, compared with 1.25 MeV gamma photons (Horowitz 1984).

The dose response for these types of radiation is also very different to that of high energy gamma irradiation. Because of the high ionization density within the track and the isolation of the tracks, the dose response tends to be linear at all doses up to saturation (McKeever and Horowitz 1990). The cause is related to the difficulty of interaction between the isolated tracks.

Fading. As with the materials used in dating, TLD materials should exhibit high stability of the signal over exposure and analysis time scales. Since these are much shorter than those encountered in dating applications, lower temperature glow peaks can be used than is usual in dating. Stabilities over several months or years are normally all that are required.

Materials. Table 7 lists several popular TLD materials along with a summary of the main properties. Several of the materials used as TLDs require elaborate annealing proe,edures in order to reuse them with a sufficient degree of accuracy. This is a disadvantage and, unfortunately, the most popular materials for personnel dosimetry (LiF based systems) all appear to require annealing.

Tab

le 7

. C

hara

cter

isti

cs o

f so

me

key

TL

D m

ater

ials

Pho

spho

rus

Zef

f G

low

A

ppro

xim

ate

Rat

io o

f T

L

peak

em

issi

on

resp

onse

°C

m

axim

um

S (3

0 ke

V)/

(n

m)

S (6

OC

o)

LiF

:Mg

, T

i 8.

14

210

425

1.3

CaF

2 (n

atur

al)

16.3

26

0 38

0 13

C

aF2

: D

y 16

.3

200

480

13

CaS

04

:Tm

15

.3

220

450

10

Li 2

B40

7: C

u 7.

4 20

5 36

8 0.

8 M

gS

i04

:Tb

10

.5

200

380

to 4

00

4 B

eO

7.13

38

0 to

220

33

0 1.

4 A

I 20

3 (c

orun

dum

) 10

.2

180

420

3.5

A120

3• 3

P20

S' M

gO

11.7

25

0 57

0 3.

5 : M

n (

glas

s)

a N

A -

no

t ap

prec

iabl

e

Sen

siti

vity

L

inea

r (i

n pe

r-un

it)

rang

e (G

y)

I 5

-10

to

I 20

to

50

10 t

o 50

16

10

to

I 60

10

to

30

8 10

to

10

40 t

o 10

0 10

to

4 3

10 t

o 0.

5 10

0 10

to

10

0.1

10 t

o 10

Fad

ing

The

rmal

5-1

0%

per

yea

r N

A

25%

per

mo

nth

7

-30

% p

er m

on

th

25%

per

2 m

onth

s N

A

7% p

er m

on

th

2%

per

4 m

onth

s 23

% p

er y

ear

Opt

ical

NA

a

Sen

siti

ve

Sen

siti

ve

30%

in

5 h

10%

in

3 h

Sen

siti

ve

Sen

siti

ve

Sen

siti

ve

Sen

siti

ve

~

v. tv ..., ::r

-(1

) .... 3 0 C- 2. 0 (1) '" ()

(1)

0 ()

(1) ;J>

"0

'E.

. n'

~ o· 0 '" - '" W


Research is in progress to develop new materials or procedures which do not require the annealing of the samples.

Mineral Applications. Fluorites (obtained in Belgium, Brazil, and India) have been found to have extremely high sensitivity in radiation dosimetry applications (Mejdahl 1978). A disadvantage is the high atomic number of CaF2 which gives rise to a significant over-response at low energies. To correct for this, suitable high Z shield materials are employed, e.g., brass-encapsulated fluorite powder has been used on a wide scale in radiation monitoring applications (Nambi et al. 1987). The activators in natural fluorite are the naturally present rare earth impurities such as Ce, Gd, Dy, Sm, Tb, etc. Fluorites are known to exhibit photo-transferred thermoluminescence properties which can be gainfully employed in UV exposure monitoring (Sunta et al. 1970).

Natural quartz has also been used effectively in radiation dosimetry applications. This material can be used in two separate ways - either by utilizing the high temperature, stable glow peaks, and undertaking a conventional dosimetry study, or by using the low temperature, thermally unstable peak at 100 DC. This latter peak has been observed to undergo an increase in sensitivity proportional to the imparted dose (the so-called pre-dose effect). By calibrating the sensitivity increase, an estimate of the absorbed dose may be arrived at. Both techniques have been employed recently in reassessing the dosimetry at the A-bomb sites of Hiroshima and Nagasaki, and in fallout dose assessment in the United States (Maruyama et al. 1987; Haskell et al. 1985).

Other Applications

Defect Characterization. The TL emission from minerals can be usefully employed in elucidating the nature of intrinsic and radiation-induced solid state defects by establishing correlations with other characteristics, such as optical absorption and electron spin resonance (ESR) (Nambi 1985; Yang and McKeever 1990; McKeever 1989). The sensitivity of TL enables as few as 109

defects to be detected in a system. While identification is not possible by TL alone, the method is very useful in establishing the thermal stabilities of the defect centers.

Geological Applications. Thermoluminescence can be correlated with a host of geological and/or geochemical aspects such as radioactive anomalies, metamorphism, alteration, tectonic influences, lithological peculiarities, paleoclimatic conditions, etc. (McDougall 1968). However, such attempts have remained largely empirical and in general have not been pursued with seriousness in recent times. An exception to this is the study of meteorites. Here, substantial information relating to the mineralogical classification, thermal histories, and orbital histories can be extracted from the TL signal (McKeever 1985; Sears 1988; Sears et al. 1989).


Summary

Thermoluminescence is a widely used experimental technique and enjoys application in a broad range of disciplines. Some of these uses have been briefly noted in this chapter. Further details can be found in the specialist texts and publications.

References

Aitken MJ (1985) Thermoluminescence dating. Academic Press, New York Haskell EN, Kaipa PL, Wrenn ME (1985) Environmental and accident dosimetry using the

Pre-Dose TL technique. Nucl. Tracks 10: 513-6 Horowitz YS (ed.) (1984) Thermoluminescence and thermoluminescent dosimetry, vols. I-III.

CRC Press, Boca Raton Maruyama T, Kumamoto Y, Ichikawa Y, Nagamoto T, Hoshi M, Haskell E, Kaipa P (1987)

U.S.-Japan joint reassessment of atomic bomb radiation dosimetry in Hiroshima and Nagasaki. Final Report DS86, 1

McDougall DJ (ed.) (1968) Thermoluminescence of geological materials. Academic Press, London, New York

McKeever SWS (1985) Thermoluminescence of solids. Cambridge University Press, Cambridge

McKeever SWS (1989) Energy-storage mechanisms and thermoluminescence processes in minerals. In: Coyne LM, McKeever SWS, Blake DF (eds) Spectroscopic characterization of minerals and their surfaces. American Chemical Society, ACS Symposium Series 145: 166-79

McKeever SWS, Horowitz YS (1990) Charge trapping mechanisms and microdosimetric processes in lithium fluoride. Radiat Phys Chern 36: 35-42

Mejdhal V (1978) Measurement of environmental radiation at archaeological excavation sites by means of TL dosimeters. PACT 2: 70-83

Mejdhal V (1986) Thermal dating of sediments. Radiat Protect Dosim 17: 219-27 Mische EF, McKeever SWS (1990) Mechanisms of supralinearity in lithium fluoride dos

imeters. Radiat Protect Dosim 29: 159-75 Nambi KSV (1985) Proc. First Natl. Seminar on Defects in Insulating Solids, RIT, Jamshed-

pur, India, 3 . Nambi KSV, Bapat VN, David M, Sundaram VK, Sunta CM, Soman SD (1987) Country

wide environmental radiation monitoring using thermoluminescence dosemeters. Radiat Protect Dosim 18: 31-8

Sears DWG (1988) Thermoluminescence from meteorites: shedding light on the cosmos. Nuel Tracks and Radiat Meas 14: 5-17

Sears DWG, DeHart JM, Hasad FA, Lofgren GE (1989) Induced thermoluminescence and cathodoluminescence studies of meteorites. In: Coyne LM, McKeever SWS, Blake DF (eds) Spectroscopic characterization of minerals and their surfaces. American Chemical Society, ACS Symposium Series 145: 190-222

Sunta CM, Kathuria SP, Nambi KSV (1970) Proc Natl Symp on Radiation Physics, BARC, Bombay, 299

Yang XH, McKeever SWS (1990) The pre-dose effect in crystalline quartz. J Phys D Appl Phys 23: 237-44

Vlasov VK, Kulikov OA (1989) Radiothermoiluminescence dating and applications to Pleistone sediments. Phys Chern Minerals 16: 551-8

Zimmerman DW (1972) Relative thermoluminescence effects of alpha and beta radiation. Radiat Effects 14: 81-92


3.5.3 Exoelectron Spectroscopy of Minerals

V.S. KORTOV

Theoretical Aspects of Exoelectron Spectroscopy

A brief account of the fundamentals of EES seems appropriate. EES, as a physical effect, is a nonstationary phenomenon in the surface which is in a nonequilibrated state. This state arises from such external effects as mechanical strain, corpuscular and electromagnetic irradiation, and sharp drops in temperature. Electron emission from such surfaces is excited, e.g., by heating or illumination.

The diversity of factors affecting the state of solid surface layers lead to a variety of models for describing EES. One commonly used is based on the band theory of solids. According to this concept, the various defects in the surface layer are most important. Growth defects, as well as those due to impurities or external agencies, are believed to form electron traps which are filled during preliminary crystal excitation. In order to leave the trap, the electron will need an activation energy which is determined by its binding energy in the trap (the energy level depth e) and the energy necessary to overcome the crystal electron affinity X (Fig. 61).

Thermally activated exoelectron emission is widely used for surface layer trap spectroscopy of dielectrics. In this method, samples are generally heated at a constant rate. In this case, emission results from thermo-ionization of traps at depths up to 100 nm. Delocalized electrons reach the conduction band. From there part of them overcome the potential barrier X and escape into the vacuum, whereas the rest can recombine with hole traps, which causes thermally activated luminescence. Electron emission is registered as characteristic curves

Fig. 61. Band cherne for a ubsurface layer of dielectric

3.5.3 Exoelectron Spectroscopy of Minerals 167

with maxima whose number and location are determined by the electron trap spectrum of the layers. For bulk traps of one type, the count rate of electrons I(T) and the form of an isolated maximum are described by the equation

I ( I: + x) I(T) = PoC exp - kT . (1)

Here Po is the frequency factor, C is the trap concentration, and I is the kinetic order.

In a number of cases the maxima for which bulk traps are responsible are registered synchronously. Peaks from adsorption traps do not have luminescent analogs. These peaks can also be described by Eq. (1), when X = O.

One of the main problems of EES is to determine I: and C(T). In this thermally activated process the activation energy E and Po can be calculated. As follows from the band model, this analysis enables us to obtain E in the interval between I: and I: + X. However, the difference between I: and E, in many cases, is not significant, for two reasons: firstly, wide-band dielectrics, as a rule, are characterized by a small value of X (I: ~ X); secondly, on excitation, the charge state of the crystal surface changes, making it easier for electrons to overcome the surface potential barrier.

Among the methods of excitation of dielectric crystals, electron bombardment is widely used because of the possibility for using electron beams with various densities and considerably different electron energies. Upon electron bombardment with energies of 1 to 1.5 keY with coefficients of secondary emission (1 > 1, in most cases a positive charge is formed on and near the sample surface. Electrons penetrating to a distance determined by their energy are captured on traps, and create a negative bulk charge, whereupon, electrical fields with a strength of 105 _106 V (cm are formed in dielectrics. Calculations show that electrons delocalized from bulk traps during thermostimulation are accelerated in the bulk charge electrical field up to the energies of several electronvolts, which significantly exceeds the X value. Moreover, these fields cause a benddown el/lo of energy bands and decrease of X, all of which promote emission of the electron into vacuum.

In EES of bulk traps, the excitation of a sample by X-rays is widely used. In this case the crystal surface, as a rule, is positively charged due to photoelectron loss in the course of irradiation. Optical absorption and ESR are used to identify the traps responsible for emission.

The trap spectrum created by adsorption centers as well as by surface structural defects, electrons with energies less than that of the crystal forbidden gap suffer excitation. In this case (1 < 1, the surface is negatively charged, and there is no potential barrier for the emission of delocalized electrons, therefore E = 1:.

Thus, over a wide temperature range when the surface has a charge and an electric field stimulating the exit of electrons into vacuum, one can find the spectrum of bulk and surface traps after calculating E. In particular, if we transform Eq. (1) by considering the heating of a sample at a constant rate J1. and


take logarithms, we obtain a relation to determine E and Po as:

dC

In~= -~-lnPo C1 kT 11 •

(2)

The concentration of active centers and its rate of change can be determined from a TSEE curve by the formulae:

C(T) =~. JI(T)dT IlY T

and

dC 1 ----·I(T)

dT IlY ,

where Y is the proportion of electrons registered in the experiment.

(3)

(4)

To determine the parameter 1, one can use a method described elsewhere. The value F(l) is calculated as the ratio of areas under the experimental curve I(T):

00

f I(T)dT F(l) = .:..:;T;"---__

fI(T)dT o

(5)

For the first order, F(l) = 0.38, while for the second F(l) = 0.52. When the parameter 1 is known, the values E and Po are calculated from Eq. (2). This method is very convenient for the calculation of TSEE curves by the minicomputers of exoelectron spectrometers.

At elevated temperatures (T ~ 600 K), relaxation of a bulk charge in the subsurface layer takes place, and to determine the energy depth of a trap found by the described method, the value E must be corrected for x. If X is unknown, the e for deep traps is determined by Seidle's formula (1962), taking into account the temperature maximum shift of thermoactivation process at various rates of sample heating. Note as well, that the determination of E from the TSEE peak analysis and e from the experiments with various heating rates permits X to be found as the difference of these values.

In conclusion, we should like to draw attention to the main analytical possibilities of EES:

1. Variations of electron energies and of vacuum for TSEE measurements enable us to identify the traps created by structural defects and adsorbates.

2. Comparison of the TSEE curves before and after irradiation with different doses makes possible the investigation of the dynamics of trap accumulation on the surface and in the subsurface layer of crystals.

3.5.3 Exoelectron Spectroscopy of Minerals 169

3. The variation in the electron excitation energy and the creation of emissionactive zones at various depths from the sample surface allows us to determine the spatial distribution of traps in the subsurface layer.

4. Thermoactivation analysis of TSEE curves permits us to determine the energy depth (1)) of traps and evaluate the potential barrier (X) on crystal surfaces.

Samples and Experimental Methods

Mono- and polycrystalline samples of phenakite, willemite Zn2Si04, beryllium orthogermanate Be2Ge04, and solid solutions of them, as well as crystals of synthetic IX-quartz have been used. The samples of phenakite and quartz were single crystal plates with sizes of 0.5 x 10 x 10 mm, cut from optically transparent natural and artificial single crystals.

According to spectral analysis AI, Mn, Ti, Ge, and Fe impurities (0.02-0.2 wt.%) were present in natural single crystals of phenakite. The artificial samples were tablets 10 mm in diameter, 1.5 mm thick obtained by high temperature solid phase synthesis from a mixture of the initial oxides using mineralizers.

TSEE changes were recorded with a linear heating rate of 20 K/min in a vacuum of 10-4-10- 7 Pa using a secondary electron multiplier as electron detector. The design of the operating chamber enabled the TSEE and TSL curves to be registered simultaneously. Prior to the measurements the samples were excited by X-rays (Fe anode, 25 kV, 10 rnA) or by electron bombardment with energies from 6 eV up to 4 keY. To obtain detailed information about the nature of emission-active centers, ESR and optical spectroscopy were used.

Since with electron excitation it is possible to trap not only primary but also secondary electrons returning to the sample with much less energy, collection of reflected current is a necessary condition to control the filling of electron traps at various depths from the crystal surface. In view of this, in our experiments with electron bombardment, the grid with positive potential was placed above the sample.

To investigate radiation surface damages quartz and phenakite samples irradiation by fast neutrons in the operating channel of a nuclear reactor have been used.

The possibilities of exoelectron spectroscopy to investigate defects in dielectrics are demonstrated for phenakite Be2Si04, its structural analogs Zn2Si04, Be2Ge04, solid solutions Be2Sil-xGex04 (x = 0 --;- 1), and IX-quartz.

The above assumption is confirmed by the results of synchronous TSEE and TSL measurements of phenakite crystals after X-ray irradiation exciting mainly bulk traps (Fig. 62). Samples were irradiated in air at 300 K. Upon subsequent heating, TSEE peaks are registered at the same temperatures as after electron bombardment. The most intense TSEE peaks at 330 and 480 K are accompanied by TSL. The coincidence of these TSEE and TSL peaks testifies that bulk

170

-; til

a. ~

-4

~ ...,


2

30 j ci

20

10

O~,---------.-------.--------.--______ -L 0 300 400 500

T.K

600

Fig. 62. TSEE (1) and TSL (2) of phenakite after X-ray excitation

electron traps of one ongm take part in the relaxation process of crystal excitation. At T > 550 K there is no such coincidence, as difficulties in TSL registration arise due to the thermal background of the crystalholder and temperature luminescence extinction. The existence of a TSL peak at 380 K which has no TSEE analog must be pointed out. It means that TSL at the given temperature arises without participation of electron traps.

Traps in the Surface Layers of Simple and Complex Oxides. To further investigate the origin of electron traps in phenakite, exoemission properties of its simple oxides were studied. The relevance of such an investigation is due to the fact that the tetrahedra [BeO 4]6 - and [SiO 4]4 -, typical of simple BeO and Si02 oxides, are also the foundation of the crystal lattice of phenakite and of other silicate minerals with beryllium content. Besides, the results of optical absorption and ESR measurements show that radiation centers in quartz and phenakite crystals are of similar origin. This allows use of data obtained for analogous traps in quartz when discussing the origin of traps in phenakite.

The results of TSEE measurements of quartz monocrystals and ceramic Be02 are given in Fig. 63a. The coincidence of TSEE maxima at 670 K attracts

2 Ceramic BeO was used in connection with monocrystal pyroelectrical properties and with the effect of spontaneous exoelectron emission without preliminary excitation which makes the strict comparison of BeO, Si02 and Be2Si04 TSEE difficult.


10

5

3

"'; III

0. 2 S

... 52

0

4

3

2

0 300 400

_I

500

T, K

\ \ I I I

3 o

'Ill

2 0.

S M

52 1 ..,

\ /'\ " // \ " '" \ ;,.._..... \ , 0

b

c

600 700

171

Fig. 63a-c. Exoemission spectra of initial simple oxides Si02 (1a), BeO (2a), and solid solutions Be2Sil-xGexOy (b), Zn2_xBexSiOy (c): 1 - x = 0; 2 - x = 0.1; 3 - x = 0.3; 4 - x = 0.5; 5-x=1

attention when BeO and phenakite TSEE curves are compared. In phenakite, as will be shown later, the above maximum is due to the existence of oxygen vacancies, whereas in BeO the anion sublattice is characterized by high stability and the formation of traps with oxygen vacancies is unlikely. Frenkel defects of beryllium atoms in interstitial and empty cation sites are the most characteristic of the BeO lattice. In particular, the TSEE maximum at 510-550 Kin BeO is usually explained by electron trap destruction on interstitial beryllium atoms. At the same time, in complex silicate structures, intrinsic defects, as a rule, are due to the distortion of the silicon-oxygen crystal sublattice. In view of all this, it should be assumed that the coincidence of TSEE maxima in BeO and phenakite at 670 K is accidental and exoemission in both compounds is due to the destruction of traps of different origin but with similar values of energy depth.


The interconnection of emission centers in quartz and phenakite crystals is more definite. When TSEE curves of the above crystals are compared, it is necessary to point out the proximity of temperature location of almost all exoemission maxima in the temperature range 300-570 K. On this basis one can postulate that the similarity of traps is responsible for TSEE processes in the crystal investigated. Existence of an electron trap in quartz with TSEE maximum at 325 K which is also characteristic of phenakite (330 K) and ascribed to the impurity [Ge04J - 5 centers should be emphasized.

To elucidate the role of anion and cation sublattice defects in the formation of traps in the crystal surface layer, the TSEE of the phenakite structural analogs BezGe04, ZnzSi04 and solid solutions based on them has been studied (Fig. 63b,c).

The TSEE curves of solid solutions BezSil-xGex04 (x = 0-1) qualitatively coinciding with those for phenakite show identical origin of emission traps in this isomorphous series. Since the Ge atom in BezGe04 is an element of the regular lattice and does not create a trap. there is no TSEE peak characteristic of phenakite at 330 K caused by the destruction of impurity germanium centers. Smooth decrease of the TSEE intensity upon gradual substitution of Si4+ by Ge4+ shows that the emission activity of solid solutions is mainly determined by the traps created in silicon-oxygen structure constituent.

When the composition of the solid solutions Znz- xBexSi04 (x = 0-0.3) is altered, the TSEE intensity remains practically constant and this essentially distinguishes them from the solutions with isomorphous substitution in the silicon-oxygen sublattice. Besides, considerable distortions of the crystal lattice due to the great difference between Znz + and Bez + ionic radii are probably the reason for active center energy depth variations which can be seen in the TSEE maxima shift in comparison with that of phenakite.

Kinetic Analysis of TSEE and TSL Processes and Determination of Trap Energy Depth in Phenakite. Certain results of calculation for the kinetic order and trap parameters in phenakite have been obtained by the method described previously (Table 8).

The difference in the kinetic order and temperature position of TSEE and TSL maxima at 330-335 K attracts attention. The second order ofTSL kinetics is most probable with the luminescence recombination mechanism on the assumption that conduction electrons participate in TSL. This fact once more confirms that the impurity [Ge04J5 - center is an electronic one. Its thermoionization causes the ejection of electrons into the conduction band and synchronous TSEE and TSL peaks.

From TSEE measurements at 590 K. when the action of a surface charge is weakened, it was found that the value of X in phenakite crystals was 0.2 eV.

Based on these calculations the energy levels for traps in the subsurface region and on the phenakite surface have been drawn up (Fig. 64). The depths of trap determining the appearance of high temperature TSEE peaks, are corrected


Table 8. TSEE- TSL kinetic parameters and energy depth of traps in phenakite

Tm, [K]

TSEE

150 180 200 330

435 480 590 670

TSL TSEE TSL

335 2 380 1 430 1 475 1

2 1

0.2 e V 0.4 eV ' 0.5 eV

0.geV [GeO,1 5-

E [eV]

TSEE

0.2 0.4 0.5 0.9

1.0 1.6 2.3 2.8

1.6 eV

2.3 eV Lum,nescence

TSL

0.8 0.7

1.5

Po [S-I]

TSEE

2 x 108

2 x 1011 1 x 1011 4 x 1012

1 X 1010 6 x 1014

6 x 1013

leV ° 2ads

Surface

TSL

2 x 107

4 X 1013

173

Fig. 64. Energy level scheme of active centers in phenakite crystals. (After Kortov et al. 1985)

by X. Those traps whose origin has been decoded with the help of ESR and optical spectroscopy are also marked in the scheme.

The TSEE maximum at 435 K, as has been stated, is connected with the presence of adsorbates on the phenakite surface. Emission, in this case, can occur just from the levels of surface defects formed by ion-radicals, for instance OR -,0;,0 -, etc. It has been shown that oxygen adsorption causes an increase of the emission peak intensity in the range 425- 445 K characteristic of many oxides. ESR has also shown that upon heating of some oxides in the temperature range of 440 K, ESR signals change their shape, which is explained by valence transformations in the adsorbed oxygen layer. Thus, our experimental results


and earlier published data permit the conclusion that the TSEE maximum at 435 K in phenakite is due to recombination processes in the layer of adsorbed oxygen covering the sample surface.

Since, as is shown in the present chapter, most emission active centers in phenakite and quartz are of identical origin, electron trap energy parameters in the surface layers of the above crystals have similar values.

References

Braunlich P (1968) Thermoluminescence and thermally stimulated current tools for the determination of trapping parameters. In: McDougall OJ (ed) Thermoluminescence of geological materials. Academic Press, London

Holzapfel G (1976) The evolution of volume concepts to describe exoelectron emission. In: Proc 5th Int Symp Exoelectron Emission and Dosimetry, Svikov, pp 19-34

Kortov VS, Zatsepin AF, Ushkova VI (1985) Exoelectron spectroscopy of traps in surface layers of phenakite and quartz. Phys Chern Mineral 12: 114-121

Marfunin AS (1979) Spectroscopy, luminescence and radiation centers in minerals. Springer, Berlin Heidelberg New York

3.6 Infrared Spectroscopy

3.6.1 Band Assignments in Infrared and Raman Spectroscopy

A.N. LAZAREV, P.F. McMILLAN, and S.W. KIEFFER

An experimental vibrational spectrum (usually infrared or Raman) consists of a set of discrete peaks or overlapping bands as a function of frequency. In most applications of vibrational spectroscopy, it is necessary to assign some or all of these to particular types of atomic vibration. This identification of experimental infrared or Raman bands with individual vibrational modes is referred to as a spectral assignment. This approach is used extensively in structural studies of minerals in which the vibrational spectra are interpreted to give information on structural properties such as bond lengths and angles, or coordination polyhedra. With suitable calibration, this information can be made quantitative. Band assignments are also often useful in modeling the thermodynamic properties of minerals, for enumerating the contributions of particular vibrational degrees of freedom to frequency regions in the phonon density of states, g(w), and for comparing the thermochemical and elastic properties of minerals within and between structural classes.

The most rigorous and complete way of assigning experimental spectra is to calculate the normal vibrational modes from first principles. Although reliable ab initio methods for doing this for complex minerals are now becoming

3.6.1 Band Assignments in Infrared and Raman Spectroscopy 175

available, they have not yet been applied to a wide range of structures or compositions. Most normal mode calculations carried out for minerals to date have been empirical or semi-empirical, with force field parameters or the crystal potential function adjusted or refined by comparison with experimental data. This process can yield much useful information, especially when data for crystals with isotopically substituted atoms are available to further constrain the calculation. However, due to the intrinsically underdetermined nature of the vibrational problem, even the best refined empirical calculations do not necessarily give a true description of the normal modes of the real crystal, and the resulting spectral interpretations can be ambiguous.

In many cases, it is also possible to use the crystal symmetry to assign the experimental vibrational spectra. For example, if only one vibrational mode appears within a given symmetry species, then the atomic displacements of that mode are completely constrained by symmetry and the spectral assignment is unique. This is the case for the infrared absorption band of periclase (MgO), which must involve motions of the Mg2 + sublattice against the anion sublattice, along the Cartesian directions, and for the single Raman active mode of diamond, which can be described as a C-C stretching vibration, with one carbon sublattice moving against the other. In molecular vibrational spectroscopy, the symmetry elements of the molecule are determined by inspection. The symmetry species of the molecular vibrations are then deduced using the algebraic methods of group theory. For crystalline materials, the vibrational displacements of each atom about its equilibrium position are examined with respect to the atom site symmetry within the unit cell, to obtain the symmetry species for the crystal lattice vibrations. If molecular species can be identified within the crystal, the symmetry species for the vibrations of each molecular group can be correlated with those for the free molecule, to analyze the effect of the crystalline matrix on the vibrations of the molecular group. For example, a free tetrahedral Si04 group has four vibrational symmetry species, including singly and triply degenerate (Ai and T 2) stretching modes, and doubly and triply degenerate (E and T 2) bending vibrations. On inclusion in an olivine lattice, the vibrational modes of the four Si04 groups within the unit cell become coupled, and the symmetry is reduced to Cs (m), removing the degeneracies and resulting in vibrational band splittings. The magnitudes ofthese splittings can be correlated with the degree of distortion of the Si04 groups in the lattice.

Group theoretical methods are well established for enumerating the number of each type of symmetry species for all of the vibrational modes of a given crystal lattice. These can then be used to predict the number of infrared and Raman active (and spectroscopically inactive) modes for the crystal. If additional modes are observed in the spectra, this can indicate the presence of impurities, or local deviations of the mineral from ideal symmetry (orderdisorder, superstructure formation). In addition, symmetry analysis combined with polarized infrared or Raman measurements can be used to determine the orientation of structural units within the lattice. For example, polarized infrared spectroscopy of minerals containing hydroxyl groups in the O-H stretching


region gives direct information on the orientation of the O-H group, which can be difficult to elucidate from diffraction experiments. Likewise, infrared and Raman spectroscopy of carbonate minerals can be used to determine the orientation and degree of distortion of the carbonate groups within the lattice.

Much useful information can be gained from approximate, qualitative interpretations of the infrared and Raman spectra of minerals. This is especially true when localized vibrational modes of molecular units or particular structural groups can be identified within the crystal spectra. For example, the O-H stretching vibrations of hydroxyl-containing minerals are usually only weakly coupled to other modes of the lattice, and the O-H stretching region can be assigned as if isolated -OH units were embedded in the crystalline matrix. This has been used extensively to study the local OH environment in hydroxyminerals, including clays, amphiboles, topaz, and micas, and nominally anhydrous minerals containing OH groups in defect sites, such as olivine, pyroxene, garnet, feldspar, and quartz.

In a similar way, the deformation mode of molecular H20, and the stretching and bending vibrations of carbonate anions are usually uncoupled from the mineral lattice vibrations, and can be easily assigned in the spectra and interpreted in terms of the geometry and local environment of these molecular groups. Although the stretching and bending vibrations of tetrahedral and octahedral Si04 or Si06 groups tend to be more strongly coupled to other lattice vibrations, it has proved useful to analyze the spectra of silicate minerals using the assumption of pure Si-O stretching or OSiO bending vibrations (and SiOSi bending in polymerized silicates) as a first approximation. In general, bands in the 1710-1200 cm -1 region are assigned to Si-O stretching motions (Fig. 65), and bands in the 300-700 cm - 1 region as OSiO or SiOSi bending vibrations (Fig. 66), and any coupling with other lattice vibrational modes in treated as a perturbation on this assignment. This type of approach has been extended to other "covalent" structural units within mineral structures, such as aluminate, titanate, borate and phosphate polyhedral units, and characteristic frequency ranges for stretching vibrations of particular MOn polyhedra have been suggested. This has proved useful in studies of cation coordinations in minerals and glasses, although the method must be applied with some care, because vibrational coupling between stretching vibrations can affect the interpretation, and considerable overlap in frequency ranges for different polyhedra can occur. In polymerized minerals and glasses containing SiOSi linkages, bands in the 400-700 cm - 1 region have been assigned to bending vibrations of the SiOSi group, and their frequency has been shown to vary inversely with the SiOSi angle. This has proved useful in studies of the densification mechanisms of silicate minerals and glasses.

Vibrational spectroscopy has been used extensively in structural studies of silicate melts and glasses, where the lack of long range order limits the information obtained from diffraction studies. In these vibrational studies, the glass spectra are often interpreted by comparison with spectra of corresponding crystalline phases, as well as assignments based on band symmetry, and results

3.6.1 Band Assignments in Infrared and Raman Spectroscopy 177

I I I T

Misc.

Stsh J 0.22

0.037 0.074 0.037 0.07~ ~ I,. .

0051 Qrtz I J -p;; 2;,154 Anar : Albt '"

2.0

..... 2.5

a t 0.159 Talc, Muse

2.75 t 0.089 ( 0.089 Trem

3.0 J 0.100 ~ 0.100 Enst. lade, Diop

0.05~ Ky tAnd ---- Silt ~ ~0.15

AnlGrls~~Py j I :~ Gam I :~ AI I i~ 4.0

0.055 t 0.0551 10.110 Zire

I 10.048 I ~0.142 I I Fors

700 800 1100 1200

Fig. 65. Summary of frequencies (in wavenumber units, cm -1) at which antisymmetric Si-O-Si, Al stretching bands occur in silicates of different polymerization. The degree of polymerization is expressed by the 0/Si ratio on the ordinate. The decimal fractions noted by each band represent the fraction of the total degrees of freedom of the unit cell for that mode. In this and in Figs. 66 and 67, the mineral abbrevations are: Stsh stishovite; Anor anorthite; Albt albite; Qrtz quartz; Talc talc; Muse muscovite; Trem tremolite; Enst enstatite; Jade jadeite; Diop diopside; Ky kyanite; And andalusite; Sill sillimanite; Sp spessartine; Py pyrope; An andradite; Gr grossular; Al almandine; Garn garnets; Zire zircon; Fors forsterite. The variolls symbols with the lines are simply to distinguish lines appropriate to different minerals. Details relating to the data shown are in Kieffer (1985)

of vibrational calculations and isotopic substitution experiments. More recently, these vibrational studies have been supplemented with data from nuclear magnetic resonance spectroscopy, which gives detailed information on the local environments of silicon, aluminium, and other nuclei within the glass structure.

Another application of vibrational mode assignments for minerals is in heat capacity modeling from model vibrational density of states functions. In this

178

2.0

2.5

.... a 2.75

3.0

4.0


I Qnz,O.037

W4Anor 0.102 wffAiJ Alb! 0.102

F10 Qnz 0.185 ?0a e; Anar O. 102 :a

Alb! 0.102

WDiop 0.067 m 5! O. 7

800

q-20%

q-10%

5

q< 10%

q-25%

900

Fig. 66. Summary of the frequencies at which symmetric Si-O-Si, Al stretching and bending modes occur in silicates of different polymerization, as expressed by the OjSi ratio. These frequences span ranges indicated by the boxes; geometric patterns distinguish different OjSi ratios. The decimalfractions in the boxes represent the fraction of degrees offreedom of the unit cell for those modes. Abbreviations are given in the legend to Fig. 198. Details relating to the data shown are in Kieffer (1985)

type of calculation, the g(w) is independent of the detailed vibrational mode assignment, but such assignments can help in enumerating the number of vibrational modes present in particular frequency ranges within the density of states (Figs. 65-67). Some thermodynamic properties, such as entropy, are particularly sensitive to the distribution of modes at low frequencies, and therefore, band assignments in this region are of great interest. These bands are often difficult to assign, and there may be significant overlap between acoustic and optic modes (Fig. 67). Although some generalities appear to exist (e.g., within a given structural class such as feldspars, there may be a nearly linear inverse correlation between cation mass and lowest optical mode frequency), there are enough exceptions to any rules yet formulated that great caution must be exercised Rigorous calculation of normal vibrational modes from first principles can be especially helpful in understanding this region of the vibrational spectrum, because simple models typically are not adequate.

3.6.1 Band Assignments in Infrared and Raman Spectroscopy

2.5

1Q. 2.75 o

o

I

<71

100

I I I

200

I

I

500

I

I 600

179

700

Fig. 67. Summary of frequencies of lattice and acoustic modes for typical minerals with different O jSi ratio. The inverted triangles represent estimates of the zone-boundary (K = 0) acoustic mode frequencies of the two shear and longitudinal modes, calculated with an assumed dispersion relation as described in Kieffer (1985). The patterned bands represent the range of frequenceies of the distortion modes observed in infrared and Raman spectra, i.e., at K = O. The dashed extension on the low-frequency end of each band is an estimate of the extension of the spectrum due to dispersion acorss the Brillouin zone. The decimal fraction gives the fraction of degrees of freedom associated with these modes, obtained by subtracting all other enumerable modes from 1.00

References

Decius JC, Hexter RM (1977) Molecular vibrations in crystals. McGraw-Hili, New York Farmer VC (ed) (1974) The infrared spectra of minerals. Mineralogical Society, London Ghose S (1988) Inelastic neutron scattering. In: Hawthorne FC (ed) Spectroscopic methods in

mineralogy and geology. Reviews in mineralogy vol 18, Mineralogical Society of America pp 162- 192

Kieffer SW (1985) Heat Capacity and entropy: systematic relations to lattice vibrations. In: Kieffer SW, Navrotsky A (eds) Microscopic to macroscopic: atomic environments to mineral thermodynamics. Reviews in mineralogy vol 14, Mineralogical Society of America, pp 65- 126

Lazarev AN (1972) Vibrational spectra and structure of silciates. Consultants Bureau, New York

McMillan PF (1985) Vibrational spectroscopy in the mineral sciences. In: Kieffer SW, Navrotsky A (eds) Microscopic to macroscopic: atomic environments to mineral thermodynamics. Reviews in mineralogy vol 14, Mineralogical Society of America, pp 9- 63

McMillan PF, Hofmeister AM (1988) Infrared and Raman spectroscopy. In: Hawthorne FC (ed) Spectroscopic methods in mineralogy and geology. Reviews in mineralogy vol 18, Mineralogical Society of America, pp 99- 159


Rossman GR (1988) Vibrational spectroscopy of hydrous compoentns. In: Hawthorne FC (ed) Spectroscopic methods in mineralogy and geology. Reviews in mineralogy vol 18, Mineralogical Society of America, pp 193-206

White WB (1975) Structural interpretation of lunar and terrestrial minerals by Raman spectroscopy. In: Karr C (ed) Infrared and Raman spectroscopy of lunar and terrestrial materials. Academic Press, pp 325-358

3.6.2 Polarized Infrared Spectra

A. BERAN

Powder infrared spectroscopy is the most common method of obtaining infrared spectra of minerals. However, important experiments in mineralogy are carried out on single crystals, and it is often useful to obtain polarized IR spectra in order to study orientation effects of vibrating groups.

The first mineralogical experiments with "glass plate"-polarized near-IR radiation date back to Merritt (1895). Today's most commonly used polarizing elements are gold wire grids on a substrate that is transparent over the range of interest. AgBr covers ca. 5000 to 300 cm - 1, polyethylene covers ca. 500 to 30 cm -1; near-IR measurements can be optimized by using LiJ0 3 substrate. IR beam condensors and microscopes as parts of the instrument expand the range of possible mineralogical applications and make it possible to obtain polarized IR spectra of small sample regions, down to measuring areas of 10/lm in diameter.

One of the criteria that must be met for the absorption of radiation is that the vector direction of dipole moment change and the direction of the electric vector of the incident polarized radiation be the same. Depending on the symmetry of the mineral, several crystallographically oriented crystal plates must be prepared, and a polarized IR spectrum can be recorded with the beam perpendicular to the double polished plate. Knowing the orientation of the electric vector of the polarized radiation relative to a crystallographically defined direction (crystal axis) of the crystal plate, the orientation of the dipole moment change of a distinct vibrating group can be determined by rotating the crystal plate. If a specific absorption band is found to be much more intense in a given position of the crystal plate than the one obtained after rotating the sample 90 degrees, then for the vibrating groups giving rise to the intense absorption, the vector direction of dipole moment change and the direction of the electric vector of the polarized radiation are largely parallel. The "dichroic ratio" of the absorption intensities in the directions parallel and perpendicular to a crystallographically defined direction is represented by the extinction coefficients ell and d. for IR radiation whose electric vector vibrates parallel and perpendicular, respectively, to the defined direction of the crystal.

Several PQli!rized spectra of lunar and terrestrial minerals in the near-IR region are presented in a review article by Burns and Vaughan (1975).

3.6.2 Polarized Infrared Spectra 181

Polarized IR spectra are particularly useful in the determination of "water" in minerals. At one extreme, these measurements can be used to deduce possible orientations of structurally incorporated OH groups and H20 molecules; at the other extreme, lack of polarization effects allows to distinguish fluid inclusionwater and included nonoriented hydrous impurities. Regarding the pleochroic behavior of OH absorption bands, calculations on the basis of the "electric dipole transition model" proposed by Strens et al. (1982) reveal excellent agreement between calculated and observed absorptions in most cases. An expression defining the variation of the intensity of OH absorption bands with the orientation of the electric vector with respect to the crystal axis is given by Rouxhet (1970). On the basis of polarization measurements, Tsuboi (1950) first reported data on the position of hydrogen in muscovite; Zemann and coworkers determined the IR pleochroism of the OH stretching frequency for several minerals (ca. 3700-3000 cm -1), starting on the determination of the OH orientation in azurite (Tillmanns and Zemann 1965). For the extensive literature on this topic the reader is referred to the review articles of McMillan and Hofmeister (1988) and Rossman (1988).

IR spectroscopy provides an extremely sensitive method for detecting trace hydrogen bonded to oxygen in the structure of various minerals. Some few examples are given here to illustrate the conclusion that can be drawn from polarized IR studies of nominally anhydrous minerals.

On the basis of the pleochroic scheme of the OH fundamentals in forsterite, it is proposed that OO(OHh and OOz(OH}z tetrahedra occur as structural elements, assuming that vacancies are on Si sites (Beran and Putnis 1983). A similar model is proposed for phenakite, with 002(OH}z and 003(OH) tetrahedra as structural elements, assuming Be vacancies (Beran 1990). From these data, it is evident that some similarities in the O-OH replacement mode exist between OH-bearing olivines, phenakites and "hydrogarnets". OH groups are also incorporated in pyroxenes. In diopside OH dipoles are partially replacing the 0(2) (zigzag) oxygens pointing to the 0(3) positions (Beran 1976; Skogby et al. 1990).

Two structurally similar minerals, danubrite and An 59 labradorite, show sharp and broad absorptions, respectively, in the region of the OR stretching fundamentais. The OH dipoles in danburite show a distinct orientation parallel to the b-axis; the OH groups in labradorite are oriented approximately perpendicular to (001) (Beran 1987). The proposed models are in accordance with bond valence calculations, showing that in both framework structures the most deficient oxygens, 05 in danburite and Oem in labradorite, are partially replaced by OH. As is a general rule, the O-H direction is approximately perpendicular to the three surrounding cations, where the donator oxygen is the top of a flat trigonal pyramid. Polarized IR studies of kyanite, sillimanite, and andalusite show these minerals to contain structurally bound OH groups. If OH is incorporated during crystallization, then Al2SiOs phase relations could be significantly affected by fractionation of hydrous components among the mineral phases (Beran et al. 1989).


Polarized IR spectra of Verneuil-grown colourless corundum, rose to deep red ruby, blue, "amethyst"- and "alexandrite"-colored sapphire reveal extremely pleochroic and narrow OH absorption bands. Assuming vacancies in the cation lattice, OH is coordinated to two AI-atoms forming "groups" of face-sharing AI2(OH)Og-double octrahedra (Beran 1991). The recognition ofa structural OH content in rutile was initially based on the observations of pleochroic bands in the IR spectrum of Verneuil-grown crystals (Soffer 1961). Beran and Zemann (1971) have established the presence of OH in rutile from natural occurrences and proposed a model for the structural OH incorporation. Regarding the pleochroic scheme of the bands, maximum absorption occurs when the electric vector of the polarized IR radiation vibrates perpendicular to the c-axis, thus, indicating OH dipoles on the O-site oriented perpendicular to the plane of the three coordinating Ti-atoms (Hammer and Beran 1991).

H20 molecules reside in the structural channels of beryl and cordierite at two different sites that can be distinguished on the basis of the polarization properties (see Rossman 1988). In nepheline, H20 molecules exist in multiple crystallographic orientations, presumably in the K-site. Two orientations of water are indicated at room temperature; when the crystal is heated to 350°C, a third orientation develops. All three types of H20 have the H-H axis perpendicular to the c-axis; the two-fold axis of the different H20 types is tilted at different angles to the c-axis (Beran and Rossman 1989).

References

Beran A (1976) Messung des Ultrarot-Pleochroismus von Mineralen. XIV. Der Pleochroismus der OH-Streckfrequenz in Diopsid. Tschermaks Min Petrogr Mitt 23: 79-85

Beran A (1987) OH groups in nominally anhydrous framework structures: an infrared spectroscopic investigation of danburite and labradorite. Phys Chern Mineral 14: 441-445

Beran A (1990) The occurrence of OH absorptions in phenakite - an IR spectroscopic study. Mineral Petrol 41: 73-79

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, Putnis A (1983) A model of the OH positions in olivine, derived from infrared-spectroscopic investigations. Phys Chern Mineral 9: 57-60

Beran A, Rossman GR (1989) The water content of nepheline. Mineral Petrol 40: 235-240 Beran A, Zemann J (1971) Messung des Ultrarot-Pleochroismus von Mineralen. XI. Der

Pleochroismus der OH-Streckfrequenz in Rutil, Anatas, Brookit und Cassiterit. Tschermaks Min Petrogr Mitt 15: 71-80

Beran A, Rossman GR, Grew ES (1989) The hydrous component of sillimanite. Am Mineral 74: 812-817

Burns RG, Vaughan DJ (1975) Polarized electronic spectra. In: Karr C (ed) Infrared and Raman spectroscopy oflunar and terrestrial minerals. Academic Press, New York, pp 39-72

Hammer VMF, Beran A (1991) Variations in the OH concentration of rutiles from different geological environements. Mineral Petrol 45: 1-9

McMillan PF, Hofmeister AM (1988) Infrared and Raman spectroscopy. In: Hawthorne FC (ed) Spectroscopic methods in mineralogy and geology. Rev Mineral, vol 18. Mineral Soc Am, pp 99-159

Merritt E (1985) Dber den Dichroismus von Kalkspat, Quarz und Turmalin fUr ultrarote Strahlen. Ann Phys Chern 55: 49-64

3.6.3 Applications of Infrared Spectroscopy 183

Rossman GR (1988) Vibrational spectroscopy of hydrous components. In: Hawthorne FC (ed) Spectroscopic methods in mineralogy and geology. Rev Mineral, vol 18. Mineral Soc Am, pp 193-206

Rouxhet PG (1970) Hydroxyl stretching bands in micas: a quantitative interpretation. Clay Minerals 8: 375-388

Skogby H, Bell DR, Rossman GR (1990) Hydroxide in pyroxene: variations in the natural environment. Am Mineral 75: 764-774

Soffer BH (1961) Studies of the optical and infrared absorption spectra of rutile single crystals. J Chern Phys 35: 940-945

Strens RGJ, Mao HK, Bell PM (1982) Quantitative spectra and optics of some meteoritic and terrestrial titanian clinopyroxenes. In: Saxena SK (ed) Advances in physical geochemistry, vol 2. Springer, Berlin Heidelberg New York, pp 327-346

Tillmanns E, Zemann J (1965) Messung des Ultrarot-Pleochroismus von Mineralen. I. Der Pleochroismus der OH-Streckfrequenz in Azurit. Neues Jahrb Mineral Mh 1965: 228-231

Tsuboi M (1950) On the positions of the hydrogen atoms in the cyrstal structure of muscovite, as revealed by the infra-red absorption study. Bull Chern Soc Jpn 23: 83-88

3.6.3 Applications of Infrared Spectroscopy to Structure and Bonding in Minerals and Glasses and to Speciation of Hydrous Components

W.B. WHITE and A.M. HOFMEISTER

The 3n - 6 vibrational degrees of freedom of an isolated molecule or the 3n - 3 vibrational degrees of freedom of a crystal can be classified according to the symmetry point group of the molecule or to the factor group of the crystal. Modes producing a change in the dipole moment of the atoms vibrating couple directly with incident electromagnetic radiation and thus are manifested as absorption bands in the the infrared. The wave functions for molecular vibrations are localized, the energy levels are sharp, so that transitions between them give narrow absorption bands with intrinsic widths on the order of fractions of a wavenumber. The wave functions for crystal vibrations are completely delocalized, leading to bands of phonon states spanning a discrete range of wavenumbers. However, the momentum associated with infrared photons is so small that IR absorption in perfect crystals is required by momentum conservation to take place only at zero wavevector, the center of the Brillouin zone (sometimes called the k = 0 selection rule). Thus, IR band widths for ordered crystal are also small, in the range of tenths to a few tens of wavenumbers.

Infrared spectra can be measured, using dispersive or Fourier transform spectrometers, by direct transmission through thin (0.5 to 5 jlm) layers, through pellets of powdered specimen embeded in a dielectric medium (typically KBr or petroleum jelly), by specular reflectance from polished surfaces, by attenuated total reflectance, and by diffuse reflectance. Spectra obtained by direct transmission and by deconvolution of specular reflectance spectra have proper line shapes. Spectra of polar materials (virtually all minerals) have distorted line shapes and shifted band positions when measured on powder dispersions.


Mineral Identification and Structural Characterization

Most minerals have a substantial number of atoms in their unit cells and thus a multi-band infrared spectrum. IR spectra, like Raman spectra and X-ray powder diffraction patterns, can be used as fingerprints for mineral identification. However, the difficulty in obtaining spectra with accurate line shapes and positions from powder or polycrystalline specimens has limited this application.

Structural entities that maintain their molecular character in the mineral, such as water, carbonate ions and SiO!- polyhedra, contribute bands to the IR spectra which strongly resemble the spectra of the free molecules or ions and thus can be identified. Molecular units with the same symmetry and similar atomic masses and bond strengths such as the triangular nitrate, carbonate, and borate ions have similar spectra and are not easily distinguished by means of IR spectra alone. In addition to these molecular or internal modes are external or "lattice" modes arising from vibrations between the molecular entities and other ions or groups that make up the crystal.

When more than one type of molecular entity exists in the crystal or when the molecular entity occupies a site with local symmetry less than that of the free molecule, lifting of degeneracy and relaxation of selection rules adds additional features to the IR spectrum. Careful analysis of these features, usually by group theoretical methods, provides information on the correlation field coupling the molecular units and the crystal filed of the local site.

Spectra of silicate minerals are dominated by bands of the Si04 tetrahedra, which shift in wavenumber depending on the degree of polymerization of the silicate structure. Si-O stretching bands occur in the range of 850-900 cm -1 for olivines or garnets with isolated Si04 tetrahedra and shift to higher wavenumber with increasing polymerization, reaching 1100-1150 cm -1 for the silica polymorphs.

Order/Disorder

When nonequivalent ions order on non-equivalent sites such as Mg/Fe ordering in pyroxenes or AI/Si ordering in feldspars, there are no changes in selection rules, and thus no new bands. However, bands in the ordered structure are usually narrower. When nonequivalent ions order on equivalent sites to form derivative structures such as ilmenite/corundum or chalcopyrite/sphalerite, the lowered symmetry relaxes selection rules and the ordered structure gives an IR spectrum with more bands than that of the parent mineral. Group theoretical method are used to predict the changes.

Polytypic stacking is ordering in one dimension. The unit cells of the polytypes are multiples of the parent cell, and thus the Brillouin zone is a fraction of the parent's Brillouin zone. New bands which appear in the spectra of polytypes can be assigned to zone boundary phonons of the parent structure made IR-active by the folding in of the parent Brillouin zone.

3.6.3 Applications of Infrared Spectroscopy 185

Disorder created by substitution of foreign ions or vacancies breaks the translation symmetry of the lattice and thus weakens of k = 0 selection rule. At low concentrations (a few atomic percent) of substitutional ions or vacancies, the effect is to broaden the IR and Raman lines. At higher concentrations, particularly of vacancies, bands become very broad (half widths in the range of 10-50 cm -1), selection rules break down, and ultimately the spectrum becomes a map of the density of states. Line broadening from substitutional disorder is only slightly temperature-dependent.

Complete solid solution is a special case. The IR peak positions of some solid solutions vary smoothly with composition along the series much like the Vegard law behavior in X-ray diffraction (one-mode behavior). In other solid solutions, bands appear that are characteristic of each endmember. These vary slightly in peak positions, but instead change relative intensity with composition along the series (two-mode behavior). Whether a given solid solution exhibits one-mode or two-mode behavior depends on relative masses of the substituting ions, and in complex structures (e.g., perovskites) some bands may exhibit one-mode behavior while other bands exhibit two-mode behavior. Theory also predicts that two-mode behavior will occur if bond strengths alone differ considerably for the two endmembers. This is observed for translations of the Si04 tetrahedron in solid solutions between grossular and pyrope (or almandine) garnet.

There also exists orientational disorder. It occur when a low energy barrier exists between alternative orientations of a molecule or ion within a structure. NHt in many structures may take on several orientations or may go into free rotation. Nonspherical ions such as Pb2+ and BP + may take on more than one position depending on the orientation on the lobe of charge occupied by the nonbonding lone pair electrons. At finite temperatures, randomizing the orientation of these species introduces a disorder and thus a line broadening of the vibrational spectra. This broadening is strongly temperature-dependent because of removal of the thermally-induced disorder at low temperatures.

IR Spectra of Glasses

The defining characteristic of natural and synthetic glasses is the absence of long range translational symmetry. However, substantial short range order remains in the form of distinct silica, germania, phosphate, or borate polyhedra making up the glass-forming network, producing vibrational spectra. Bands are broad, typically 30-100 cm - 1, but shift systematically with degree of polymerization of the network and can be used to estimate the numbers of bridging and nonbridging bonds in the individual tetrahedra. The low wavenumber "lattice" modes are usually lost when crystals are melted to form glasses, but very low wavenumber (50-200 cm -1) bands have been assigned to the vibrations ofalkali or alkaline earth ions within the cage formed by the network-forming polyhedra.


Hydrous Phases in Minerals

The small mass of the hydrogen atom places its vibrations at much higher frequencies than the lattice vibrations in silicates and the strongly polar nature of the O-H bond results in efficient absorption of infrared energy. These facts make it possible to observe minute quantities of OH and H20 in a mineral or a glass that is nominally anhydrous as well as in hydrated minerals. Such information is useful in that speciation and concentration of trace amounts of hydrogen in a mineral is pertinent to physical properties of a material such as rheology and response to radiation as well as to phase transformations and chemical reactions.

Symmetry analysis shows the bent H20 molecule will have two stretching modes (symmetric and asymmetric) and one bend. These have been observed in the 3200-3600 and 1600 cm -1 regions, respectively, with the exact positions depending on the host. Existence of a combination mode near 5200 cm - 1 is strong proof that H20 is present. Sharp bands are expected and observed for water oriented within the mineral structure, whereas broad bands are related to liquid water (i.e., fluid inclusions consisting of more than a few hundred molecules). One site for OH in a structure yields one sharp IR band. Minor variations in the orientation will add breadth to this one peak. Thus, differentiating multiple speciation of OH from H20 often requires information on the combination region.

Concentrations are determined by using Beer's law

A = e·t·C, (1)

where the extinction coefficient e is a constant determined from a standard with known hydrous content. A is the absorption coefficient, t is thickness, and C is concentration. The IR measurements should be taken from a single crystal or a thin film rather than from KBr powder because of the hydro scopic nature of the later. General calibration curves are available, but accurate results require standards of similar chemistry and structure.

Recent studies of nominally anhydrous minerals have documented the presence of OH in pyroxene, amphibole, kyanite, rutile, olivine, and p-spinel (Fig. 68); whereas H 20 is present in garnet and feldspar and as micro-inclusions in diamond. Hydrogarnet substitutions 4H = Si have been demonstrated.

Inferences on Bonding

Vibrational frequencies vary with the square root of bond force constants. For ionic crystals, the force constants usually vary roughly with the inverse cube of interatomic distances. Because the classifications and selection rules for infrared absorption are determined by the crystal structure, isostructural families of compounds have similar spectra. Minimum sets of force constants can be

3.6.3 Applications of Infrared Spectroscopy

0.4~----.....,..,==----------...,

Thickness = BOllm

0.3 W () Z ... « u; m .. II: 0.2 0 U) m «

g 0.1

po. ..

O.O-'---.----~---____r----__l

3200 3400 3600 3800

WAVENUMBERS, cm-1

187

Fig. 68. IR absorption spectra polycrystalline Fo90 (bottom) and a section of the same specimen which was transformed to beta spinel (top). Thickness is about 80 jJm. The p-phase region contains 113 ppmH(wt) . The IX-phase region contains 85 ppm H(wt). The sharp bands in the p-phase at 3330 and 3616 cm - 1 are consistent with crystallographic predictions of two sites for OH. (Young et al. in press)

extracted through normal coordinate analysis and these can be used to rationalize relations between bond character and structure in a relatively crude way.

Chemical and isotopic substitution in a mineral provide a means for determining what factors influence bond strengths. For garnets, the lattice constant and cation mass are most important. This structure is a network of various edge sharing polyhedra, so that changes in the bond length of anyone cation influences the bond lengths and electron sharing among all other bonds. In general, bond length should be the more important factor because most minerals consist of polyhedra that expand and compress independently. Other factors are compressibility of the solid and whether the substituting ions have spherical electronic configurations or not (e.g., fluoride perovskites).

Temperature-Dependent IR Spectroscopy

Reduction in temperature can do two things to a mineral spectrum: one is to sharpen the peaks in response to changing population levels. The other is to shift the bands in response to a decrease in the bond lengths and hence strengths.

Pressure-Dependent IR Spectroscopy

Increasing the pressure likewise shifts the peak positions, in response to changes in the structure, but does not change the population of the energy levels. Broadening occurs, but the origin of this phenomenon is uncertain. Phase transformations are readily observed through changes in peak positions and intensity with pressure. Pressure shifts are helpful in assigning atomic motions to IR bands. Anharmonicity of the solid is inherent in the non-zero values that


have been measure for the mode Gruneisen parameters

Yi = - dln(vJ/dln(V) (2)

of olivine and various spinels. Lower values are obtained for bands connected with Si than for bands associated with other cations, which is consistent with the incompressibility of the Si-O bond.

Calculation of Thermodynamic and Elastic Properties

Combining a Kieffer-type model with complete band assignments of vibrational spectra provides for accurate calculation ofCy and S (within 0.5%). For garnets, this is sufficient to indicate that calorimetric data on natural grossular is elevated by 5% over the true value, probably due to hydrous component. Spectroscopic calculations of thermodynamic properties of MgSi0 3 ilmenite and of germanates have been successfully used to calculate phase boundaries. The pressure dependence of the modes in olivine has been used to establish the dependence of thermal expansivity on pressure.

Bulk modulus and its pressure derivatives can be calculated from vibrational frequencies and their pressure derivatives. The accuracy of the calculation is generally about 5%, depending on the structure of the mineral and the completeness of the spectral data.

References

Aines RD, Rossman GR (1984) Water in minerals? A peak in the infrared. J Geophys Res 88: 4059-4071

Allen FM, Buseck PR (1988) XRD, FTIR, and TEM studies of optically anisotropic grossular garnets. Am Mineral 73: 568-584

Chopelas A (1990) Thermal properties of forsterite at mantle pressures derived from vibrational spectroscopy. Phys Chern Mineral 17: 149-156

Farmer VC (ed) (1974) Infrared spectra of minerals. Mineralogical Society of London, 525 pp Gervais F, Blin A, Massiot D, Chopinet MH, Naudin F (1987) Infrared reflectivity spectro

scopy of silicate glasses. J Non-crystal Sol 89: 384-401 Hawthorne FC (ed) (1988) Spectroscopic methods in mineralogy and geology. Rev Mineral 18,

698 pp Hofmeister AM, Chopelas A (1991) Vibrational spectroscopy of end-member silicate garnets.

Phys Chern Mineral 17: 503-526 Hofmeister AM, Xu J, Mao H-K, Bell PM, Hoering TC (1989) Thermodynamics of Fe-Mg

olivines at mantle pressures: mid- and far-infrared spectroscopy at high pressure. Am Mineral 74: 281-306

Kieffer SW, Navrotsky A (eds) (1985) Microscopic to macroscopic. Rev Mineral vol 14, Mineral Soc Am, Washington DC 427 pp

McMillan P, Akaogi M, Ohtani E, Williams Q, Nieman R, Sato R (1989) Cation disorder in garnets along the Mg3A12Si3012-Mg4Si4012 join: an infrared, Raman and NMR study. Phys Chern Mineral 16: 428-435

Merzbacher CI, White, WB (1988) Structure on Na in aluminosilicate glasses: a far-infrared reflectance spectroscopic study. Am Mineal 73: 1089-1094

3.7 Raman Spectroscopy in Earth Sciences 189

Ross NL, Navrotsky A (1988) Study of the MgGe0 3 polymorphs (orthopyroxene, clinopyroxene and ilmenite structures) by calorimetry, spectroscopy and phase equilibria. Am Mineral 73: 1355-1365

Rossman GR, Smyth JR (1990) Hydroxyl contents of accessory minerals in mantle eclocites and related rocks. Am Mineral 75: 775-780

Skogby H, Bell DR, Rossman GR (1989) Hydroxide in pyroxene: variations in the natural environment. Am Mineral 75: 764-774

Scott JF (1974) Soft-mode spectroscopy: experimetnal studies of structural phase transitions. Rev Mod Phys 46: 83-128

Turrell G (1972) Infrared and Raman spectra of crystals. Academic Press. London, 384 pp Young TE, Green HW, Hofmeister AM, Walken D (in press) Infrared spectroscopic investiga

tion of OH in p-(MgFe)2Si04 and coexisting olivine: implications for mantle evolution and dynamics. Phys Chern Min

3.7 Raman Spectroscopy in Earth Sciences

1. DUBESSY, R.Y. ORLOV, and P. McMILLAN

In the last decade, Raman spectroscopy has been increasingly applied in Earth Sciences including identification of phases and determination of their composition, structural studies of minerals, melts, fluids, and glasses, and access to thermodynamic variables. The aim of this chapter is to give a very basic introduction to the theory of Raman spectroscopy and to review its main applications in the Earth Sciences.

Fundamental in Raman Spectroscopy

Spontaneous Raman Scattering (SRS). When a substance is irradiated by monochromatic radiation with energy Eo (frequency vo, such that Eo = hvo), the light scattered by the substance consists of radiation with the same energy Eo, the Rayleigh scattering (10- 3 of the incident intensity) and weak spontaneous Raman scattering (SRS) ( < 10- 6 of the incident intensity). The SRS consists of radiation energies Eo ± E j (frequencies Vo ± vJ This frequency shift of the Raman lines with respect to the exciting radiation results from an energy exchange between the exciting radiation and the vibrational levels of the material. The different energy exchange processes involved in the different kinds of light scattering are given in Fig. 69 and a Raman spectrum of quartz is illustrated in Fig. 70.

The SRS is linked to the dipole moment p(l) induced in the matter by the oscillating electric field E vector of the exciting radiation according to the equation: p(1) = C(E, where C( is the polarizability tensor of the material. C( is a second rank symmetric tensor and its derivative with respect to a vibrational mode coordinate describes the intensity as well as the polarization state of the Raman line.

190

E1

Eo

~

>. .... ro

~

i Voo

Val

IR R SRS ASRS FL RRS

ANTI -STOKES

465

1: :0 (l'- quartz .... ro

-1000

127

-127 206

-500 o 500

RAMAN SHIFT (cm- i )


virtual state

Fig. 69. Energy diagram of different types of light-matter interactions Eo, E I , E2: electronic levels, vo.o and VO•I : energy of vibrational levels for the first electronic level Eo: I R infrared absorption spectroscopy; R Rayleigh scattering; SRS Stokes spontaneous Raman scattering. ASRS Anti-Stokes spontaneous Raman scattering; Fl fluorescence. RRS Resonance Raman scattering

STOKES

1000

Fig. 70. Stokes and Anti-Stokes Raman spectra for IX-quartz

It is difficult to predict the Raman activity of vibrational modes from first principles, but this can be easily done by inspection of the symmetry properties of the molecule or crystal using the method of group theory. This method allows the determination of the number and symmetry types of infrared and Raman active vibrational modes and gives a prediction of their polarization properties.

The energy supplied Eo by the incident radiation on the sample has been assumed to be smaller than the energy Eel required to excite the material to an excited electronic level. However, when Eo > Eel' intense fluorescence occurs and completely hinders the weak Raman signal. This is often the case in transition metal and rare earth-bearing minerals. Near infrared sources of


excitation with lower Eo, or methods of nonlinear optics and time-resolved methods (picoseconds) are used to avoid or reject fluorescence.

Instrumentation. A Raman experiment consists of a light source and a system which measures the intensity of the scattered radiations as a function of its wavelength. The typical light sources for linear Raman experiments are continuous wave argon and krypton ion lasers. These lasers offer 35 usable radiations from near UV to the near infrared boundary. Dye lasers give the possibility of continuously tuning the wavelength of the radiation over all the visible spectrum. The dye is usually pumped by the higher power lines of the CW argon laser. Dye lasers are often used in resonance Raman experiments.

As the Raman signal is weak compared to the intensity of the incident laser radiation, the detection must be highly sensitive and the spectrometer must efficiently reject the exciting radiation. Two different kinds of instruments are now available: dispersive spectrometers and Fourier transform interferometers.

In dispersive devices, the scattered radiation is separated as a function of wavelength by a sequence of holographic gratings (2, 3, or 4). In the case of a single-channel detection, the intensity of each wavelength is successively focused on the photocathode of a photomultiplyer tube by rotation of the gratings. In the case of a multichannel detection, the dispersed radiations are incident on an intensified photodiode array or a CCD detector. This permits a reasonably large portion of the spectrum around 300~500 cm -1 width to be studied, and results in better sensitivity, less time-consuming experiments, use of lower laser irradiance for fragile samples, but lower spectral resolution compared to single channel detection.

In the Fourier transform method, using a Michelson interferometer, the time autocorrelation function of the signal is recorded and then Fourier transformed into the spectrum as a function offrequency. The main disadvantage is that lines below 50 cm - 1 are difficult to observe even with the new filter technology. However, this techniques permits the use of near infrared excitation, (1060 nm line of the Nd~ Yag laser), which essentially eliminates fluorescence. Other advantages include high spectral resolution and high speed for obtaining the spectra. In addition, it is now possible to combine infrared and Raman spectroscopies in the same instrument. This technique will certainly find wide application in spectroscopy of strongly fluorescent minerals.

The incident laser beam can be focused through an optical microscope to give a spatial resolution of near 1 /-lm. Such devices are the Raman microprobes. The same objective collects the Rayleigh and Raman radiations in a near backscattering geometry. Scanning of the laser spot across the sample may provide a 2D image of Raman active vibrational mode and thus of a compound.

Raman experiments can be carried out over wide ranges of temperature, from liquid helium temperature to near 3000 K (C02 laser heating experiments), and at pressures up to tens or hundreds ofGPa in the diamond anvil cell. Other cell designs are used for studies of liquids and fluids at high temperatures and pressures.


Comparison with Infrared Absorption Spectroscopy. Vibrational properties are also commonly studied using infrared spectroscopy. In general, infrared and Raman spectra obtained for a given sample provide complementary data. However, in many applications, Raman spectroscopy presents a number of advantages over infrared techniques. The energy of interest in vibrational spectroscopy is approximately 10-4000 cm - 1. In an infrared experiment, this range requires light with wavelengths of 100 Jim to 2.5 Jim. It is not yet possible to cover this entire range with a single detector, unlike in Raman spectroscopy. Usually, the spectrum from around 200 to 4000 cm -1 can be easily obtained in a single infrared experiment, but obtaining data below 200 cm - 1 requires increasingly greater effort. Secondly, due to the wavelength of infrared light, it can be difficult to focus infrared beams for micro-infrared spectroscopy. However, visible light can be easily focused to approximately 1 Jim using conyentional optics, and micro-Raman spectroscopy is a much simpler and better-developed technique. Third, water is a common solvent of interest in geochemical experiments. However, water has such strong absorption across the entire infrared range that it is difficult to obtain infrared spectra of species dissolved in water. In contrast, water has a very weak Raman signal, and species dissolved in aqueous fluids or liquids can be easily studied in situ by Raman spectroscopy. Finally, the infrared spectra of solids generally consist of broad bands which are only partly due to absorption. The analysis of such spectra can be difficult, and in the case of powders, the spectra can be dependent on the size and shape of the powder particles. In general, this problem does not arise in Raman spectroscopy, where the dimensions of the powder particles are large compared with the wavelength of the incident light beam.

Applications in Earth Sciences

Phase Identification and Analysis. In some cases, Raman spectroscopy is the simplest and even the only method to identify a mineral, specially when included in another transparent and nonfluorescent mineral. Furthermore, there are isomorphous minerals with close dimensions of the elementary cell for which X-ray phase analysis is of low efficiency.

Micro-Raman spectrometry is the most popular technique for the identification of gases of the C-O-H-N-S system inside fluid inclusions (C02, H2S, CO, S02' N2, O2, H2, CH4 , C2H6 , C3HS) and the estimation of their mole fraction after calibration. These gas analyses are useful for the identification of fluid mixing and unmixing processes, chemical equilibrium and disequilibrium, and f02 and fS2 estimations of circulating paleofluids.

The dependence of the Raman shift of the v 1 line of CH4 on pressure can also be used for the determination of gas pressure inside the inclusion. 13C02 molecules have been identified but isotopic compositions with accuracies comparable to those obtained by traditional techniques in isotope geochemistry


have not yet been achieved. SOi - is the only poly atomic ion identified with a detection limit around 2 x 10 - 3 molal. Monoatomic ions are indirectly identified by the Raman spectrum of the salt hydrate which nucleates on cooling the electrolyte solution.

Because of the nondestructive character of such analysis, this method is also used in gemmology.

More generally, micro-Raman has been increasingly applied to Material Sciences. In addition, Raman remote sensing methods have been developed for the control of environments, especially atmosphere and ocean waters.

Structural Studies. The study of polarized Raman spectra of crystals with respect to the crystal orientation gives valuable information on the symmetry of vibrational modes, and hence the symmetry of the structural groups (for example, distortion of CO~ - in carbonate minerals from trigonal to planar symmetry), and their orientation (for example, the orientation, of CO~ - ions in apatite and of water molecules within channels in natural cordierites).

The Si-O stretching frequency depends on the polymerization state of the tetrahedral Si04 and on the coordination state of silicon. This has been used extensively in structural studies of silicate glasses and melts. The O-H stretching frequency is highly dependent on the OH environment and the presence or absence of hydrogen bonding and Raman spectroscopy is an efficient tool for locating the hydrogen atoms in OH-bearing minerals.

The mechanical quality of piezoelectric quartz is directly related to their OH and H2 0 content. Raman bands at 3520 and 3590 cm -1 have been assigned to these OH ... OH- hydrogen bonded structures. By contrast, the wide nonpolarized band between 3000 to 3600 cm - 1 corresponds to randomly oriented H 2 0 molecules.

The identification of the stacking period on a poly type can be obtained with accuracy as good as that achieved by X-ray diffraction and has been applied for the understanding of Zn-blende formation. The stage of graphite intercalation compounds can be determined by the number (1 or 2) and the frequency of the stretching vibrations of carbon atoms in aromatic layers.

In ionic solid solutions, the frequency of vibrational modes varies monotonously with the change of the composition whereas the intensity of line varies weakly. This one-mode type of oscillations makes it possible to determine the composition of the solid solution. However, there are reverse situations for which the intensity of lines varies strongly whereas their frequency is roughly constant with composition. This is a two-mode type of oscillations and it is not possible to distinguish between composition or exsolution effects by inspection of the spectrum.

Raman spectroscopic studies have placed a major role in developing our current understanding of water structure. Raman spectrum of liquid water at room pressure and temperature is interpreted in terms of hydrogen-bonded H 20 molecules. Raman studies of water over wide ranges of temperature and pressure have shown the variations of hydrogen-bonded structures with temper-


ature and pressure. The modification of water structure by dissolved electrolytes was also extensively studied.

Phase Transitions. Classic applications of Raman spectroscopy in solid state physics and chemistry deal with the study of displacive phase transition through observation of soft modes. A soft mode is an anharmonic vibration whose atomic displacements mimic the displacive changes required by the phase transitions. For a thermally driven phase transition the soft mode frequency varies with temperature, reaching zero at the transition temperature Te. The rx-f3 transition in quartz is of this type and the broad band at 207 cm - 1 decreases rapidly in frequency as temperature increases. However it has been shown recently by Raman spectroscopy and many other methods that the rx-f3 quartz phase transition does not proceed directly but passes through a series of incommensurate phases between 844 and 846 K.

There is current intense interest in the possibility of soft mode behavior linked to displacive phase transitions in MgSi0 3 perovskite which is thought to be a major constituent phase of the Earth mantle. The dependence of vibrational frequencies on pressure and temperature, Gruneisen parameters, are important thermodynamic quantities. Study of diamond has established constancy of those parameters up to 700 kbar and has shown the stability of the tetrahedral configuration of carbon at very high pressure.

Order-Disorder. Vibrational spectra are sensitive to the degree of order of a material, with broadened bands being observed for more disordered phases. The simplest case of order-disorder study concerns graphite. Well-crystallized graphite exhibits a strong Raman band at 1580 cm - 1. The structure of carbon compounds can be related to the extension and planarity of the aromatic layers and their stacking. With increasing disorder by comparison with graphite, the 1580 cm -1 band broadens and moves to higher frequency and new band appears at 1350 cm - 1. The theoretical treatment allows the interpretation of the relation between the correlation length La calculated from X-ray data and the Raman spectra, which gives a measure of the degree of order of the carbon mineral. This method has been used to monitor the graphitization of natural carbons by natural or experimental heating, and the alteration of initially strictly ordered graphite.

The rearrangement K-Na disordering in anorthoclase occurs over a wide temperature interval and has been identified by the widening oflines at 515 and 475 cm -1 of the Si-O network. Systematic changes in the spectra have been observed with decreasing Si-AI order in alkali feldspars and cordierite.

Orientation disordering associated with the rotation of complex ions is easily identified. By contrast, X-ray diffraction suggests incorrectly a center of symmetry, since this method averages over all the static distribution.

Thermodynamic Properties. Heat capacity of crystals can be calculated from the dispersion curves of acoustic and optic modes over all the Brillouin zone. However, the knowledge of phonons at only the center of the Brillouin zone


spectra is sufficient enough for a good prediction of solid heat capacity with the Kieffer's model (Kieffer and Agoskov, Chap. 6.5.2 this Yol.). The optic modes are known from the Raman and infrared spectra. In Brillouin spectroscopy, the incident laser light interacts with acoustic vibrations and other low energy excitations, and is scattered inelastically with a very small change in energy. These very small changes in energy are measured with an interferometer, rather than an energy-dispersive spectrometer as used in Raman spectroscopy. The acoustic wave velocities measured in a Brillouin scattering experiment are directly related to the elastic properties of the sample.

In polyatomic ionic aqueous solutions, the variations of equilibrium constants with pressure and temperature can be directly measured by Raman spectroscopy, thus giving the enthalpy (AH OR) and the reaction volume variation (AY\).

Other Raman-Related Laser Spectroscopies

As already noted, there are a great number of Raman-related laser spectroscopies which are well known in physics, chemistry, biochemistry and bio-physics, but have not yet had wide application in the Earth Sciences. Three of these which are likely to provide the most useful information are resonance Raman, electronic Raman, hyper-Raman, and coherent anti-stokes Raman spectroscopies.

Resonance Raman Spectroscopy (RRS). In normal scattering experiments, the incident laser energy is chosen to lie far from any electronic transitions of the sample in order to avoid absorption and fluorescence. However, if the laser energy nearly coincides with an electronic absorption frequency, it can cause a large intensity enhancement of Raman active modes associated with the absorbing center, or chromophore (Fig. 69). This technique has been used to study the color-producing species in ultramarine (synthetic blue lazurite), identified as S2' and S; complexes.

Electronic Raman Spectroscopy (ERS). ERS occurs on dipole-forbidden transitions between electronic states of equal parity. Most of the studies deal with the study of the crystal field splitting of transition metals (3 dn) or rare earths (4 fn). F or instance, it has been established that y4 + replacing Zr4 + in zirconium decreases local symmetry down to Cz and carries out tunnel transitions at frequencies of 7 x 1011 between two potential minima spaced by 0.15 A. The ERS intensity lines are two to four orders of magnitude lower than in normal Raman experiments. This technique requires equipment for obtaining cryogenic temperatures.

Hyper Raman Spectroscopy (HRS). HRS is an nonlinear optic three photon process which requires high-power lasers such as ruby or Nd-Yag lasers. In HRS experiments, two incident laser photons, with Vo frequency, combine


within the sample through non linear optical mixing and results in weak hyper Raman scattering with frequencies 2vo + Vi and thus situated in a shorter wavelength interval range where fluorescent radiations are absent. The selection rules of HRS are different from those of SRS, and all infrared active modes are active in HRS experiments. This can be extremely useful for obtaining infrared frequencies for samples or under conditions where the infrared spectra are difficult to measure. For example, the low frequency ferroelectric soft mode of perovskites, normally only active in the far-infrared, is observed in HRS. In addition, HRS has been used to separate the contribution of transverse and longitudinal modes in Si02 glass and liquid.

Coherent Anti-Stokes-Raman Scattering (CARS). CARS is also a nonlinear optic technique using one laser with constant VI frequency and a tunable laser with frequency V2 such as VI - V2 = Vi' where Vi corresponds to a vibrational frequency of the sample. Under these conditions, a coherent strong emission occurs from the sample at a frequency Vas = 2VI - V2 = VI + Vi' In principle, there is no need for a spectrometer in CARS experiment. The spectral resolution is very high (down to 0.00 1 cm - 1). It is clear that in CARS useful signal and fluorescence does not overlap. This method only applies to transparent and homogeneous media such as gases, flames, plasma and transparent and perfect crystals, with very narrow Raman line widths.

References

Anderson A (1971) The Raman effect vols 1 and 2. Marcel Dekker, New York Berg B, Vallode M, Martinez G (1984) Raman scattering investigation of the rx-p transition

and of the incommensurate phase in quartz. J Phys C Sol State Phys 17: 167-171 Buback M, Crerar DM, Koplitz L (1987) Vibrational and electronic spectroscopy ofhydroth

ermal systems. In: Ulmer GC, Barnes HL (eds) Hydrothermal experimental techniques. Wiley, pp 333-359

Dele-Dubois MZ, Dhamelincourt P, Schubnel HJ (1980) Etude par spectroscopie Raman d'inclusions dans des diamants sapphirs et emeraudes. Rev Gemmol 11-13; 64: 13-16

Chase DB (1986) Fourier-transform Raman spectroscopy. J Am Chern Soc 108(24): 7485-7488 Dhamelincourt P (1982) Instrumentation and recent application in micro-Raman spectro

scopy. Microbeam Anal 17: 261-269 Dubessy J, Poty B, Ramboz C (1989) Advances in C-O-H-N-S fluid geochemistry based on

micro-Raman spectrometric analysis of fluid inclusions. Eur J Mineral 1: 517-534 Geilikman MB (1982) Mechanisms of polytype stabilization during the wurtzite-sphalerite

transition. Phys Chern Mineral. 8(1): 2-7 Gardiner DJ, Graves PR (1989) Practical Raman spectroscopy. Springer, Berlin Heidelberg

New York Ghose S (1985) Lattice dynamics phase transitions and soft modes. Rev mineral 14: 127-163 Hemley RJ, Bell PM, Mao HK (1987) Laser techniques in high-pressure geophysics. Science

237: 605-612 Irish DE, Brooker MH (1976) Raman and infrared spectral studies of electrolytes. In: Clark

RJH, Ester REM (eds) Advances in infrared and Raman spectroscopy, vol 2, pp 212-311 Lyons KB, Sturge MD, Greenblatt M (1984) Low-frequency Raman spectrum of ZrSiO 4: V4 +.

An impurity-induced dynamical distortion. Phys Rev B 30(4): 2127-2132

3.8.1 Principles, Technique, Applications in Mineralogy 197

McMillan PF (1989) Raman spectroscopy in mineralogy and geochemistry. Annu Rev Earth Planet Sci 17: 255-283

McMillan PF, Hofmeister AM (19!!8) Infrared and Raman spectroscopy. In: Hawthorne FL (ed) Reviews in mineralogy, vol 18. Spectroscopic methods in mineralogy and geology, pp 99-159

Orlov R Yu, Guseva EW (1989) Raman spectroscopy in mineralogy and material science. Izv AN SSSR, Ser Geol 4: 84-95

Pasteris ID, Wopenka B, Seitz lC (1989) Practical aspects of quantitative laser Raman microprobe spectroscopy for the study at fluid inclusions. Geochim Cosmochim Acta 52(5): 979-988

Salje E (1986) Raman spectroscopy investigation of the order parameter behavior phase transition and evidence for Na-K site ordering. Phys Chern Mineral 13(5): 340-346

Turrell G (1972) Infrared and Raman spectra of crystals. New York, Academic Press Vogt H (1982) Coherent anti-stokes Raman scattering and hyper-raman scattering. In:

Cardona M, Giintherodt (eds) Light scatering in solids II. Springer, Berlin Heidelberg New York, pp 277-327

3.8 Electron Paramagnetic Resonance (EPR)

3.S.1 Principles, Technique, Applications in Mineralogy

J.A. WElL, Y. DUSAUSOY, and S.L. VOTYAKOV

Electronic paramagnetic resonance (EPR), sometimes called electronic spin resonance (ESR), was discovered in 1945 by Zavoisky. EPR, a spectroscopic method used in various fields of biology, chemistry, mineralogy, and physics is based on the resonant absorption of microwaves by paramagnetic atoms, molecules, ions, and free radicals in dielectric solids, liquids, and gases. Very high sensitivity (1013 spins in 1 cm3) for identification of trace amounts of paramagnetic centers is characteristic of this nondestructive analytical method. In mineralogy, EPR allows the study of paramagnetic impurities as trace elements and of defects formed by natural or artificial irradiation, in crystalline as well as amorphous materials. These species often are evident as color centers. The main results are: identifying, locating and measuring the concentration of impurities in the compound structure, determining local crystal electric and magnetic fields and consequently local distortion around each center, suggesting chargecompensation processes, revealing atom motional effects, and identifying different types of radiation defects. Many applications have been proposed, relating: to the crystal chemistry of trace elements, radioactivity effects, thermal history and dating of rocks, to nucleation and crystallization processes, to exploration geology, and to characterization of coals and lunar samples. Recent investigations are characterized by combined application of various spectroscopic methods: EPR with optical absorption, luminescence, and nuclear magnetic resonance (NMR).


Basic Principles and Experimental Arrangement

EPR depends on the presence of partly filled electron shells, i.e., presence of unpaired electrons. The spin magnetism of electrons cancels when they are paired. In the simplest case, that ofthe free electron, the magnetic moment vector is P.e = - gePeS where ge is the spectroscopic factor of the free electron (without orbital angular momentum), equal to 2.00232; Pe is the electronic (Bohr) magnet on equal to 0.9274 x 1023 JT- I and S is the (unitless) electronic spin operator vector. The unit of magnetic fields is the Tesla (1 T = 10 Gauss).

Electronic Paramagnetic Resonance. If a static external magnetic field (flux density) B is applied, a weak magnetic interaction called the Zeeman effect can be described by the energy (Hamiltonian) operator:

.Yf = - P.e . B = gePeB· S = gePeBSz, (1)

where Sz is the projection of Son B (the z axis) and takes on 2S + 1 values, given by M = - S, - S + 1, ... , + S. For instance, in the case of one unpaired electron (M = ± 1/2), there is a small but essential difference, given by the Boltzmann distribution, in the spin populations of the ground (E _ 1/2) and excited (E+ 1/2) energy levels. An alternating magnetic field BI (frequency v) can be used to stimulate transitions between adjacent energy levels. If the spin sample is placed in such a field (applied perpendicular to the field B), it comes into resonance (r) when the magnetic-field magnitude B or the frequency v has been adjusted so that:

hVr = EM-I - EM = gePeBr, e.g., M = + 1. (2)

The resulting absorption curve in this simplest case shows a single line (Fig. 71a-c).

EPR Technique. In practice, adjusting v is more difficult than varying B. Accordingly, the resonance condition is attained by scanning the magnetic field strength B until the resonance signal (value Br) is obtained. The transition is excited using BI radiation having a constant frequency v produced by a microwave source, e.g., a klystron (Fig. 72). Typical EPR spectrometers operate at the following frequencies:

Frequency 9.5 GHz wavelength 3.2 cm, called X band

Frequency 35 GHZ wavelength 0.86 cm, called Q band.

A few spectrometers operating near 100 GHz, at very high magnetic fields, have now been constructed.

The basic spectrometer arrangement is shown in Fig. 72: an electromagnet gives a homogeneous scannable magnetic field B, a source yields monochromatic microwaves adjustable within a limited frequency range (centered at v) and power range, a resonating cavity holding the sample is located between the poles


E A

'-----+----·8 '---"'--~I---""-- 8

a b

.M d8

'---.£--+---8

c

Fig. 71a--c. a Magnetic energy splitting for S = 1/2, showing transition at constant frequency and varying magnetic field: b Absorption line, c Absorption line first-derivative

Fig. 72. A schematic representation of a basic EPR spectrometer arrangement

of the magnet. The cavity stores the microwaves and thereby sets up oscillating field B1 perpendicular to B; the sample is located at the maximum of this excitation field. The resonance signal is detected by a variation of the current induced in a detection crystal. The klystron/cavity and the cavity/detector are interconnected by waveguides or by coaxial cables (Fig. 72). Often the sample is held at cryogenic temperatures in order to optimize the sensitivity and resolution. Narrow lines are optimal to achieve these.

To minimize perturbations of the signal due to noise and baseline shifts, the resonance line is B-field modulated (say, sinusoidally at 100 kHz) with an amplitude small compared to its width. With use of synchronous detection, the detected signal appears as the derivative dA/dB of the absorption line amplitude with respect to the static magnetic field (Fig. 71).

Various sophisticated versions of EPR can be utilized. For instance, combined EPR and NMR (called ENDOR = electronic nuclear double resonance) gives enhanced spectral resolution when nuclear spin effects are present. In recent years, pulsed microwave EPR spectrometers have become available, allowing Fourier-transform techniques (FT-EPR) to be applied. This type of instrument will enable rapid time-resolved measurements, correlated coupling to


pulsed lasers and other excitation sources, as well as multi-scanning (computer storage) to improve signal-to-noise ratios.

Another advance has been the development of EPR imaging techniques. These will allow determination of the distribution of paramagnetic impurities within mineral samples.

Electronic Paramagnetic Resonance in Crystals

As shown above, the resonance spectrum of a free. electron consists of one absorption line. In solids, the spectrum is more complicated due to the local electric field, to the coupling of electronic with nuclear magnetic moments, and to the interactions between each electron magnetic moment with the magnetic fields generated by other electrons (ions with more than one unpaired electron).

The g Matrix. In a crystal, the g factor deviates from ge, corresponding to a pure spin moment and takes on a variable value according to the orientation of the crystal relative to the magnetic field B. This g anisotropy is due to the localsymmetry crystal field, which introduces an effective energy separation .1 between the ground and electronic excited states of the species studied. The g value is given roughly by g = ge - cA./ .1, where .A. is the spin-orbit coupling constant and c is a coefficient the value of which depends on the type of orbitals involved in the coupling. The orders of magnitude of .1 and .A. typically are 10 000 and 10 cm -1 respectively. The term cA./ .1 can be considered as a perturbation of the ge value. Then the interaction between Band S is given by .Yt' = PeBogoS, where the 3 x 3 matrix g substitutes for the scalar value ge' Basically it contains six independent parameters, three being geometrical (defining its principal-axis coordinate system) and three principal values which relate to energy ratios as described above. This matrix can be used to calculate the g factor at any orientation of B relative to the crystal.

In many crystal structures, the same paramagnetic species occurs at several orientations. Such symmetry-related centres will give rise to separate EPR spectra, at line positions described by symmetry-related g-matrices, which may, however, superimpose at some angles of B relative to the crystal.

The Hyperjine and Superhyperjine Structure. In the presence of a nuclear magnetic moment neighboring the electronic magnetic moment, a weak interaction arises between these two moments and produces a (usually) small further splitting of the spin energy levels and hence of the EPR spectrum, the so-called hyperfine structure (HFS). If one or more nuclear spins are on the surrounding ligands rather than on a central ion, the interaction gives rise to the so-called superhyperfine structure (SHFS). These hyperfine effects in the EPR spectrum yield detailed information about the unpaired electron distribution, i.e., on or near what atoms it is located, and allows identification of these atoms. The hyperfine splittings occur even at B = O.


Each nuclear magnetic moment is expressed by Po = goPoI where I is the nuclear spin operator, Po the nuclear magnet on (which is 1836 times less than Pel, and nuclear g-factor go depends on the nucleus considered (see tabulations). The nuclear magnetic moment Po is thus much weaker than the electronic moment Pe and hence its effect can be considered as a small additive perturbation Bo of the magnetic field seen by the electron spin. The interaction is given by:

(3)

where

N

.Yt' A = S·Ao·I + L S·Aj·Ij. (4) j-l

Here Ao is the primary hyperfine (Aj the superhyperfine) matrix, the component values of which depend on the orientation, distance (and magnitude) of the nuclear dipole relative to the electronic spin. If the Zeeman effect due to B is much larger than the electronic-nuclear spin interaction, then each energy level of the electronic spin will be given by: E(M,mI) = gPeBM + AMmb where the nuclear quantum number mI takes on the 21 + 1 values - I, - I + 1, ... , + I. The appropriate selection rule (ILiMI = 1, AmI = 0) gives the single-nucleus resonance condition:

or Br = hvrfgPe - AmdgPe. .. . (5)

Here A is the hyperfine splitting parameter derivable from the matrix A, and varies with crystal orientation. Thus every line is split into (usually) 21 + 1 lines called the hyperfine structure of the EPR spectrum. (Fig. 73). The relative intensity of the HFS sets is a function of the natural abundance of isotopes with nonzero nuclear spins. For example: see the hyperfine structure from 53Cr (I = 3/2, 9.5% natural abundance) in Fig. 74. In the case of N > 1 nuclei with nuclear spins, the spectrum is accordingly more complicated.

E

M

1 "2 ~ ........ ~

, .... 1 ............

-"2

m 1 o

-I

-I o 1

I I

:: l-------m-~ Fig. 73. Magnetic energy level splitting for an ion with S = 1/2 and I = 1

202

Cr 3+

(FS) t Cr3+_ 27AI3+

( HFS)

B~

I,', I, . , ,

53Cr3+ _ 27A1 3+

(SHFS)


Fig. 74. EPR spectrum of forsterite at 9.48 GHz and 295 K. It shows a fine-structure (FS) line of Cr3 +; HFS quartet of 5 3Cr3 + and HFS sextet of Cr3 + ;Z 7 AI3 + and SHFS of a 53Cr3+ /27 Al 3+ pair center. (Bershov et al. 1983)

The Fine Structure in the Case S > 1/2. If the paramagnetic center has only one unpaired electron, the EPR spectrum shows one absorption line, possibly split into a hyperfine multiplet. For a total electronic spin larger than 1/2, the spectrum is more complicated, exhibiting so-called fine structure (plus possibly hyperfine structure). The electrons interact through spin-orbit and spin dipolar magnetic coupling. These interactions introduce a new term S· D· S (where D is the electronic quadrupole 3 x 3 matrix) into the general expression of the spin Hamiltonian:

Yf = PeB·g·S + S·D·S + ... + YfA' (6)

yielding Zeeman energies + fine structure + hyperfine structure. Even if the magnetic field B is zero, for a center (say, a transition-metal ion)

with S larger than 1/2, the crystal field splits the 2S + 1 electronic spin-energy levels, partly or completely. The splitting, often called the zero-field splitting (ZFS), depends on the crystal-field symmetry and on the odd or even number of electrons (Kramers or non-Kramers ions). In all cases, any remaining degeneracy is fully removed by the Zeeman interaction (Fig. 75). Since the Zeeman levels now are no longer equally spaced, a fine structure of the spectrum appears, with theoretically 2S allowed absorption lines (I.dMI = 1), in the absence of hyperfine structure. However, the absorption-line number can be modified for various reasons: if the zero-field splitting is much larger than the microwave photon energy hv, then only transitions within each Kramers doublet ( ± 1/2; ± 3/2; ... ) can be observed for modest B. If the transition probability between any two levels is very low, then the absorption line can disappear into the background noise. For many ions, e.g., Fe3 +, Cr3 +, Mn 2 + ... , "forbidden" transitions (I.dMI

3.8.1 Principles, Technique, Applications in Mineralogy

I 1 I 1 I

+1 +3+5 I -2-2 -2 1

8=0

FREE ION

+1 -2

NON-CUBIC 8=0 8-

6

5

4

3

2

CRYSTAL FIELD ZEEMAN EFFECT

203

Fig. 75. Theoretical crystal-field splitting and Zeeman effect for an ion with S = 5/2

= 2,3) can also be observed. Figure 76 shows the energy levels of an Fe3 + center in substitution for AI3+ in kyanite, including the allowed IL1MI = 1 and forbidden IL1MI > 1 transitions, and the corresponding EPR spectrum.

Fine Structure and Local Environment. The fine structure often is calculated as a summation of fine-structure components I'B::'O::' where the energy parameters B::' in the spin Hamiltonian describe the contribution made by all the atoms of the crystal to the crystal field and reflect the charge distribution (symmetry and magnitude) around the magnetic center. Here n ranges over all even nonnegative integers, while m = - n, - n + 1, ... , + n. The 0::' spin operators serve to take into account the geometric environment. Then the spin Hamiltonian can be written:

(7)

The first sum, over the O~ operators, represents an alternative way of writing the term S· D· S introduced in Eg. (6). The primary deviation from local cubic symmetry is given by the B~ value. For lower symmetry, the other parameters B~, B; 1, B; 2 contribute to the fine structure. All these second-order parameters take into account the contribution of first, second, ... neighboring atoms which decreases (in principle) as r- 3 where r is the distance between the paramagnetic center and the atoms. By contrast, only the atoms of the coordination polyhedron (formed by the nearest neighbors) contribute effectively to the fourthorder B~ parameters, since they vary as r - 5 or faster. Hence the determination of the latter parameters contributes to the knowledge of the immediate local symmetry, i.e., of the geometric configuration of the central polyhedron.

Various theoretical estimations of ZFS parameters have been done, among them a number of publications devoted to the so-called superposition model (Newman and Urban 1975; Lehmann 1979-1980). This empirical model is based on the assumption that the ZFS is determined by distortions of the first coordination sphere alone and that it can be expressed as a sum of contributions


6 E

5

+5 Ic==~:1::~~----~L---lr----~---"2 1.00r f ~ 0.25[=j:t~t==:=~=+==t===[==~4 2 0

it -1.20 r--='F=r=::::::t::::::--+--.J~ 3 I I I I I

I ~ M I = 3: 2: 2: I I Ii I I ---7-,--~-----i---~---~----

dA dB

o

I I I I I I I I I I I I I I I I I I I I I I I I

I

5 10

B

B(KG)

15

Fig. 76. Energy-level diagram for Fe3+ in a kyanite single crystal showing the allowed ILIMI = 1 and forbidden (ILlMI > 1) transitions and the X-band EPR first-derivative spectrum.

(Dusausoy et al. 1990). An extraneous line is also visible

from the nearest-neighbor atoms. Correlations of ZFS with site distortions of Mn2 +, Fe3 +, Cr3 + centers in different hosts were attained, so that site positions of impurity ions and different ligands can be obtained.

Determination of Site Occupancy. The study of the angular dependences of the EPR transitions relative to the crystal axes and coordination-polyhedron symmetry axes allows identification of the site (when more than one type is possible) of paramagnetic ions, e.g., Fe3 +, Cr3 +, in cubic oxides, in garnets, in sillimanite and in cassiterite, and Nb 3 +, Yb3 +, V4 + in zircon crystals. In the case where the coordination polyhedron of the paramagnetic center has a very low symmetry (e.g., triclinic), a possible way to locate the paramagnetic ion in the structure is to determine the orientations of the pseudosymmetry axes of the fourth-order parameters B:r' of the spin Hamiltonian and compare these with the orientations of the pseudosymmetry axes of the crystal field calculated for each polyhedron in the structure. This was done for example, in the case of Fe3 + in feldspars, and in kyanite (Fig. 76).

Applications of EPR in Material Science and in Mineralogy

EPR spectra can be observed in almost all solid samples, liquids and gases containing 3d-Sd and 4f-Sf ions, certain atoms and molecules with unpaired electrons (H, NO, NH;, etc.), and electrons or holes localized at "point" defects


(yielding 0-, SiO~-, CO2, PO!-, etc.) in crystals. For almost all paramagnetic centers, the lower-limit detectable concentration typically is about 0.0001 % in modern EPR spectrometers, while the upper limit (0.1-1 %) is determined by the extent of EPR line broadening due to the interaction between the paramagnetic ions (defects) which occurs with increase in their concentration.

EPR experiments can provide detailed information about the oxidation state of impurity ions, as well as their structural position and type of charge compensation, and yield structural models of various point defects. Examples of such investigations in solid dielectrics, semiconductors, ferroelectrics, laser crystals, and luminophors have been reviewed.

EPR studies of natural and synthetic minerals have been carried out in almost all classes of oxides, hydroxides, sulfides, halides, sulfates, phosphates, and silicates, chiefly on two paramagnetic center types: paramagnetic impurity ions and paramagnetic radiation defects. Here are some selected examples:

Quartz is a common mineral in almost all igneous rock; it can also be found in sedimentary and metamorphic rocks. Crystals are mass-produced for various technical applications. A series of quantitative studies have been performed on paramagnetic defects (involving Ag, AI, Cu, Fe, Ge, H, P. Ti) in this material. A complex charge-transfer chemistry has been found, involving transport not only of electrons but also of interstitial ions M + (where M = H, Li, Na, Ag, Ag2).

Numerous paramagnetic species, such as [AI04/M] + (an 0 - center) and [Ti04/M]O (a Ti3+ center), have been characterized. Another type of common defect is typified by the oxygen-vacancy species known as the E'l center, in which an unpaired electron is localized at a Si ion adjacent to an oxygen vacancy. Some ofthe centers induce coloration of the mineral, and for some species the EPR and optical absorption spectra have been correlated. Their origin in some cases is due to natural radioactivity and these can be used as indicators of radioactive sources and dosage. When their concentration increases linearly with radiation time, then dating of recent faults has been proposed to be feasible. Correlation between the EPR spectra of natural quartz and facies of rock metamorphism type of granitoids and pegmatites has been studied. The spectroscopic characteristics of granitoids, pegmatites, kimberlites, and sedimentary rocks have been used as possible indicators of their genesis. Aluminum traces in quartz have been proposed as an indicator for the temperature offormation. Attempts to correlate gold content with EPR signals in adjacent quartz crystals have been reported.

In feldspars, paramagnetic ions (Fe, Cr, etc.) and numerous radiation defects have been observed, such as AI-O - -Si, AI-O - -AI (in albite, sanidine, microcline, orthoclase, anorthoclase, scapolite, labradorite), Si-O - _X2+ (in amazonite, labradorite, oligoclase, bytownite), AI-O- _X2 + (in microcline), NHj, and N 2 -. These studies show correlation between the EPR spectra and color, thermal history, and natural irradiation. The petrological application of EPR spectroscopy of feldspars from pegmatites and granitoids of the Ukraine were studied by Matyash et al. (1981).

The EPR studies and microprobe analyses on cassiterites and synthetic doped Sn02 led to an explanation of the atom replacement and charge


compensation when Fe3 + substitutes for Sn4 +. Four such paramagnetic centers were encountered: center "Sd2": 2Sn4 + --+ Fe3+ + Nb5 +; center "Sd3": 2Sn4 + --+ Fe3+ + Ta5 +; center "I": Sn4 + + 0 2 - --+ Fe3+ + OH- (with OH- in the coordination polyhedron); center "SD 1" has the same composition but with OH- at longer range. The powder EPR spectra of cassiterite from various different mines can be interpreted on the basis of one of the three spectrum types shown in Fig. 77.

Numerous publications have been devoted to EPR investigations of accessory minerals such as zircons, local structural defects associated with rareearth ions, d-orbital transition elements, and oxygen vacancies. A great variety of impurity ions and radicals was observed by EPR in apatite and members of its group, in beryl, and in scheelite. Recent EPR studies of single-crystal diamonds have led to elucidation of nitrogen-atom impurity centers probably associated also with oxygen, and of substitutional nickel-atom centers.

Ice, a mineral commonly found at temperatures below 273 K, exhibits EPR spectra of 0- radicals and of H02 after y-irradiation. Atomic hydrogen and hydroxyl radicals have also been studied in this medium.

EPR investigations of carbonate minerals are very useful in solving the sedimental petrology problem of Mn2 + concentration and its distribution between nonequivalent positions within the dolomite structure, as produced by the conditions of the mineral crystallization and evolution.

St AGNES Sd I

o 2 3

B(KG) -

I I

4

Fig. 77. Powder EPR spectra of cassiterites from: Saint Agnes, England (Fe20 3 0.32 wt. %); Elmeki, Niger (Fe20 3 0.88, Nb 20 s 0.79wt. %); Penonta, Spain (Fe2 0 3 0.50, Nb2 0 s 0.82, Ta20 S 0.35 wt. %). Lines from other centers are also visible. (Ruck et al. 1988)


EPR studies, featuring modern pulsed techniques, have revealed important details of paramagnetic species held in (or on) zeolites. For instance, the studies of AgO and ofCu2 + in such hosts have been informative from both structural and catalytic aspects.

EPR studies of various fossil fuel sources have disclosed, besides organic free radicals, various metal-ion species Mn2+, Fe3+, and porphyrin-bound V02 +.

Many EPR studies of glasses have been carried out. For example, detailed analysis has revealed that the 9.5 GHz EPR powder/glass spectrum of Fe3+ in silicon dioxide has a strong peak at ca. 165 mT and is spread over ca. 750 mT, with all features dominated by the ZFS parameters rather than by the Zeeman (g) parameters. The interpretation the spectra of glasses, powders, and amorphous minerals is generally based on spectral simulation achieved by statistically superimposing the anisotropic spectra arising from all crystal orientations of a given center. Certain d- and f- transition elements in noncrystalline samples (silicate, phosphate, borate glasses) were studied in order to identify the nature of ions or molecular complexes such as CrO~ - and VOl- impurities, radiation defects, to determine the coordination site, to correlate EPR spectra with glass composition, and to monitor the nucleation mechanism.

In summary, all the studies mentioned indicate that EPR spectroscopy is an important practical tool, useful to provide information about the nature of the point defects in minerals, for understanding the luminescence and optical absorption properties of minerals, revelation of their crystallization conditions, radiation and annealing history, etc. connected with geochemistry and petrology.

References

Abragam A, Bleaney B (1970) Electron paramagnetic resonance of transition ions. Clarendon Press, Oxford

Bednarek J, Plonka A (1987) Single-crystal electron spin resonance studies on radiationproduced species in ice Ih. J Chern Soc Faraday Trans 1 83: 3725-3735; 3737-3747

Bershov LV, Gaite JM, Hafner SS, Rager H (1983) Electron paramagnetic resonance and ENDOR studies of Cr3 + -A13 + pairs in forsterite. Phys Chern Miner 9: 95-101

Calas G (1988) Electron paramagnetic resonance. In: Reviews in mineralogy. Spectroscopic methods in mineralogy and geology vol 18, Mineral Soc Am, Washington DC pp 5l3-571

Che M, Fraissard J, Vedrine JC (1974) Application of electron paramagnetic resonance and nuclear magnetic resonance to the study of silicates and clays. Bull Groupe Fr Argiles 26: 1-53

Dusausoy Y, Babkine J, Gaite JM, Hafner SS, Rager II (1990) Localisation par RPE des ions traces Fe3+ dans la structure disthene. Reunion de la societe Fran~aise de Mineralogie, Rennes, 3-8 Septembre

Graham WRM (1987) Recent progress in the study of metals in fossil fuel sources by EPR. In: Weil JA (ed) Electronic magnetic resonance of the solid state. Canadian Society for Chemistry, Ottawa pp 323-330

Griscom DL (1990) Electron spin resonance. Glass Sci Technol 4B: 151-251 Herve A (1985) La resonance paramagnetique electronique. In: Calas G (ed) Methodes d'etudes

spectroscopiques des mineraux. Soc Franc Mineralogie et Cristallographie, pp 3l3-389


Isoya J, Kanda H, Norris JR, Tang J, Bowman MK (1989) Fourier-transform and continuouswave EPR studies of nickel in synthetic diamond: site and spin multiplicity. Phys Rev B 41: 3905-3913

Kevan L, Narayana M (1983) Electron spin echo studies of the location and adsorbate interactions of paramagnetic metal species in zeolites. ACSSS 218 (Intrazeolite Chemistry) New York pp 283-299

Kliava Ya G (1988) EPR spectroscopy of disordered solids. Zinatne, Riga Krasnobaev AA, Votyakov SL, Krokhalev V Ya (1988). Spectroscopic properties of zircon and

its geological applications. Nauka, Moscow Lehmann G (1979,1980) Correlation of zero-field splittings and site distortions. Phys State Sol

(B) 92: 417-424; 99: 623-633 Lloyd RV, Lumsden DN, Gregg JM (1985) Relationship between paleotemperatures of

metamorphic dolomites and ESR-determined Mn(II) partitioning ratios. Geochim Cosmochim Acta 49: 2565-2568

Low W (1968) Electron spin resonance - a tool in mineralogy and geology. Adv Electronics Electron Phys 24: 51-108

Marfunin AS (1979a) Physics of minerals and inorganic compounds. Springer, Berlin Heidelberg New York. Also: (1979b) Spectroscopy, luminescence and radiation centers in minerals. Springer, Berlin Heidelberg New York

Matyash IV, Litovchenko AS, Bagmut NN, Proshko V Ya (1981) Radiospectroscopy of feldspars. Naukova Dumka, Kiev

McKinney TM and Goldberg IB (1989) Electron spin resonance. In: Rossiter BW and Hamilton JF (eds) Physical methods of chemistry. Volume 3B, Chap 4 (2nd ed). John Wiley, New York

McWhinnie WR (1985) Electron spin resonance and nuclear magnetic resonance applied to minerals. In: Berry FJ and Vaughan DJ (eds) Chern. Bonding Spectrosc Miner Chern Chapman and Hall, London, pp 209-249

Michoulier J, Gaite JM (1972) Site assignment of Fe3+ in low-symmetry crystals. Application to NaAlSi30 s. J Chern Phys 56: 5205-5213

Morton JR, Preston KF (1987) Landolt-Bornstein numerical data and functional relationships in science and technology. In: Fischer H (ed) New Series, Group II, vol 17a. Springer, Berlin Heidelberg New York, pp 577

Newman DJ, Urban W (1975) Interpretation ofS-state ion EPR spectra. Adv Phys 24: 793-844 Newton ME, Baker JM (1989) 14N ENDOR of the OKI centre in natural type Ib diamond.

J Phys Condens Matter 1: 10549-10561 Petrov I, Agel A, Hafner SS (1989) Distinct defect centers at oxygen positions, in albite. Am

Mineral 74: 1130-1141 Ruck R, Dusausoy Y, Gaite JM (1988) Electron paramagnetic resonance of a new Fe3 + centre

in cassiterite. Bull Mineral 111: 143-147 Ruck R, Dusausoy Y, Nguyen Trung C, Gaite JM, Murciego A (1989) Powder EPR study of

natural cassiterites and synthetic Sn02 doped with Fe, Ti, Na and Nb. Eur J Mineral 1: 343-352

Sherbakova MYa (1981) Electron and hole centers in scheelite crystals according to EPR data. Molecular spectroscopy and x-ray analysis of minerals. Nauka, Novosibirsk pp 87-128

Sherbakova MYa, Distanova AN, Teleshev AV, Dovgel VN, Minin NA, Radionova RB (1985) EPR investigation of quartz from various types of granitoids. Nauka, Novosibirsk Geol Geophys 9: 89-96

Speit B, Lehmann G (1982) Radiation defects in feldspars. Phys Chern Mineral 8: 77-82 Solntsev VP (1981) Nature of color centers and EPR in beryl and chrysoberyl. Problems of

theoretical and genetic mineralogy. Nauka, Novosibirsk pp 92-140 Stevens KWH (1952) Matrix elements and operator equivalents connected with the magnetic

properties of rare-earth ions. Proc Phys Soc Lond A65: 209-215 Taylor PC, Baugher JF, Kriz HM (1975) Magnetic resonance spectra in polycrystalline solids.

Chern Rev 75: 203-240 Vassilikou-Dova AB, Lehmann G (1987) Investigations of minerals by electron paramagnetic

resonance. Fortschr Mineral 65: 173-202

3.8.2 Electron Nac1ear Double and Multiple Resonance 209

Wei! JA, Buch T, Clapp JE (1973) Crystal point group symmetry and microscopic tensor properties in magnetic resonance spectroscopy. Adv Magn Reson 6: 183-257

Wei! JA (1984) A review of electron spin spectroscopy and its application to the study of paramagnetic defects in crystalline quartz. Phys Chern Mineral 10: 149-165

Wei! JA (1993) A review of EPR spectroscopy of the point defects in !X-quartz: the decade 1982-1992. In: Helms CR, Deal, BE (ed) The physics and chemistry ofSi02 and the Si-Si02

interface. 2. Plenum Press, New York, pp 131-144

3.8.2 Electron Nuclear Double and Multiple Resonance

l.R. NIKLAS, A.B. BRICK, and I.-M. SPAETH

Electron paramagnetic resonance (EPR) is a well established tool for the characterization and determination of the microscopic structure of paramagnetic defects in solids. There are, however, limitations of its usefulness arising from its limited power to resolve ligand hyperfine interactions which provide, for example, valuable information about the lattice site of the defect. Another problem is, that particularly in natural crystals such as minerals, different types of defects with overlapping EPR spectra are present simultaneously, making an analysis often very difficult if not impossible.

These problems can be overcome by observing nuclear magnetic resonance (NMR) transitions of the different neighbor nuclei of the defect rather than the hyperfine EPR transitions. Because of sensitivity problems the NMR transitions cannot be measured directly but they can be detected indirectly by observing intensity changes of simultaneously induced EPR transitions under special experimental conditions. Thus the high resolution of NMR is combined with a better sensitivity gained by this quantum transformation. This technique is called electron nuclear double resonance (ENDOR). In a very simple first order treatment of the spin Hamiltonian, ENDOR lines are observed for nuclei with nuclear spin I = 1/2 at frequencies fENDOR:

(1)

where ms is the electronic spin quantum number, Wshf the ligand hyperfine interaction energy of the nucleus corresponding to the ENDOR line, gI the nuclear g-value of this nucleus, JiK the nuclear magnet on, and Bo the static magnetic field. The electron spin of the defect can be inferred from the number of ENDOR lines for each neighbor nucleus due to the different values ofms' and the chemical nature of the nucleus giving rise to fENDoR can be determined from gI by measuring the dependence of fENDoR on Bo. For nuclear spins I > 1/2 a quadrupole interaction [not shown in Eq. (1)] causes splittings of the ENDOR lines and yields additional information on electric field gradients. The symmetry of neighbour shells and thus the structure of the defect follows from a detailed analysis of the dependence of fENDoR on the direction of the magnetic field Bo

210

Y1

X


Fig.78. ENDOR spectrum of defects in sodium beta alumina generated by

X ionising radiation. The labels X, YI, Y2, and Z indicate lines for different defect species. (After Barklie et al. 1980)

o 10 20 30

a:: If) UJ

a

a:: If) UJ

I ...... UJ

b

a:: If) UJ

I ...... UJ

c 320

Endor frequency (MHz)

340 360

Magnetic field ( m T)

Fig. 79. a Integrated EPR spectrum of defects generated by ionizing radiation in KMgF 3 doped with Fe3 +. b Part of the Fe3 + EPR spectrum as measured by ENDOR-induced EPR. c F-center EI-EPR spectrum. (After DuVarney et al. 1980)

relative to the crystal lattice. One example of an ENDOR spectrum is shown in Fig. 78 for defects in sodium beta alumina generated by ionizing radiation. The lines are unusually broad due to lattice imperfections. The different labels X, YI, Y2, and Z indicate ENDOR lines belonging to different types of defects. The different defects can be distinguished by using the ENDOR-analog to excitation spectroscopy in optical luminescence. By measuring the intensity of an ENDOR line as a function of the magnetic field Bo (formally as in an EPR experiment), one obtains exclusively the EPR (absorption-) spectrum of the defect species giving rise to this ENDOR line. This is called ENDOR-induced EPR (EI-EPR). An example is shown in Fig. 79 for defects produced by X-irradiation at 300 K in KMgF 3 doped with Fe3 +. Figure 79a shows the (integrated) conventional

3.8.3 EPR: Improvement of Experimental Technique 211

EPR spectrum, Fig. 79b part of the Fe3+ EPR spectrum obtained by EI-EPR using an Fe3+ ENDOR line for the measurement, and Fig.79c shows the F-center EI-EPR spectrum using an F-center ENDOR line. Both EPR spectra are superimposed, the F-center EPR spectrum is not even visible in conventional EPR (Fig. 79a). Overlapping ENDOR spectra due to different types of defects can be separated by electron nuclear triple resonance (double ENDOR). Also spatial resolution of ENDOR spectra is possible. The indirect detection ofNMR can be carried even further by optical detection of EPR (ODEPR) which in turn is used to detect NMR. This optical detection of ENDOR (ODENDOR) opens the possibility to link optical spectra of defects to their structure as obtained from the ENDOR analysis.

References

Abragam A, Bleaney B (1970) Electron paramagnetic resonance of transition ions, Calrendon Press, Oxford

Barklie RC, Niklas JR, J-M Spaeth (1980) J Phys C Sol State Phys 13: 1745 Bauer RU, Niklas JR, Spaeth J-M (1983) Phys State Sol b 119: 171 DuVarney RC, Niklas JR, Spaeth J-M (1980) phys state sol 97: 135 Feher G (1959) Phys Rev 114: 1219; 1249 Ishchenko SS, Brik AB (1987) Fiziol Tverd Tela Leningrad 29: 3481 Kevan C, Kispert W (1976) Electron spin double resonance spectroscopy. Wiley and Sons, N.Y Niklas JR, Spaeth J-M (1980) Phys. state sol bIOI: 221 Seidel H, Wolf HC (1968) In: Fowler WB (ed) Physics of color centers. Academic Press, N.Y. Spaeth JM (1988) Experimentelle Technik der Physik 36: 257 Spaeth JM (1989) Spec Period Rep 11B: 89 Spaeth JM (1990) In: Rossiter BW, Hamilton JF (eds) Physical methods in chemistry, vol 5,

Chap 6 Spaeth JM, Lohse F (1990) J Phys Chern Sol 51: 861 Spaeth JM, Koschnick F (1991) J Phys Chern Sol (in press) Spaeth JM, Niklas JR, Bartram RH (1991) Multiple magnetic resonance spectroscopy for

structural analysis of point defects in solids. Springer, Berlin Heidelberg New York

3.8.3 EPR: Improvement of Experimental Technique

Y A.S. LEBEDEV

During the past decade (1980-1990), both the philosophy and the technique of EPR experiments have been essentially revised in view of their new applications. These achievements were based partly on entirely new physical ideas (such as the use of spin-selected reactions for indirect detection of EPR), or on new technical inventions (such as the construction of very small "personal" spectrometers). However, most of the recent achievements have combined new physical ideas with new technical possibilities, resulting in the development of the pulsed EPR technique, multifrequency EPR, EPR imaging, and other techniques.


In the modern pulsed EPR technique the most important role belongs to the new type of EPR cavities -loop gap resonators, which have a higher concentration of microwave energy and a shorter time response. With such cavities and modern electronics and signal processing, an impressive set of new pulsed EPR experiments became possible, such as Fourier-transform EPR, pulsed electronnuclear double resonance in solids, longitudinal detection ofEPR and EPR with vector jumps of the magnetic field and, two-dimensional EPR spectroscopy. The new pulsed methods permit the detection of short-lived centers (10- 9-10- 8 s), and for stationary centers they greatly increase the volume of information obtained in experiments, as direct measurement of relaxation times T 1 and T 2,

and of small hyperfine and dipole-dipole interactions has become possible. The next significant improvement in EPR technique is connected with the

development of multifrequency EPR, especially of EPR at very high frequencies (VHF). For more than 30 years, practical application of EPR was carried out only at the 10 GHz frequency range. Later, however it was proved that an increase in detection frequency up to 100-150 GHz is followed by a definite increase in spectral resolution and sensitivity. For solid state samples it is significant that VHF-EPR spectra of power or amorphous materials may provide a volume of information as large as single crystal studies. Comprehensive analysis of spectra recorded at very different frequencies (from VHF to 0.1-1.0 GHz) leads to a further improvement in the accuracy of structure identification and investigation. Sometimes VHF-EPR may be advantageous as a possibility to use extremely small probes, from 1-10 picogramms.

For applications in material science and mineralogy, the most promising appear to be EPR imaging, or EPR tomography. Spectra detection in a nonuniform magnetic field with known configuration makes EPR frequency dependent on spatial coordinates. After proper computer processing (reconstruction of projections or similar procedures), the spatial distribution of paramagnetic centers within the sample may be restored with linear resolution up to 10-100 Jl in practice and up to 1-10 Jl in theory. Similar approaches make it possible to focus the EPR effect on some small part of a sample, giving rise to the technique of EPR-microscopy.

This indirect-detected EPR uses a more or less standard technique for the excitation of EPR transitions, but detects the thus induced change in optical radiation, luminescence, photoconductivity, or rate of chemical reaction. Indirect detection of EPR excels by its unique sensitivity, spectra being sometimes detected from only 10-100 centers in a sample. As a comparison the ordinary 10 GHz baRd EPR sensitivity is 101°_1011 (centers per 1 G of line width), while for VHF-EPR (150 GHz) it is 107-108 •

In addition to these improvements in EPR technique, which greatly broaden its field of application, there has also been a tendency to make ordinary EPR spectrometers more compact and flexible. Special compact magnets and modern personal computers were used to construct special "personal" EPR spectrometers small in size and weight, having at the same time good sensitivity and spectral resolution. Being relatively cheap, and also transportable, these

3.9 Nuclear Magnetic Resonance (NMR) Spectroscopy 213

"personal" spectrometers may be used not only in stationary laboratories but also for out-of-building control, geological expeditions, etc.

References

Bowman MK (1990) Fourier transform electron spin resonance. In: Kevan L, Bowman MK (eds) Modern pulsed and continuous-wave electron spin resonance

Froncisz W, Hyde JS (1980) The loop-gap resonator: a new microwave lumped circuit ESR sample structure. J Magn Reson 47: 515-521

Gorcester J, Millhauser GL, Freed JH (1990) Two-dimensional electron spin resonance. Ibid, pp 119-194

Grupp A, Mehring M (1990) Pulsed ENDOR spectroscopy in solids. Ibid, pp 195-230 Hoch MJR (1981) Electron spin resonance imaging of paramagnetic centers in solids. Sol State

Phys 14: 5659-5666 Lebedev YaS (1990) High-frequency continuous-wave electron spin resonance. Ibid,

pp 365-404 Schweiger A (1990) New trends in pulsed electron spin resonance methodology. Ibid,

pp 43-118 Yakimenko OYe, Smirnov AI, Lebedev YaS (1990) EPR imaging of structurally heterogenious

media. Appl Magn Reson 1: 1-19

3.9 Nuclear Magnetic Resonance (NMR) Spectroscopy

R.J. KIRKPATRICK

Stages of Development; Wide-Line and High-Resolution MAS NMR

NMR spectroscopy provides information about the structure and dynamic behavior of atoms and molecules in solids, liquids, and gases. The NMR phenomenon was discovered in 1946, following Electron Paramagnetic Reson- \ ance in 1944, and preceding Nuclear Quadrupole Resonance in 1949. Between 1951 and 1952, when the first work on natural minerals was done, and about 1980, only wide-line (continuous wave, CW) NMR methods were generally available for solids. In this technique, the magnetic field is continuously varied and the absorption of energy by atomic nuclei is observed. For many gases and liquids, in which rapid molecular tumbling causes averaging of line-broadening interactions, these techniques provide quite high-resolution spectra, and NMR quickly became the method of choice to investigate the structure of organic molecules and other solution species. For solids, however, such line narrowing does not occur, and the peaks in CW spectra are typically very broad.

Despite this limitation, CW NMR methods can provide much important insight into the structure and dynamical behavior of solids. Of particular mineralogical interest are results for electric field gradients at 27 Al in many aluminosilicates (e.g., Ghose and Tsang 1973), order-disorder in, e.g., feldspars, spinels, and ambligonite, B-coordination in crystalline and glassy borates and


borosilicates, and the structure and dynamical behavior of protons in clay minerals, zeolites, and other hydrous minerals (Fripiat 1980).

In about 1980, the simultaneous development or wide-spread use of highfield superconducting magnets, the pulse-Fourier-transform method of data acquisition, and magic-angle spinning (MAS) allowed greatly increased spectral resolution and signal/noise ratios. Since then there has been an explosive growth in the application of MAS and pulse-Fourier-transform NMR methods in solidstate science with many hundreds of published papers. Computer automated spectrometers that are relatively easy to operate are now widely available, and NMR is on its way to becoming a standard tool in mineralogy.

This revolution in instrumentation provided the opportunity for several developments of direct interest to mineralogy.(1} It became possible for the first time to directly investigate the local (nearest -neighbor, NN, and next-nearest neighbor, NNN) structure of disordered crystals and amorphous phases through the chemical shift and to compare these results with those for average structures obtained by diffraction methods. (2) It became possible to obtain reasonably high-resolution spectra of nuclei with a quadrupole moment, greatly increasing the number of nuclei that can be usefully probed (Table 9). (3) It became possible to obtain useful, and in many cases high-resolution, spectra at both low and high temperatures (4 K to 1200 0c) for materials of mineralogical interest, allowing investigation of many structural phase transitions and even high-temperature melts. (4) It became worthwhile to undertake quantum chemical calculations of NMR chemical shifts for solids, and such calculations are now underway.

Over the past decade a large data base for the MAS and pulse-Fouriertransform NMR behavior of minerals has been accumulated, beginning with 29Si in zeolites and other silicates and aluminosilicates. Some data are now available for representatives of most groups of silicate and aluminosilicate minerals, especially those that show AI, Si disorder on the tetrahedral sites, such as feldspars and micas. In addition, there is some work on borates, phosphates, and carbonates, and a large body of data for glasses of mineralogical interest.

High-resolution NMR spectroscopy of solids is still in a period of rapid growth with respect to spectroscopic methods, technology, and applications. In the future there will be expanded use of techniques and empirical correlations already developed to investigate questions of mineralogical interest, increased quantum chemical understanding of the NMR behavior of minerals (and therefore deeper understanding of the observations), and improved technology (including higher magnetic fields and such techniques as dynamic angle spinning).

Principles and Innovations

The NMR experiment investigates phenomena related to transitions between nuclear spin energy levels for which the degeneracy has been lifted by an external


Table 9. Some nuclides of potential use in NMR studies of solids

Nucleus Readily Spin Quadrupole Natural Frequency observed moment abundance MHz

(l0- 24cm2) (11.7 T)

H-1 Yes 1/2 99.985 500 H-2 Yes 1 0.0028 0.015 76.8 Li-7 Yes 3/2 -0.03 92.58 194.3 Be-9 Yes 3/2 0.0512 100 70.3 B-1O Yes 3 0.074 19.58 53.7 8-11 Yes 3/2 0.0355 80.42 160.4 C-13 Yes 1/2 1.1 125.7 N-14 Yes 1 0.016 99.6 36.1 N-15 Yes 1/2 0.37 50.7 0-17 Yes 5/2 - 0.026 0.037 67.8 F-19 Yes 1/2 100 470.4 Na-23 Yes 3/2 0.14 100 132.3 Mg-25 Yes 5/2 N.D. 10.1 30.6 AI-27 Yes 5/2 0.149 100 130.3 Si-29 Yes 1/2 4.7 99.3 P-31 Yes 1/2 100 202.4 S-33 No 3/2 - 0.064 0.76 38.4 CI-35 Yes 3/2 - 0.0789 75.5 49.0 K-39 Yes 3/2 0.11 93.1 23.3 Sc-45 Yes 7/2 - 0.22 100 121.5 Ti-49 No 7/2 N.D. 5.5 28.2 V-51 Yes 7/2 -0.04 99.76 131.4 Cu-63 Yes 3/2 0.16 69.1 132.5 Zn-67 Yes 5/2 0.15 4.1 31.3 Ga-71 Yes 3/2 0.112 39.6 152.5 Ge-73 Yes 9/2 -0.2 7.8 17.4 Se-77 No 1/2 7.6 95.3 Br-79 Yes 3/2 0.33 50.5 125.3 Rb-85 Yes 5/2 0.27 71.25 48.3 Sr-87 No 9/2 0.2 7.0 21.7 Y-89 Yes 1/2 100 24.5 Zr-91 No 5/2 N.D. 11.2 46.7 Nb-93 Yes 9/2 -0.2 100 122.2 Mo-95 Yes 5/2 0.12 15.7 32.6 Ag-109 Yes 1/2 48.18 23.3 Cd-ll3 Yes 1/2 12.26 110.9 In-115 Yes 9/2 1.14 95.72 109.6 Sn-119 Yes 1/2 8.58 186.4 Te-125 No 1/2 6.99 158.0 Cs-133 Yes 7/2 -0.003 100 65.6 Ba-137 No 3/2 0.2 11.3 55.6 La-139 Yes 7/2 0.21 99.9 70.6 Yb-l71 No 1/2 14.3 88.1 W-183 Yes 1/2 14.4 20.8 Pt-195 No 1/2 33.8 107.5 Hg-199 Yes 1/2 16.8 89.1 T1-205 Yes 1/2 70.5 288.5 Pb-207 Yes 1/2 22.6 104.6


magnetic field. Thus, nuclides with nonzero nuclear spins (Table 49) can in principle be examined. Transitions between these nuclear spin energy levels correspond to the radio-frequency range. For instance at the relatively large magnetic field of 11.7 Tesla (11 700 Gauss), the frequency for 29Si is 99.3 Mhz, the frequency for 23Na is 132.3 MHz, and that for 1H is SOO Mhz.

This frequency, v, is given by the following relationship

where h is Planck's constant, H is the magnetic field at the nucleus, Pn is the nuclear magneton (the elementary value of the magnetic moment of nuclei), and gn is the nuclear g-factor ( = JlI; where Jl is the magnetic moment of the nuclide of interest and I is the spin quantum number for the nuclide of interest).

The classical (nonquantum mechanical) description of NMR considers the precession of the new nuclear magnetic moment of all the individual nuclei of a given nuclide in a sample around the direction of the applied magnetic field. The frequency of precession, the Larmor frequency, is equal to the frequency of the radiation necessary to cause spin sublevel transitions.

The NMR behavior of atomic nuclei under these conditions is analogous to the EPR behavior of electrons, which undergo transitions between electron spin sublevels. The nuclear magneton, Pn, however, is a factor of 1836 less than the Bohr magneton, p, for the electron. Thus, NMR frequencies are of the order of 103 less than EPR frequencies.

The NMR frequencies are orders of magnitude less than those of nuclear gamma resonance (Mossbauer) spectroscopy, because NMR observes transitions between spin sublevels of the nuclear ground state arising from the applied magnetic field, whereas NGR observes transitions between the nuclear ground and excited states without a magnetic field necessarily present.

Nuclides can be divided into three types with respect to their spin quantum number and NMR behavior. Those with spin I = 0 have no magnetic moment and yield no NMR signal. Those with I = 1/2 have two spin energy levels and behave as magnetic dipoles. These are of great interest to NMR of solids, because MAS produces quite narrow peaks for them. Examples include 1 H, 29Si, and 31 P. Those with I > 1/2 have both a dipole moment and a quadrupole moment. Interaction of this quadrupole moment with the electric field gradient at the nucleus normally causes significant peak broadening even under MAS. Using powerful superconducting magnets and high power spectrometers, however, it is now possible to obtain quite good MAS and static Fourier-transform spectra of these nuclides, which are much more common than I = 1/2 nuclides. Examples include 170, 23 Na, 27 AI, 133CS, and many others: Table 49).

One of the most important reasons that NMR spectroscopy is chemically/structurally useful is that the magnetic field at the nucleus, H, is typically shielded slightly from the applied magnetic field, Ho, by the electrons in the vicinity of the nucleus. This shielding is a tensor property and is given by the relationship

H = Ho(1 - u),


where (Jis the shielding tensor. This shielding gives rise to the chemical shift, which is the parts per million (ppm) fractional difference of the resonance frequency of the observed nucleus from that of an experimentally useful standard,

This chemical shift is extraordinarily sensitive to differences in local structure and bonding arrangements and is the main NMR parameter that has been used in the past decade to investigate minerals. It is sometimes also possible to determine the entire chemical shielding tensor (described by the chemical shift anisotropy, CSA).

The CSA and several other nuclear interactions, including the interaction of the dipole moments of the nuclei, cause the peak width for solids to be very broad under ordinary conditions. The Hamiltonians that describe these interactions, however, all contain terms involving the factor (3cos28 - 1). This term goes to zero if 8 = 54.7°. Because of the behavior of this term, if the sample is physically spun on an axis at 54.7° to the orientation of the applied magnetic field (the "magic-angle"), the time-average value of these interactions goes to zero, and the broad peak collapses to one nearly as narrow as for solutions (Fig. 80). Typical peak widths for I = 1/2 nuclides in well-ordered solids are of the order of 1 ppm, and in addition there are the so-called spinning sidebands spaced at the spinning speed in frequency units. Magic-angle spinning, then, yields the narrowest peaks for sollids and is usually the technique of choice for spin I = 1/2 nuclides and often for quadrupolar nuclides.

For nuclides with spin I> 1/2, interaction of the quadrupole moment of the nucleus, eq, with the electric field gradient (crystal field gradient) at the nucleus, eQ, gives rise to a Hamiltonian which is not fully averaged by MAS to second order. Thus, such nuclides in environments with a nonzero electric field gradient give rise to complicated peak shapes that depend on the spin I and the shape and magnitude of the electric field gradient. This interaction is described by the quadrupole coupling constant, e2qQ/h, a measure of the magnitude of the quadrupole interaction (directly proportional to the electric field gradient for a given nuclide) and the asymmetry parameter, '1, a measure ofthe deviation of the electric field gradient from cylindrical symmetry. Both of these parameters can be determined from continuous wave and pulse-Fourier-transform NMR.

For minerals, the NMR chemical shift most often provides information about the static, local structure of the material examined. That is, (1) different chemical shifts are due to different structures and compositions in the NN and NNN environment only, and (2) the structure observed is the time-average over all positions occupied by the nucleus due to normal lattice vibrations (i.e., the static structure). This occurs because the frequencies of these vibrations are orders of magnitude greater than any chemical shift differences they cause.

It is also possible, however, to obtain significant information about the dynamic behavior of a material through NMR measurements. Under the right circumstances, this dynamical information can be obtained via measurement of various NMR relaxation times or through line shape analysis.

218

A

B

c

ppm


Fig. 80 A-C. Comparison of 29Si static (C) and MAS (A, 8) NMR spectra of synthetic Ca3SiOs. The static spectrum in C and the MAS spectrum in 8 are on the same scale, and the MAS spectrum in A is on an expanded scale. Note the greatly increased resolution in the MAS spectra compared to the static spectrum. This phase has nine crystallographically distinguishable Si sites and yields seven 29Si NMR peaks. The ones at - 69.1 and - 73.5 ppm are accidentally degenerate. The peak at - 71.4 ppm is from Si in impurity Ca2Si04 • (Courtesy of X.-D. Cong)

There are two basic types of NMR relaxation times. The spin-lattice (longitudinal) relaxation time, T 2, is a measure of the rate of return of an excited nuclear spin system (all individual nuclei of a given nuclide) to equilibrium. The spin-spin (transverse) relaxation time, T l' is a measure of the rate ofloss of phase coherency in the nuclear spin system. T 1 measurements have been the most common for inorganic solids. For I = 1/2 nuclides, T 1 relaxation can only occur if there are fluctuations in the magnetic field at the nucleus occurring at the Larmor frequency. It appears that for most solids such relaxation occurs by interaction with unpaired electrons (paramagnetic centers), although there have been few systematic studies of minerals. For nuclides with I > 1/2, T 1 relaxation can also occur by fluctuations of the electric field gradient at the nucleus at the Larmor frequency. Because the power spectrum of the lattice vibrations of solids normally has a component at these relatively low frequencies, the T 1 values for quadrupolar nuclides in minerals are generally much less than those of I = 1/2


nuclides. Relaxation rate measurements are especially important for investigating structural phase transitions, because the frequencies of soft lattice modes near such transitions are often very low and the intensity of the vibrational power spectrum in the NMR frequency range is greatly enhanced near the transition temperature (Blinc 1981; Rigamonti 1984).

Line shape analysis can sometimes be useful for investigating atomic motion, because fluctuations in the chemical shift occurring at frequencies within about an order of magnitude of the differences in the chemical shifts between environments can have an interpretable effect on the peak shape. For instance, if a nuclide occurs in two environments (sites) with different chemical shifts, and if chemical exchange is occurring between those two sites, the observed spectrum varies greatly with the frequency of the exchange. At exchange frequencies much less than the peak separation there is no effect of the spectrum, at exchange frequencies of the order of the peak separation there is one broad peak, and at exchange frequencies much higher than the separation motional averaging causes only one peak with a chemical shift at the abundance-weighted averaged position. This phenomenon has provided information about such diverse issues as atomic motion in melts and the behavior of adsorbed cations on clays (Farnan and Stebbins 1990; Weiss et al. 1990, and references in both).

Useful introductions to NMR theory and practice are given by Abragam (1961), Schlicter (1978), Farrar and Becker (1971), Becker (1980), Fukushima and Roeder (1981), Harris (1983), Fyfe (1984), Akitt (1983), Gerstein and Dybowski (1985), Wilson (1987) and Sanders and Hunter (1987). A useful introduction to theory, experimental practice, and the results for silicates is provided by Engelhardt and Michel (1987). Oldfield and Kirkpatrick (1985), Kirkpatrick (1988), and Stebbins (1988) provide shorter introductions and summaries.

Summary of Applications: Nuclides and Minerals

Of the approximately 40 nuclides of potential interest in NMR studies of minerals, only a few have been widely used (see Engelhardt and Michel 1987; Kirkpatrick 1988; Stebbins 1988, for many examples and references to original papers).

1. Continuous wave NMR spectra have been obtained for 1 H, 2H, 7Li, 9Be, 11 B, 19F, and 27 Al in a variety of minerals, mostly before 1980. These results provided much useful information about electric field gradients in minerals and about AI-site occupancy.

2. 29Si has been the most important nuclide for MAS NMR studies of minerals. There is a large body of data for silica polymorph, zeolites, feldspars, clay minerals, micas, nepheline, cordie rite, leucite, opals, glasses, gels, and the full range of alumino silicate phases. The correlations between the 29Si NMR chemical shifts and tetrahedral polymerization, NNN site occupancies, cationoxygen bond strengths, and mean Si-O-Si bond angle per tetrahedron are now


relatively well understood. Tetrahedral and octahedral Si can be readily distinguished.

3. 27 Al MAS NMR can readily distinguish 4-,5-, and 6-coordinated Al and has provided significant information about site occupancies in many aluminosilicates, especially clays, zeolites, and glasses. Aluminophosphates have also been observed.

4. 23Na has been observed in many crystalline and amorphous Na-aluminosilicates and has provided information about especially the local symmetry of the Na-environments.

5. 31 P has been observed in many phosphate minerals and glasses. 6. II B has been observed in many borate and borosilicate crystals and

glasses. NMR readily distinguishes 3- and 4-coordinated B and can provide information about the bonding symmetry of the B03 triangle.

7. 170 is potentially one of the most useful nuclides in mineralogy, because oxygen is the most abundant anion in nature. Unfortunately, 170 has a natural abundance of only 0.037% and cannot be observed at this level in solids. There have been several studies of synthetic, isotopically enriched phases, and newly developed techniques of gas exchanging natural samples may allow more extensive use. For some materials it is possible to distinguish bridging from nonbridging oxygens and to obtain information about the O-site symmetries.

8. 13C is the workhorse nuclide of organic chemistry and has been used in studies of carbonate minerals, oil, coal, and soils.

9. 19F has been studied extensively by continuous wave methods but can be difficult to observe under MAS due to the large homonuclear dipole broadening and high resonance frequency. There have been some studies of silicates and glasses.

10. 133CS has been used to study the structure and dynamical behavior of adsorption sites on clays.

Specific Applications

NMR methods are capable of addressing a wide range of structural and dynamical questions concerning the behavior of solids. Applications to minerals include the following.

1. The nature of Si, Al order/disorder in alumino silicate minerals has been reconsidered over the last decade due to the data provided by, primarily, 29Si MAS NMR. Because this technique provides information about the NN and NNN structure, for many phases it is now possible to define the state of local Si, Al order on the tetrahedral sites and to compare it to the results from diffraction measurements. Diffraction methods yield only the average structure for the bulk sample, and provide information about site occupancies of Si, Al disordered phases only through such parameters as average cation-oxygen bond distances and diffraction peak intensities. NMR, on the other hand, can often provide direct, quantitative values for the fractions of Si-sites with different numbt:rs of


Al NNN. Calculations based on these data can then provide quantitative models for the nature of the disorder. Such work has been done for many minerals, including zeolites, feldspars, cordierite, micas, clay minerals, nepheline, leucite, and majorite garnet. Similarly, 27 Al MAS NMR has been used to determine tetrahedral/octahedral Al ratios in clay minerals and 31 P MAS NMR has been used to investigate site occupancies in phosphates.

One important result of this work is the recognition of the significance of the scale of observation on interpretations of structure and atomic order. To have a reasonably complete structural picture of a phase (excluding defects), it is necessary to know the average structure determined from diffraction measurements, any domain structure on the scale of tens to hundreds of as determined from TEM observation, and local order/disorder on the NN and NNN scale as determined by NMR or some other spectroscopic method such as EXAFS/ XANES or NGR.

2. Because of their importance as molecular sieves and catalysts, zeolites have been studied more extensively than any other group of minerals. There are more than 40 natural and 100 synthetic zeolites, including gallosilicates, aluminophosphates, and silicoaluminophosphates. Representatives of all types have been investigated by NMR. Different aspects of zeolites that can be investigated by NMR include the structure of the framework e7 AI, 29Si, and 170), the structure and dynamical behavior of exchangeable cations e3Na, 7Li, 2osTI, and 133CS) and hydrocarbon molecules e Hand 13C), water molecules and hydroxyl groups eH), and the nature of the pores e 29Xe).

3. Investigation of clay minerals has also been especially useful, because of the fine grain size, complex interlayers and potential tetrahedral and octahedral disorder of these phases. Interlayer water has been studied extensively by 1 H and 2H NMR (see Fripiat 1980 for an introduction) and interlayer cations by 27 Al and 133CS NMR. Tetrahedral disorder has been investigated by 27 Al and 29Si NMR and octahedral disorder by 21AI NMR.

4. Because detection of NMR signal does not depend on the presence of an ordered crystal structure, it has been used extensively to investigate the structure of glasses and other amorphous materials. Unfortunately, for most amorphous phases the NMR peaks are much broader than for crystals, due to static structural disorder. For alkali silicate glasses, it is sometimes possible to determine the fraction of 29Si in sites with different polymerizations (QD sites, where n is the number of bridging oxygens per tetrahedron, n = 0-4), but for alkaline earth silicate glasses and aluminosilicate glasses only a single, broad peak is normally observed. The peak maximum, however, varies systematically with composition. For peraluminous glasses and gels, 27 Al MAS NMR can distinguish 4-,5-, and 6-coordinated AI. For some glasses 170 MAS NMR can distinguish bridging and non bridging oxygens. NMR has also provided significant insight into the stability of potential nuclear waste glasses and the mechanisms of their reaction with water.

One of the most novel recent applications of NMR to glasses has been the observations of changes of the chemical shift and relaxation times for glasses


and melts with temperature through the glass transition (Farnan and Stebbins 1990, and references therein). This work has provided significant new information about the nature of the glass to melt transition and the mechanisms of viscous flow in alumino silicate melts.

5. NMR spectroscopy is a powerful tool to investigate crystalline phases undergoing structural phase transitions (SPTs). Most NMR work with SPTs has involved either measurements of changes in relaxation times or quadrupole coupling constants with temperature through the transition, and there is a very large literature on this subject (Blinc 1981; Rigamonti 1984). Recently, MAS NMR at elevated temperatures has shown that it is possible to characterize SPTs and the structure of the incommensurate phases that are often associated with SPTs via the chemical shift (Phillips et al. 1991), and additional studies are underway.

Recent Technical Developments

In addition to MAS, many other NMR spectroscopic techniques have been developed in recent years. The most important of these for mineralogy has been proton cross-polarization with MAS (1 H CPMAS, Sindorf and Maciel 1983; see Yannoni 1982, for an introduction). In this experiment the spin systems of both the protons and the nuclide to be observed (e.g., 29Si) are simultaneously excited and nuclear spin allowed to transfer from the protons to the observed nuclide. This experiment greatly enhances signal from individual observed nuclei near protons and allows discrimination among peaks from hydrous and anhydrous phases. It can also allow qualitative estimates of differences in, e.g., 1 H_29Si distances for different Si sites in the same phase.

Other recently developed techniques include the following.

VAS NMR CRAMPS GASPS SEFT SECSY SHECOR SEBBORD

: Variable angle spinning NMR : Combined rotational and multiple pulse spectroscopy : Gated spin echo pulse sequence : Spin echo Fourier transform : Spin echo correlation spectroscopy : Selective heteronuclear correlation spectroscopy : Spin echo broad band off resonance decoupling.

In addition, there are many newly developed two- and three-dimensional NMR techniques. These involve Fourier transformation over two or more variable time parameters in the data acquisition sequence and can provide information about, for instance, the connectivity of sites. Such techniques have been little used in mineralogy but have significant potential.

Quantum Chemical Calculation of NMR Parameters

To date, most applications of NMR spectroscopy to minerals has involved empirical interpretation of the spectra based on simple structural or chemical


concepts. For instance, for 29Si there are well established correlations between chemical shift and NN coordination, NNN environment (e.g., tetrahedral polymerization and the ligands to a Si-tetrahedron), cation-oxygen bond strengths, electronegativity, various bond angles, and to a lesser extent, distances.

However, the NMR chemical shift and chemical shift anisotropy are reflections of the local bonding environment, that is, the local distribution of electrons. For instance, the observed changes of the 29Si chemical shift of the Si02 polymorphs are not caused by the changes in mean Si-O-Si bond angle per tetrahedron, they are only correlated to them. The actual cause of the changes in chemical shift are the changes in electron distribution from structure to structure or site to site. Changes in electron distribution can be investigated via quantum chemical calculations, and it is possible to obtain the NMR parameters from such calculations. The chemical shifts must be calculated at a very high level, because both the electronic ground states and excited states must be included. The quadrupole coupling constant and the asymmetry parameter require only calculation of the ground state levels but are affected by longer range interactions than the chemical shifts. Calculations of chemical shifts for silicates have recently been undertaken by Tossell and Lazzeretti (TosseIl1991, and references therein) and are now providing insight into the origins of chemical shift variations in these materials. As computational capabilities increase, future calculations are likely to not only continue to provide such insight but also to enable the use of NMR parameters in understanding better the general issue of electron distribution in minerals.

References

Abragam A (1961) The principles of nuclear magnetism. Clarendon, Oxford, 599 pp Akitt JW (1983) NMR and chemistry, an introduction to the Fourier-transform multi-nuclear

era. 2nd ed, Chapman and Hall, London, 263 pp Becker ED (1980) High resolution NMR, theory and application. 2nd ed. Academic Press,

New York, 354 pp Blinc R (1981) Magnetic resonance and relaxation in structurally incommensurate systems.

Phys Rep 79: 331-398 Engelhardt G, Michel D (1987) High resolution NMR spectroscopy of silicates and zeolites.

Wiley, New York, 485 pp Farnan I, Stebbins JF (1990) High-temperature 29Si NMR investigation of solid and molten

silicates. J Am Chern Soc 112: 32-39 Farrar TC, Becker ED (1971) Pulse and Fourier transform NMR: introduction to theory and

methods. Academic Press, New York Fripiat JJ (1980) Applications of NMR to the study of clay minerals. In: Stucki JW, Banwart

WL (eds) Advanced chemical methods for soil and clay minerals research. NATO Adv Studies Inst Ser C, Vol C63. Reidel, Dordrecht

Fukushima E, Roeder SB (1981) Experimental pulse NMR, a nuts and bolts approach. Addison-Wesley, Reading MA, 519 pp

Fyfe CA (1984) Solid state NMR for chemists. CRC Press, Guelph, Ontario Gerstein BC, Dybowski CR (1985) Transient techniques in NMR of solids. Academic Press,

New York, 295 pp Ghose S, Tsang T (1973) Structural dependence of quadrupole coupling constant e2qQ/h for

27AI and crystal field parameter D for Fe+ 3 in aluminosilicates. Am Mineral 58: 748-755


Harris RK (1983) Nuclear magnetic resonance spectroscopy. Pitman Books, London Kirkpatrick RJ (1988) MAS NMR spectroscopy of minerals and glass. In: Hawthorne FG (ed)

Spectroscopic methods in mineralogy and geology. Reviews in mineralogy Vol 18, Min Soc Am, Washington, DC

Lippmaa E, Magi M, Sarno son A, Tarmak M, Engelhardt G, (1980) Structural studies of silicates by solid-state high-resolution 29Si NMR spectroscopy. J Am Chern Soc 103: 4889-4893

Lippmaa E, Magi M, Samoson A, Tarmak M, Engelhardt G (1981) Investigation of the structure of zeolites by solid-state high-resolution 29Si NMR spectroscopy. J Am Chern Soc 103: 4992-4996

Oldfield E, Kirkpatrick RJ (1985) High-resolution nuclear magnetic resonance of inorganic solids. Science 227: 1537-1544

Phillips BL, Kirkpatrick RJ, Thompson JG (1991) 29Si magic-angle spinning NMR spectroscopy of the ferroelastic to incommensurate transition in Sr2Si04. Phys Rev-B-Condemned Matter 43: 13280-13284

Rigamonti A (1984) NMR-NQR studies of structural phase transitions. Adv Phys 33: 115-191 Sanders JKM, Hunter BK (1987) Modern NMR spectroscopy. Oxford Univ Press, Oxford,

308 pp Schlicter CP (1978) Principles of magnetic resonance, 2nd edn. Springer, Berlin Heidelberg

New York SindorfDW Maciel GE (1983) 29Si NMR study of dehydrated/rehydrated silica gel using cross

polarization and magic-angle spinning. J Am Chern Soc 105: 1487-1493 Stebbins JF (1988) NMR spectroscopy and dynamic processes in mineralogy and geo

chemistry, In: Hawthorne FC (ed) Spectroscopic methods in mineralogy. Reviews in mineralogy Vol 18, Min Soc Am Washington DC

Tossell JA (1991) Calculation of the effect of deprotonation on the Si NMR shielding for the series Si(OH)4 to SiO:-. Phys Chern Mineral 17: 654-660

Weiss CA, Kirkpatrick RJ, Altaner SP (1990) Variations in interlayer cation sites of clay minerals as studied by I33CS MAS NMR spectroscopy. Am Mineral 75: 970-981

Wilson MA (1987) NMR Techniques and applications in geochmeistry and soil chemistry. Pergamon Press, Oxford

Yannoni CS (1982) High resolution NMR in solids: the CPMAS experiments. Accts Chern Res 15: 201-208

3.10 Nuclear Quadrupole Resonance (NQR)

IN. PENKovand D. BRINKMANN

NQR spectroscopy yields experimental access to the hyperfine interaction between the electric quadrupole moment eQ of nuclei with spin larger than 1/2 and the electric field gradient (EFG) tensor Vij present at the nuclear site. The transition frequencies lie in the range up to 1000 MHz. NQR can elucidate various problems concerning the composition, the real structure (electronic and crystalline) and other properties of condensed matter over a wide range of temperature and pressure.

The NQR parameters studied are: the quadrupole interaction constant, eQVzJh, the asymmetry parameter, ", and the spin-lattice and spin-spin relaxation times, T 1 and T 2, respectively.

NQR spectra have been obtained from the following nuclei lOB, 14N, 33S, 35,37CI, 63,65CU, 69,71Ga, 75As, 79,81Br, 113In, 121,123Sb, 1271, 201Hg, and 209Bi.

3.10 Nuclear Quadrupole Resonance (NQR) 225

The most important nuclei for mineralogical studies are As, Sb, Bi, and Cu. NQR studies have been performed in the following minerals:

7S As: arsenic As, domeykite CU3As, realgar AsS and its Se analog, dimorphite AS4S3, auripigment AS2S3 and its Se analog, smitite AgAsS2, lorandite TIAsS2, lautitev CuAsS, proustite Ag3AsS3, arsenolite and klodetite AS20 3; 121, 123Sb: antimony Sb, stibnite Sb2S3 and its analog, chalcostibite CuSbS2, miargyrite AgSbS2, skinnerite Cu3SbS3, bournonite CuPbSbS3, stephanite AgsSbS4, frankeite Sn3PbsSb2S14' senarmontite and valentinite Sb20 3, nadorite PbSb02CI, servantite SbSb04; 209 Bi: bismuth Bi, bismuthine Bi2S3, wittichenite CU3BiS3' bismite and sillenite Bi20 3, eulitine Bi12Si020; 63,6SCU: cuprite Cu20, tenorite Cu20 chalcosine Cu2S, kovelline CuS, klockmannite CuSe, chalkopyrite CuFeS2, cubanite CuFe2S3' bornite CuSFeS4,

delafossite CuFe02 and its analogs CuAI02 and CuGa02, skinnerite CU3SbS3' and wittichenite Cu3BiS3.

As a starting point in structural studies, the NQR method may be used to determine the number of crystallo-chemically nonequivalent atomic sites, especially in the case of supercells (for instance in chalco sine, klockmannite), and the space group symmetry.

The method is helpful in the study of impurity states and their spatial distributions. It has been established that in layer structures impurities are mainly localized in the Van der Waals space close to the coordinationally unsaturated atoms. In coordinational structures (e.g., arsenolite-senarmontite) the impurities are distributed over the basic atomic (molecular) sites and this causes a stabilization of the structure (senarmontite), a lowering of symmetry, and variations in their properties.

The electronic structure of typical coordination complexes of Cu(l), the role of valency and lattice contributions to the total EFG at the Cu nuclei has been determined on the basis of their NQR spectra and calculations of the EFG. Polarity and covalency bonds in pyramidal groups RS3 (R: As, Sb, Bi) in sulfides and sulfosalts have been estimated. The role of donor-acceptor bonds has also been shown.

In the field of technical and technological mineralogy, NQR investigations have been performed in dielectrics (auripigment, realgar, etc.), ferroelectric semiconductors (proustite, pyrargirite, stibnite), semiconductors (chalcosine, CuAI02, and its analogs), superconductors (covelline, klockmannite), magnetics (chalcopyrite, cubanite), materials for opto-electronic devices (bismite, sillenite, etc.). The NQR spectra and relaxation times are typical for the respective compounds.

Hitherto unknown phase transitions have been found in stibnite, smithite, skinerite, wittichenite, proustite, and pyrargirite. Other features of the lattice dynamics such as mobility of atoms, atomic groups, electrons, and other charge carriers have been studied. The possibility of investigating the chemosorption ability of minerals (stibnite, bismuthine, auripigment) has been demonstrated. It has been established that these minerals (especially stibnite) adsorb heavy metal


cations (CU2 +, Ag+, Pb2 +, etc.) which playa role as activators of flotation by the sulfhydril preparation. The maximum adsorption of cations is observed in the region of the phase transition (300 K for stibnite). The chemical mechanism which fixes the mineral particles to the flotoreagents has been discussed.

Genetic (or typomorphous) features of minerals are reflected in shifts of the NQR frequencies. The shifts are usually due to various irregularities in composition and structure.

The NQR method is an excellent instrument for quality control of synthetic analogs of minerals in new techniques. The properties of a substance can be predicted from the features of the NQR spectrum. For example, in chalcopyrite type minerals, NQR reveals the presence of a local magnetic field which is characteristic for anti-ferromagnetic structures. Ferroelectric activity in chalcogenides is monitored by the observation oflarge EFGs at the As and Sb nuclei due to the presence of RS3 trigonal groups with a stereochemically active pair of un-shared s electrons.

Since NQR spectra are highly specific for the various nuclei, the NQR method can be used for a fast qualitative and quantitative analysis of the polymineral ores.

NQR has played an important role in the study of the new class of hightemperature superconductors such as the YBa2CuXOy (x = 3 or 4; y = 6.5 - 7 or 8) compounds since NQR provides information on static and dynamic properties on an atomic scale. Topics which are investigated are temperature and pressure dependence of the NQR frequencies, determination of EFG and Knight shift tensors, and relaxation times. Especially Knight shift and relaxation studies are apt to understand the Cu spin dynamics and their role in the mechanism of superconductivity.

References

Abdullin RS, Kal'chev VP, Penkov IN (1987) Investigation of copper minerals by NQR: crystallochemistry, electron structure, lattice dynamics. Phys Chern Mineral 14: 258-263

Briukmann D, Mali M (1994) NMR-NQR Studies of high-temperature super conductors. In: NMR basic principles and progress, vol 31. Springer, Berlin Heidelberg New York

Brinkmann D (1992) Probing the electronic structure of Y-Ba-Cu-O superconductors by copper NQR/NMR. Z Naturforsch 47a

Lucken EAC (1969) Nuclear quadrupole coupling constants. Academic Press, London

3.11 Muon Resonance

3.11 Muon Resonance. Application to the Study of the Hydrogen Atom Position in Quartz

J.A. WElL

227

The positive muon, Mu +, can be obtained in beam lines by utilizing certain nuclear reactions. It tends to pick up an electron, e -, in chemical materials to form muonium, Muo, which is a light (ca. 1/9) isotopic form of atomic hydrogen, HO. It will react chemically in much the same way as HO, to form bonds with ordinary atoms. The muon (and hence also Muo and its compounds) is highly unstable, with a mean lifetime t1/2 of ca. 2.2 x 10-6 s. Thus work with muonic chemical and mineralogical systems is technically demanding, but quite feasible.

Since the muon has a nuclear spin I = 1/2 and an associated magnetic moment gllP,J (where gil = 2.002332 and magneton PIl = lelh/(2mll) = 4.485222 x 10- 26 JT- 1) magnetic resonance phenomena (e.g., NMR for diamagnetic muonium compounds, and EPR for paramagnetic ones) are feasible. However, utilizing detection of the positron e + emitted during decay of the muon, and the tendency for the e + to come off along the direction of muon spin (and magnetic moment), yields an even more sensitive tool. Internal and externally applied magnetic fields align the spin moments (muonic and/or electronic), cause Larmor precession(s), and affect the e + counting rates in specific directions. The muon spin rotation (jlSR) technique, elaborated in various ways for example by combining it with magnetic resonance technologies, has given valuable insights into the nature of species in gases, liquids, and solids.

Muonium formed in crystalline Si02 as well as in fused quartz was first reported in 1958, and by 1990 has been the subject of more than 30 publications. Si02 turns out to be a favorable medium for jlSR studies. In IX-quartz below 120K, Muo is trapped in complex cavities within the structure, in complete analogy with HO, as ascertained by measuring its anisotropic hyperfine coupling. The jlSR technique for the former and EPR for the latter are in complete agreement, for example in finding that both atoms are compressed, with their electron clouds shrunk inward by the surrounding ° and Si atoms as manifested by enhanced Fermi contact terms (compared to the free-space atoms).

Above 120K, Muo is detrapped and diffuses in the quartz crystal, primarily along the open c-axis channels. Here jlSR can follow the species, measure its motionally averaged hyperfine parameters, whereas (continuous-wave) EPR fails. The temperature dependence of the jl + spin-polarization relaxations gives information about the Muo motions. It is to be expected that muonium analogs of the various other hydrogenic centers known to occur in IX-quartz will be studied. Similar work in various other minerals should also be feasible.

In fused quartz, the glass structure influences the Muo hyperfine coupling in a random manner, causing relaxation of the jl + spin polarization. Surface sites and their physicochemical behavior have also been shown to be amenable to jlSR study.


In analogy with muonium, positronium (PSO = e + e -) also is detectable in quartz, decays with a half-life tI/2 of ca.1O- 9 s, by measuring the ')I-rays released. The dependence of tI/2 on the chemical surroundings gives valuable physical information, for instance on radiation damage and impurity centres present.

References

Brewer JH, Fleming DG, Spencer DP (1981) Jl+e- hyperfine interactions in quartz crystals. In: Kauffman EN, Shenoy GR (eds) Nuclear and electron spectroscopies applied to materials science. Elsevier, Amsterdam, pp 487-493

Chappert J, Grynszpan RI (1984) Muons and pions in materials research. North-Holland, Amsterdam, Chap 10

Cox SFJ, Symons MCR (1986) JlSR spectroscopy on free radicals: a complement to ESR spectroscopy. Hyperfine Interact 32: 689-706

Patterson BD (1988) Muonium states in semiconductors. Rev Mod Phys 60(1): 69-159

CHAPTER 4

Remote Sensing Methods: Visible, Infrared, and Microwave

230 Chapter 4. Remote Sensing Methods

Remote Sensing Methods: Visible, Infrared, and Microwave

B. CERVELLE

The basis of electromagnetic remote sensing is the recording by sensors placed on board aircraft or satellites of analogic or digital data, proportional to the intensity of an electromagnetic beam - visible, infrared, microwave ranges -reflected, emitted, or back scattered from the surface of the Earth or of a planet. Where the source of illuminating beam is the Sun, remote sensing is said to be passive and works in the visible and infrared ranges. An example of active remote sensing mode is given by a microwave "radar" beam - centimetric wavelengths - emitted from the vector, aircraft, or satellite, laterally to the tracks, and backscattered by the ground toward an antenna on board the same vector (Synthetic Aperture Radar SAR). Therefore, the detection of the returned echos and their analysis lead to deriving information about surface or subsurface properties.

Two main properties of the remote sensing devices characterize their efficiency: spatial and spectral resolutions. Because the number of photons coming from an object and reaching the detector is limited, the beam can be split either in many spectral bands, or in many pixels, but not both. Therefore, remote sensing sensor capabilities result from a trade-off between spatial and spectral resolutions. For example, the Earth observation satellite Spot has only three spectral channels (XS1, XS2, and XS3 "bands") but a ground resolution of 20 m (multispectral mode), or 10 m (panchromatic mode), whereas the Landsat Thematic Mapper (TM) has seven spectral channels for a ground resolution of 30m.

Therefore, remote sensing methods consist mainly in the processing of data acquired in each spectral bands to detect, classify, map, or identify objects thanks to their "spectral signatures".

Spectral Signatures of Mineral Objects in Remote Sensing

Spectral signatures are the radiometric manifestation of a target which may be a phenomenon or an object. In Earth sciences, the target can be the geochemical composition or the moisture content of a soil, the presence of a mineral at the surface, or the degree of alteration of a rock. For example, the spectral signature of the divalent Ni content (Fig. 81) in a serpentinite is the ratio R(710)/R(500), where R is the diffuse reflectance of this rock for the wavelength indicated (nm). This ratio characterizes the intensity of the absorption band due to the crystal field transition 3 A2 -+ 3T 1 (F) of Ni2 + in the tetrahedral sites of the mineral. The equivalent of this Ni spectral signature at the satellite level would be XS2jXSl for the Spot channels, and TM3/TMl for the Landsat bands.

Remote Sensing Methods 231

CONFIDENCE

98.4 %

6

2 3 4 5 6

N i 2+ content (weight %)

Fig. 81. Laboratory spectral signature of Ni-content, ratio between reflectance at 710 nm, and reflectance at 500 nm. Numbers refer to chemical analyses published by Maquet et al. (1981) for lateritic ores from New Caledonia

Spectral characteristics for remote sensing purposes have been determined by laboratory measurements. Unfortunately, most of the present Earth observation satellite have spectral bandwidth too broad (some 100 nm) to permit retrieving the basic spectral characteristics of the minerals as determined in the lab. On the one hand, satellite spectral data are qualitatively rather poor, on the other, they are quantitatively rich because the number of pixels observed is huge. Therefore, the methods to exploit these data are mainly statistical and their utilization depends on the spectral region for the optical (visible and infrared) spectrum, and on the physical characteristics of the microwave beam in SAR remote sensing.

Visible and Infrared Remote Sensing. The optical spectrum, i.e., visible (VIS) and infrared (IR), contains several regions where the gas of the atmosphere absorbs the Sun's radiations, for example between 1.3 and 1.5 JLm, and between 1.8 and 2.0 JLm. Therefore, it is necessary to distinguish different domains on both sides of each opaque atmospheric filter (Fig. 82).

- From 0.4 to 1.1 JLm (VIS and near infrared - NIR), diffuse reflectance of minerals, inversely proportional to the absorption, is chemically controlled by the contents of first-row transition elements. On the one hand, charge transfers occurs in the UV, between two atoms or ions, for example between Fe3 + and oxygen, or between Fe3 + and Fe2 +. The intense absorption bands they generate extend into the VIS. On the other hand, crystal field absorption is responsible


ATMOSPHERIC lRANSMISSION EMl1TANCE

~ 100 \

~ ~~

g

~

~~~ __ ~ ____ ~~~~~~~~ __ ~~~~~ __ ~~o 0.2 0.4 0.6 0.8 1.0 2.0 I 4.0 6.0 8.0 10.0 20.0

1 ~ 1 N:A,: Wrn[ MWm I". 1 Fig. 82. Out-of-atmosphere solar irradiance spectrum (dashed line) and ground black-body (300 K) emittance spectrum (dotted line) compared with the transmission spectrum of the Earth's atmosphere. (After Schott 1989)

for most of the reflectance maxima in transparent minerals, especially the silicates. The spectral signatures depend on the nature of the transition elements, their coordination (number and symmetry), and therefore the metal-ligand distances.

Since iron, with its different valencies, represents practically 98% of the metals at the Earth's surface, Fe2 + and Fe3 + are the cations mainly responsible for the VIS-NIR spectral characteristics of the minerals pixels observed by remote sensing.

- From 1.1 to 2.5 Jim, i,e., the short wave infrared domain (SWIR), the atmosphere has two opaque regions centered around 1.4 and 1.9 Jim, due to the absorption bands of water. However, the vibrations of the OH ions between 2.0 and 2.5 Jim have a strong influence on the spectral behavior of hydroxyl minerals: therefore, the phyllosilicates have a low reflectance in this domain by comparison with their reflectance in the atmospheric window around 1.6 Jim. The ratio R(1.6 Jim)/R(2.2 Jim) allows the characterization of the outcropping altered rocks, the mother rocks having a much higher reflectance. The channel Landsat TM 7 (2.08-2.35 Jim) has been chosen for this purpose.

- From 3 to 5 Jim, called the middle wave infrared (MWIR), the light reaching the satellized sensors has two components: (1) one part is the reflexion of the sunlight by the surface; (2) the other part, due to the thermal agitation, is emitted by the surface (Fig. 82).

Simultaneous phenomena make difficult the interpretation of data recorded during the day for geological purposes. On the other hand, the images recorded at night can display forest fires, hydrocarbure flames, lava flows etc.


- From 5 to 8 JIm, the atmosphere absorbs totally the sun light. Civil satellites have no channel in this domain.

- From 8 to 14 J1.m, called the thermal infrared (TIR), many meteorological satellites (Meteosat, GOES 1-5, NOAA 6-7, Nimbus 7), and the Earth observation satellite Landsat TM (channel 6), have the capability to detect the emittance of the ground. Phenomena involved in this TIR domain of remote sensing are relatively different from what happens in the VIS-NIR and SWIR regions.

Thermal Infrared Remote Sensing. In the thermal region of the optical spectrum, i.e., from 8 to 14 JIm, the detected signal, emitted by the ground illuminated by the sun, is mainly a function of the temperature of the ground. Where the reflectances in the VIS to SWIR are only dependent on the surface, emissivities in the TIR are influenced by the volume of the rocks, so that mineral objects which may have the same reflectance in the VIS can be differentiated through their different emittances, because the thermal inertia is a function of physical properties in volume: density, porosity, and hydro scopic state.

Therefore, remote sensed thermal data allow the differentiation of limestone, dolomite, granite, and volcanic rocks. From a mineralogical point of view, Si04

tetrahedras have their fundamental vibrations in these spectral regions, and the distance between these tetrahedras influences the location of the absorption bands. For example, quartz has a minimal emittance at 8.5 JIm, whereas it is at 10 J1.m for olivine. Recent spectroscopic experiments in the 8-12 J1.m have made it possible to develop a model for the lithological identification of igne rocks. Based on the crystal chemistry of silicates minerals, the model utilizes data recorded by remote sensing systems with broad bands, which can be traded off against the requirement of high spatial resolution for lithological identification.

The interpretation of the Heat Capacity Mapping Mission (HCMM) data has shown that the soils and even the vegetation do not mask completely the underlying rocks, which is a great advantage for the observation of the temperate climate zones. HCMM data have allowed the study of many phenomena that exhibit different levels of thermal capacity: differentiation between dry and moist soils over the daily solar heating cycle, mapping of the thermal anomalies associated with turbidity of waters, etc. Therefore, although the spatial resolution ofTIR sensors cannot be as high as that ofVIS-SWIR sensors, TIR remote sensing data are complementary to those obtained in the other parts of the optical spectrum. This is also the case for the microwave remote sensing methods.

Microwave Remote Sensing Methods for Mineral Surfaces. Microwave remote sensing with SAR sensor of Seas at, SIR-A, B, C, or with ERS-l and Radarsat, works in an active mode, the beam illuminating the ground being emitted and controlled from and by the satellite. Since the waves are not absorbed by the clouds and by the vegetation, it is an all-weather, night-and-day remote sensing method, which represents a great advantage for very cloudy regions and highly vegetated areas. The intensity ofthe backscattering depends on the physical and


morphological properties of the terrain: topography, roughness of the surface, and dielectric constant of the mineral components. The influence of these different parameters is still under investigation.

Ground Roughness Effect

Following the Rayleigh criterion, the transition between a rough surface and a smooth surface relative to the wavelength A. of the incident waves is given by the relation:

h = A./(8cos e),

where e is the incidence angle and h the mean height of irregularities at the surface. For A. = 24 cm, the wavelength used for the Seasat sensor, and for e = 38°, the transition as been shown experimentally to be for h = 4 cm.

Where the roughness of the mineral surface (soil, sand, etc.) is less than 4 cm, the microwave beam is specularly reflected. Unless the surface is conveniently oriented, the reflected beam does not return to the antenna, and the object appears very dark. On the contrary, it looks bright with a suitable orientation. Therefore, the SAR image of such a surface, a sandy desert for example, is very contrasted.

Effect of the Incidence Angle

Essentially based on Doppler effect, back scattered SAR microwaves are very sensitive to the geometry of the terrain. Angles of incidence are comprised between 20° and 60° relative to the vertical. Seasat and SIR data processing has shown that a low incidence angle (20-30°) enhanced the effect of topography, while high incidence angles (40-60°) are more dependent on the roughness. This can be explained by the fact that for diffusing surfaces - h > A./(8 cos e) - the backscattered signal is not a strong function of the incidence angle e, while for specular reflectance surfaces - h > A.j(8 cos e) - the signal reflected is very strongly dependent on e.

Effect of the Physico-chemical Properties of the Surface

For a given wavelength, incidence angle, and polarization state of the microwave beam, the back scattered signal is a function of the averaged dielectric constant of the target (Fig. 83). This means that the moisture content is a very important factor. For example, a microwave beam reaching a smooth, dry, and porous sandy surface will be partly backscattered, and partly refracted at the interface air/sand. The refracted beam may be transmitted across the sand, and then backscattered by a more compact or humid underlying lithology. The

Remote Sensing Methods

G

1.60

0.80

3.00 4.00 DIELECIRIC CONSTANT

5.00 £'

235

Fig. 83. Gray-scale values G of a radar SIR-A image as a function of the dielectric constant E' (real part) of the imaged soils

penetration depth can be calculated, as a function of the magnetic permittivity (close to 1 for the rock containing low ferromagnetic elements) and of the complex dielectric constant of the medium. For very dry sands, the depth penetration can reach 6-7 m, property which is very useful for water prospecting in desert regions.

Obstacles to the Utilization of Remote Sensing Methods for Mineralogical Purposes

From the laboratory extraction of spectral signatures to their utilization in processing remote sensing data, there are several obstacles: the effect of the atmospheric filter, the modifying role of the alteration layer of the rocks and their contamination by dust and micro-organisms, the effect of roughness (centimetric scale) and of the topography (metric scale) on the radiometry of the surface, taking into account the geometry of the sun's illumination.

- The Earth's atmospheric filter influences differently the radiometric values for each channel: atmospheric corrections are applied to the primary satellite data, in order to transform out-of-atmosphere reflectance values into groundsurface values. The correction models are rather simple, and use no ancillary data, such as the atmospheric composition at the moment of acquisition of the data by the satellite sensor. They correct only for H20 vapor and gas Rayleigh scattering, the main uncertainty being due to the actual concentration of dust level in the atmosphere.

- Petrologists need fresh rock sections for their recognition. This observation is also true for remote sensing goals. Nevertheless, field spectroradiometry has demonstrated that the reflectance spectra recorded on outcropping surfaces of the rocks (the "patinas"), reveal absorption features in good correlation with the rock crystal-chemistry: the band numbers and their location in wavelength coincide, although the patina spectra have lower ordinal resolution.


- Where a blackish-brown alteration layer is masking the nature of the rock, particularly in desert areas ("desert varnish"), this layer, opaque in the visible range because formed by Fe and Mn oxi-hydroxides, prevents to record informations about massive rock geochemistry. Whereas the spectra are featureless in the 400-1500 nm range, it has been noted that the desert varnish-induced absorption decreases as the wavelength increases, so that the spectra exhibit well marked band series in the SWIR (Fig. 84) and TIR. This conclusion is particularly useful for the recognition of silicate and carbonate formations on the numerical images remotely sensed on Earth.

- Pixel heterogeneity is the general case when data are remotely sensed from the space on the surface of the Earth. Radiometry of such pixels results from the weighted addition of optical properties of the different outcropping mineral components. Advanced research is in progress to finalize computed intelligent information extraction from reflectance spectra recorded on mineral mixtures.

- Optical remote sensing works generally in a passive mode, i.e., the illumination geometry is not man-controlled. The variation of the incidence angle of the sunbeam combined with the irregularities (roughness) of the surface and the relief (topography) of the land leads to strong modification of the basic spectral signatures. Models have been proposed, but they require simplifications and assumptions.

- Multispectral data acquired by sensors such as Landsat or Spot concerns channels approximately 100 nm wide, so that only broad spectral signatures can be recorded. The new tendency is to acquire data from space in many narrow

CALCITE (ref.)

2.34

...•.•..... , : °0

9..1. ./ .... h" ...........

10

1.91 2.00 2.19 2.34

Fig. 84. Mid-infrared spectrum (1.8-2.5 Ilm) of the outcropping surface (CI) of a rock covered by desert varnish. A spectrum of the fresh rock (calcitic) is given for comparison. Wavelengths of the absorbance maxima (i.e., reflectance minima) are pointed out in Ilm


spectral channels - hyperspectral remote sensing - which should allow in the near future the extraction of much more detailed information about mineral composition of the planetary surfaces.

- Spectral signatures in space environments other than Earth should be reassessed taking into account new environmental parameters: pressure, temperature, different composition of the atmosphere if any, sun illumination angle, and so on.

Conclusion

Hyperspectral imaging radiometers, able to sample the optical spectrum over more than 200 spectral 10 nm-wide channels, are now utilized on the field and from aircraft to test the spectral signatures models developed at the laboratory level. The next step should be reached with space borne imaging spectrometers, allowing the remote discrimination or identification of individual mineral species, the main difficulties being the unavoidable trade-off between spectral and spatial high resolutions, and the necessity of atmospheric corrections for each narrow spectral channel taking into account the local meteorological conditions.

Such developments will lead to achieving remote sensed imaging spectrometry on the surface of the Earth or of any planet or asteroid, and to deducing from the interpretation of their spectral features and crystal-chemical properties of their surfaces.

Comparison between microwave and optical remote sensing methods shows that both techniques provide very different but complementary information. On the one hand, SAR sensors provide cloud-free data very sensitive to surface roughness and topography, and to moisture content. On the other hand, VIS and IR remotely sensed data are more sensitive to surface crystal-chemistry.

ReferenCdi

Blom RG, Elachi C (1981) Spaceborne and airborne imaging radar observations of sand dunes. J Geophys Res 86: 3061

Bothorel A. Cervelle B, Chorowicz J, Tamain G, Alem EM (1984) Spectral signatures (visible, near-infrared) of rocks and ores. Application to remote sensing of three types of ore bodies from S. Morocco, Mod Geol 8: 277-294

Cervelle B, Moelo Y (1990) Advanced microspectroscopy. In: Vaughan DJ, Jambor JL (eds) Mineral Assoc can, Ottawa, pp379--408

Clark RN, Roush TL (1984) Reflectance spectroscopy: quantitative analysis techniques for remote sensing applications. J Geophys Res 89: 6329-6340

Elachi C, Brown WE, Cimino JB et al (1982) Shuttle imaging radar experiment. Science 218: 996-1003

Goetz AFH, Vane G, Solomon JE, Rock BN (1985) Imaging spectrometry for Earth remote sensing. Science 227: 1147-1153

Hapke B (1981) Bidirectional reflectance spectroscopy. 1: Theory. J Geophys Res 86: 3039-3054


Hapke B, Wells E (1981) Bidirectional reflectance spectroscopy. 2: Experiments and observations. J Geophys Res 86: 3055-3060

Huguenin RL, Jones JL (1986) Intelligent information extraction from reflectance spectra: absorption bands positions. J Geophys Res 91: 9585-9598

Hunt GR (1979) Near infrared spectra (1.3-2.4 /lm) of alteration minerals. Potential for use in remote sensing. Geophysics 44: 1974-1986

Hunt GR, Salisbury JW, LenhoffCJ (1970-1976) Visible and near-infrared spectra of minerals and rocks:

I. Silicates minerals. Modern Geology, 1970, 1, pp 283-300 II. Carbonates. Modern Geology, 1971,2, pp 23-30

III. Oxides and hydroxides. Modern Geology, 1971,2, pp 195-205 IV. Sulphides and sulphates. Modern Geology, 1971 3, pp 1-14 V. Halides, phosphates, arsenates, vanadates and borates, Modern Geology, 1972, 3, pp

121-132 VIII. Intermediate igneous rocks. Modern Geology, 1973, 4, pp 237-244

IX. Basic and ultrabasic igneous rocks. Modern Geology, 1976,5, pp 15-22 XI. Sedimentary rocks. Modern Geology, 1976, 5, pp 211-217

XII. Metamorphic rocks. Modern Geology, 1976, 5, pp 219-228 Khale AB, Goetz AFH (1983) Mineralogical information from a new thermal infrared

multispectral scanner. Science 222: 24-27 Khale AB, Madura DO, Soha JM (1980) Middle infrared multispectral aircraft scanner data:

analysis for geological applications. Appl Opt 19: 2279-2290 Lynn DW (1984) In: Remote Sensing Soc (ed), Satellite remote sensing: review and preview.

UK, p 41. Proceedings of the 10th Anniversary International Conference of the Remote Sensing Society, Reading, UK.

McCauley JF, Schaber GG, Breed CS, Grolier MJ, Haynes CV, Issawi B, Elachi C, Blom R (1982) Subsurface valleys and geoarcheology of the eastern Sahara revealed by Shuttle Radar. Science 218: 1004-1020

Maquet M, Cervelle B, Gouet G (1981) Signature of Ni2+ and Fe3+ in the optical spectra of limonitic ore from New Caledonia: application to the determination of the nickel content. Mineral Depos 16: 357-373

Moore RK (1983) In: Am Soc Photogrammetry Manual of remote sensing, RN Colwell (ed) vol 1, pp 399-538

Mustard JF, Pieters CM (1989) Photometric phase functions of common geologic minerals and applications to quantitative analysis of mineral mixture reflectance spectra. J Geophys Res 94: 13619-13634

Pinty B, Ramond D (1987) a simple bidirectional reflectance model for terrestrial surfaces. J Geophys Res 91: 7803-7808

Pinty B, Verstraete MM, Dickinson RE (1989) A physical model for predicting bidirectional reflectances over bare soil. Rem Sens Environ 27: 273-288

Podwysocki MH, Segal DB, Abrams MJ (1983) Use of multispectral scanner images for assessment of hydrothermal alteration in the Marysvale, Utah, mining area. Econ Geol 78: 675-687

Salisbury JW, Hapke B, Eastes JW (1987) Usefulness of weak bands in mid-infrared remote sensing of particulate planetary surfaces. J Geophys Res 92: 703-710

Salisbury JW, Walter LS, Vergo N (1989) Availability of a library of infrared (2.1-25.0 /lm) mineral spectra. AM Mineral 74: 938-939

Short NM, Stuart LM Jr (eds) (1982) "HCMM anthology" In: NASA Scientific and Technological Information Branch SP 70Ul, Washington DC

Sunshine JM, Pieters CM, Pratt SF (1990) Deconvolution of mineral absorption bands: an improved approach. J Geophys Res 95: 6955-6966

Switzer P, Kovalick NS, Lyon RJP (1981) Estimation of atmospheric path radiance by covariance matrix method. Photogramm Eng Rem Sens 47: 1469-1476

Walter LS, Salisbury JW (1989) Spectral characterization of igneous rocks in the 8 to 12 /lm region. J Geophys Res 94: 9203-9213

CHAPTER 5

Microprobe Analysis

240 Chapter 5. Microprobe Analysis

5.1 Electron Probe Microanalysis

S.l.B. REED and I.M. ROMANENKO

Electron Microbeam Techniques

Modern electron microbeam instruments are represented by the electron microprobe (EMP), the scanning electron microscope (SEM), the transmission electron microscope (TEM), and the Auger microprobe.

The technique of electron probe microanalysis (EPMA) involves measuring the intensity of characteristic X-rays generated by electron bombardment, from which the chemical composition can be determined. The sample consists of a polished section, either a petrological thin section or an opaque mineral mount. The composition of individual mineral grains or even variations within a single grain can be determined, with a spatial resolution of approximately 2,um.

The SEM is closely related to the EMP, but is mainly used for the study of sample surfaces with high spatial resolution (e.g., 5 nm) using back scattered and secondary electron signals emitted as the beam is scanned across the surface. Although chemical analysis is not the main purpose of the SEM, it is possible with the addition of an X-ray spectrometer (usually of the energy dispersive type). Quantitative analysis is more difficult to achieve than in the EMP, however, mainly because of poor control of sample geometry and beam current.

The TEM is used for the study of samples prepared as thin foils (typically < 200 nm thick in the case of silicates) or small particles on thin films. Very high

spatial resolution (down to below 1 nm) can be obtained. The X-ray signal may be detected by an energy dispersive spectrometer, allowing semi-quantitative analysis with spatial resolution of perhaps 30 nm. This technique is known as analytical electron microscopy (AEM). Electron diffraction patterns giving crystallographic information can also be obtained.

The spatial resolution of various micro-analytical techniques is compared in Fig. 85.

Principles of Electron Probe Microanalysis

The classical electron microprobe (Castaing 1951) consists of:

- electron gun and column to generate and focus electron beam onto sample, - specimen stage with precise x, y, and z movements, - optical microscope for viewing and selecting areas on sample for analysis, - at least one wavelength dispersive (WD) X-ray spectrometer and possibly an

energy dispersive (ED) spectrometer.

The WD spectrometer is a monochromator in which a single X-ray wavelength is selected by Bragg "reflection", the wavelength being changed by


100 -f-----,

80

c o 60 ...., ::J o Ul Q) .... 40 r- - I

......, Q Q)

o 20

I

I I

-- XRF Microprobes - - PIXE Microprobes

- - - - LMA (loser microprobe) -------- EPMA ---------- SIMS

o _:_:_:_:_::1... _____ _

o 10 20 30 40 Lateral resolution, J1.m

50

Fig. 85. Spatial resolution of different microprobe techniques

241

varying the angle of reflection. Different wavelength ranges are covered by crystals of different lattice spacing. In the case of the ED spectrometer the energy-resolving capability of the lithium-drifted silicon detector is utilized, the X-ray spectrum being obtained by sorting the detector pulses electronically according to their height.

The energy ofthe electron beam is typically in the range 5 to 30 keY and the diameter of the focused beam is 1 J1.m or less. The emitted X-ray intensity is dependent on the electron accelerating voltage and the current in the beam (typically 10-100 nA). Quantitative analysis is carried out by comparing characteristic X-ray intensities generated in the sample and in standards, using identical instrumental conditions. The standards may be pure elements or compounds of known composition. "ZAF" corrections are required to allow for the effects of X-ray absorption, the penetration and back scattering of incident electrons and X-ray fluorescence occurring within the sample, all of which are dependent on composition.

Instruments

The first commercial EMP, the Cameca MS-85, was introduced in 1958 and was succeeded by the following models manufactured by Cameca (France): MS-45, Camebax-MBX, -Micro, -Microbeam, and SX-50. Other commercial instruments include the Microscan 1, 5, and 9 (Cambridge Instruments, England);


JXA-3, -5A, -50A, Superprobe-733, -8600 (JEOL, Japan); EMX, SEM-XMA, SEMQ (ARL, USA); MAC-400 (USA); MAR-2, -3, -4 (USSR). A comparison between the Cameca MS-85 and SX-50 models is as follows:

1. X-ray take-off angle: MS-85 - 16°, SX-50 - 40 0; 2. spectrometers: MS-85 - 2 WD., SX-50 - up to 6 WD. or 5 WD. and 1 ED. 3. beam diameter for probe analysis: MS-85 - about 5/lm, SX-50 - 0.1 /lm; 4. computers: MS-85 - none, SX-50 - 2 microprocessors and 1 computer.

It may be noted, however, that the technical advances evident in this list are associated with a factor of 10 increase in cost.

Only two "classical" EMP instruments are currently manufactured: these are the Cameca SX-50 and JEOL 8600, both of which are available in several versions, costing from $400k to $8ook (US) approximately. Many different models of SEM are available. These are usually considerably cheaper and can be fitted with X-ray spectrometers for analytical applications. However, as noted above, they are less satisfactory for quantitative analysis than the "true" EMP instrument.

Cathodoluminescence or CL (light emission stimulated by electron bombardment) can be observed through the optical microscope in the EMP and is of considerable interest in mineralogy. CL detectors which can be mounted on SEMs are also available. Recent research shows the possibility of detecting rare earth elements (REE) at concentrations down to 1 ppm in phases such as zircon, monazite and xenotime using the new CL spectrometer designed by the Analytical Center of Mechanbor (Leningrad, USSR), mounted on a CAMEBAX-Micro EMP instrument (Zamorjanskaja et al. 1987).

Analytical Signal in EPMA

The analytical signal in EPMA is the net intensity of characteristic X-ray lines from sample and standards, after background (b.g.) subtraction. In WD analysis background is estimated in various ways: (1) by measuring b.g. intensity on both sides ofthe line; (2) by measuring b.g. on one side, applying a correction for slope; (3) by measuring b.g. on "blank" samples, with a correction for the atomic number difference between the blank and the sample. In ED analysis background is usually removed either by means of a "top-hat" digital filter or by fitting the shape of the continuum using a mathematical model.

Spectral interferences are uncommon when WD spectrometers are used, because of their high resolution. Overlap between peaks is frequent in ED analysis, however, and special spectrum deconvolution and fitting procedures are used to deal with this problem.

"Light" elements (Z < 10) have X-ray lines of energy below 1 keY which can be detected by WD spectrometers with synthetic multi-layer structures designed to diffract long-wavelength X-rays, or by ED detectors which have either a very thin window or none at all. Overlap problems in this region of the spectrum are

5.1 Electron Probe Microanalysis 243

accentuated by the presence of Land M lines of heavier elements. Also quantitative analysis for light elements is considerably more difficult than for "ordinary" elements, owing to severe absorption and the effect of chemical bonding on line positions and shapes. However, considerable progress in this field has been made in recent years.

Limits of Detection

The lower limit of detection is one of the most important characteristics of EPMA and other analytical methods. It can be defined as the concentration equivalent to three times the standard deviation of the background count. The equation:

LD = 3SBR- 1C

gives the limit of detection (LD) as a function of the relative s.d. of the background (SB) and the peak to background ratio R for a weight fraction C of the element concerned. In principle SB can be reduced indefinitely by increasing the counting time, but the practical limit is '" 0.Q1. The peak to background ratio (R) is ultimately limited by the finite natural width of the X-ray lines but in practice is determined by the resolution of the spectrometer, values of around 1000 being typical for pure elements. Figure 86 shows empirical limits of detection as a function of atomic number.

Standards

Standards for quantitative microprobe analysis should be:

1. homogeneous with respect to major analyzed elements; 2. stable under electron bombardment and in air; 3. capable of receiving a high quality polish.

The following types of standard are commonly used for rock-forming mineral: (1) stoichiometric end-members of solid-solution series with known minor elements; (2) pure oxide crystals; (3) other compounds or pure metals in cases where stable oxidized compounds do not exist. For rare earths the following standards may be used: (1) crystals of pure oxides, phosphates, aluminates and gallium garnets; (2) silicate and boro-silicate glasses. For sulfides, arsenides, and tellurides synthetic compounds are usually used. For platinum-group minerals pure metals and alloys are employed.

Problems in Electron Probe Microanalysis of Various Mineral Groups

Rock-Forming Minerals. An important problem in EPMA by WD spectrometry is the difference in valence state of elements in samples and standards,

E Q. Q

cO 0

:.;::; U <1J

+" <!)

0

"-0

+"

'E :.:J

244

1200

1000

800

600

400

200

0 0 20

/ ..... -I' /

'"

, /

40

Atomic number, Z

/ /

I

/

I I

Chapter 5. Microprobe Analysis

I I

I I

)'

I

+-+-+-+-+ K-lines (JXA-SA data) x - - L-lines (JXA-SA data) .............. M-lines (JXA-SA data) • __ ... Birks et 01. ('absolute' ratios)

.. _--- ...

60 80 100

Fig. 86. Limits of detection ofEPMA (K-, L-, and M-lines) including 'absolute' values (for ideal spectrometer of infinite resolution)

which causes displacement of characteristic lines especially for C, F, Na, Mg, AI, and S. Thus it is undesirable to use sulfide standards for the analysis of S in silicates and preferably only sulfates should be used (e.g., SrS04, BaS04)'

Sulfides, Tellurides, and Platinum Group Minerals. The most important problems in the EPMA of sulfides, tellurides, and platinum group minerals are connected with interferences between lines and measuring background intensities. Attention to the ZAF correction procedures is also needed. For example, if Bi2S3 is used as the standard for Bi, low Bi contents are obtained.

Rare Earth Minerals. In analyzing REE minerals the main problem is the true measurement of background and line intensities in the presence of interferences and line overlaps. Note that the Mz lines are much more intense suggested by table ( ~ 1 %): see Fig. 87.

Energy Dispersive EPMA. Energy dispersive EPMA (Fig. 87) allows the analysis of major elements (from 0.1 wt. % upwards) in rock-forming minerals (Reed and Ware 1975). In most cases it is quicker than WD analysis, though the minimum detectable concentration is higher. It can also be applied to sulfides, platinum group minerals, and REE minerals. Recent developments in elec-


~

0 :2 '--N :2

0 ~

0 0:::

25

20

15

\ \

\

10 \

5

\ . \ \

-Be window, 40 degrees -Thin window, 30 degrees -Be window, 30 degrees

/ /

/ ,/

-.,......-=---------~

04n~rrrrITTI~~ITITTITI~rrrrITITTITI~rrrrITTInnTrrITIT~ 60 65 70 75 80 85 90 95

Atomic number, Z

245

Fig. 87. Mz/Ma = f(Z); Eo = 10 keV energy dispersive. Be and thin windows take-off angle = 30 and 40 degrees

tronics for ED spectrometers enable faster spectrum acquisition so that analysis times as short as 10 s (for major elements) are possible (Reed 199Q).

References

Birks LS, Seebold RE, Grant BK, Grosso JS (1965) X-ray yield and line/background ratios for electron excitation. J. Appl. Phys. 36: 699-702

Castaing R (1951) Ph D thesis, Univ. Paris Dunham AC, Wilkinson FCF (1978) Accuracy, precision and detection limits of energy

dispersive electron microprobe analysis of silicates. X-ray Spectrom 7: 50-56 Heinrich KFJ (1981) Electron beam X-ray microanalysis. Van Nostrand Reinhold, New York Heinrich KFJ, Newbury DE, Myklebust RL, Fiori CE (eds) (1981) Energy dispersive X-ray

spectrometry. NBS Special Publication, 604 pp Jarosevich E, Nelen J A, Norberg J A (1980) Reference samples of electron microprobe analysis.

Geostand Newslett 4: 43-48 Keil K (1973) Application of the electron microprobe in geology. In: Andersen CA (ed)

Microprobe analysis. Wiley, New York, pp 184-239 Long JVP (1977) Electron microprobe analysis. In: Zussman J (ed) Physical methods in

determinate mineralogy. Academic Press, London, pp 273-341 Reed SJB (1990) Quantitative ED analysis at high count rates. In: Michael JR, Ingram P (ed):

Microbeam analysis - 1990. San Francisco Press, San Francisco, pp 181-184 Reed SJB (1992) Electron microprobe analysis. US Dept Commerce, Washington (2nd ed)

Cambridge Univ Press, Cambridge


Reed SJB, Ware NG (1975) Quantitative electron microprobe analysis of silicates using energydispersive X-ray spectrometry. J Petrol 16: 499-519

Zamorjanskaja MY, Zamorjanski AN, Yajinshenker IA (1987) Optical spectrometer attached to electron probe microanalyzer. Pribory & Technika Experimenta 4: 161-163

5.2 Trace Element Microanalysis by Proton-Induced X-Ray Emission (PIXE): The Proton Microprobe

D.S. WOOLUM

Introduction

The proton microprobe (proton probe) is based on the same principles as the electron microprobe (electron probe), except that protons are used to excite the atoms in the sample and induce the emission of their characteristic X-rays. Because optimal excitation for most elements typically occurs for protons of 1 to 4 MeV energy, proton micro probes are accelerator-based instruments. MeV protons are more difficult to focus than 10-30 keV electrons; nonetheless, micrometer-diameter beam spot sizes are achievable. In general, the spatial resolution of the proton probe will depend on a number of factors: the horizontal resolution on the beam spot size and the depth resolution on factors like the proton beam energy, the sample composition, sample thickness, and the elements being analyzed; for trace elements with X-ray energies greater than about 5 ke V residing in silicate matrices, the primary factor is the proton beam penetration, which is determined by the proton energy. In fact, adjustment of the beam energy is occasionally used to achieve desired depth resolutions, although this technique is limited in practice. After a point, in decreasing the beam energy to improve the depth resolution, one loses sensitivity, as ionization crosssections decline with proton energies below a few MeV. In practice, the spatial resolution in thick samples is about ten to several tens of microns and is thus somewhat poorer than that of the electron probe.

The primary advantage of the proton microprobe over the electron probe is its greater signal-to-noise ratio. While routine electron microprobe analyses allow for analyses down to hundreds of ppm, minimum detection limits for the proton range down to a few ppm. This greater sensitivity is due to the lower X-ray backgrounds involved in the proton case.

Cookson [1987: in: Watt and Grime (1987)] has outlined the development of the proton-induced X-ray emission (PIXE) technique, starting with initial efforts to produce micro beams using specially designed collimators and culminating in more recent successes involving electrostatic, conventional magnetic, and superconducting solenoidal focussing elements. The majority of PIXE applications have been in the general areas of air pollution studies and biological science, due to the advantage of dealing with samples having a low Z (atomic number)

5.2 The Proton Microprobe 247

matrix, where X-ray backgrounds and interferences are minimal. Furthermore, much PIXE work has involved the analysis of thin-target samples, where the calculation of concentrations is uncomplicated by the need to take into account the degradation of the primary beam energy with depth in the sample or to make matrix corrections arising from the depth-dependent X-ray production and absorption. Nevertheless, a number of laboratories around the world have developed facilities and software for thick-target analysis of mineral and hard rock samples. These developments are the subject of this chapter.

Useful overviews of PIXE and its many applications can be found in Johansson and Campbell (1988), in the five Proceedings of the International Conferences on PIXE and its Analytical Applications (1977, 1981, 1984, 1987, 1990), in Watt and Grime (1987), and in Cabri and Chryssoulis (1990).

Instrumentation

Only relatively small accelerators are required to produce the proton beam energies and currents required. Typically, a single-ended Van de Graaff accelerator or a Tandem Van de Graaff accelerator is used, and proton energies up to 4 to 6 Me V and external beam currents up to several tens of microamperes are achievable with such accelerators. Energies in the range 2-3.5 MeV are normally used for mineralogical trace element work, since these optimize the ionization cross-sections for a very broad range of elements and yet are not energetic enough to generate very significant numbers of high energy photons from nuclear reactions. A bending magnet is used to direct the proton beam to the PIXE beam line, normally a dedicated line and only one of at least several at the accelerator. This is usually followed with energy regulating slits; if the energy of the beam varies, the deflection of the beam by the bending magnet varies, and the energy regulating slit senses the imbalance of the beam currents striking opposing sides of the slit aperture. It produces a signal proportional to the imbalance, which is used to provide feedback to the accelerator terminal in order to stabilize the beam energy.

The PIXE beam line is maintained at relatively high vacuum (at or better than about a micro-Torr), particularly for microbeam PIXE, in order to minimize beam defocusing and to avoid undue contamination of the system. Magnetic and/or electrostatic steerers are used to guide the beam to the target chamber and normally quadrupole magnets are used to focus the beam. The final focusing element used to produce the microbeam varies from facility to facility.

In some cases, specially designed collimators are used to produce the order of micrometer-sized beams, but care must be taken to minimize the collimator scattering which produces beam halos that can excite grains adjacent to the target mineral. Such an arrangement usually results in the loss of significant beam current, since all the beam except that portion incident on the 1-10 micrometer aperture is stopped in the collimator. As a rule of thumb, sample


currents useful for trace element work are those in excess of a few tens of a nanoampere. Further, the collimator is sputtered by the beam, enlarging the aperture and degrading its contours, and must be replaced periodically.

In some cases, quadrupole doublets or triplets are used for the final focusing. This arrangement significantly enhances the beam current densities and can produce micrometer-sized beams with careful design (Watt and Grime 1987), but working distances (focal lengths) are relatively long (the order of meters).

In a few cases, superconducting solenoidal lenses are used for the final focusing. Due to the large magnetic fields achievable (up to about 8 Tesla) working distances are relatively short (the order of a few tens of centimeters) and superior beam current densities in micrometer-sized beam spots are achievable (Maggiore 1980, 1981). On the other hand, such final focusing elements have some disadvantages. They produce large stray fields in the target chamber (magnetic samples must be well-secured and magnetic materials eliminated in the construction of the target chamber); they require liquid nitrogen and helium, which are costly, and periodic shut downs (once or twice a day typically) to fill the cryogenic dewars. Furthermore, these dewars encasing the superconducting solenoid are relatively bulky and can impose significant spatial constraints on the design of the target chamber. Finally, in some situations, the beam current densities may be high enough to degrade the sample and require special sample preparation/mounting; our experience at the Los Alamos PIXE facility indicates that for 20-50 micron spot sizes most rocks appear able to take 50 nanoampere proton currents. The rare cases where sample damage was noted was not due to sample melting but to a softening and flow in the epoxy mounting material.

Rarely, the proton beam is extracted from the beam line via a thin exit window or a differential pumping port, and collimated. Such arrangements are suitable when larger (a few tens of micrometer-sized) beams are sufficient for the task at hand and when only higher energy (greater than about 3 keY) X-rays are of interest, since the low energy X-rays are strongly attenuated in the atmosphere in centimeter-sized path lengths.

In virtually all cases, the induced X-rays are detected using an energy dispersive [typically a lithium-drifted silicon, or Si(Li)] detector, which has a large acceptance angle for efficient collection of the relevant X-rays (2-30 keY, typically) and can achieve a moderate energy resolution (about 145 eV for the 5.90 keY Mn K-alpha line). Such detectors have the advantage of being capable of simultaneous multi-element analysis but have severe count rate limitations in PIXE due to dead time and pulse pile-up effects. Important future developments in PIXE will most likely see the design and fabrication of efficient wave length dispersive crystal spectrometers specially designed to maximize energy resolution (to resolve interfering peaks) while maximizing X-ray collection efficiency.

For microbeam PIXE it is important to optically view the sample in order to find and position the targeted mineral grains in the proton beam. Specially designed and standard microscope systems have been installed, depending on the resolution desired and the working distance constraints. Optical microscopic documentation of the sample normally involves normal incidence optical


viewing, and samples appear very different and distorted even when the incident light is not normal to the sample surface. Thus, in order that selected grains can be confidently identified and positioned for analysis, ideal viewing systems for PIXE provide normal incidence optical viewing. When viewing is less than ideal, major element X-ray signatures in the collected spectrum can be used to verify the identity of the phase exposed to the beam, as well as to check for the possibility that the beam is sampling an inclusion or another contiguous grain in the section.

Target stages usually allow for the mounting of multiple samples, to minimize sample change (primarily pump-down) times and, thus, to minimize throughput. These may be manually or computer-driven, and of a variety of types (e.g., piezoelectric, stepper motor, micrometer, etc.) with widely varying spatial resolution and reproducibility (usually of the order of sub-micrometer dimensions to 10 Ilm). A wide variety of schemes exist for monitoring the integrated beam current dose delivered to the sample which must be determined in order to perform quantitative analyses. These will be noted in more detail below.

Analytical Considerations

Access to electron microprobes is much more readily obtained, and, in general, electron microprobes are capable of routine analyses of major and minor elements, so, normally, PIXE is used only when trace element analyses are required. For trace element analyses, sample preparation is a serious concern. As in electron probe work, in order to be able to assume simple irradiation geometries, samples must be appropriately polished to obtain a plane surface. In addition, sample preparation procedures must be designed so as to avoid the possibility of introducing contaminants in the cutting, impregnating (if applicable), polishing, coating (to make the sample conducting), and mounting of the sample.

Normally, samples are prepared as thick polished sections. However, when grain sizes are comparable to, or smaller than, the spatial resolution, in order to minimize the possibility of having underlying subsurface grains contributing to the analysis of a grain, one can prepare a thin section (less than the order of 100 Ilm) which allows for transmitted light observation so three dimensional effects can be checked directly. Alternatively one can prepare a doubly-polished thin section of a thickness comparable to the grain sizes, but the sample thickness must be known for quantitative analysis (usually infinite thickness is assumed in quantitative analyses of geological samples: see below). Such special sample preparation procedures may also be required, even when grain sizes are large, when exceptionally fine spatial resolution is required, as for determining shallow diffusion profiles for example.

In order to partially obviate the count rate limitations of the Si(Li) X-ray detectors, absorbers are normally used in conjunction with the detectors. The


absorbers are typically thin metal foils. The absorber composition and thickness is chosen to selectively absorb major and minor element X-rays compared to the trace element X-rays, which are normally of higher energy. X-rays generated in the sample can induce secondary fluorescence in the absorber, so the absorber must be of high purity so that the X-rays generated in it do not contribute to sample spectrum in regions used for the trace element analysis. The absorbers must also be uniform in thickness, without imperfections (e.g., pinholes) and of accurately known thickness. Because of the exponential attenuation of the X-rays in the absorber, small uncertainties in the absorber thickness can result in relatively large uncertainties in the relative attenuation of the lower energy X-rays.

In the electron probe, the primary source of background is the Bremsstrahlung background from both the electrons in the primary beam and the secondary electrons generated in the sample by the beam. In the proton probe, there still is Bremsstrahlung background due to the secondary electrons generated in the sample, but the Bremsstrahlung from the primary proton beam is much smaller. (The intensity of Bremsstrahlung radiation emitted by a charged particle is inversely proportional to the mass of the particle, and since the mass of the proton is roughly 2000 times the mass of the electron, the Bremsstrahlung background from the primary beam is orders of magnitude less in the case of the proton probe.) This is the primary reason for the superior signal-to-noise ratio, and thus sensitivity, for PIXE when compared to electron probe microanalysis.

During a conventional PIXE analysis, X-rays detected by the Si(Li) detector are stored in a multi-channel analyzer. Characteristic X-ray peaks are superimposed on the broad Bremsstrahlung background in the X-ray spectrum collected. The energy of the X-ray peaks identifies the elements contained in the sample, and the integrated peak counts in an isolated peak is proportional to the concentration of that element in the sample. Figure 88 shows a typical X-ray spectrum from a particularly challenging sample; it is a thick homogenized sample of a carbonaceous chondrite meteorite, which contains virtually all the elements in their "cosmic" proportions. This spectrum illustrates the difficulties arising from the use of energy-dispersive X-ray detectors like the Si(Li) detector. With only about 150 eV energy resolution, there are severe overlaps of X-ray peaks, K-alpha and K-beta lines in the transition metal region of the spectrum. Rather elaborate computationally intensive data analysis procedures must be developed to first fit background and then deconvolve the overlapping peaks in the residual background-subtracted spectrum.

Mineral/rock PIXE analyses Fig. 88 usually involve the use of thick targets where (1) the variation of the proton energy (and thus the X-ray production) with depth in the sample and (2) the self-absorption of X-rays produced below the surface are important factors. Converting the integrated X-ray peak counts to concentrations primarily requires knowledge of the relevant atomic parameters (e.g., energy-dependence of the ionization cross-sections, fluorescence yields, X-ray mass absorption coefficients), the irradiation geometry (angle of incidence of the beam and takeoff angle to the X-rays toward the detector), the

5.2 The Proton Microprobe

10'.------------------,

Fe

5

Ni

Zn

10 ENERGY, KeV

ORGUEIL

15

251

Fig. 88. The 3 Me V PIXE spectrum of the Orgueil carbonaceous chondrite meteorite. On this scale, the visible peaks are K-line doublets: elemental labels indicate the K-alpha peak. Thick Al absorbers highly suppress the major element X-rays, including Fe and Ni. The minor peaks below Fe, in order of decreasing energy, are Ni and Fe escape peaks, the pulser peak (used to make dead time corrections, if needed) and an Al secondary fluorescence peak from the absorber

concentrations of the major and minor matrix elements (to determine the X-ray absorption character of the matrix), the energy dependence of the detector efficiency, and the thickness and composition of any absorber used in conjunction with the detector.

There are two basic methods for determining absolute concentrations. One involves the use of standards, and the other involves absolute or relative calibrations of yields, without the use of standards. In the case of electroninduced X-ray fluorescence, the X-ray production is complicated in that the electrons are multiply scattered and follow erratic paths as they are slowed down. By contrast, MeV protons follow essentially straight line paths that are easily modeled. The physics of the processes involved is well understood and the values of the required atomic parameters are also well known in the case of PIXE. Thus, it is possible to calculate absolute concentrations without the use of standards, and this is frequently done in PIXE.

Standardless analyses can be made by measuring yields of the trace element~ relative to an independently known element in the sample. Normally, this is a major or minor element that can readily be determined by the electron probe, where the homogeneity of the distribution of the reference element can also be checked. The closer in atomic number the reference element is to the atomic numbers of the trace elements being analyzed the less sensitive is the analysis to uncertainties in the absorber thickness. We (Burnett et al. 1988, 1989), using the Los Alamos-developed software (Duffy et al. 1987) reported an accuracy of ± 10% for heavy elements at 50 ppm, or larger, concentrations in a 15-min

PIXE analysis. Improvements at lower levels are possible with improvements in the physical description of the background and in correcting for sum peaks. It appears realistic to believe that the effective analytical levels can be pushed down to around 10 ppm with ± 10% accuracy. This has already been demonstrated with longer count times (Burnett et al. 1989).


Standardless analyses can also be made by using the major matrix elements to generate an absolute calibration curve giving the X-ray yield per incident proton (or per unit time-integrated beam current: e.g., per microcoulomb incident charge accumulated in the proton irradiation) as a function of Z. Checks of this method have yielded accuracy estimates that range up to 10-15%, depending on the matrix (Clayton 1981), and discrepancies in the yields were systematic (independent of atomic number) and were attributed to small inaccuracies in the charge integration and detector geometry, not to uncertainties in the atomic data.

Less commonly, it is possible to prepare synthetic standards of the same (or compositionally similar) matrix, or to use single element standards. Unless the matrix of sample and standard are nearly identical this involves the additional consideration of the production and absorption in the standard and requires accurate beam dose determinations to compare the standard and sample irradiations for absolute concentration calculations.

Comparison with Other Techniques: SIMS, SXRF

In Synchrotron Radiation X-ray Fluorescence (SXRF), high energy synchrotron radiation photons from a Ge V energy electron storage ring are used to induce X-ray fluorescence. Very intense synchrotron radiation fluxes are available. A beam line at the Brookhaven National Lab NSLS (National Synchrotron Light Source) is at an intermediate stage of development and currently uses a Si(Li) detector. On paper (Gordon 1982), when wave-length dispersive detectors (crystal spectrometer detectors) are used. SXRF should eventually have better sensitivity than PIXE, but right now the two techniques appear to have comparable analytical performance.

The biggest difference between SXRF and PIXE is in sample preparation requirements set by the very different depth responses. For SXRF, the X-ray production cross-sections do not decrease significantly with depth, and with a Si(Li) detector depth response is determined almost entirely by X-ray absorption. For Y K-alpha, the absorption half-thickness is around 300-400 Jlm in silicates. (For PIXE it is more like a few tens of micrometers in silicates.) Thus, for microanalysis using a Si(Li) detector doubly polished samples of accurately known thickness are required. It is desirable that sections be free-standing since even an ultra-pure backing to the sample section scatters beam photons and enhances background levels. Furthermore, SXRF has the disadvantage that diffraction peaks are present in the spectrum when white light incident synchrotron photons are used in the analysis of crystals. This can be eliminated, or altered to eliminate the interference with the X-ray peaks of interest, by rotating the sample. It should not be a problem once higher resolution detectors are used or when monochromatic excitation of the sample is performed. One additional difference between SXRF and PIXE as they are currently used in geochemical analyses, is the shape of the backgrounds. PIXE backgrounds generally peak at


about 5 keY and drop rapidly at higher energies. SXRF backgrounds are primarily due to scattering of the incident photon beam and they peak at higher energies. Compared to SXRF, PIXE signal-to-noise ratios are better at higher energies (above about 10 keY) and worse at lower energies (below about 5 keY.

The ion microprobe [Secondary Ion Mass Spectrometry (SIMS)] principles, instrumentation, and analytical considerations for mineral analysis are summarized, for example, by Cabri and Chryssoulis (1990). SIMS normally uses a 10-20 keY ion beam focussed to 10-50 microns to sputter the sample. Secondary ions produced in the sputtering process are analyzed with a mass spectrometer. SIMS has better sensitivity than PIXE, can analyze at the ppm level, and can perform isotopic analyses. Compared to PIXE, the SIMS technique involves strong matrix effects, in general (Ray and Hart 1982). These can be minimized with strong energy filtering (Crozaz and Zinner 1986), with a sacrifice in sensitivity, but sometimes standards that are closely matched in chemical composition to the samples being analyzed are required. At times this is merely an annoyance, and at times a very difficult requirement to meet (e.g., the case of some exotic phases).

As in many similar comparisons, it is true here that no one trace element technique can be viewed as the do-all and be-all for all applications. Given a well-defined scientific problem, care must be exercised in the choice of a technique to use in its solution. Or, alternatively, careful consideration is advised when selecting problems to work on, given the availability of a particular technique.

References

Burnett DS, Woolum DS, Benjamin TM, Rogers PSZ, Duffy CJ, Maggiore C (1988) High precision thick target PIXE analysis of carbonaceous meteorites. Nucl Inst Meth B35: 67-74

Burnett DS, Woolum DS, Benjamin TM, Rogers PSZ, Duffy CJ, Maggiore C (1989) A test of Carbonaceous chondrite elemental abundance smoothness. Geochim Cosmochim Acta 53: 471-481

Cabri LJ, Chryssoulis SL (1990) Advanced methods of trace element microbeam analysis. In: Jambor JL, Vaughan DJ (eds) Advanced microscopic studies of ore minerals Vol 17. Mineralogical Association of Canada, Ottawa

Clayton E (1981) Uncertainties in theoretical thick target PIXE yields. Nucl Instr Meth 191: 567

Crozaz G, Zinner E (1986) Scanning electron microscopy 1986/II, SEM Inc, Chicago, pp 369-376

Duffy CJ, Rogers PSZ, Benjamin TM (1987) The Los Alamos PIXE data reduction software. Nucl Inst Meth B22: 91-95

Gordon B (1982) Sensitivity calculations for multi-element trace analysis. Nucl Inst Meth 204: 223-229

Johansson SAE, Campbell JL (1988) PIXE: A novel technique for elemental analysis. Wiley, Chichester, UK

Maggiore CJ (1980) Materials analysis with a nuclear microprobe. Scanning electron microscopy/1980/I, SEM Inc, Chicago, pp 439-454

Maggiore CJ (1981) The Los Alamos nuclear microprobe with a superconducting final lens. Nucl Inst Meth 191: 199-203


Proceedings of the International Conferences on PIXE and its analytical applications (1977-1990) Nucl Instr Meth 142 (1977); 181 (1981); B3 (1984); B22 (1987) and B49 (1990)

Ray G, Hart SR (1982) Quantitative analysis of silicates by ion microprobe. Int J Mass Spectr Ion Proc 44: 231-255

Watt F, Grime GW (eds) (1987) Principles and applications of high energy ion microbeams. Hilger, Bristol, UK

5.3 Nuclear Microprobe and Microscopic Analysis

P. TROCELLIER

MeV ion beam techniques are now currently applied in numerous scientific fields like metallurgy, biology and medicine, geology, archeometry, solid state physics, and environmental sciences, as described in the Proceedings of the Eleventh Ion Beam International Conference, held in Balatonfused (Hungary) in July 1993.

It is relatively easy to produce MeV ion beams with a diameter in the range 1 mm-l00 J.lm (milliprobe configuration) by means of adjustable aperture collimators. This type of beam is sufficient to investigate small specimens and to obtain local informations on their chemical composition. However, it was necessary to wait for the beginning of the 1970s when J.A. Cookson and his group in Harwell designed the first magnetic quadrupole lenses device to focus high energy positive ions, and built the first nuclear microprobe in the world. They succeeded to focus down to 4 J.lm width a 3-MeV proton beam and obtained, using Rutherford back scattering spectrometry, the first scanned image on a 400-mesh copper grid deposited on a carbon backing.

At the beginning of the 1990s, nuclear microprobe is coming to its technical maturity. More than 60 instrumental facilities are now running all around the planet. Nuclear microprobe offers to the scientist a wide variety of spectroscopic methods to overcome the mystery of both structure and elemental composition of materials at the micrometer scale. This property appears to be a determining advantage in comparison with other microanalysis techniques.

A Single Device for a Dozen Microanalysis Methods

As a positive ion microbeam (protons, deuterons, helium-3 or helium-4 ions, or heavier ions) impinges a solid surface, lots of atomic and nuclear interactions occur leading to the production of various secondary radiations, as shown in Fig. 89.

Most of these physical processes offer the basis for a microscopic analysis method:

- induced X-ray emission, labeled as PIXE when protons are used (see Chap. 5.2 by D. S. Woolum);

5.3 Nuclear Microprobe and Microscopic Analysis

SECONDARY ELECT RONS

ATOMIC INTERACTIONS

PHOTONS ISIPS)

INCIDENT ION BEAM

SCATTERED PARTICLES IRBS and CCM I

RECOIL PARTICLES IERDAI

DIRECT OBSERVATION PARTICLES OF

ISEM)

{

CHARGEO

NUCLEAR REACTIONS GAMMA-RAYS INRA and PI GE I

TRANSMITTED IONS ISTlM,IMTI

IFS S)

255

NUCLEAR INTERACTIONS

Fig. 89. Schematic description of atomic and nuclear interactions resulting from the bombardment of a solid surface with a MeV ion microbeam

- secondary electron microscopy (SEM) sensitive to surface relief more than to chemical composition; sputter induced photon spectroscopy (SIPS) as developed by I.S.T. Tsong and co-workers at ke V energy; forward scattering ion spectrometry (FSS);

- Rutherford backscattering spectrometry (RBS) and elastic recoil detection analysis (ERDA) generally induced by MeV helium-4 ions;

- channeling microscopy often coupled with RBS, sensitive to crystalline disorder; nuclear reaction analysis (NRA) induced by protons, deuterons, helium-3 ions or heavier particles such as nitrogen-1S ions, between 100 eV and 8 MeV; induced gamma-ray emission, labeled as PIGE when MeV protons are used; scanning transmission microscopy (STIM) and ion microbeam tomography (IMT), based on the collection of transmitted particles.

Analytical Capabilities and Performances

The main characteristics and relevant analytical performances offered by nuclear microprobe techniques are described in Table 10.

PIXE and PIGE are generally combined for major, minor, and trace element determination purpose in solids. From lithium to uranium, minimum detectable

Tab

le 1

0. C

hara

cter

isti

cs a

nd p

erfo

rman

ces

of in

vest

igat

ion

tech

niqu

es d

evel

oped

in

nucl

ear

mic

ropr

obe

anal

ysis

(N

MA

) N

v.

0'\

Met

hod

Con

figu

rati

on

App

lica

tion

P

erfo

rman

ces

fiel

d In

cide

nt

Em

erge

nt

Pro

bing

dep

th

Dep

th

MD

L

reso

luti

on

(wt.

ppm

) (n

m)

PIX

E

2-4

MeY

p

K,

L,

and

M

Tra

ce e

lem

ent

100

nm t

o 5

to 1

00 (

thic

k ta

rget

) X

-ray

s fr

om

dete

rmin

atio

n se

vera

l jlm

0.

1 to

10

(thi

n fil

m)

Na

to U

PIG

E

1-5

MeY

p

Gam

ma-

rays

T

race

ele

men

t 1

to 1

00 jl

m

1 to

200

0 fr

om L

i to

Zn

dete

rmin

atio

n Is

otop

ic r

atio

RB

S 2

-4 M

ey4 H

e S

catt

ered

T

hin

laye

r st

udy

0.01

to

3 jlm

10

to

50

100

to 1

000

4He

Hea

vy e

lem

ent

in l

ight

sub

stra

te

FS

S

3-4

MeY

p

Sca

tter

ed p

at

Lig

ht e

lem

ents

Ij

lm

20

1000

fo

rwar

d an

gles

pr

ofil

ing

ER

DA

2

-4 M

ey4 H

e R

ecoi

l lh

H

iso

tope

0.

1 to

0.5

jlm

(R

) 25

(R

) 10

to

100

prof

ilin

g 1

to 1

0 jlm

(T

) 40

(T

)

NR

A

0.5-

2 M

eYp

C

harg

ed

2H t

o 34

S 0.

01 t

o 1

jlm

5

to 5

0 0.

01 t

o 10

00

(")

::r

0.1-

2 M

eYd

pa

rtic

les

or

isot

ope

(RR

) (R

R)

0.5-

3 M

eY 3

He

gam

ma

rays

di

stri

buti

on

1 to

10

J.lm

10

0 to

500

~

'1:1 n .... yo

~

<S . .... 0 '1:

1 .... 0 cr'

(1) » ::s e:...

'<

i!'. '"

5.3 Nuclear Microprobe and Microscopic Analysis 257

limits vary respectively, from 1 to 100 weight ppm for PIXE and from 1 to 2000 weight ppm for PIGE between 2 and 5 MeV in thick targets.

Considering its kinematics, RBS is well adapted for heavy element (Z 2': 40) measurement in light matrix (Z ~ 20) and/or for thin layer characterization. In crystalline samples, RBS can be coupled with ion channeling spectroscopy in order to determine precise atom location of the species out of their usual sites, and to identify impurities or compositional anomalies. Therefore, it is possible to image the lateral variations in crystal structure and atom position combining microbeam scanning and high depth resolution channeling. This technique, called channeling contrast microscopy (CCM), was first proposed by J.e. McCallum and co-workers in 1983. It has led to numerous mapping applications in the semiconductors field. An extensive review of elastic collision based methods was given by W.K. Chu and M.A. Nicolet in 1978.

In the case of thin targets, the detection of the scattered ions can be performed in forward angles (FSS). Hence, the scattering spectrum gives information about the depth distribution of the major components across the film.

ERDA is essentially devoted to hydrogen and deuterium microdetermination. Two geometrical configurations can be considered. The first operates in reflection (R) and leads to a probing depth of a few tenth of micrometers, depending on the helium-4 ion energy. The second corresponds to transmission collection (T) and gives a probing depth up to 7 lim in thin targets at 3 MeV, as demonstrated by J. Tirira and co-workers in 1990.

Direct observation of nuclear reactions induced by:

- protons: (p, alpha) and (p, gamma) type; - deuterons: (d, p) and (d, alpha) type; - helium-3 ions: eHe, p) and eHe, alpha) type; - heavier ions such as 15N

is specific of light element isotopes from 1 H to 34S. Sometimes, nuclear crosssections exhibit strong and narrow resonances (RR) in function of the incident ion energy. Hence, it enables elemental depth profiling to be undertaken over several micrometers. The most famous resonance example corresponds to the nuclear reaction: lHe 5N, alpha gamma) 12C at 6.385 MeV, as proposed by W.A. Lanford (1986) for hydrogen depth profiling. The surface resolution will in this case be better than 5 nm.

Scanning transmission ion microscopy (STIM) is a technique that permits the imaging resolution of the nuclear microprobe to be improved by at least one order of magnitude. Particularly interesting for thin films investigation, for example biological tissues, STIM is based on multiple event energy loss and scattering through the specimen, these are dominated by major elements of the matrix. Both processes can be used to image the density variations of the specimen within the scanned area. Ion microbeam tomography (IMT), using protons higher than 4 MeV energy, produces cross-sectional views of materials with a high imaging resolution in a manner similar to STIM. Material density is derived from the energy losses of incident ions, primarily due to interactions

258 Chapter 5 .. Microprobe Analysis

with electrons. Several oral or poster contributions were given on these topics during the Second and the Third International Conferences on Nuclear Microprobe Technology and Applications, respectively held in Melbourne in February 1990 and in Uppsala in June 1992, notably by H. Lefevre, M. Cholewa, respectively G. Bench, and A. Pont au.

Instrumentation

Most of the nuclear microprobe facilities operating in the world are based on an electrostatic accelerator: single-ended or tandem Van de Graaffwith a maximal spot in the whole beam area. This spot is afterwards focused down to the micrometer level by means of magnetic quadrupole lenses, electrostatic lenses, or superconducting solenoid. In 1987, F. Watt and G. Grime gave a detailed review of Me V ion beam focusing procedure.

The analysis chamber is equipped with a target holder driven by stepping motors. An optical microscope, often coupled with a CCD video camera, permits location of the part of interest on the sample surface. Energy dispersive X-ray detector (Si Li), germanium HP crystal and silicon surface barrier detectors are setted in the rear part of the chamber near the sample under investigation. The microbeam scanning is controlled by electromagnetic coils or electrostatic plates placed before or after the lenses.

All the signals available from the detection devices are converted into pulse height amplitude spectra. The data-handling system, connected to the microprobe, drives the microbeam scanning and the multi parameter data acquisition when elemental maps have to be performed. It also ensures the result display, the data storage onto high memory capacity units, and the mathematical treatment of the spectra, using adapted algorithms.

Microscopic analysis

Nuclear microprobe is now established as a viable analytical technique that offers complementary investigation methods to more conventional microanalysis tools as described in Table 11.

Biological and medical sciences appear to constitute the most extensive research field for nuclear microanalysts since 1974. This fact is essentially due to the high lateral resolution and the relevant minimum detectable limit that can be achieved, especially in PIXE mode. Biological applications of NMA essentially deal with trace element transport in tissues in order to study metabolic mechanisms and elemental desequilibrium. Nevertheless, these studies imply constraints such as specific specimen preparation and loss of integrity during ion bombardment. All these points were fully illustrated by R.D. Vis in 1985 (Vis 1986).

The following paragraph will be concentrated on reviewing applications in the field of mineral materials without the intention of cataloguing all works but rather of identifying the specific contribution of the nuclear microprobe.


Table 11. Comparison of microscopic techniques

Technique Elements Lateral Analyzed Depth Sensitivity resolution volume resolution (wt. ppm) (/lm) (Jim3 ) (m)

Electron microprobe Z ~ 12 0.5 0.5 100 (EMA)

Ion microprobe Z~l 0.05 0.001 0.005 0.001 to 10 (SIMS,IMA)

Laser probe Z~3 0.5 0.1 0.05 to 10 (LPMS, LAMMA)

Nuclear microprobe Z~l 0.3 1 to 10 0.005 to 0.2 0.Q1 to 100 (NMA)

Synchrotron X-ray Z~3 10 5 0.1 fluorescence (SXRF)

Metallurgical applications of the nuclear microprobe essentially concern the determination of light element isotope from 1 H to 28Si with the exception of 2°Ne. It has involved the measurement of changes in the composition of materials as these can strongly influence their properties. Physicochemical phenomena such as oxydation, corrosion, segregation, precipitation, mechanical durability, and phase formation were being studied. In these cases, NRA and PIXE are the most common used analytical methods. They give simultaneously access to the distributions of major matrix components and of the principal "light contaminants".

Microbeam applications in solid state physics and semiconductors technology have shown a rapid development since the beginning of the 1980s. Microcircuit structures with lateral dimensions on the micrometer scale have been investigated using RBS and channeling spectrometry. Thin films composition, crystal growth, impurity redistribution and atom location of dopants constitute the most classical works. The development of imaging tools such as CCM has considerably enhanced the interest of semi-conductors community in nuclear microprobe capabilities, as reviewed by I.S. Williams and coworkers in 1988.

Nuclear microprobe has only been applied for a few years in archeological research. PIXE and PIGE are in this case the principal experimental methods carried out. Specific instrumental arrangements such as external ion beam devices have been designed for this purpose. A wide variety of materials are investigated: glass, clay, stone, alloy, noble metal, bone, tooth, paper, paintings, etc. Topics such as chemical alteration, aging, authentification, protection and restauration, trade and social evolution in ancient communities, and manufacturing processes have been explored. A synthetic review of nuclear microprobe applications in archeology has been given at the Pont a Mousson workshop in 1985.


Geological materials are often complex heterogeneous associations of minerals containing a wide variety of elements in different chemical state. Nuclear microprobe analysis (NMA) appears to be an alternative to electron microprobe analysis (EMA) with energy dispersive X-ray spectroscopy (EDXS) and to scanning secondary ion mass spectrometry (IMA or SIMS), because PIXE, PIGE, NRA, and RBS modes can be combined. Special attention must be paid in PIXE investigation due to numerous X-ray peak overlaps from a wide range of minor and trace elements and pile-up peaks from major elements such as calcium and iron.

Specific applications of NMA in geology could be classified in five principal research axes: cosmic materials, ore minerals, volcanic phenomena, alteration mechanisms, and energy resources studies.

Work on cosmic materials (meteorites, cosmic dust, and lunar rocks) is aimed at different purposes:

- elemental fractionation between different mineral phases using microPIXE; - light element abundance by means of NRA; - origin of stellar carbon with 12C(d,p)13C reaction; - superficial hydrogen distribution in lunar rocks by resonant nuclear reaction

to evaluate solar wind effects; - local isotopic ratio measurements using NRA and PIGE.

Analyses of ore minerals are based on trace element determinations whereby minimum detection limit (MDL) is considered as the most restrictive parameter. Generally, MDL less or equal than 10 weight ppm is required for diagnostics concerning transition series, rare earth, and platinum group elements.

The application of nuclear microanalysis to volcanic samples has been considered on the one hand for trace element research in which PIXE and PIGE modes give fruitful results, and on the other hand for magmatic mechanism study. In this case NRA and ERDA provide relevant insight about the composition and the physical properties of the volatile phase trapped within volcanic minerals. Solid and fluid inclusions offer a new research axis for several nuclear microanalysis groups. The main objective in this case is to collect data on both thermodynamics and kinetics parameters of mineral genesis.

Alteration mechanisms of geological materials concern, for example, submarine basalts and glasses or terrestrial natural oxides and sulfides.

Recent investigations have been undertaken on mineral deposits for energy resources exploitation such as coals, oil shales, and oil field rocks in order to study:

- trace element chemistry; - interaction between liquid phase and mineral surface; - diffusion processes through the host medium.

The ultimate purposes are, in fact, progress in both knowledge of the deposit formation mechanisms and extraction procedures.


Perspectives

One of the most interesting developments in nuclear microprobe technology during the past 5 years is the improvement in lateral resolution. For example, F.W. Martin (Worcester, 1990) has obtained a 0.1 Jl.m diameter microbeam of 1.06 MeV protons using an achromatic lens doublet. Efforts are being made to minimize chromatic, spheric, and parasitic aberrations of electromagnetic quadrupoles, as reviewed recently by G. Legge (1989).

Moreover, progress in ionization processes helps to increase the stability and the brightness of ion sources. The immediate benefit of these improvements concerns the examination mode based on transmission imaging, where specimen mapping can be achieved with a resolution about 0.1 Jl.m, using microbeam intensity as low as 10 - 15 A.

Many efforts are actually dedicated to the study of the secondary effects of nuclear microprobe investigation on solid media in order to predict its ultimate capabilities. It includes energy transfer mechanisms, temperature rising, insulating materials exploration, and beam-induced damage.

Jointly to the microbeam improvements, the availability of both more and more performant detectors and data handling systems gives raise to the accuracy of the informations obtained. The sophistication of mathematical codes required for spectrum interpretation and the development of computing devices for data presentation strengthen the quantitative character of nuclear microprobe results.

From a solid state physics point of view, nuclear microprobe analysis is now capable to help in scientific understanding of grain boundary and interface phenomena, combining several of its spectroscopic modes.

References

Cookson JA, Ferguson ATG, Pilling FD (1972) Proton microbeams, their production and use. J Radioanal Chern 12: 39-52

Chu WK, Nicolet MA (1978) Backscattering spectrometry. Academic Press, New York Lanford W A (1986) Ion beam analysis of glass surface: dating authentification and conserva

tion. Nucl instr meth B14: 123-126 Legge GJF (1989) Microprobe analysis in ion beams for materials analysis. In: Bird JR,

Williams JS (eds) Ion beams for materials analysis. Academic Press, Sydney Martin FW, Goloskie R (1990) An achromatic proton microbeam of 1.06 MeV energy and

0.1 micron width. Proc 2nd Int Conf Nuclear Microprobe Technology and Applications, Melbourne (Australia), February 5-9, 1990, Nucl. Instr. meth. B54 (1991) 64

McCallum JC, McKenzie CD, Lucas MA, Rossiter KG, Short KT, Williams JS (1983) Channeling contrast microscopy: application to semiconductor structures. Appl Phys Lett 42: 827-829

Proc Ion Beam Analysis in the Arts and Archaeology Workshop, Pont a Mousson, February 18-20, 1985 (1986) Nucl instr meth B14: 1-167

Proc 11th Ion Beam Analysis Conf, Balatonfund (Hungary), July 5-9, 1993 (1994). Proc 3rd Int Conf Nuclear Microprobe Technology and Applications, Uppsala (Sweden) 8-12, June 1992. Nucl Instr Meth B77 (1993) Nucl Instr Meth B85: 1-942

262 Chapter 5. Microprobe Artalysis

Proc 2nd Int Conf Nuclear Microprobe Technology and Applications, Melbourne (Australia), February 5-9, (1990) Nucl instr meth B54: 1-440

Tirira J, Trocellier, P, Frontier JP, Trouslard P, Costantini JM, Mori V (1990) 3D hydrogen profiling by elastic recoil detection analysis in transmission geometry. Nucl instr meth B50: 135-139

Tsong 1ST, Yusuf NA (1978) Absolute photon yields in the sputter induced optical emission process. Appl Phys Lett 33: 999-1001

Watt F, Grime GF (1987) High energy ion microbeams. Hilger, Bristol Vis RD (1986) The proton microprobe: applications in the Medical Field. CRC Press, Boca

Raton Williams JS, McCallum JC, Brown RA (1988) Ion microbeam applications in semiconductors.

Nucl instr meth B30: 480-485

CHAPTER 6

Electron, Acoustic, and Tunneling Microscopy of Minerals

264 Chapter 6. Electron, Acoustic, and Tunneling Microscopy of Minerals

6.1 Electron Microscopy of Minerals

H.-R. WENK, A.C. McLAREN, a.M. PENNOCK, and V.A. DRITS

There are basically two distinct types of electron microscope. The transmission electron microscope (TEM) is the analog of the transmission light microscope and provides information about the internal structure of a specimen which is thin enough to be transparent to the electron beam. The scanning electron microscope (SEM) provides information about the surface structure of the specimen. Nowadays, both types of microscope are commonly equipped with an energy dispersive X-ray detector for measuring the intensities of the characteristic X-rays which are emitted by the specimen and hence its chemical composition.

In this chapter the structure, characteristics, and uses of both types of electron microscope will be reviewed briefly, together with the origin of image contrast and some characteristic examples of the use of electron microscope in mineralogical research.

6.1.1 Fundamentals of TEM and HRTEM

A.C. McLAREN

Basic Construction of the Transmission Electron Microscope and Imaging M odes. The modern transmission electron microscope consists of an electron gun and a number of electromagnetic lenses, all within an evacuated column. The electron beam is accelerated by a high voltage and then focused by two condenser lenses onto the thin specimen. A number of specimen holders for specific applications are available, but the minimum requirement for crystallographic studies is a double-tilt holder that allows the specimen to be tilted about two axes which are at right-angles to each other and to the optic axis of the microscope. The first condenser contains a fixed aperture, while the second condenser is equipped with a series of interchangeable apertures of different sizes. There are usually five imaging lenses (objective, diffraction, intermediate, and projectors 1 and 2) and the final image of the object is formed on a fluorescent screen which may be viewed directly or through a pair of binoculars. The image can be recorded on a photographic film by tilting the screen out of the beam. Another series of interchangeable apertures of various sizes can be located in the back focal plane of the objective lens. The image is focused by the objective lens and the magnification controlled by the intermediate and projector lenses. The electron diffraction pattern which is formed in the back focal plane of the objective lens can be projected onto the fluorescent screen (or photographic film) by removing the objective aperture and suitably adjusting


the power of the diffraction and lower lenses. The area of specimen from which the diffraction pattern is derived can be selected by means of an aperture located in the image plane of the objective - selected area diffraction (SAD). Smaller areas can be achieved with convergent beam electron diffraction (CBED), in which the incident electron beam is focused to a fine spot on the specimen. If the convergence angle is appropriately chosen, the diffraction pattern consists of an array of nonoverlapping disks which (for specimens greater than ca. 50 nm thick) exhibit contrast which can provide information about crystal symmetry which is not available from a normal SAD pattern.

Electrons interact very strongly with solids, so the specimen must be less than about 200 nm thick in order to be usefully transparent to 100 kV electrons. Observations at the limit of instrumental resolution require the specimen thickness to be about 10 nm at 100 kV. Thicker specimens can be used at higher accelerating voltages. Microscopes with accelerating voltages up to 400 kV are currently available commercially.

There is no universal technique for preparing thin specimens from bulk samples. Metal specimens are usually prepared by electropolishing techniques and chemical polishing has been used successfully for Si, for example. Crushed fracture fragments are often used for high-resolution lattice imaging where only very small thin areas are needed due to the high magnification ( > 500 000 x ) being used. For most silicates and other minerals ion (or atom) bombardment of 25-30 Jlm thin sections (previously prepared by standard petrological methods) is now the preferred technique.

The fundamental optical principles of image formation by the objective lens in a transmission electron microscope are the same as those in a light microscope and were first formulated by Ernst Abbe in 1873. According to Abbe, the image of an illuminated object is the result of a twofold diffraction process. First, the Fraunhofer diffraction pattern of the object is formed in the back focal plane of the lens. The light waves, of course, travel beyond this plane and arrive at the image plane, where they overlap and interfere to form a magnified image of the object. The nature of the image depends upon the relative amplitudes and phases of the waves which pass through the aperture in the back focal plane. If a large number of diffracted waves pass through the aperture in the back focal plane of the objective lens of an optical microscope, it can be assumed that the image is a reasonably faithful representation of the object. However, mainly because of the aberrations of the magnetic lenses used in electron microscopes, the number of beams used to form an image must usually be restricted by the aperture in the back focal plane of the objective lens, and this, of course, sets a limit to the resolution achievable. The various imaging modes used in TEM depend upon how many and which beams are used to form the image.

The most commonly employed imaging mode is bright field (BF), in which only the central beam, usually called the transmitted beam, is allowed to pass through the aperture in the back focal plane of the objective lens. Image contrast is observed if, for whatever reason, the diffracted beams originating from different parts of the specimen vary in intensity. If many strong diffracted beams


are excited, then an extremely complex BF image may be formed. Therefore, it is usual, especially when observing images due to crystal defects, to orient the specimen to produce two-beam conditions, that is, there is only one strongly diffracted beam in addition to the transmitted beam. Dark-field (DF) images form when only this strongly diffracted beam is allowed to pass through the objective aperture. To obtain a good quality DF image, the incident beam is tilted by means of an electromagnetic device in the illumination system, so that the diffracted beam travels along the optic axis.

Since these imaging modes use only the transmitted beam or a single diffracted beam, the resulting images contain only information on a scale which is coarse compared with the spacing between crystallographic planes. In order to produce high-resolution lattice images of the planes in a crystal, it is necessary to allow at least two beams to pass through the aperture in the back focal plane of the objective. The detailed interpretation of electron microscope images produced using any of these operating modes requires as complete an understanding as possible of the diffraction process.

Origin of Image Contrast in the TEM

The wavelength associated with a beam of electrons accelerated through 100 kV is 0.0037 nm. Thus for d-spacing ofthe order of 0.2 nm, the Bragg angles () are of the order of 10 - 2 radian.

Because the Bragg angle (}s is proportional to 1/d(hkl), it is usual to express the conditions for diffraction in terms of the reciprocal lattice and the Ewald sphere construction. In reciprocal space, the Bragg reflecting plane (hkl) is specified by the vector g which is normal to (hkl) and of magnitude 1/d(hkl). Since the radius 1/A. of the Ewald sphere is large compared with 1/d(hkl) for 100 kV electrons, many diffracted beams are simultaneously excited when a monochromatic electron beam is incident upon a single crystal. However, it is usually possible experimentally to orient the crystal with respect to the incident beam so that only one diffracted beam has an intensity Ig which is comparable to the intensity 10 of the transmitted (or straight-through) beam. This situation is known as the two-beam approximation and we consider the images produced under these conditions.

The DF image is determined by Ig and reveals those parts of the specimen which give rise to that beam. The BF image is determined by 10 and if there is no absorption

(1)

and the BF and DF images are complementary. It is clear therefore that contrast will arise if different parts of the specimen are diffracting with different intensity.

10 and Ig can be calculated as a function of depth z in the crystal and deviation LJ(} from the exact Bragg angle () using the kinematical theory or the


dynamical theory of electron diffraction. Without absorption, the intensities calculated by the dynamical theory are

10 = cos2naz + (s/a)2 sin2naz

1 . 2 Ig = --2 sm naz

(atg)

where

1 a = - J 1 + (stg)2

tg

and

s = (2/,1)L10 sinO.

(2)

(3)

(4)

It can be seen that the intensities oscillate (out-of-phase) with z with a period l/a. At the exact Bragg angle s = ° and, from Eq. (3), the period of oscillation is tg the extinction distance which is given by

tg = nV cosO/ ,1F g' (5)

where V is the volume of the unit cell and F g the structure factor for the reflection g = hkl. For example, tg is of the order of 500 nm for g = 020 in the feldspars with 100 kV electrons.

When absorption is included, the F g are complex and

1 1 i -=--,+-,-,-. tg tg tg

Similarly,

a = a' + ia"

and the expressions for 10 and Ig become

10(a) = exp ( - Jlz){[cosh u + (s/a') sinh U]2 - (a\) - 2 sin2na'z}

Ig(a) = exp( - Jlz){(a\)-2 [sinh2 u + sin2na'z]},

where Jl is a linear absorption coefficient and

u = na"z

l/a" = a't't" g g

a' ~ -!- J 1 + (st~)2 tg

(6)

(7)

(8)

t: is called the absorption length and t~/t~ is of the order of 0.1. If there is no absorption Jl = 0, a" = 0, u = 0, t: --+ 00, cosh u = 1, sinh u = 0, a' = a = l/tg J 1 + (stg)2, and t~ = tg.

Io(a) and Ig(a) are plotted as a function of z for s = ° in Fig. 90a. It follows that a wedge-shaped specimen will exhibit thickness fringes whose spacing will


1.0

Ig(al

o~---+----~----~--- z

10 (al

a

b

'---=---'----'-----.L-_z

-3

-s t' g

I' 9

-2

21' 9

31' 9

Fig. 90. a Io(a) and I.(a) as a function ofz calculated from the two-beam dynamical theory of electron diffraction with abosrption, for s - O. tJt~ = 0.05 and t~ - t~. (Hirsch et al. 1965). b Io(a) and Ig(a) as a function of st~ calculated from the two-beam dynamical theory of electron diffraction with absorption. The crystal thickness is 3t~. t~/t~ = 0.1 and t~ = t~. (Amelinckx 1965)

decrease from a maximum at the exact Bragg angle as s increases. Due to absorption, the fringe contrast decreases with increasing specimen thickness.

In Fig. 90b, Io(a) and Ig(a) are plotted as a function of st~ for a crystal of constant thickness. Thus, a bent crystal will exhibit bend contours, the bright regions of which correspond to those parts of the crystal for which s is close to zero. If there is negligible absorption, the corresponding BF image will be complementary. However, it will be noted that with absorption, Io(a) is asymmetrical with respect to s, while Ig(a) is symmetrical; that is

Io(a) ( + s);6 Ig(a) ( - s)

but

Ig(a) ( + s) = Ig(a) ( - s).

This is because in Ig(a), s always appears as S2, but in Io(a), s appears as sin the product term 2(s/a')cosh u sinh u. Thus, due to absorption, the contrast at a


bend contour is symmetrical about s = 0 in a OF image, and asymmetrical in a BF image. For thick crystals, the oscillations in both the BF and OF images tend to fade out; at t ~ 10t~ they are no longer observed.

Crystal defects are made visible in TEM images by contrast mechanisms which are closely related to the origins of bend contours and thickness fringes. If the crystal planes in the neighbourhood of a defect are bent, then the diffracted intensity from the distorted region close to the defect will be different from that ofthe surrounding undistorted crystal and contrast will arise in a way similar to a bend contour. Since the diffracted intensity varies with depth in the crystal, the contrast from the defect will depend on its position in the thin foil. Even though a defect may produce no bending of the atomic planes, it may become visible because it modifies the normal variation of diffracted intensity with depth in a perfect crystal.

In order to illustrate these ideas, consider first a specimen consisting of two parts, I and II, both perfect crystals of the same material and of uniform thickness but of different orientation, joined along a planar boundary which is normal to the plane of the specimen and parallel to the incident electron beam, as shown in Fig. 91a. Suppose the objective aperture is positioned so that it allows only the transmitted electron beam to pass through, (i.e., a BF image is formed) and that only part I is diffracting strongly. Consequently, the transmitted beam from part II will be strong compared with he transmitted beam from part I, and in the BF image part II will appear bright and part I relatively dark. The contrast arises because parts I and II are diffracting differently and because there is an aperture in the back focal plane of the objective lens. If a OF image were formed using the strong diffracted beam from part I, then the image would be of opposite contrast to the BF image.

If the boundary between parts I and II were inclined to the plane of the specimen and only part I was diffracting then the wedge-shaped region of part I would give rise to thickness fringes. If both part I and II were diffracting (either with the same g and different s, or different g and different s) then the image of the boundary would again be a fringe pattern, but not a simple superposition of the thickness fringes from the two overlap'ping parts. Fringe patterns are, in general, produced by any inclined boundary and are analogous to thickness fringes. However, the fringe pattern actually produced depends upon the nature of the boundary (stacking fault, twin boundary, etc.) and the diffracting conditions. In principle, the contrast observed under different diffracting conditions can be used to determine the nature of the boundary, such as the fault vector of a stacking fault, for example. The same general discussion also applies to a specimen consisting of two parts I and II, which have different structures and compositions, the boundary between them being either coherent or incoherent.

Close to a dislocation, the crystal planes are highly distorted, as can be seen in the sketch of an edge dislocation shown in Fig. 91b. For simplicity this distortion can be considered as a rotation of the planes normal to the page -clockwise on the right-hand side of the dislocation core and anticlockwise on the left-hand side. Now suppose that well away from the dislocation the incident


II

a T(I1)

b

b

Fig. 91. a Schematic diagram showing the manner in which the aperture in the back focal plane of the objective lens gives rise to contrast in a bright field (BF) image of a specimen consisting of two differently oriented crystal grains. (McLaren 1991). b Schematic diagram of an edge dislocation of Burgers vector b. Away from the core, the beam is incident on the planes at a glancing angle O. = OD + LlO. Thus due to the bending of the planes near the core, the angle of incidence on the right-hand side is O2 < 01 and on the left-hand side it is 03 > 01•

(McLaren 1991)


electron beam is at an angle () to the planes shown that is slightly greater than the exact Bragg angle ()B; i.e., () = ()B + A(). Thus, the diffracted intensity from the region close to the dislocation on the right-hand side will be higher than the background intensity, and in a BF image the dislocation will be seen as a dark line on the right-hand side of the true position of the dislocation core.

It is important to realize that not all planes in the neighbourhood are equally distorted. For the edge dislocation shown, the reflecting planes are normal to the Burgers vector b and the contrast is strong. However, the planes normal to the dislocation line (i.e., parallel to the page) are undistorted and, in an image formed with these planes, the dislocation will be out-of-contrast. This condition for invisibility can be expressed mathematically by g. b x u = 0 where g is the reciprocal lattice vector of the operating reflection and u a unit vector parallel to the dislocation line. For a pure screw dislocation (b parallel to u) all planes parallel to bare undistorted in an elastically isotropic crystal and such a dislocation will be out-of-contrast for all reflections which satisfy the condition g. b = O. These conditions form the basis of the determination of the Burgers vector of a dislocation from the contrast observed with different reflections.

Inclusions or precipitates which strain the crystal matrix are made visible in a similar way. Small voids (or gas bubbles) which have negligible strain fields can be observed by several contrast mechanisms, such as phase contrast (in out-offocus images) and by structure factor contrast.

High Resolution Transmission Electron Microscopy

HRTEM is now generally taken to mean an imaging mode in which at least two beams are allowed to pass through the objective lens aperture. Images produced in this way show the periodicity of the lattice and, for this reason, they are often referred to as lattice images. To produce a two-dimensional lattice image, the crystal is oriented so that the incident electron beam is aligned accurately parallel to the zone axis of interest. Because of lens aberrations, it is often advantageous to limit the number of beams used to form the image and this of course sets a limit to the practical spatial resolution (Abbe theory).

An image obtained under optimum operating conditions of defocus etc. is essentially a map of the change density in the crystal projected onto the image plane - the darker regions corresponding to regions of higher charge density. However, this intuitive interpretation ofHRTEM images, though often possible, may be misleading, especially if the image detail is on a scale which is finer than the theoretical point-to-point resolution (ca. 0.3 nm) of the electron microscope. Consequently, it is now usual to compare the images observed under known conditions with images which have been computer-simulated for model structures using the Cowley-Moodie n-beam dynamical theory of electron diffraction.

HRTEM has been used most successfully in mineralogical research for identifying domains and intergrowths, even if they are only a few unit cells in extent.


Recent Developments in Transmission Electron Microscopy

During the past decade a number of important developments have been made in both TEM and SEM. Some of these developments have been applied to mineralogical problems, while others appear to be potentially useful.

Improvements in lens design have made high-resolution lattice-imaging a virtually routine operation. Consequently this technique (together with complementary computer facilities for simulating images for a range of operating conditions, specimen thickness, etc.) is now making an important contribution to mineral structure determination using very small samples or microdomains within a crystal (Buseck et a!. 1988). The commercial development of highvoltage (e.g., 400 kV) microscopes has also increased the resolution currently attainable.

Changes in lens design have also made it possible to focus the incident electron beam to a very fine spot (as small as 2 nm) on the specimen so that chemical analysis via the emitted X-rays, using energy dispersive X-ray (EDX) detectors, can be carried out from correspondingly small volumes of specimen. With the associated computer facilities, good quality quantitative chemical analysis is now routine (Williams 1984). However, a significant limitation of EDX spectrometers is that the efficiency of the Li-doped Si detector with the standard 811m thick Be window, falls off dramatically for elements oflow atomic number Z; thus only X-rays from Na (Z = 11) and heavier atoms can be detected. The need in materials science for both qualitative and quantitative analyses of light elements has stimulated the development of windowless EDX detectors (Williams 1984) and Electron Energy Loss Spectrometers (Egerton 1989), both of which are now commercially available.

At or near the exact Bragg angle, electron channeling parallel to the reflecting planes takes place (the Bormann Effect - see McLaren 1991). As a consequence, the intensities of the peaks of the emitted characteristic X-rays depend on the orientation of the crystal with respect to the incident electron beam. This effect is the basis of the relatively new technique of ALCHEMI -Atom Location by Channeling Enhanced Microanalysis (Spence and Tafto 1983). This technique has been applied to the determination of site occupancies of minor elements in olivenes (Tafto and Spence 1982) and to measurements of the degree of AljSi ordering in K-feldspar (McLaren and Fitz Gerald 1987), for example. Some difficulties in interpretation have recently been overcome by a better theoretical understanding of the electron channeling mechanism and, hence, the usefulness of ALCHEMI in mineralogy should increase. An excellent practical guide to the technique has been published recently by Otten (1989).

Perhaps the most notable development in electron diffraction techniques is the advent of micro diffraction, particularly convergent beam electron diffraction (CBED). This development is again a consequence of the ability of modern condenser lenses to focus the beam to a fine spot on the specimen. With such a convergent beam, each Bragg reflection gives rise to a disk whose size is determined by the range of angles in the cone of the illumination. For a thin

6.1.2 Scanning Electron Microscopy and Image Formation 273

specimen, the disks are featureless, but for thicker specimens in which there is strong dynamical interaction of the electron beams, each disk shows a complex variation in intensity. Because CBED patterns are derived from a very small volume of specimen, Kikuchi lines are much clearer than those in SAD patterns, which tend to wash out because of buckling of the specimen within the relatively large sampled volume. The symmetry of the contrast within the central disk in CBED and the symmetry of the pattern as a whole (which usually exhibits highorder Laue zones) is highly sensitive to crystal orientation. When the incident beam coincides with a low-order zone axis, for example, the symmetry of the CBED pattern reflects the symmetry of the crystal with great precision. Consequently, CBED patterns can be used to determine the point and space groups of the crystal. Eades (1985) has given a concise account of how to obtain CBED patterns and interpret the main features observed. More extensive reviews, including the use of CBED for determining crystal symmetry elements, have been published by Williams (1984) and Steeds (1979). McLaren and Fitz Gerald (1987) have used CBED to determine the presence, or lack, of a mirror plane in K-feldspars, and hence the crystal symmetry (monoclinic or triclinic).

The use of SAD as an analytical tool suffers from the fact that the patterns are normally recorded on photographic films which are subsequently measured and analyzed by hand. Consequently, the information contained in a diffraction pattern is not immediately available to the operator during the experiment. Although techniques for acquiring SAD data using the computer facilities available on an analytical electron microscope have been developed (Carr 1982; Hagemann 1982), this facility is not generally available, in spite of its potential usefulness in mineralogy.

A significant aspect of all the developments discussed above is that they are available (or potentially available) in a single instrument, thereby making transmission electron microscopy one of the most powerful techniques for the study of the structure and chemistry of minerals.

6.1.2 Scanning Electron Microscopy and Image Formation

G.M. PENNOCK

The design and operation of the electron gum, condenser and objective lenses in an SEM are similar to a TEM. The objective, or final lens, focuses the beam onto the sample surface and the beam is moved or scanned across an area. At any instant, signals being generated by the beam can be collected by appropriate detectors and one of them used to control the intensity on a cathode ray tube (CRT). The CRT is scanned synchronously with the incident beam so that an image of the area as it is scanned is built up on the CRT (rather like a TV image). The magnification of the image is simply the ratio of the area scanned to the area


on the CRT. The accelerating voltages used in SEM range from about 1 kV to 30 kV, beam currents from 1 pA to 10 nA and beam diameters from about 5 nm to 1 J1.m at the specimen. The normal specimen holder allows the specimen to be tilted and translated; usually a wide range of movement towards and away from the final lens is available giving afree working distance from about 5 to 30 mm between the sample surface and the final lens.

At any instant in time, the primary (incident) electrons interact with the sample in a volume, V, whose general shape and size are shown in Fig. 92a. The interaction volume and depth of penetration X of the electrons below the sample surface are a maximum when the beam is normal to the sample surface and increase as the average atomic number and density of the sample decrease and as the beam energy increases. The shape of the interaction volume is pear shaped for lower atomic number materials and more hemispherical for heavier elements. Estimates for the penetration depth of the beam at normal incidence and 20 kV for aluminum, copper and gold are 4.0, 1.5 and 0.9 J1.m respectively.

The electrons interact with the sample by both elastic (i.e., without loss of energy) and inelastic scattering. A fraction of the incident electrons are absorbed while others are scattered out of the specimen. The sampling volume is that fraction of the interaction volume that is emitted from the sample and contributes to the signal (Fig. 92b).

Backscattered electrons are produced by elastic interaction of primary electrons with atomic nuclei. Consequently, the number (i.e., intensity) of back scattered electrons increases with increasing atomic number Z, and has a cosine angular distribution about the incident beam direction for normal incidence (Fig. 92c). Many primary electrons lose energy by inelastic processes before becoming backscattered and consequently the energy distribution of back scattered electrons ranges from zero energy up to the energy of the incident beam. Tilting the sample increases the intensity of the backscattered electrons and causes the angular distribution to become asymmetrically peaked away from the beam; the contrast from differences in the average Z is reduced at grazing incidence and high contrast topographic images are obtained. Although back scattered electrons from a medium atomic weight specimen are emitted from an area of diameter ~ 1-2 J1.m, the more intense part of the emitted signal, which occurs at the point of impact of the primary beam on the sample surface, can be selectively displayed on the CRT giving an image with a resolution ~ 25nm.

Inelastically scattered incident beam electrons lose energy by producing (1) secondary electrons, (2) Bremsstrahlung or continuum X-rays, (3) characteristic X-rays, (4) Auger electrons, and (5) infrared, visible, and ultraviolet radiation.

Secondary electrons have a low energy (::;; 50 eV) and are emitted from a thin surface layer ( ~ 5 nm for metals and ~ 30 nm for insulators). The majority are weakly bound, valence electrons which are ejected from the atom after excitation by the incident beam, leaving the atom in an ionized state. Secondary electrons have an angular cosine distribution about the incident beam at normal

6.1.2 Scanning Electron Microscopy and Image Formation

Low Voltage

High Voltage

a

b

Low Atomic Num ber

BEAM

High Atom ic Number

BEAM

l~s'm',' + V v

275

Fig. 92. a The interaction volume (V) and depth of penetration (X) of the electron beam in low and high atomic number materials at low and high beam voltages. b Schematic diagram showing sampling volumes for Auger (a) secondary (b) and back-scattered (e) electrons signals. The X-ray and cathodoluminescence sampling volumes (d) are of the same order as the interaction volume for the incident beam. c Angular distribution of electrons (circle) emitted from the surface of a sample with a beam at normal incidence. The position of the solid state detector (SS) and scintillator-photomultiplier detector (SP), with its collection solid angle Q and the take-off angle 'P, are shown. d Schematic diagram showing the origin of trajectory topographic contrast in a scintillator-photomultiplier detector. Secondary electrons (SE) are deflected by the positive bias into the detector but faster, high energy backscattered (BS) electron trajectories are not affected. Absorption of the emitted electrons by the sample also affects the contrast - for example, the cavity appears dark

incidence which is not affected by surface tilt, as occurs with back scattered electrons. The number of emitted secondary electrons increases approximately with the secant angle of the surface tilt of the sample causing topographic contrast in images. Secondary electrons are sensitive to surface fields caused by voltage or magnetic differences. The spatial resolution for secondary electron images is of the order of the beam diameter.


BEAM

SS ! ~,..........

SS

c ----.;;;;;;; ........... ----Samp le

BEAM

SE

d

Fig. 92 Continued

Valence band electrons which are excited by the incident beam into the conduction band can alternatively lose their energy by emission of infrared, visible, and ultraviolet radiation, known as cathodoluminescence. CL signals are sensitive to defects and impurities which introduce new electron energy levels between the valence and conduction bands of the crystal. If inner shell, tightly bound electrons of an atom are ejected by the incident beam, the ionised atom can lower its energy by filling the vacancy with an electron from a higher atomic

6.1.2 Scanning Electron Microscopy and Image Formation 277

shell and, at the same time, giving off energy in the form of either an Auger electron or a characteristic X -ray.

Incident beam electrons which are not emitted may leak through the specimen to earth producing what is often referred to as a specimen current (SC) whose magnitude is essentially the difference between the incident electron current and the emitted electron current. The SC signal can also be used to generate an image in the usual way. If the specimen is positively biased to 50 V, the secondary electrons are prevented from leaving the surface and the SC image and the back scattered image are complementary, i.e., they are of opposite contrast.

At angles near the Bragg angle, the incident beam electrons channel between the reflecting planes in the crystal (the same phenomenon is observed during anomalous transmission in TEM). The channeling effect is only significant in the thin surface layer, about 5 nm thick. Generated signals are altered by this channeling and provide a means of examining crystal orientation, usually using back-scattered electrons. The different orientations in flat, polished polycrystalline materials cause a low level contrast in the image as a result of the different channeling behavior in the differently oriented grains. If the beam is rocked through a large solid angle about a stationary point on the sample surface a selected area channeling pattern is observed from which crystal orientations can be determined. The resolution is about OS in 5 p.m areas.

Two types of electron detector are used. The solid state detector (SS) is a semiconductor device which generates a current when struck by an electron, the current being proportional to the energy of the electron. It is placed directly below the pole piece of the final lens, a position which maximizes the detection efficiency by ensuring a high task-off angle and a large solid angle of collection. This detector has a good response to high energy backscattered electrons, but is relatively insensitive to the low energy secondary electrons. A scintillatorphotomultiplier (SP) type of detector is commonly used to collect back scattered and secondary electrons and convert them into light photons. A positive bias (~ + 250 V) on the front of the detector increases the collection of the low energy secondaries including those emitted from the surface but traveling away from the detector (Fig. 92d). The trajectory of high energy back scattered electrons are not affected by the bias and only those backscattered electrons emitted from the surface in the "line of sight" of the SP detector are collected; some of the electrons (both back scattered and secondary) may be absorbed by the sample. These trajectory effects also cause topographic contrast in the image. Application of a negative bias (~ - 250 V) prevents the secondaries from reaching the detector but has no effect on the collection of the higher energy fraction of the backscattered electrons; only line of sight back scattered electrons enter the detector, giving a high contrast, topographic image. Cathodoluminescent signals are collected by a photomultiplier. X-rays emitted by the specimen are usually collected by a solid state Si(Li) detector; the output signal is proportional to the collected X-ray photon energy so that a spectrum of X-ray intensity as a function of energy can be obtained and characteristic X-ray signals


used for chemical analysis. This technique is also used in TEM. It should be noted that not all instruments are fitted with a complete range of detectors.

After collection and amplification the detector signal is adjusted using video processing to give a range of image intensity from black to white. These processes do not add to the information in the signal but make the image more easily assimilated by the eye. For instance, a selected level in the signal can be suppressed and the remaining signal amplified to give a suitable image intensity, which is essential in the study of channeling contrast. The signal can be differentiated to enhance regions where sharp contrast occurs, adding definition (but not resolution!) to the image. Nonlinear (gamma) ramping of the signal enhances either the black or the white levels and has the effect of flattening the contrast in images; it is often used to enhance regions of low signal information, such as from a cavity.

The final variable which affects the image in SEM is the scan speed and the number of lines in the frame scan. The number of lines on the monitor need not exceed the resolution of the eye which is ~ 0.1 mm. Image integration can be used to reduce noise in beam sensitive materials which require fast scanning.

SEM instruments have a large depth of focus, F, given by 0.2/MO( mm where M is the magnification and 0( is the divergence angle of the beam; 20( ~ D/H, D being the final lens aperture diameter and H the distance between the final lens aperture and the sample. F ~ 20 pm for M = 1Ooo, H = 10 mm and D = 200 pm. Depth of focus is increased at long working distances but these

conditions decrease the resolution since the beam diameter is large at the sample surface. Bearing such compromises in mind, the optimum spatial and/or atomic number resolution which can be obtained mainly depends on the quality of the electron-optics, the contrast developed in the specimen/detector system and the sampling volume of the signal from within the specimen. The contrast level is often the most significant limitation to the resolution and this is very specimendependent.

Specimen preparation is generally minimal with SEM. Samples are fixed to the specimen stage in an orientation which will optimize the generation and collection of the signal of interest. The samples need to be conducting to avoid charge build-up and possible thermal damage. Gold is often used as a coating when topographic information is required, as it gives a good secondary electron signal. The coating needs to be thin (about 10-20 nm), continuous, and show a minimum of structure. Carbon is used when information about the composition is required, as with back scattered electron or X-ray signals, as the gold layer absorbs and obscures the signals from the sample. Other alternatives to coating include operating at low accelerating voltage, using the backscattered electron signal or operating at a fast scanning rate and integrating the signal. Channeling contrast and other weak contrast techniques require samples with a good, damage-free surface polish.

Technical developments now commercially available include, for example, the environmental SEM, the liquid nitrogen cold stage and on-line computer analysis of channeling patterns. These techniques are potentially useful, for

6.1,3 Applications of Transmission Electron Microscopy 279

example, in identifying fluid inclusions and in determining the preferred orientation of crystal grains in a rock

6.1.3. Applications of Transmission Electron Microscopy

H.R, WENK

The main use of the electron microscope in mineralogy is to investigate defects which may be introduced during growth, during phase transformations or during deformation, There is a large literature which has been reviewed in monographs (e,g" Wenk 1976, 1980; White 1985; McLaren 1991; Buseck 1992), We will give a brief overview of some of these applications by describing the main crystal defects of interest, their geometry and origin and illustrate TEM characterization with examples, for convenience mainly from our own resear~h, Defects affect physical and chemical properties, For example, mechanical properties can only be understood in terms of the types and arrangement of dislocations, Defects also influence the stability of the crystal and modify the reactivity, especially important for fluid-rock interactions, Structural defects can often be used to infer the history of a mineral which underwent changes during cooling or during deformation,

Another application of the TEM is the direct determination of crystal structures which has become possible through high resolution electron microscopy (e,g" Buseck et aL 1988), Again a few examples will highlight some of the possibilities and recent developments,

Lattice defects are normally classified in terms of their geometry, Point defects are single "mistakes" in the regular and periodic positioning of atoms, These take the form of a missing atom or vacancy, a foreign atom, or an interstitial atom introduced between normal lattice sites (Fig, 93a), As individuals, point defects are difficult to observe, but when numerous they affect macroscopic properties (e,g" chemical composition and electrical resistivity),

• •••••• • •••••• ••••••• • •••••• ••••••• • •••••• SF • •• • •• ••••••• • •••••• • • • • • • • • • • ,L • •• • •••••• • ••• • • • • ••••• • •••••• • • ••• • •• • ••••• • •••••• • ••• • • • • ••••• • ••••••

(0) (b) (c)

Fig. 93a-c. Basic types of crystal defects, a Point defects: interstitial atom and vacancy, b Line defect: edge dislocation, c Planar defect: stacking fault (SF)


With line defects or dislocations (Fig. 93b) the perfect lattice is disrupted and displaced along a line. Other planar defects can occur (e.g., twin boundaries, stacking faults, antiphase boundaries) which are characterized by displacement over a whole plane of atoms. Linear and planar defects cause substantial local distortions of the crystal structure and this strain can be imaged with the transmission electron microscope, making use of diffraction and brightfield and darkfield imaging.

Deformation Microstructures

Dislocation microstructures are indicative of the strain history and of deformation mechanisms. The TEM has been used widely to investigate dislocations in minerals. Along with illustrations of examples it is appropriate to review very briefly some basic principles of dislocation theory (for more extensive treatments, see e.g., Hull 1965; Barber 1987).

When single crystals are subjected to shearing stresses, plastic deformation occurs on one, or more, sets of crystallographically defined slip planes (hkl), which each have a specific slip direction [uvw]. Large discrepancies between the observed strengths of crystals and theoretical strengths calculated on the assumption that slip occurred instantaneously right across slip planes, led Taylor (1934) and others to introduce the concept of dislocations. Figure 94 shows a slip plane, over part of which slip has occurred. Slip has transposed the upper part of the crystal over the lower part by one unit cell in the direction of the vector b, so that the lattice perfection is restored across the slipped area except along the dislocation line AB, which defines the limits of the slip. The vector b is called Burgers vector. At A where the dislocation line is perpendicular to b the lattice distortion is effectively caused by an extra vertical plane and such a distortion is called an edge dislocation. At B, where the dislocation line is parallel to b, the lattice planes perpendicular to it are distorted and form a continuous helix, so that the dislocation is a screw. It is apparent that the character of a dislocation and of the strain field around it changes as its orientation changes with respect to b. In general, a dislocation will have edge and screw components. Slip of the whole upper part of the crystal in Fig. 94 over the lower part occurs when the dislocation glides right across the slip plane. Ultimately all the bonds across the plane have been broken, but the process occurs only in the vicinity of the dislocation, so that relatively low stresses are required to cause slip and to effect small changes in crystal shape.

For a quantitative characterization of deformed crystals it is necessary to determine the slip plane (hkl) and the Burgers vector [uvw]. Lattice strain along the dislocation line can be imaged with the TEM. Figure 95 illustrates dislocations appearing as lines with terminations in experimentally deformed dolomite. Dislocations are concentrated in a slip plane that can be identified by trace analysis by tilting the specimen into various orientations and identifying the corresponding orientation of the crystal with diffraction patterns. As outlined

6.1.3 Applications of Transmission Electron Microscopy 281

Fig. 94. Propagation of a dislocation on a slip plane (hkl) partway across a crystal. A portion of the crystal above the slip plane has moved one lattice unit in the direction of the Burgers vector b over the lower portion of the crystal, thereby creating a dislocation line. The parts of the dislocation with edge (A) and screw character (B) are shown

.€l

1105 0114 • • •

1011 • • •

• • •

. , 1;.- I , I ' .

~. . / ' . :' ''''' ~

.~. _"'\l1lI 25'0 nm

~ •

'i4

'. Fig. 95a-d. Contrast experi· ments to image dislocations in a single crystal of dolomite, deformed experimentally at 600 0c. The mottled contrast is due to incipient decomposition (Wenk et al. 1990). a -Symmetrical (zone axis) electron diffraction pattern, < - 4 - 591) zone. b Same as a

but tilted into a two-beam condition with reflection g = 01-14 preferentially oper-

ating. c Image under diffraction conditions b in bright field mode. d Image under the same conditions with g = 01-14 operating in dark

field mode


above the visibility and contrast of dislocations depends critically on the diffraction conditions and two-beam conditions are most informative. Figure 95a shows a symmetrical (zone axis) diffraction pattern from dolomite (obtained from an area 1 Jl.m diameter); in Fig. 95b the specimen has been tilted into a "two-beam" condition with only the strong O1I4 reflection operating in addition to the main beam. By placing an aperture around the latter, a brightfield (bJ.) image is obtained (Fig. 95c), in this case showing some dislocations in strong contrast. If the main beam is deflected to direct the diffracted beam through the aperture, we obtain a darkfield (dJ.) image (Fig. 95d) in which the contrast is exactly reversed if the specimen is extremely thin, so that all the electrons suffer no loss in energy in passing through it (elastic scattering).

The Burgers vector b of dislocations can be determined by means of contrast experiments, using two-beam conditions. Defects are generally not visible if the nermal to the diffracting plane hkl (the corresponding diffraction vector g) is parallel to the Burgers vector. In this precise orientation the specimen appears to be a perfect crystal viewed in the direction of the electron beam. As described in section 1 pure screw dislocations are invisible if g' b = 0, i.e., if the Burgers vector [uvw] lies in the diffracting planes (hkl). If zero contrast is obtained for two separate reflections, then the Burgers vector is in the direction defined by the intersection of the two planes. In general, g' b x u = 0 where u is a unit vector in the direction of the dislocation line. This criterion is often neglected because, at least for low order reflections, the contrast contribution is small. The example shows how a slip system (hkl) [uvw] can be identified with the TEM.

In general, Burgers vectors correspond to lattice vectors but sometimes the magnitudes of the Burgers vectors are only a fraction of a lattice vector. Such dislocations are called partial dislocations, and as they propagate they leave behind a stacking fault (Fig. 96a) or, in ordered crystals, an antiphase boundary (Fig. 96b), planar defects across which the lattice is faulted and displaced. The glide of a second partial dislocation, trailing behind and terminating the fault, restores lattice perfection. Since the strain energy of a dislocation increases with the magnitude of the Burgers vector, it is often energetically favorable for dislocations to propagate as two partials rather than a unit dislocation such as for basal slip in dolomite (Fig. 97a) .

••••••• • • • • • • • .0.0.0 . ••••••• • •••••• .0.0.0. ~~-~~-~~-. !_._ !T· • • • SF __ ._~T. 0.0. APB ....... - ....... - .0.0. 0.-••••••• • •••••• .0.0.0. ••••••• • •••••• .0.0.0. • • • •••• • •••••• .0.0.0.

(Q) (b) (c)

Fig. 96a-c. Propagation of edge dislocations under a shearing stress on a slip plane (dashed lines). a Unit dislocation. b Partial dislocation with stacking fault (SF). c Partial dislocation in an ordered crystal with an antiphase boundary (APB)


1 pm , .--1 •

Fig. 97a-d. Dislocations in minerals. a DF micrograph of partial dislocations and stacking faults (with fringes, on 0001 planes) in dolomite deformed at 420 °C. b BF micrograph of dislocations which are concentrated in deformation bands in dolomite deformed at 600 dc. Two slip systems (c = 0001 and fl = 10- 12) are viewed edge-on, the third one (f3 = 01-12) is inclined. c Dislocation microstructures with dense tangles and f = {01- 12} twins in dolomite deformed at 500 dc. d Dislocation network and subgrains, indicative of recovery, in naturally deformed quartz from a mylonite. (a- c Barber et al. 1981, d Wenk et al. 1983)


Climb Force · . . . .. ~ ..... . • ••••• • ••••• • ••••• • ••••• __ ~!T_·_·_· __ •••••• • • • •• •• ~ __ -_ ~ ! T!. ! ~ • •••••• • •••••• • •••••• • •••••• • •••••• • •••••• • •••••• • ••••••

a

•••••• - ~ ~ ~~ ~ _0_ 1 · ..... . · ..... . · ..... . •• • • • •• t -.-;;.1;;;-• •••••

b

c Fig. 98a-c. Dislocation climb to a different slip plane by diffusion of vacancies (a) can eliminate dislocations of opposite sign (b) and produce regular low energy arrays such as edge dislocations stacked on top of each other (c)

More information about deformation mechanisms is obtained by investigating interaction of dislocations. At low temperatures, the movement of dislocations is restricted to the particular slip planes for the crystalline material and dislocations are often concentrated in slip or deformation bands (Fig. 97b). If more than one slip system is present, cross-slip of dislocations may occur and can lead to dense tangles of dislocations (Fig. 97c) and concentrations of strain energy which can be reduced if dislocations are able to move out of their slip planes. This is achieved at higher temperatures by diffusion of vacancies to the core of the dislocations (Fig. 98a) which causes them to climb. This additional degree of freedom enables dislocations of opposite sign to move closer together and disappear upon combining (Fig. 98b). Dislocations ofthe same sign arrange themselves into arrays (Fig. 98c) which appear often as networks with a lower total energy and therefore fairly stable against annealing. While slipbands are typical of low temperature deformation, the significance of regular networks, loops and dipoles is that they are associated with thermal activation and diffusion, such as exist in deformed high grade metamorphic rocks.

Transformation Microstructures

Many minerals undergo phase transformations during cooling. At high temperature they are often structurally disordered and chemically homogeneous. At low temperature they may be ordered with a different symmetry or a crystal which is homogeneous at high temperature may separate into domains of different composition. Both types of transformations may introduce characteristic planar defects which can be analyzed with the TEM.


: : :~: : : E)E)e~

~::: a

.A 08

-A8

.0. 0 • 0

o • 0 • 0 •

o • 00. 0

.0. 0 • •

o • 0 • 0 •

b

APB

Fig. 99. a A disordered crystal (AB) begins to order (A, B) in two regions. b Upon coalescence the ordered domains form an APB

Ordering Transitions. A rearrangement of atoms from a disordered to an ordered state is often accompanied by a loss in symmetry. This loss in symmetry can be due to either compositional ordering (such as the AI-Si distribution in feldspars) or positional ordering (such as the arrangement of Si-tetrahedra in a and f3 quartz). Once ordering commences from a number of independent sites in a disordered crystal (Fig. 99a) the ordered regions, which are called domains, will continue to expand until they impinge on each other (Fig. 99a, b). The domains will either coalesce into a single domain (this occurs when the atomic sequences at the interfaces are perfectly "in step") or will be separated by a boundary. An ordered crystal can consist of one domain or of many domains. The number and size of domains is dependent upon the cooling history.

Domains are separated by boundaries and these boundaries can be classified by the symmetry operation which brings two unlike domains to coincidence. Interfaces between domains which can be brought to coincidence by a translation are called antiphase boundaries or APBs (Figs. 99b and 100a); boundaries between domains related by an inversion operation are inversion boundaries (Fig. 100b); if a lattice rotation brings two domains to coincidence we call them twin boundaries (Fig. lOOc). The type of variant depends on the symmetry change of the space group during the phase transition. Boundaries have characteristic contrast when imaged with the TEM (e.g., Amelinckx and Van Landuyt 1976). Of importance is the background shade on either side of the boundary, the fringes accompanying the boundary and the conditions for visibility. Some basic rules are summarized in Table 12.

Transformation boundaries may be curved or straight, depending on their strain energy and the conditions pertaining and subsequent to the transformation. Two types of APBs can be recognized: an APB is conservative if the displacement vector R is parallel to the fault and the surrounding lattice is merely geometrically distorted (Fig. lOla); conversely, an APB is nonconservative if R does not lie in the fault plane. In the latter case the APB can be considered as if a layer of material has been inserted or removed (Fig. 101b), which changes the local composition. This is particularly significant if APBs are periodic because it can represent bulk nonstoichiometry.

Periodic planar defects may be the result of phase transformations (e.g., in intermediate plagioclase, Fig. 102a) or growth (e.g., in J dolomite, Fig. 102b). Commonly, periodic APBs are expressed in diffraction patterns: if faults are in


a

Fig. l00a-c. Depending on the change in space group symmetry during ordering, ordering may produce a APBs (in this case from a displacive transformation), b inversion boundaries (l B) and c twin boundaries (TB). For twin boundaries the shear vector increases with distance from the twin plane

Table 12. Contrast on planar interfaces

Type of interface Fringe Background SAD Extinction condition

Translation: APB n Fringe Same No effect g·R = n

BF& DF (unless periodic) All main reflections symmetrical

SF IX Fringe Same No effect g·R = n BF symmetrical (unless periodic) DF not

Rotation: grain boundary Thickness fringes Different 2 patterns No extinction

twin boundary Thickness fringes Generally: different 2 patterns g·R = 0 edge-on view: same 1 row in

common

Inversion IX-like fringe Same in BF No effect If projection is ("black-white" different in DF centrosymmetric

contrast)

rational lattice planes and repeat in multiples of a unit cell dimension (e.g., substituting every third Mg layer by a Ca layer in dolomite) superstructure reflections appear in rational positions and can be indexed (Fig. l02a). If the modulation is incommensurate (such as in intermediate plagioclase, Fig. 102b) satellites appear which are symmetrical about the main reflections and do not have rational indices. If the APBs are not strictly periodic, streaking is observed


a ••• 0 0 0 b • • • 0 0 0

• • • o 0 R't. • • • • • 0 0 0

000 • • • • • • • • .00 0 0 0 0

000 • • • • • •

• • 0 0

• • • • 0 0

• •

• 0

IR 0

•

Fig. IOla,b. If the displacement vector R lies in the fault plane the APB is called conservative (a); if it does not lie in the fault plane it is nonconservative (b) and appears as if an extra layer has been inserted or removed

Fig. I02a,b. Periodic planar defects. a Commensurate periodic 0001 APBs in domains of b dolomite (Van Tendeloo et al. 1985). b Incommensurate APBs in metamorphic intermediate plagioclase (Wenk and Nakajima 1980). Dark field images and corresponding diffraction patterns are shown

in the diffraction pattern (in the case of b-dolomite streaking is in addition to the superstructure reflections, Fig. l02a).

Feldspars are an excellent example to illustrate planar defects. Consider, for instance, calcic plagioclase feldspar. At high temperature the space group is that of albite, CT, and the AI-Si distribution is disordered. At low temperature Al and Si are ordered to avoid AI-O- AI bonds and the space group is II with a unit cell twice as large (Fig. 103a, inset). Upon cooling, various sites in the disordered structure will acquire preferences to be either Al or Si and around these nuclei, ordered regions with alternating AI-Si will grow, creating ordered domains.

288

e Si

OAI

Chapter 6. Electron, Acoustic, and Tunneling Microscopy of Minerals

----m

Fig. I03a,b. Al- Si ordering in the tetrahedral framework of feldspars produces a b-APBs in lunar anorthite (courtesy of W. Muller) and b albite and periciine twinning in microciine (McLaren 1984). The ordering pattern across the faults is schematically illustrated in insets

These domains can be brought to coincidence by a 1/2b translation and the boundaries separating them are APBs (Fig. 103a). In potassium feldspar the C2/m (monoclinic sanidine) -+ CI (triclinic microcline) transition involves a rotation variant and produces twin boundaries (Fig. 103b).

Another example is perovskite, which is cubic at high temperature, transforms first to a tetragonal and then to an orthorhombic structure through displacements of oxygens from special to general positions. Finally the structure changes from a centrosymmetric to a noncentric geometry. Figure 104a shows resulting defects such as two types of twin boundaries (121 for cubic to tetragonal and 101 for tetragonal to orthorhombic), APBs, and inversion boundaries, the latter on a very fine (modulated) scale.

Depending on the nature of the transition, changes are sluggish (most chemical ordering which requires diffusion) or rapid and reversible (such as small atomic displacements). The r:x-{J transition in quartz at 573 °C can be observed in situ in a TEM equipped with a heating stage and proceeds through an intermediate stage with twins (Fig. 104b).


Fig. 104. a APBs (APB), (101) and (121) twin boundaries, and inversion boundaries visible as a modulated structure (M) in niobian perovskite. It is assumed that most of these planar defects formed during phase transformations (Hu et al. 1992). D is a dislocation. b Displacive rx-fJ transition in quartz observed in situ in the electron microscope at 573 °C. The transformation is associated with Dauphine twinning, which is influenced by the dislocation microstructure. (Barber and Wenk 1991)

Exsolution. So far we have discussed interfaces between domains of similar structure and composition. Another type of domain forms during exsolution from a chemically homogeneous phase. Classical examples occur in pyroxenes (pigeonite/enstatite-augite) and perthite: (albite-microcline/sanidine). There is a slight mismatch in the lattice parameters of the two chemically different phases, so that the interphase interface is not strictly coherent and dislocations occur in the interface to accommodate the latticl~ mismatch. The interfaces are generally parallel to the surface of least lattice distortion, resulting in regular lamellar structures. When coarse scale (1-10 mm) lamellar structures occur, they are often interpreted as the result of heterogeneous nucleation and growth of the equilibrium phases (Fig. 105a). On the other hand, very fine lamellar structures with diffuse interfaces (Fig. 105b) are attributed to spinodal decomposition below the equilibrium solvus, at relatively high undercooling.

The cooling history can often be inferred from the morphology of the microstructures which form during phase transformations. Take for example an alkalifeldspar of intermediate composition. The microstructure (Fig. 106a) is compared with the phase diagram (Fig. 106b). A first exsolution into Na- and Krich semicoherent lamellae occurred at high temperature (Abl). During ordering of sanidine, a second exsolution occurs on a finer scale with coherent lamellae of albite (Ab2) and monalbite (Mb2). The monoclinic symmetry of monalbite is preserved due to strain from neighboring orthoclase (Or). Orthoclase shows fine modulations due to incipient twinning during conversion to triclinic microcline.


Fig. IOSa,b. Exsolution in pyroxenes. a (001) ex solution lamellae in lunar pigeonite. They probably nucleated on a subgrain boundary which crosses the picture. b Fine modulated tweed structure in lunar pigeonite, indicative of spinodal decomposition. (Champness and Lorimer, 1971)

The chemical composition can be measured in TEMs equipped with energy dispersive X-ray detectors. With a STEM which adds the high beam localization of a SEM to a TEM, quantitative analyses can be performed on areas as small as 20nm.

Growth Defects

All natural crystals contain growth defects, and to the extent that they remain unmodified by other processes, these defects are a record of growth. Particularly as crystals grow from aqueous solutions at low temperature where kinetics are unfavorable, various types of lattice defects are introduced into the structure. These include dislocations, stacking faults, sequential stacking disorder, twins, and point defects. Dislocations having a screw component are important since they often facilitate growth. Such dislocations can provide useful information about growth directions. Growth bands, reflecting local variations in composition or impurity content, mark past positions of growth surfaces and contribute to the overall growth microstructure. They are well known in metallic systems where growth from an impure melt leads to constitutional supercooling with consequent periodic fluctuations in growth rate and impurity concentration (Hurle 1962) and also occur in aqueous systems. In low-temperature rhombohedral carbonates domains with different local composition and structure have been attributed to changes in composition of the aqueous solution during growth (Fig. lO2a and lO7a). Slight differences in composition (and hence interplanar spacing) or in orientation of lattices in adjacent sectors may lead to the formation of stacking faults. Stacking disorder with a nonideal repeat

6.1.3 Applications of Transmission Electron Microscopy

b 900

800

~ 700 o ~ E500 ::t

soo

1.00

300

1 monoclinIc feldspar

," "': .• -'""'--, ,,:, ,- ..

o 10 20 30 ~o ~O 50 70 80 90 100

mole ",,0

291

Fig. l06a,b .. Exsolution and ordering in perthite. a DF micrograph of mesoperthite from Sri Lanka with coarse lamellae of twinned albite (Abl) and orthoclase (Or), and a second generation of fine Na-rich lamellae of albite (Ab2) and monalbite (Mb2) which exsolved during ordering of sanidine. Orthoclase shows a modulated tweed structure due to incipient twinning. (Courtesy of C. Evangelikakis). b Phase diagram of alkalifeldspar. (Champness and Lorimer 1976)


. - ." . ----" -.-- . . -.-. 7A 14A

...

Fig. I07a-c. Heterogeneous phases in minerals which originated during growth at low temperatures. a Fe-rich dolomite with domains of y-dolomite. Left side low magnification DF image, highlighting domains of the y superstructure (see weak and diffuse reflections in diffraction pattern). Right side atomic resolution [0001] image with different cation ordering patterns in dolomite (D) and y dolomite (C). Image simulations are inserted. (Wenk et al. 1991). b Mixed layer illite-smectite (7/ 14A) and R 1 illite/smectite, both replacement products of illite in clay. (Jiang et al. 1990). c Fibrous crystals of opal CT in cherts from the Monterrey Ftn. The diffraction pattern illustrates presence of both cristobalite c and tridymite (t)


sequence of a given structural unit in a crystal is particularly common in clay minerals (Fig. 107b) and metastable silica minerals in sedimentary rocks (Fig. 107c). Stacking disorder is predominant if the growth direction is parallel to the stacking direction of interest.

Structure Determination

With improvement in electron microscope design, point-to-point resolution sufficient to resolve interatomic spacings has been obtained in many materials. Although individual atoms may be resolved under favorable conditions, determination of the crystal structure from the images is seldom straightforward. Traditionally, two-dimensional images are obtained for a range of thickness and focus and a series of experimental images is then compared with contrast calculations based on the multibeam dynamic theory. A qualitative match of experimental and simulated images has been considered as good evidence that the structure model is correct. It is often advantageous to reconstruct the periodic structure by digitizing experimental HREM images (Fig. 108a), Fourier transforming them and then back-transforming them from structure factors determined at reciprocal lattice points. This removes nonperiodic noise and improves image quality (Fig. 108b). The simulated and experimental [100] zone axis image of staurolite at Scherzer focus and a thickness of 100A show good agreement (Fig. 108). This technique was applied to minerals and the study of biopyriboles and clay minerals have been excellent examples.

Unfortunately, such a procedure does not enable us to assess the quality of the model structure quantitatively. Also there is no way to analytically derive a model from the image. Furthermore, the image corresponds to a 2d projection of the structure in which information is invariably lost due to the superposition of

Fig. lOSa-c. Images of the structure of staurolite as viewed along [001]. lEOL ARM-lOOO operating at 800 kV, Scherzer focus. 8 Experimental image in thin area ( < 40 A). b Image reconstructed using structure factors from the computed Fourier transform of 8. c Multibeam dynamic contrast calculation assuming the structure of staurolite, microscope conditions as described above and specimen thickness of 30 A. (Downing et al. 1990)


---x

I , y

a b

S i

d

AI O.SAI Si Fe o


atoms at different levels in the unit cell in the beam direction. With resolution of 1.5 A, as it is available in some microscopes, it should be possible to resolve single atoms which are typically spaced 2-3 A, if one could obtain a threedimensional deconvolution of the information contained in high-resolution images. The basic necessary information is not contained in a single image but in a combination of images which view the structure in different directions. Biophysicists have developed such a method to investigate the structure of proteins.

This so-called electron crystallography resembles an X-ray structure determination in which the crystal structure is reconstructed by a Fourier synthesis from amplitudes and phases of diffracted waves obtained directly from TEM images. Information on phases and amplitudes from several images, representing different views of the specimen, is combined and then Fourier transformed to a three-dimensional map of the Coulomb potential of the crystal which resembles the electron density.

In the case of proteins, resolution is mainly limited by the sensitivity of organic materials to damage by the electron beam. The same method has recently been applied to determine the structure of the mineral staurolite. In this case, a much higher point-to-point resolution can be obtained. However, a condition for 3d electron crystallography is that dynamic scattering is minimal and that the phase object approximation is valid which is only true for very thin crystals. With 39 structure factors from five projections and d-spacings larger than 1.6 A a 3d potential map of staurolite was calculated (Fig. 10ge). Since all atoms are located near z = 0 and z = 0.25 in the unit cell and because of assumed orthorhombic symmetry, most of information is contained in these two xy Coulomb potential sections which are compared in Fig. 109a-d with the structure model. There is an excellent correspondence between the experimental potential map and the crystal structure:, with all atoms, including oxygens, clearly resolved. The method has great potential for determination of unknown crystal structures, particularly where homogeneous regions exist only in submicron-sized domains as is the case in many minerals. It is estimated that 3D structure determinations should be possible on small domains only ten unit cells wide and should not only resolve atom positions but also site occupancies. This structural analysis has also been approached with channeling enhanced microanalysis which is, however, much more limited. Another application of electron crystallography is space group determination from structure factor phases in cases where evidence from diffraction patterns is ambiguous. In this context the convergent beam diffraction method ought to be mentioned, which also can be used for symmetry determinations on very small volumes.

Fig. 109. a,b Sections displaying the Coulomb potential of staurolite as derived from the 3d structure reconstruction with 39 structure factors from 5 different images. Sections at z = 0 and z = 0.25 (units) are shown, each corresponding to a slice 0.2 A thick. c,d The corresponding crystal structure indicating cation and oxygen positions. e Three-dimensional view of the Coulomb potential of staurolite. (Downing et al. 1990)


6.1.4. Applications of Scanning Electron Microscopy

G.M. PENNOCK

Knowledge of the size and shape of grains and pores in rock aggregates is important in gaining an understanding of the mechanical behaviour of the lithosphere and in metamorphic processes and ore genesis. The large depth of focus and topographic detail which can be attained with SEM is essential in this type of study. For instance, the pore topography along a grain interface in a deformed quartz aggregate has been studied. The aggregate was isostatically hot pressed at 1200 K and 300 Mpa confining pressure and 200 Mpa pore water pressure. The nature of the microstructure of the interface and the pore topography vary according to the grain size and the deformation conditions. This information is useful in determining fluid-rock interactions and permeability in deforming grain aggregates.

The back scattered signal in conjunction with energy dispersive X-ray spectroscopy (EDS) provides a powerful means of examining and identifying the shape, distribution and composition of phases present in a polished sample. The average atomic number contrast in a picritic basal can be observed using the back scattered electron signal. The background glassy matrix region has the lowest average atomic number and showed some devitrification into fine scale regions of higher average atomic number. The blocky, central grain is lighter than the matrix and therefore of higher average atomic number. EDS was consistent with the large grain being olivine altered along the cracks into serpentine of a lower average atomic number than the olivine. Compositional zoning of the olivine is evidenced from the increase in average atomic number towards the grain edge. The more faceted phase had an average atomic number generally lower than the olivine and EDS was consistent with the phase being clinopyroxene. The needle-like laths showed the highest average atomic number had an EDS spectra consistent with that of illmenite.

Recent interest in the high pressure phase behaviour of wiistite has led to a study of the pressure dependence of the elastic moduli using ultrasonic interferometry on single crystals. The perfection of the crystals were checked using the Laue X-ray technique and electron channeling. Low magnification electron channeling pattern can be obtained using the scanning transmission mode (STEM) in a TEM instrument. The material was a wiistite boule in which the elastic moduli data showed anomolous isotropic behaviour. The top lefthand corner shows a star arrangement of bands of the [100] cube zone axis. A discontinuity in the bands is seen to run E-W across the middle of the micrograph. Analysis of other patterns showed discontinuities of more than 10 degrees across several subgrain boundaries which accounted for the isotropic behaviour. It is interesting to note the effect differentiation of the signal had on the channeling pattern. The channeling band from one of the cube axis of the

6.1.4 Applications of Scanning Electron Microscopy 297

upper subgrain was parallel to the E-W scanning direction. In consequence, signal differentiation did not enhance th,e contrast and this band is therefore extremely weak compared to the other bands in the channeling pattern. Channeling patterns and backscattered Kikuchi patterns are increasingly used in texture analysis to determine crystal orientations in polycrystalline aggregates.

References

Ahn JH, Peacor DR (1986) Transmission and analytical electron microscopy of the smectiteto-illite transition. Clays Clay Mineral 34: 165-170

Amelinckx S, Van Landuyt J (1976) Contrast effects at planar interfaces. In: Wenk HR (ed) Electron microscopy in mineralogy. Springer, Berlin Heidelberg New York, pp 68-112

Barber DJ (1987) Dislocations and microstructures. In: Wenk HR (ed) Preferred orientation in deformed metals and rocks: an introduction to modern texture analysis. Academic Press, Orlando pp 148-182

Buseck PR (Ed) (1993) Minerals and reactions at the atomic scale: Transmission electron microscopy. Rev Mineral 27

Buseck P, Cowley J Eyring L (eds) (1988) High resolution transmission electron microscopy and associated techniques. Oxford Univ Press, New York 645 pp

Downing KH, Hu Meisheng, Wenk HR, O'Keefc: MA (1990). Resolution of oxygen with the TEM: 3d-electron crystallography of staurolitf:. Nature 348: 525-528

Drits VA (1986) Electron diffraction and high resolution electron microscopy of mineral structures. Springer, Berlin Heidelberg New York

Glaeser RM (1985) Electron crystalography of biological macromolecules. Annu Rev Phys Chem 36: 243-275

Goldstein n, Newbury DE, Echlin P (1981) Scanning electron microscopy and X-ray microanalysis. Plenum Press, New York

Hull 0 (1965) Introduction to dislocations. Pergamon Press, Oxford Hurle DTJ (1962) Mechanisms of growth of metal single crystals from the melt. Pergamon

Press, New York Loretto MH (1984) Electron beam analysis of materials. Chapman and Hall, London McLaren AC (1991) Transmission electron microscopy of minerals and rocks. Cambridge

Univ Press, New York Schwarzer RA, Weiland H (1988) Texture analysis by the measurement of individual grain

orientations - electron microscopical methods and application on dual-phase steel. Text Microstruct 8/9: 551-577

Spence JCH, Tafto J (1983) ALCHEMI: a new te(:hnique for locating atoms in small crystals. J Microsc 130: 147-154

Steeds JW (1979) Convergent beam electron diffraction. In: Hren n, Goldstein n, Joy DC (eds) Introduction to analytical electron microscopy. Plenum Press, New York, pp 387-422

Taylor GI (1934) Plastic deformation in crystals. Proc R Soc Lond 145: 362-404 Thomas G, Goringe MJ (1979) Transmission electron microscopy of materials. Wiley, New

York Von Heimendahl M (1980) Electron microscopy of materials, an introduction. Academic

Press, New York Wenk H-R (ed) (1976) Electron microscopy in mineralogy. Springer, Berlin Heidelberg New

York Wenk HR (1979) Elektronia Mikroskopia B Mineralogii. MIR, Moscow, 541 pp (in Russian) White JC (ed) (1985) Short course in applications of electron microscopy in the Earth Sciences.

Mineral Assoc Can, Fredericton


References for Figures

Barber DJ, Wenk H-R (1991) Dauphine twinning in deformed quartzites: implications of an in situ TEM study of the rr.-p phase transformation. Phys Chern Mineral 17: 492-502

Barber DJ, Heard HC, Wenk H-R (1981) Deformation of dolomite single crystals from 200-800°C. Phys Chern Mineral 7: 271-286

Champness PE, Lorimer GW (1971) An electron microscopic study of a lunar pyroxene. Contrib Mineral Petrol 33: 171-183

Champness PE, Lorimer GW (1976) Exsolution in silicates. In: Wenk HR, Champness PE, Christie JH (eds) Electron microscopy in mineralogy. Springer, Berlin Heidelberg New York, pp 174-204

Jiang WT, Peacor DR, Merriman RJ, Roberts B (1990) Transmission and analytical electron microscopic study of mixed-layer illite-smectite formed as an apparent replacement product of diagenetic illite. Clays Clay Mineral 38: 449-468

McLaren AC (1984) Transmission electron microscope investigations of the microstructures of microlines. In: Brown WL (ed) Feldspars and Feldspathoids. Nato ASI Series, Reidel, pp 373-409

Meisheng H, Wenk H-R, Sinitsyna D (1992) Microstructures in natural perovskites. Am Mineral 77 (in press)

Van Tendeloo G, Wenk H-R, Gronsky R (1985) Modulated structures in calcian dolomite: a study by electron microscopy. Phys Chern Mineral 12: 333-341

Wenk H-R, Nakajima Y (1980) Formation and decomposition of APB-structures in calcic plagioclase. Phys Chern Mineral 6: 169-186

Wenk H-R, Barber DJ, Reeder RJ (1983) Microstructures in carbonates. In: Reeder RJ (ed) Carbonates: mineralogy and chemistry. Reviews in mineralogy. Min Soc Am 11: 301-361

Wenk H-R, Meisheng H, Lindsey T, Morris W (1991) Superstructures in ankerite and calcite. Phys Chern Mineral 17: 527-539

6.2 High Resolution Acoustic Microscopy

U. BELLER

In recent years, scanning acoustic microscopy has become a powerful tool for material characterization on the micrometer scale. The advantage of this technique is the ability of ultrasound to penetrate materials which are opaque to light and electron beams. Therefore it is possible to detect pores, cracks, and inclusions located beneath the surface.

Generally, there are two different types of scanning acoustic microscope, one acting in transmission mode and the other in reflection mode. The latter is the most commonly used and its principles are briefly described here.

Scanning acoustic microscopy uses sound waves in the range of a few hundred Mhz up to a few Ghz. The high frequency sound field is excited by means of special piezoelectrical transducers. The transducer is sputtered directly onto the rear side of the lens (Fig. 110).

The plane acoustic waves emerging from the transducer pass the lens towards the spherical cavity ground on the front side. They are focused onto the object by using a coupling fluid, water (sound velocity: 1500ms- 1) being the

6.2 High Resolution Acoustic Microscopy

Ac oustic 'Waves

Acoustic An tireflection LaY 5.!er-+-_~

output to Amplif ier and Displ ay

ledronic Switch

Tran sducer

Sap phire - Lens

Medium

Obje.!!..kt~-::::;::::::7=,b#1i;':;;;;~?=~~~-r-r

Fig. 110. Principle of scanning acoustic microscopy in the reflection mode

299

most frequently used. Part of the acoustic energy interacting with the object is remitted, picked up by the lens, and guided towards the transducer, which transforms the reflected ultrasound into an electrical signal of the same frequency. The receiver transforms this signal representing one image point into a gray level on the video display and its value represents the signal intensity. Accordingly, the brightness on the monitor depends on the reflected ultrasound, energy, which is due to the local acoustic properties of the specimen. By scanning the lens across an area in two orthogonal direction, a two-dimensional image is built up point by point.

Two basic requirements are important for the design of the acoustic lens: the high velocity of sound and the acoustic absorption. In the high frequency range up to 2 Ghz, single crystals of sapphire are a suitable material. The velocity of sound along the c-axis in sapphire is about 11 000 ms - 1. To reduce the energy loss as a result of the large impedance difference at the interface lens-fluid, an antireflection layer is sputtered onto the cavity of the lens. Because of the significant effect of this energy loss on image quality, the radius of the curvature is kept very small (less than 1 mm).

In material characterization, the scanning acoustic microscope is used for imaging grain structures, e.g., unetched metals and for detection of microcracks, e.g., in quartz grains in granite. Other applications are the investigation of integrated circuit component in microelectronics. Furthermore, for the characterization of thin-film properties, the scanning acoustic microscope is an important instrument.


References

Addison RC, Somekh M, Rowe JM, Briggs GAD (1987) Characterization of thin-film adhesion with the scanning acoustic microscope. SPIE 768: 275-284

Briggs A (1985) An introduction to scanning acoustic microscopy (Microscopy handbooks 12). Oxford Univ Press, Oxford

Hollis RL, Hammer R (1980) Defect detection for micro-electronics by acoustic microscopy. In: Ash EA (ed) Scanned image microscopy. Academic press, pp 155-164

Kushibiki J, Maehara H, Chubachi N (1982) Measurement of acoustic properties for thin films. J Appl Phys 53: 5509-5513

Nikoonhad M (1984) Recent advances in high resolution acoustic microscopy. Contemp Phys 25: 129-158

Lemons RA, Quate CF (1979) Acoustic microscopy. In: Mason WP, Thurston RN (eds) Physical acoustics. Academic Press, London, pp 1-92

Quate CF, Atalar A, Wickramasinghe HK (1979) Acoustic microscopy with mechanical scanning - a review. Proc IEEE 67: 1092-1114

6.3 Scanning Tunneling and Atomic Force Microscopy

A.V. ERMAKOV and S.V. TITKOV

In 1982, G. Binning and H. Rohrer (Nobel Prize in physics in 1986) developed an essentially new type of a microscope with the atomic resolution to study solid surfaces. The operation principle of this microscope is based on the quantum tunneling effect. It occurs when a metallic needle is brought close (less than 20 A) to the surface of conducting sample under investigation and an electric tunneling current is induced by an applied voltage.

In tunneling microscopes the probing tungsten needle has a tip which ideally consists of a single atom. It is scanned across the sample surface by piezoelectric translators which, under the action of an applied voltage, either expand or contract, thereby causing the needle to move along the X, Y, and Z directions. An automatic controller maintains a constant value of the tunneling current by raising and lowering the position of the needle whose resulting movement duplicates the surface relief of the sample under study. The amount of the needle shift is determined by the voltage at the piezoelectric translator, which is processed by the computer and presented either on a video display or on a plotter. By moving the tip along a number of parallel lines, a three-dimensional image of the surface structure is obtained.

The method of tunneling microscopy achieves its unique resolution due to the existence of a very strong exponential dependence of the tunneling current (I) on the distance between the probing needle and the surface under study (s): I '" Exp ( - 2KS), where the decay constant is expressed as K = J2m<jJ /h2, and <jJ is the electron output.

The maximum resolving power of a tunneling microscope in the plane of the sample surface is 2 A and 0.02 A normal to the surface. This implies that it is

6.3 Scanning Tunneling and Atomic Force Microscopy 301

possible to study structural elements with dimensions of a few hundredths of an atomic diameter.

The dimensions of the field of sight in this method are determined by the structure of piezoelectric translators responsible for the scanning. The dimensions of the field of sight range from tens of A to hundreds of Jl.m and can easily be adjusted electronically.

Very recently, scanning tunneling microscopy has revolutionized surface science. This method allows the following characteristics of surfaces to be obtained: (1) their atomic relief showing the arrangement of individual atoms; (2) the local distribution of the potential barrier height for electrons; (3) the local distribution of the density of the electron energy states in the surface layer of the sample (tunneling spectroscopy).

Among the major advantages of the method of tunneling microscopy are the following features. In a tunneling microscope the image is produced without destruction or any damage to the sample being studied (it is not even necessary to deposit a thin metal film onto the sample surface as in the case of electron microscopy). Using this method, investigations can be performed not only in a vacuum but in air or in a liquid medium as well. In the tunneling microscope the sample experiences no dehydration or radiation damage. Besides, there is no limit to the size of the samples to be studied.

However, as follows from the description of the operational principle of the tunneling microscope, this method does: not allow the investigation of samples having no electrical conductivity.

To study dielectrics, G. Binning and his co-workers developed a method known as atomic-force microscopy in 1986. This method makes use of the interatomic (intermolecular) interaction forces taking place between the surface being examined and the probing needle separated by a distance of 1 to 100 A. In the atomic-force microscope the resolution in the direction normal to the surface is 1-3 A; in the transverse direction it is Jl 0-15 A, which is close to the resolution of the tunneling microscope.

The method of the tunneling microscopy has found a wide application in microelectronics, since the influence of the surface structure upon the operational characteristics of crystals is especially high. Investigations of semiconducting materials (Si, GaAs), physical (friction and superconductivity), chemical, and biological processes have been performed with the aid of the tunneling microscope.

Application of tunneling and atomic-force microscopy to minerals is at its beginning. It was found that hematite and galena (001) surfaces, exposed by fracturing in air or under oil at room temperature, do not reconstruct and have the same atomic structure as the equivalent plant in the bulk. However, the nanometer-scale morphologies of these supposedly simple and flat surfaces are highly complex and consist of undulations which have dimensions of approximately 100 to 400 A in diameter. In addition to undulations, relief on hematite is seen in the form of pits and ridges with a vertical size of several tens of Angstroms. Cleavage steps, 30-50 A high, are readily apparent on the galena


surface. STM images of pyrite have shown, that atomic positions on a pyrite growth surface do not correspond to those expected for a simple termination of the bulk pyrite structure; it is likely that a surface oxidation product was imaged.

The AFM results suggest that the {O 1 O} surface of albite, exposed by fracturing, exhibits very small pits and cleavage steps, which are high energy reactive sites. This method has recently been used to image also muscovite, montmorillonite, and illite surface.

Thus, tunneling and atomic-force microscopy provide a unique means for studying the atomic structure and the nanometer-scale morphology of surfaces which playa key role in mineral dissolution, sorption at the solution-mineral surface, mineral surface reactions (for example, in ore deposit formation), and fracture propagation.

References

Baratoff A, Binning G, Fuchs H, Salvan F, Stoll E (1986) Tunneling microscopy and spectroscopy of semiconductor surfaces and interfaces. Surf Sci 168: 734-743

Binning G, Rohrer H (1986) Scanning tunneling microscopy, IBM J Res Dev 30: 355-369 Binning G, Rohrer H, Gerber Ch, Weibel E (1982) Surface studies by scanning tunneling

microscopy. Phys Rev Lett 49: 57-61 Binning G, Quate CF, Gerber Ch (1986) Atomic force microscope. Phys Rev Lett 56: 930-933 Drake B, Prater CR, Weisenhorn AL, Gould SAC, Albrecht TR, Quate CF, Cannell OS,

Hansma HG, Hansma PK (1989) Imaging crystals, polymers, and processes in weater with the atomic force microscope. Science 243: 1586-1589

Eggleston CM, Hochella MF Jr (1990) Scanning tunneling microscopy of sulfide surfaces. Geochim Cosmochim Acta 54: 1511-1517

Golovchenko JA (1986) The tunneling microscope: A new look at the atomic world. Science 232: 48-53

Hansma PK, Tersoff J (1987) Scanning tunneling microscopy. J Appl Phys 61: RI-R23 Hartman H, Sposito G, Yang A, Manne S, Gould SAC (1990) Molecular-scale imaging of clay

mineral surfaces with atomic force microscope. Clays Clay Mineral 38: 337-342 Hochella MF Jr, Eggleston CM, Elings VB, Parks GA, Brown GE Jr, Wu ChM, Kjoller K

(1989) Mineralogy in two dimensions: scanning tunneling microscopy of semiconducting minerals with implications for geochemical reactivity. Am Mineral 74: 1233-1246

Hochella MF Jr, Eggleston CM, Elings VB, Thompson MS (1990) Atomic structure and morphology of the albite {O 1 O} surface: an atomic-force microscope and electron diffraction study. Am Mineral 75: 723-730

Sharp TG, Zheng NJ, Chang CS, Tsong 1ST, Buseck PR (1989) Scanning tunneling microscopy studies of galena (00 1) cleavage surfaces (Abstr) EOS 70: 1394

CHAPTER 7

Recent Developments in Analytical Methods in Mineralogy

304 Chapter 7. Recent Developments in Analytical Methods in Mineralogy

7.1 General Overview of the Methods of Analysis of Minerals, Rocks, Ores, and Materials

P.I. POTTS

As in other branches of science, analytical techniques for the chemical characterisation of geological materials have undergone a revolution over the past 30 years. The introduction of a succession of new and increasingly powerful techniques over this period has resulted in (1) a substantial increase in the range and sensitivity with which elements may be determined on a routine basis; (2) a considerable increase in analytical productivity in terms of determinations per day; and (3) a reduction in the manpower required to generate such data. This revolution has been accompanied by a progression away from essentially manual procedures involving chemical manipulations, as used in classical and rapid schemes of analysis, to instrumental techniques capable of automated operation with the minimum of operator intervention.

A direct consequence of the introduction of new analytical technology is the new areas of scientific endeavor that can now be exploited. This link can clearly be traced, for example, from the successful development of high resolution germanium gamma-ray spectrometers for instrumental neutron activation analysis in the 1970s and the consequent expansion in rare earth element geochemistry. Similarly, the 1980s were the decade of isotope geochemistry based on the development of high precision, high sensitivity thermal ionisation mass spectrometry instrumentation, and associated sample preparation procedures. At this distance in time, it is more difficult to recognize the contribution of earlier techniques, now considered in a routine and less glamorous light. Notable amongst these, is X-ray fluorescence spectrometry, which initiated widespread interest in trace element geochemistry by offering the capability of determining trace elements such as Rb, Sr, Y, Zr and Nb to ppm detection limits on a routine basis.

Despite the contribution made by novel techniques in the advancement of geochemical studies, there is a continuing requirement for the routine analysis of geological samples for major and trace elements. Indeed, there is a certain irony in the contrast in philosophies used to select samples for analysis between contemporary and earlier geochemical investigators. In earlier years, laboratories generally had a very limited capability for the analysis of large numbers of samples by wet chemical procedures. An examination of samples in thin section using an optical microscope would often then be used as the basis for selecting samples for chemical analysis. In modern geochemical studies, all samples collected from a field area are likely to be submitted for routine major, minor, and trace element analysis. Indeed, these routine geochemical data are then often used as the basis for selecting samples to be analyzed by more specialized techniques used, for example, in the determination of the rare earth elements or isotope ratio measurements.

7.1 Methods of Analysis of Minerals, Rocks, Ores, and Materials 305

Modern geochemical laboratories are, therefore, equipped to determine a comprehensive range of major and trace elements on a routine basis with a large annual throughput of samples. Except where economic or social factors dictate otherwise, these criteria largely preclude the use of manual chemical methods of analysis in favor of techniques capable of simultaneous multi-element analysis. As will be seen from the next sections, there is significant overlap between the analytical capabilities of modern instrumental techniques. Furthermore, despite the potential of inductively coupled plasma-mass spectrometry, no single technique can be considered to be fully comprehensive. However, certain combinations oftechniques complement each other to a large degree in offering extensive coverage of major and trace elements routinely required by modern geochemical studies. It is convenient to summarize the capabilities of these complementary techniques by considering them as follows.

Inductively Coupled Plasma-Atomic Emission Spectrometry (ICP-AES) Plus Flame Atomic Absorption Spectrometry (AAS)

For many laboratories with a strong tradition in classical and rapid schemes of analysis, a natural progression arose with the introduction first of AAS and then of ICP-AES. Both these instrumental techniques are based on solution analysis and therefore take advantage of skills and expertise in the dissolution of samples. To offer a comprehensive analytical capability, up to three sample dissolutions procedures may be required. In the determination of the major elements, a lithium metaborate fusion is normally used so that on dissolution of the fused mixture, silica is retained in solution. Trace elements may also be determined on this solution; however, a separate sample attack based on HFperchloric acid is often used during which silica is expelled by fuming off as the tetrafluoride with a consequent beneficial reduction in the total salt content of analyte solutions. The ICP-AES technique does not have sufficient sensitivity to determine the rare earth elements without an additional ion exchange preconcentration procedure. Atomic absorption spectrometry is required particularly in the determination of Rb, and to some extent K and Na, elements that cannot otherwise be determined with adequate sensitivity by ICP-AES.

As well as comprehensive coverage of the major and trace elements of geochemical interest, the ICP-AESjAAS combination is capable of determining selected light elements such as Li, Be, and B at trace levels. Furthermore, with the appropriate matching of sample and standard solutions, these techniques are relatively insensitive to matrix effects. However, the precision of determinations, particularly of the major elements, is not as high as can be achieved by XRF, for example. Furthermore, either efficient procedures for the batch dissolution of samples or a degree of automation is essential for sample preparation to achieve an adequate sample throughput. Particular attention must be paid to ensure the complete dissolution of samples containing resistant phases such as chromite, zircon, and selected rare earth element phases.


X-Ray Fluorescence Spectrometry (XRF) Plus Instrumental Neutron Activation Analysis (INAA)

Dissolution of samples is avoided in this combination of techniques, both of which involve the analysis of samples in a solid form. Two preparation schemes are normally used for XRF: glass discs for the determination of the major elements and compressed powder pellets for the trace elements. Samples for INAA are normally dried and sealed in polyethylene vials for irradiation without further preparation. In view of the high precision of results, XRF is arguably the technique of choice for the determination of the major elements, provided accurate correction models are used to compensate for interelement matrix effects. Furthermore, the XRF technique is capable of determining to high sensitivity several geochemically important trace elements (including Rb, Sr, Y, Zr, Nb, Th) which can pose difficulties in some competitive techniques that are based on solution analysis. To complement these determinations, INAA is capable of determining to high sensitivity a range of rare earth elements as well as other key trace elements including Ta, Hf, and Th. In earlier years, INA A did not always have the highest reputation in the reliability of analytical results, mainly because calibration normally depends on the simultaneous irradiation of a single "standard" rock. These difficulties have been largely overcome as the reliability of trace compositions in geochemical reference materials has improved. Disadvantages of this combination of techniques is the inability of determining low atomic number elements, some of which (including Li, Be, B, and F) have important geochemical applications. Careful appraisal of the accuracy of calibration techniques is also necessary in view of the normal practice of relying on reference materials as calibration standards. Furthermore, INAA is not completely self-contained, access to a nuclear reactor being required for the activation of samples.

Inductively Coupled Plasma-Mass Spectrometry

This technique is the one that comes closest to satisfying a universal geochemical panacea in the provision of geochemical data. Not only are there very few elements that cannot be determined satisfactorily (a few suffer mass interferences from molecular constituents of the plasma), but also the instrumentation can be tuned to give uniformly high sensitivity across the entire mass range. Routine detection limits are normally in the sub-ppm range. With such a high specification in instrumental performance, attention has naturally been focused on what some would see as the Achilles heel of the technique, sample dissolution and solution chemistry. In this respect, ICP-MS suffers no greater disadvantage than other solution techniques except in its extended capability for the direct determination of selected trace elements that may be present in resistant mineral phases. In view of the relatively short time interval over which ICP-MS instrumentation has been in routine practical operation, it is too early to evaluate fully the capabilities of this technique in comparison with those listed

7.1 Methods of Analysis of Minerals, Rocks, Ores, and Materials 307

above. However, although ICP-MS is unlikely to be effective in the routine determination of the major elements, the technique has already demonstrated its high analytical productivity in the direct determination of trace elements, including the rare earth elements, without the necessity for a preliminary ion exchange separation.

Thermal Ionization Mass Spectrometry (TIMS)

Although ICP-MS instrumentation is capable of isotope ratio determination, the technique cannot currently match the high precision ofTIMS essential in studies of RbjSr, NdjSm, and Pb-U-Th isotope schematics. In this respect, TIMS has no rivals, and coupled with the appropriate sample preparation facilities, is the key instrument in laboratories undertaking isotope geochemistry studies. The technique also has the capability of determining trace elements, notably the rare earth elements (REE), using isotope dilultion techniques. These determinations require extensive solution chemistry to separate the elements of interest from other major and trace elements, and in the case of the REE, into subgroups of adjacent elements. If chemically purified samples are not presented for analysis, isobaric interferences are likely to prevent accurate determination. Although isotope dilutions were regularly performed in many TIMS laboratories, these studies have become less popular (except for samples containing very low REE abundances) in view of the extensive chemistry and instrument time required in relation to the requirements of competitive techniques.

Future Developments

No one can predict with certainty the future course of analytical geochemistry. However, it is possible to discern trends that have already influenced the developments in this field. One clear trend is the progressive improvement in sensitivity that has arisen from the introduction of a succession of new techniques. At one time, the parts-per-million level appeared to represent limit to routine determinations. However, now analysts have their horizons on the subparts-per-billion level. A further trend has been in improved methods of automation, not only in instrument operation, but also in sample preparation procedures to reduce man-power costs and enhance analytical productivity.

Although it is all too easy to become blinded by the technology of potential analytical developments, it is important to realize that new analytical procedures are only likely to find widespread use if one of two criteria are satisfied. Either the technique must be capable of determining existing suites of elements with higher analytical productivity in relation to running costs, or alternatively be capable of promoting new geochemical applications by a substantial improvements in sensitivity. The latter criterion may be satisfied by the development of new techniques for the determination of geochemically important trace elements at the ppb and sub-ppb level or by the direct determination of


"difficult" trace elements that cannot often be included in existing schemes of analysis (including, perhaps, elements such as the halides, As, Sb, Se, Te, Hg, TI, Bi, and the platinum-group elements). Since the geochemical interpretation of these data may depend in part on their mineralogical behavior, it seems likely that equal importance will be placed on the development of microanalysis techniques for the trace element determinations in individual mineral grains at high spatial resolution. Whatever the outcome, it is unlikely that the more traditional techniques described in the next sections will be supplanted in routine geochemical applications.

7.2 Classical and Rapid Methods

P.J. POTTS

For the first half of the 20th century, the only techniques available for the quantitative analysis of rocks and minerals were the so-called classical techniques. In their original form, these techniques were based on multiple chemical separations to isolate individual analytes in a chemically purified state for determination by mass using a balance. A large number of variations existed in the precise scheme adopted by individual analysts. However, in brief overview, a classical separation consisted of the following stages:

1. Fusion of the sample (1 g) with anhydrous sodium carbonate; after cooling, the melt was dissolved in hydrochloric acid.

2. The resultant solution was boiled and evaporated to dryness to dehydrate silicic acid to insoluble silica. Soluble chlorides were taken up in hydrochloric acid solution, leaving a residue of silica. The small porportion of silica that was also taken into HCI solution was recovered by a second evaporation.

3. The hydrochloric acid solution was neutralized with ammonia, so causing the precipitation of insoluble hydroxides of iron, aluminium, titanium, and phosphorus (the R20 3 group).

4. The precipitation of calcium (and strontium) as insoluble oxalates. 5. The precipitation of insoluble phosphates of magnesium and manganese. 6. Determination of sodium and potassium in the filtrate.

For each group of elements separated, detailed procedures existed to isolate individual components for determination gravimetrically. These procedures were designed to make allowances for the small fraction of analyte carried over from the previous stage or interfering elements co-precipitated with the current one. Clearly, from a modern point of view, such procedures suffer from a number of serious limitations. The principle analytical criticism is that gravimetric determinations do not measure an element-specific signal (in the sense that an atomic emission line is characteristic of an individual element) and depend to a large extent on the reliability of the separation procedure and the skill of the

7.2 Classical and Rapid Methods 309

analyst. Furthermore, gravimetric measurements have limited sensitivity; only the major elements and a few minor elements may be determined satisfactorily, but not the array of trace elements that are so essential in modern geochemical analysis. Further limitations stemmed from practical considerations. To undertake routine classical analysis, the analyst required a high degree of skill and motivation and was capable of determining the full suite of major elements in only a few samples each week. A further difficulty arose from the lack of internationally recognized reference samples. Such samples did not become available until the first interlaboratory analysis program was organized by the US Geological Survey in the late 1940s, involving the distribution of G-1 and W-1. It was not until data from this program were fully analyzed that the significant interlaboratory bias that could affect the accuracy of the results was fully appreciated.

Some of the limitations of classical procedures were ameliorated by the introduction of several important analytical developments that improved both element specificity and the sensitivity of determinations and led to the evolution of "rapid" schemes of analysis. The most important of these innovations are as follows.

Flame Photometry

The alkali metals were difficult to determine reliably by gravimetric procedures. However, an important development was made with the introduction of flame photometry instrumentation. This instrumentation requires the sample solution to be atomized in a flame, the energy of which is sufficient to excite electrons in outer atomic orbitals of the alkali metals (especially Na and K). Subsequent deexcitation causes the emission of characteristic atomic lines that may be measured using a simple filter photometer. Early instruments were designed with a total consumption burner in which the sample solution was sucked straight into the flame, producing a rather noisy signal. Subsequently, more sophisticated designs were developed in which the sample solution was first nebulized and an aerosol of solution droplets passed into the flame for atomization. Modern flame photometers may be designed with optical monochromators and can be used for the sensitive determination of Na and K, often using Li as an internal standard in the analysis of geological samples.

Element Specific Complexing Reagents

A further refinement arose from the development of element-specific chelating reagents, notably ethylenediaminetetraacetic acid (EDT A) and its derivatives. More specifically, Ca and Mg were difficuh elements to determine gravimetrically by classical procedures. However, EDT A is a complexing agent with six


points of coordination that can react with Ca 2 + and Mg2 + ions, forming a cagelike complex. These elements could, therefore, be determined readily by titrating an aliquot of the sample solution with EDT A using a suitable indicator to show the end-point of the reaction, without the necessity of performing a detailed group separation. The normal procedure was to determine calcium on an aliquot of the sample solution (prepared by an acid digestion in which silica was fumed off) by titration with EDT A under strongly alkaline conditions (to precipitate Mg and Mn as hydroxides) using an indicator such as murexide. Magnesium was then determined by difference by titrating a second aliquot of the sample solution with EDT A at a pH of 10, conditions under which the combined content of Ca and Mg can be determined. A correction for Mn must be applied to the result.

Colorimetric Analysis

The third major innovation involved the development of element specific complexing reagents that permitted the direct determination of individual analytes from the intensity of the colour formed by the corresponding complex. Such measurement are made at a characteristic wavelength using a simple optical spectrometer. A wide range of such colorimetric reagents have been developed for the determination of major, minor, and trace elements.

A typical example of a colorimetric technique is the determination of total iron in a sample using 1, 1 O-phenanthroline. Hydroxyammonium chloride is added to the sample solution to convert all the iron to the ferrous oxidation state. Ferrous iron is then complexed with 1,10-phenanthroline in the presence of a suitable buffer agent, forming an intense red ferrous-(1,10-phenanthroline)~ + complex. The intensity of this complex is measured at 508 nm using an optical spectrometer and results compared with those from suitable standard solutions.

Current Status

Rapid techniques survived in many laboratories into the 1960s but were then largely superceded in routine schemes of analysis by the introduction of flame atomic absorption spectrophotometry. This latter technique offers much higher element specificity so that schemes of analysis could be developed for the determination of all the major elements from one sample solution. Furthermore, instrumentation was also capable of selected trace element determination on a routine basis. The geochemical importance of rapid schemes of analysis has, therefore, declined substantially. However, some of the techniques developed for rapid schemes still find use, particularly in the following areas: (1) "Low tech" laboratories that have access to minimum instrumental resources; (2) field/exploration laboratories set up to determine a limited range of "exploration"

7.3 Atomic Absorption Spectrometry 311

elements; (3) in the determination of '''difficult'' trace elements that are not readily amenable to modern instrumental techniques; and (4) in the analysis of mineral separates where only a small mass of material is available and a determination against "absolute" elemental standards is required.

Recommended reading

Jeffery PG (1975) Chemical methods of rock analysis. Pergamon Press, Oxford Maxwell JA (1968) Rock and mineral analysis .. Wiley-Interscience, New York Potts PJ (1987) A handbook of silicate rock analysis. Blackie, Glasgow, Chap 2, pp 47-76 Wilson AD (1955) Determination of ferrous iron in rocks and minerals. Bull Geol Surv GB 9:

56-58

7.3 Atomic Absorption Spectrometry

P.l. Pons

Although the principles of atomic absorption have been known for almost two centuries, it was not until the mid-1950s that instrumentation was developed by Walsh and independently by Alkemade and Milatz, demonstrating the practical analytical application of this technique. As in almost all optical spectroscopic techniques, atomic absorption instrumentation consists of three principle components: a light source, an atom cell, and an optical monochromator. However, these components possess some specialized features, as discussed further below. The principle of atomic absorption measurements depends on the absorption of light by the outermost electronic orbitals of atoms. Optical photons of characteristic wavelength are capable of exciting such outer electrons to a higher electronic state. The analytical signal is thus simply the proportion of light of characteristic wavelength that is absorbed by a sample presented for analysis in an atomized form. Despite the simplicity of this process, there were two main stumbling blocks to the development of practical analytical instrumentation. The first arose from the excitation mechanism which can only occur over a very narrow wavelength band (typically < 0.01 nm) and is not, therefore, achieved effectively using a conventional "white light" source. The second lay in the development of an effective means of atomizing the sample. These problems were overcome as follows.

The light source used in conventional atomic absorption instrumentation is the hollow cathode lamp. This consists of glass/quartz envelope containing a hollow cathode, the inner surface of which is coating with a compound of the element of interest. The glass envelope is filled with typically neon gas at low pressure (about 5 torr). When an electric current is induced to pass between anode and cathode, a low pressure discharge is created within the hollow cathode. This discharge causes a small amount of material to be atomized from


the hollow cathode and then excited, resulting in the emission of the atomic spectrum of the coating element. This atomic spectrum, therefore, contains intense emission lines lying at precisely the correct wavelength to excite atoms of the specified elements. However, an important limitation is clearly the necessity of changing lamps if a different element is to be determined. To some extent, this limitation has been mitigated by the development of multiple element hollow cathode lamps (capable of emitting the spectrum of two, or exceptionally, up to four compatible elements) and multiple lamp turrets, which on some instruments can be exchanged under automation. However, atomic absorption spectrometry is essentially a single element technique and therefore suffers some disadvantage in analytical productivity compared with more recent multielement techniques.

Atomization in conventional instrumentation is achieved by preparing samples as solutions which are drawn into a nebulizer before being passed into a flame. The function of the nebulizer is to convert the solution into an aerosol cloud of droplets which on entering the flame suffer evaporation of the solvent (normally water), dissociation into simple inorganic species, and finally atomization. The flame therefore serves as an atom cell into which the sample is continuously aspirated for the duration of measurement of the atomic absorption signal (5-10 s). The flame must burn at a sufficiently high temperature to ensure complete atomization of the sample and modern instrumentation usually has the option of air-acetylene (2450 0c) and nitrous oxide-acetylene (3200 0c) flames, the latter to promote the atomization of the more refractory elements. To maximize sensitivity, the burner has a slotted design, the long axis of which is aligned along the optical axis of the spectrometer.

The final component is the optical monochromator used to measure the proportion of radiation from the hollow cathode lamp that is absorbed when a sample solution is aspirated into the flame. Since atomic absorbance can only occur over an extremely narrow wavelength range (corresponding to the line width of the selected atomic emission line from the hollow cathode lamp), the monochromator can be designed with a relatively low resolution specification (0.2-0.7 nm band pass) without compromising performance ("white" radiation within this band pass represents a minute fraction of the atomic emission line intensity).

The analytical signal, the absorbance (A), is related to concentration using a derivation of Beer's law:

A = 10glo(Po/P) = abc

Where Po and P are the power of the incident and transmitted beam respectively, a is the absorptivity (or absorption coefficient), b is the path length of the atom cell (flame), and c is the concentration of the absorbing species.

When the first developed, atomic absorption revolutionised geochemical analysis, which at that time was mainly undertaken by rapid wet chemical techniques (qv). In particular, AAS facilitated the determination of a range of major and trace elements to very high element specificity with interference effects that were much simpler to control compared with those encountered in

7.3 Atomic Absorption Spectrometry 313

contemporary techniques. The high element specificity arises from the excitation of an element with a selected line from its own atomic emission spectrum. Because of the narrow band-pass over which excitation can occur, interferences caused by the unwanted excitation of competing elements are virtually unknown. The interference effects that are observed are mainly associated with the flame atomization mechanism and fall into two categories. First, there is sufficient energy in the flame to cause partial ionization of elements such as the alkali metals. Since ions have a different absorption spectrum from that of atom species, ionization causes the atomic absorption signal to be suppressed by an amount that often depends on the concentration of coexisting elements in the analyte solution. This interference is normally suppressed by the addition to the sample solution of an element that is more readily ionized than the analyte, caesium being most widely used in this role. The second category of interference concerns the difficulty in dissociating some compounds, particularly those that are refractory in nature, containing strong analyte-oxygen bonds. These elements are normally determined using the hotter nitrous oxide-acetylene flame. However, to promote complete atomization, it is often necessary to add a releasing agent to the sample solution. Lanthanum is often used for this purpose, since this element forms a strong La-O bond and so competes for excess oxygen in the flame, promoting the atomization of refractory oxide species.

Conventional flame atomic absorption is, therefore, capable of analyzing geological samples for the major elements (Na, Mg, AI. Si, K, Ca, Ti, Mn and Fe) and a selection of trace elements, including Li, Be, V, Cr, Co, Ni, Cu, Zn, Rb, Sr, Ba, and Pb. Calibration is usually undertaken against aqueous standards in which solution composition is matched with that of samples. Sample dissolution can be achieved by acid attack involving HF /perchloric followed by volatilization of Si or alternatively an HF /nitric digestion in an enclosed bomb followed by dilution with boric acid to suppress the activity of HF (Si retained in solution). Some interest has also been shown in the lithium metaborate fusion procedure followed by dissolution in nitric acid.

As well as showing high element specificity, atomic absorption instrumentation is relatively cheap, rugged, and does not require sophisticated laboratory services. However, the lower analytical productivity, greater susceptibility to matrix effects, and more restricted element range compared with competing solution techniques (ICP-AES and ICP-MS) have caused a decline in the importance of AAS for routine geochemcial analysis in research laboratories. Despite these limitations, this instrumentation retains high popularity, particularly in service and field laboratories, where simplicity and reliability are more important than unnecessary sophistication.

Graphite Furnace Instrumentation

Following the proposals of L'vov, a second form of atomizer was introduced in atomic absorption instrumentation and became widely available during the 1970s. In this form, the pneumatic nebulizer and flame atomizer are replaced by


a graphite tube furnace. The sample is now injected into the furnace as a discrete aliquot of solution (usually 10 to 40 Ill). An electric current is programmed to pass through the graphite tube, the temperature of which is precisely controlled to sustain the following cycle of sample decomposition:

Drying. Sample solution is first dried at typically 110°C to evaporate the solvent.

Charring or Ashing. The inorganic residue is heated to typically 600-1200 °C to break down the matrix into simple inorganic salts with the elimination of smoke that might interfere with the atomic absorption measurement.

Atomization. The furnace temperature is rapidly increased by up to 1000°C above the charring temperature to atomize the residue. During this cycle, the transient atomic absorption signal is measured for comparison with matrix matched standards.

Clean Out. The temperature of the furnace is finally increased by a further 200°C to remove all traces of the analyte and is then cooled to allow another cycle to begin.

The precise temperature for the charring and atomization cycles depend on the characteristics of both analyte and sample matrix. Reliable analysis often involves the use of "matrix modifiers", added to the analyte solution to prevent premature atomization of the analyte during the charring cycle and/or incomplete vaporization during the atomization cycle.

A second innovation is essential for successful graphite furnace AAS - that of an accurate background correction procedure. Whereas this is a rarely used optional facility in flame AAS, background correction is essential in almost all graphite furnace AAS measurements to allow for nonanalyte absorption arising from light scatter and nonspecific molecular absorbance. Three methods of background correction are available (deuterium arc, Smith-Hieftje and Zeeman). Each has its advantages and disadvantages, the deuterium arc being relatively simple and able to accommodate most background interference, Smith-Hieftje being effective for a range of nonrefractory elements, and Zeeman being universally applicable though the most expensive in terms of additional instrumentation. All background correctors causes some reduction sensitivity; the advantage of the latter two techniques being that they can correct for "structured" backgrounds (i.e., any nonlinear variation in the magnitude of the background within the band-pass of the monochromator).

Although such instrumentation is more sophisticated, graphite furnace atomic absorption offers the advantages of much improved detection limits compared with flame AAS, for some elements approaching those attainable by ICP-MS. However, in respect of geological applications, matrix interference effects preclude measurements unless the sample is first subjected to a chemical separation. The technique has, therefore, a specialized role in geological appli-

7.4 Inductively Coupled Plasma - Atomic Emission Spectroscopy 315

cations and has been used, for example, in the determination of the rare earth elements as well as gold and the platinum-group elements after fire assay separation.

References

Bernas B (1968) A new method for decomposition and comprehensive analysis of silicates by atomic absorption spectrometry. Anal Chern 40: 1682-1686

Buckley DE, Cranston RE (1971) Atomic absorption analysis of 18 elements from a single decomposition of aluminosilicate. Chern Geol 7: 273-284

Ebdon L (1982) An introduction to atomic absorption spectroscopy. Heydon, London Slavin W (1982) Atomic absorption spectroscopy, the present and the future. Anal Chern 54:

685A-694A Potts PJ (1987) A handbook of silicate rock analysis chap 4: Atomic absorption spectrometry,

Blackie, Glasgow, pp 106-152 Sen Gupta JG (1985) Determination of the rare earths, yttrium and scandium in silicate rocks

and four new geological reference materials by electrothermal atomization from graphite and tantalum surfaces. Talanta 32: 1-6

7.4 Inductively Coupled Plasma -- Atomic Emission Spectroscopy

J.G. CROCK and P.H. BRIGGS

Since its recent introduction in the early 1970s, inductively coupled plasma -atomic emission spectroscopy (ICP-AES) has become an important technique for the analysis of geochemical materials for their trace, minor, and major element content (Fassel and Kniseley 1974). In this technique, element excitation is achieved by a plasma sustained by the interaction of ionized argon gas inductively coupled to a radiofrequency energy field. An inductively coupled plasma is an attractive spectral source because of its high temperatures, as much as 10 000 K, optical transparency, and long-term stability. About three-fourths of the common elements can be determined by this technique, with lower limits of detection in the range of 0.5 to 5 jJ.g/g. This technique is noted for linearity of response, often covering four to five ordt:rs of magnitude for most elements, and relative freedom from matrix affects, whkh often plague other spectroscopic and classical methods. The technique offers excellent measurement precision, usually from 0.1 to 3% relative standard deviation, and has good accuracy. ICP-AES is a rapid multi-element technique, by which 30 to 40 elements commonly can be determined within about 2 min.

Several recently published books and articles on the ICP-AES technique explain in detail the theory and chemistry of the plasma and the analysis of geological materials. Montaser and Golightly (1987) present the theory and chemistry of the plasma, as well as its general application to analytical chemistry. A handbook of ICP-AES analysis (Thompson and Walsh 1989) describes


the complete and partial analyses of geological materials, water analysis, and the analysis of environmental samples. Winge et al. (1985) present an excellent atlas of ICP-AES spectral information for most elements. This atlas aides in selecting the proper analytical wavelength to maximize sensitivity in a given matrix while minimizing potential spectral overlap interferences. Walsh and Howie (1986) present a current review of the application of ICP-AES to the geological sciences. Church (1981) demonstrates the utility of the ICP-AES for multielement analysis of geological materials. Lichte et al. (1987) present a condensed handbook of the routine ICP-AES analysis of geochemical materials. Reviews of the current literature on ICP-AES analyses of geological and inorganic materials are published annually or biennially in journals such as the Journal of Analytical Atomic Spectrometry and Analytical Chemistry.

ICP-AES Components

The typical ICP-AES instrument has four major components, which include the sample introduction system, sample excitation system, spectrometer, and dedicated computer. The sample introduction system includes the sampling pump, pneumatic nebulizer, and spray chamber. An aqueous sample is pumped by a peristaltic pump to a pneumatic nebulizer that creates a fine aerosol in the spray chamber where size discrimination of the mist occurs. Common types of nebulizers are Babington, Meinhard, and cross-flow. The aerosol mist, generally having droplets less than 10 Jim in diameter, is transported by argon gas from the spray chamber to the torch assembly with an efficiency of usually less than 3%, with the balance going to waste. An auto sampler is often used with the sampling pump to perform many routine analyses.

The second component is the sample excitation system. A high-temperature, argon plasma is produced and sustained by the interaction of an induced magnetic field from a radio frequency generator and a flow of argon. The torch consists of three concentric quartz tubes, and was developed independently by Greenfield and Fassel (Thompson and Walsh 1989). The sample aerosol is injected into the central tube of the torch, which, in turn, punches a hole into the bottom of the plasma. As the sample aerosol passes through the plasma, it is desolvated and the particles are atomized, ionized, and excited, resulting in the emission of characteristic radiation, typically in the ultraviolet spectrum.

The third component of the system is the spectrometer, either a simultaneous polychrometer or a scanning, sequential monochrometer. Light emission passes through an entrance slit onto a diffraction grating that separates the light into its component wavelengths and focuses them on the focal curve. In the polychrometer system, exit slits positioned precisely along the focal curve isolate the analytical wavelengths. A photomultiplier tube is mounted behind each exit slit and converts the light intensity into an electronic signal. Simultaneous systems typically have 15 to 50 of these fixed slit-photo multiplier tube combinations. The advantage of these systems is the speed of analysis, 1 to 2 min for 30 to 50 elements. Only 1 to 2 ml of solution is required. Drawbacks to the polychrometer

7.4 Inductively Coupled Plasma - Atomic Emission Spectroscopy 317

systems are that they are very expensive and lack alternative wavelength selection due to their fixed channels. In contrast, the scanning, sequential monochrometer system usually has a computer-controlled stepper motor that rotates the grating, moving the spectrum across a single exit slit. This monochrometer system allows great flexibility in choosing the analytical wavelength for the element or elements of interest. This versatility is important when differing sample matrices or nonroutine samples are encountered. It also offers the ability to do quick semi-quantitative or qualitative scans of unknown samples. Because the instrument must move rapidly to each predetermined wavelength position, these instruments tend to have slower analysis time and require as much as 15 ml of sample solution when several elements are determined. The monochrometer systems are less expensive, about one third the cost of a polychrometer system.

The fourth component of all modern ICP-AES systems is the dedicated computer and associated software. This component's basic functions include instrument control, background and interelement interference correction, calculations of element concentrations from stored calibration curves, data storage, and report writing. Many of the advances in ICP-AES analysis in recent years have been in the data manipulation area by using faster and larger-capacity mini-computers.

ICP-AES Analysis

The most common ICP-AES technique is to analyze a geological material as an aqueous solution, although there are techniques in which the elements of interest are extracted into an organic solvent that is aspirated directly (for example, Motooka 1988). Limited success has been reported with the direct nebulization of a slurry power or direct insertion of samples using a graphite cup (Thompson and Walsh 1989). The most promising solid sampling technique may prove to be laser ablation of a solid sample into the plasma (Thompson et al. 1989).

Sample dissolution is usually the most tedious, time-consuming, and limiting factor of the analysis. A multi-acid digestion, combining hydrofluoric, hydrochloric, nitric, and perchloric acids at low temperatures and pressures (Crock et al. 1983) is the most common dissolution method. Most of the common rockforming silicate minerals can be dissolved by this method. The advantages of acid digestion are the ease of the method, use of large samples (as much as 2 g, although 0.2 g is more common), low reagent blanks, and lower total dissolved salts in the analytical solution. With an acid digestion the final dilution factor commonly is less than 100, allowing many elements to be determined at or near their crustal abundance. A disadvantages of acid digestion is the volatilization of some elements, such as silicon and boron as fluoride compounds. Commonly, a higher pressure, closed-vessel "bomb" digestion and addition of a complexing agent for the excess fluoride is used to avoid volatilization of silicon and boron.


Many minerals are resistant to routine acid digestions and require a more rigorous digestion. These minerals include spinel, beryl, tourmaline, zircon, monazite, niobates, tungstates, topaz, and cassiterite. These minerals can be completely dissolved by the proper choice of a sinter or fusion digestion procedure. A sodium peroxide sinter will dissolve most resistant minerals. For example, boron and silicon are routinely determined in tourmaline by ICP-AES following a sodium peroxide sinter in a zirconium crucible at 445C. Lithium metaborate, sodium and/or potassium hydroxide, sodium carbonate, and the alkali persulfates are commonly used as fusion reagents. There are drawbacks to the use of fusions or sinters. They introduce a much higher total salt content into the analytical solution, which can clog the nebulizer, spray chamber, and torch assembly. These fusions and sinters also tend to have higher reagent blanks. A larger dilution factor is used because of the smaller sample size (to to 100 mg is common) in a larger final solution volume. The final dilution factor is commonly 200 to 400, making the determination of some trace elements impossible without a subsequent separation and preconcentration. Also, at least one element common to the reagent is not determinable, such as lithium and boron from a lithium meta borate fusion, the most common fusion reagent.

To minimize matrix and potential spectral problems and to bring the analytical concentration into the range of ICP-AES, chemical separation and preconcentration are commonly used. An example is the analysis of geological materials for their rare-earth-element (REE) content at or below their chondri tic abundance levels after a lithium metaborate fusion or an acid digestion and subsequent separation and preconcentration of the REE by sequential-acid ion chromatography (Crock et al. 1986). This method has a final dilution factor as small as 5 (1 g sample in 5 ml final solution). When only 5 ml of analytical solution is available, the polychrometer system is advantageous. If there is an adequate amount of solution, as in analyzing manganese nodules (Fries and Lamothe 1984), a monochrometer system is advantageous because of its versatility in choosing analytical wavelengths that minimize spectral interferences.

Summary

ICP-AES has few equals for a rapid, precise, and accurate analysis of geological materials, including minerals and whole rocks for their major, minor, and trace element content. Proper standards, either single-element solutions or digested geological reference materials, an understanding of the matrix and elements of concern emission spectra, a proper, complete digestion, and modern instrumentation are required for the analysis of geological materials by ICP-AES.

References

Church SE (1981) Multi-element analysis of fifty-four geochemical reference samples using inductively coupled plasma-atomic emission spectrometry. Geostand Newslett 5: 133-160

7.5 X-Ray Fluorescence Analysis 319

Crock JG, Lichte FE, Briggs PH (1983) Determination of elements in National Bureau of Standards' Geological Reference Materials SRM 278 Obsidian and SRM 688 Basalt by inductively coupled argon plasma-atomic (:mission spectrometry. Geostand Newslett 7: 335-340

Crock JG, Lichte FE, Riddle GO, Beech CL (1986) Separation and preconcentration of the rare earth elements and yttrium from geological materials by ion-exchange and sequential acid elution. Talanta 33: 601-606

Fassel VA, Kniseley R (1974) Inductively coupled plasma-optical emission spectroscopy. Anal Chern 46: 1110A-1l20A

Fries T, Lamothe PJ (1984) Determination of rare-earth elements, yttrium, and scandium in manganese nodules by inductively coupled argon-plasma emission spectrometry. Anal Chim Acta 159: 329-336

Lichte FE, Golightly DW, Lamothe PJ (1987) Inductively coupled plasma-atomic emission spectrometry. In: Baedecker PA (ed) Methods for Geochemical Analysis. US Geol Surv Bull 1770: BI-BIO

Montaser A, Golightly DW (eds) (1987) Inductively coupled plasmas in analytical atomic spectroscopy. VCH Publishers, New York, 660 pp

Motooka JM (1988) An exploration geochemical technique for the determination of preconcentrated organometallic halides by ICF'-AES. Appl Spectrosc 42: 1293-1296

Thompson M, Walsh IN (1989) A handbook of inductively coupled plasma spectrometry, 2nd edn. Blackie, London, 316 pp

Thompson M, Simon C, Bret L (1989) Calibration studies in laser ablation microprobeinductively coupled plasma atomic emissiol1l spectrometry. J At Spectrosc 4: 11-16

Walsh IN, Howie RA (1986) Recent developments in analytical methods: uses of inductively coupled plasma source spectrometry in applied geology and geochemistry. Appl Geochem 1: 161-171

Winge RK, Fassel VA, Peterson VJ, Floyd MA (1985) Inductively coupled plasma-atomic emission spectrometry; an atlas of spectral information. Elsevier, New York, 584 pp

7.5 X-Ray Fluorescence Analysis

V.P. AFONIN

X-ray fluorescence (XRF) is one of the most important tools in multielement determinations in a wide variety of minerals, rocks, and ores. XRF determinations of petrogenic elements are extensively used in exploration of mineral deposits and geochemical studies.

Instrumentation

The two available types of XRF spectrometers, wavelength dispersive spectrometer (WDS) and energy-dispersion spectrometer (EDS), differ in their mode of selecting the analytical signal.

The characteristic spectrum lines are selected in WDS via diffraction on the monocrystal plate with suitable spacing. WDS spectrometers are designed with either one scanning channel or many fixled channels. Spectrometers are the latest designs, and have flexible packaging, which makes it possible to install the desired number of fixed and scanning channels.


Energy-dispersion spectrometers have been designed using the latest achievements in the production of Si(Li) detectors for X-ray with high (140 ev) energetic resolution. Compact X-ray tubes (15-200 W), isotopic sources, and secondary radiators are employed as excitation sources. Isotopic sources can effectively excite the K-series of heavy element radiation. Si(Li) detectors require constant cooling by liquid nitrogen.

Data Reduction

The conversion of X-ray fluorescence intensity into chemical composition data is a comparatively complex process. In XRF, the value of the analytical signal of the i-element is presented as

(1)

where Cj is the concentrations of the i-element in the samples and F j the function, which depends on the composition of the samples, conditions of excitation, and fundamental physical parameters (absorption coefficients, fluorescence fields, etc.) of atoms in the substance to be analyzed. Function F in Eq. (1) can be calculated theoretically, in which case we have the so-called fundamental parameter equations, which are rather complex. To simplify calculations, it is possible to approximate the function F and obtain the socalled IX correction equations:

Cj = COj (l + ~>jjCj)' j

where

(2)

The "influence coefficients" lXij account for the effect of one element on another in the sample. IX-Coefficients can be estimated empirically or calculated from the fundamental theory. The IX-correction method [Eq. (2)] and many other variants are available as software for most commercial, computer-automated XRF spectrometers.

Metrological Characteristics

In XRF analysis of rocks and minerals, the errors of sample preparation and analytical signal recording are the major sources of random error. The reproducibility of XRF analysis results lies between 0.2 and 2%. The precision determined with standard samples (0.5-5%) showed XRF to be one of the most precise methods in analytical chemistry.

7.5 X-Ray Fluorescence Analysis 321

The main disadvantage, limiting wide application of XRF for analysis of mineral composition, consists in the rellatively high detection limit.

The best detection limits by XRF (1 ppm) are attained for elements with the atomic number Z = 20-40. Within the: range of minor Z, the detection limit increases up to 10 or 100 ppm due to a low yield of fluorescence and strong absorption of long-wave radiation, whiile in the region of large Z, the detection limit increases up to 10 ppm due to low efficiency of excitation, insufficient operating voltage, and insufficient resolution of the spectrometer for short wavelength.

Petrogenic elements (Na, Mg, AI, Si, P, S, K, Ca, Ti, Mn, and Fe) and Clarke contents of the elements Sc, V, Cr, Co, Ni, Cu, Zn, Ga, As, Rb, Sr, Y, Zr, Nb, Mo, Sn, Ba, Za, Ce, Nd, Sm, Pb, Th, and U in rocks can be successfully determined by XRF. When analyzing ores and minerals, the range of elements to be determined by XRF varies from Z = 6 to Z = 92.

Further advances in XRF analysis concern new sources of X-ray fluorescence excitation.

Of major importance during the last years has been the development of synchrotron-induced X-ray emission, which combines XRF and the new synchrotron X-ray source. The availability of electron storage rings to produce intense beams of synchrotron radiation has provided a new tool for X-ray analysis. Today, there are more than 50 electron storage rings worldwide, including also those for synchrotron radiation research. Due to its unique properties, like high intensity, tunable tmergy, and high degree of polarization, synchrotron radiation (SR) has been used for many purposes in science and technology. The high brightness of SR produced by storage rings has considerably enhanced its analytical capability. Recent research has shown that SR is a powerful analytical tool for trace element analysis by XRF. Trace elements at part-per-million levels or below can easily be detected.

The most promising perspective of SR is the use of focused beams for scanning samples with high spatial resolution. Different approaches for highresolution X-ray microprobe have been used, as a pair of concave spherical mirrors coated with tungsten-carbon multilayers or a grazing-incidence mirror system with two mirrors of elliptical cylinder shape. Such a system is applied to condense spatially synchrotron X-ray beams. A 10 /lm x 10 /lm focal spot of 10-17 keY X-rays and an intensity of 108 photons per second have already been obtained. The primary beam is restric:ted to a small area and opens up the possibilities of analyzing extremely small samples or local analysis of bulk specimens, and allows trace elements mapping with spatial resolution of the order of 10 /lm. The absolute minimum detection limit is 10- 12 g. The advantages of the X-ray microprobe over the electron microprobe include better analytical sensitivity and the possibility of obtaining information for chemical state analysis. X-Ray probe does not require samples to be placed under high vacuum; X-rays deposit much less energy in the sample than do electrons in the electron microprobe method.


Another way to improve the detection limit of XRF analysis consists in using the effect of total reflection of X-rays. Introducing, additionally to the reflector substrate, another reflector in the primary beam eliminates the high-energy part of the primary beam.

This effects a reduction in the background which results in detection limits in the 0.1 ppb or ng region. Total reflection XRF (TXFR) is most suited for liquid samples, especially if only a few microliters of a valuable substance are available, and concentrations down to the parts-per-billion level have to be detected. Since the penetration depth of TXRF is in the nanometer region, the technique becomes also a sensitive method for elemental analysis of surfaces within the thickness range below 5 nm and with detection limit down to 1011 atoms/cm 2

(10- 14 monolayer). The commercially available compact totally reflecting instrumentation offers

high stability and has recently emerged as a powerful analytical method. TXRF can be applied, combined with synchrotron radiation, to study depth profiles near the surface of samples.

References

Bertin EP (1978) Introduction to X-ray spectrometric analysis. Plenum Press, New York Hannaker P, Haukka N, Sen SK (1984) Comparative study of ICP-AES and XRF analysis of

major and minor constituents in geological materials. Chern Geol 42: 319-324 Heinrich KFJ , Newbury DE, Mykblebust RL, Fior CE (eds) (1981) Energy dispersive X-ray

spectrometry. US Nat! Bur Stand, Spec Publ, Gaithersburg Hutton JT, Elliot SM (1980) An accurate XRF method for the analysis of geochemical

exploration samples for major and trace elements using one glass disc. Chern Geol 29: 1-11 Sutton SR, Rivers ML, Jones KW, Smith JV (1988a) X-ray fluorescence microprobe analysis.

In: Synchrotron X-ray sources and new opportunities in the Earth Sciences. Argonne National Lab Techn Report, pp 93-112

Sutton SR, Rivers ML, Smith JV, Jones KW (1988b) Advances in geochemistry and cosmochemistry: trace elements microdistributions with the synchrotron X-ray fluorescence microprobe. In X-ray microscopy, vol 56. Springer, Berlin Heidelberg New York, pp 438-441

7.6 Neutron Activation Analysis

CHR. KOEBERL

Neutron activation analysis (NAA) with (mostly thermal) neutrons is a useful method for the simultaneous determination of about 25-40 major, minor, and trace elements in small geological samples. The method allows the highly selective determination of these elements [especially of the rare earth elements (REE), Sc, Cr, Co, Rb, Sb, Rb, Hf, Ta, W, Th, U, and others] in the ppm and ppb range without (instrumental NAA-INAA) or with (radiochemical NAA-RNAA)

7.6 Neutron Activation Analysis 323

chemical treatment. NAA was first proposed by von Hevesy and Levi in 1936, but had to await the development of suitable electronic equipment before being useful and an analytical technique. NAA is based on the "activation" of rock samples (usually powders) with neutrons. During the irradiation with neutrons, the naturally occurring stable isotopes of all elements that constitute the rock and mineral samples are transformed into higher-mass radioactive (unstable) isotopes by neutron capture reactions, e.g., 23Na + n -+ 24Na + Yprompt. The (very weak) prompt gamma radiation is sometimes used as analytical signal, e.g., for the determination of boron. The activated nucleus then decays with a characteristic half-life t1/2 with the following equation:

24Na (t1/2 = 14.96 h) -+ 24Mg + p- + y.

There are, of course, several other decay schemes. Some radionuclides decay by p decay only, but most nuclides emit gamma-rays at one or more characteristic energies (in the range between about 60-2000 keY). The amount of the radioactive nuclide is then determined by measuring the intensity of the characteristic gamma-ray lines in the gamma spectrum. For these measurements, a gamma-ray detector and specialized electronic equipment is necessary. Because irradiated geological samples contain radio nuclides of different halflifes, the gamma spectra change with time and it is therefore necessary to count at various time intervals after the end of the irradiation (early for rather shortlived isotopes, and later for long-lived isotopes, after short-lived isotopes have decayed and therefore reduced the background). After taking different half-life corrections, decay times, counting times, and other correction factor into account, the results are quantified by comparison with synthetic or natural standards which have been irradiated and counted together with the samples.

Irradiation

The most commonly used source for neutrons for the irradiation of samples is a nuclear reactor. Research reactors, such as the TRIGA Mark II type, are frequently used for activation purposes. Neutrons emitted during the fission of 235U (which is used in most reactors :and decays spontaneously by nuclear fission) are far too energetic to sustain a chain reaction or to be useful for neutron activation analysis. They have to be moderated by, e.g., water, which leads to a loss of kinetic energy and to a different neutron energy spectrum. The resulting neutron energies can be subclassified into three groups: (1) thermal neutrons (0.001-0.5 eV), (2) epithermal neutrons (0.5 eV-I0 keY), and (3) fast neutrons ( > 10 keY). The neutron flux falls rapidly from thermal to higher neutron energies. A high neutron flux is essential for an interaction between the reactor neutrons and the target nuclei. The probability of such an interaction is called the neutron capture cross-section (IT) and is different for different isotopes. The larger the cross-section of an isotope (measured in barn = 10- 24 cm2), the more neutron activation reactions take place, and the "easier" this isotope can


be determined by NAA. Irradiation with epithermal neutrons is advantageous in cases where a nuclide has a high I/u ratio (I = resonance activation integral), which means that these isotopes have much higher cross-sections by resonance capture of epithermal neutrons, while the interfering elements have not. The irradiation is performed by filtering out thermal neutrons with a cadmium or boron (as boron carbide or boron nitride) filter. Many elements, e.g., As, Br, Rb, Sr, Mo, Sb, Ba, Ta, U, are better determined by ENAA (epithermal neutron activation analysis).

The build-up of radioactivity during the irradiation follows an exponential curve, depending on the half-life of each isotope. The number of decays per second (activity A) induced in an isotope at the end of an irradiation is: A = NucJ>[1 - exp( - In 2tirr/t1/2)]' where N is the number of atoms, u is the neutron cross-section, cJ> is the neutron flux, tirr is the duration ofthe irradiation, and t1/2 is the half-life of the product isotope. After the end of the irradiation, each isotope starts to decay with its own half-time, following the equation At = Aoexp( - In 2/t1/2), where At is the activity (in decays per second) at the time t, and Ao is the initial activity.

Decay

The radionuclides formed by neutron capture decay into (usually stable) daughter nuclides by one (or a combination) of the following ways: (1) beta ( + or - ) or beta and gamma ray, (2) electron capture, (3) gamma ray (isomeric transition), (4) internal conversion, (5) alpha particle. About 90% of all nuclides formed by neutron capture undergo beta-decay, which is almost always associated with gamma-ray emission (only a few isotopes, e.g., 32p, 35S, and 204Tl, are pure beta emitters), e.g., 24Na -+ 24Mg + p- + y. Often the radioactive isotope does not decay immediately into the stable daughter product-it first forms a metastable daughter product by beta emission, which then decays into the ground state under gamma emission. It is these gammas which constitute the analytical signal. Because natural samples are a mixture of numerous different elements and isotopes, many isotopes form at the same time, but decay with different half-lives. Therefore, the resulting activity and gamma spectrum of a sample changes with time. Shortly after the end of the irradiation, the spectrum is dominated by gamma lines of short-lived isotopes, as well as by the high (beta) background from the decay of these isotopes.

Measurement

The instrumentation for measuring gamma spectra comprises three main parts: (1) the gamma detector, (2) the electronic signal processing and the amplification part, and (3) a multichannel analyzer (for storage of all measured signals) and a computer system for data evaluation. The development of new detectors


have been instrumental in an improvement of INAA techniques over the past decade. More than about 20-25 years ago, only sodium iodide (NaI) scintillation crystals (containing about 0.1 % Tl as an activity impurity) were used for gamma measurements. Gammas interact by the photoelectric effect, the Compton effect, or pair production, creating an excitation in the crystal. The deexcitation creates a light flash; the gamma energy is proportional to the number of light pulses which are measured by a photomultiplier. These detectors are highly sensitive, but have poor energy resolution (typically measured as the full width at half-maximum, FWHM, in keY, at the 1332 keY line of 60Co). They are therefore not suitable for multielement INAA.

The newer semiconductor-type detectors are indispensable for INAA work. Germanium is chosen as the most suitable for gamma detectors. Gamma rays interact by creating an electron-hole pair (i.e., ionization) - creating holes in the valence band and putting electrons in the conduction band. Two general types of Ge crystals exist, depending on the type of impurities present. Ge containing and excess of electron donor atoms is called n-type (excess of negative charges in its conduction band), while Ge containing excess of electron accepting atoms is called p-type (excess of positive holes in its valence band). Because of these impurities, Ge also conducts in absence of ionizing radiation. To compensate for unwanted impurities, lithium atoms are drifted into the Ge crystal, forming an n-type layer (about 1 mm thick) at the crystal surface. Lithium atoms then migrate into the crystal to compensate p-type impurities, leaving an undrifted p-type core. By applying a high-voltage bias between the lithium diffusion layer and the p-type core, the electrons migrate to the positive voltage on the n-side and the holes migrate to the negative voltage on the p-side and create an electrical signal proportional to the amount of gamma energy absorbed in the depletion region. The main advantage of these detectors is a good energy resolution, typically about 2-3 keY at l!332 keY, which is about 30-100 times better than NaI(Tl) detectors. A disadvantage was their relatively low efficiency compared to NaI(Tl) detectors. These detectors have to be kept at the temperature of liquid nitrogen (77 K) all the time, first to ~void damage to the lithium layer, and second to minimize electronic noise.

However, modern detector technology has resulted in the growth of Ge crystals with very low impurity levels « 1010 impurity atoms/cm3). Such material is known as hyperpure or high-purity germanium (HpGe) and does not need charge compensation by lithium drifting. Recent improvements in crystal growth technology have allowed the construction of latge-volume (high efficiency) detectors with very good energy re:solution (around 1.6 keY at 1332 keY), although large detectors are quite expensive. These detectors can be warmed up to room temperature without damage if not in use, but during measurements they have to be cooled. A variation of these detectors is a planar detector (a flat Ge cylinder with small active volume) which is used for the measurement of the low energy part (50-300 keY) of the spectrum.

The voltage pulses from the detector are processed in a preamplifier, a spectroscopy amplifier, and an analog-digital converter (ADC). The preampli-


fier collects and amplifies the electric charge created by the Ge detector, by using a (cooled) field effect transistor (FET). The output signal is a fast rising (50 ns) voltage signal which then exponentially decays with a time constant of about 50 p.s. The height of the voltage step is proportional to the gamma energy. The spectroscopy amplifier performs pulse shaping and baseline restoration and pile-up correction in order to produce "clean" pulses for further processing. The output pulses are near-Gaussian in shape with an amplitude that is proportional to the gamma energy. Recent very fast amplifiers save peak-shaping time by cutting off the second half of the Gaussian peak after detection of the peak maximum. These amplifiers are called "gated integrators". The signal is then fed into an ADC, which converts the analog (voltage) signal into a digitized signal, which may then be stored (as a single event) in a channel of the multi-channel analyzer (MCA). Earlier, dedicated MCAs with histographic memories (because of the need to store 106-108 events per channel) were used, while nowadays larger and fast computer systems are used. Each interaction of the detector with a gamma quantum gives rise to the complex amplification and conversion steps, and results at the end in one single event which is sorted into the memory. Many gamma ray interactions with the same energy yield many events in the same channel, producing a gamma peak in the resulting gamma spectrum.

Data Reduction

Modem MCA memories (or software-simulated MCAs) contain 4096 or 8192 channels of data, spanning the range of about 60-2000 keY. Because of limitations in the detector resolution and some electronic noise, one gamma energy line spreads over more than one channel, ideally forming a Gaussian peak (defined by more than three channels) in the spectrum. The evaluation of gamma spectra involves several steps, such as locating peaks in the spectrum, determining the peak energies and areas, and calculating the elemental concentrations. Because of the large amount of data, these procedures are performed in a set of sophisticated computer programs. Once a peak has been detected by peaksearch algorithms, the peak area can be determined by various methods, most of which require a clean and well-resolved peak. Some programs chose a mathematical function to describe individual peak shapes and can subtract nonlinear peak backgrounds. It is especially important to have peak-fit programs that are able to resolve partially overlapping peaks. This ability, together with a good detector resolution, often allows multielement INAA in cases where before RNAA had to be done. If peaks are too close to each other to be resolved by either method, spectral interference data can be calculated from synthetic standards. An interference-free line of the interfering isotope is then taken to calculate a peak area for the interfering line from the standard data.

To allow comparison of the photopeak areas in the samples with standards, a series of corrections have to be applied. All sample and standard peak areas of all nuclides are recalculated to a common time (decay correction because of


different half-lives and different measurement times), as well as to an equal measurement time (this correction takes into account that for unequal measurement times, isotopes with shorter half-lives have lower count rates at the end of the measurement than at the beginning). Another important correction concerns possible flux variations in the irradiation position between individual samples and standards. Therefore, a flux monitor (e.g., a Au-doped Al flux wire) has to be included in the irradiation vial. A correction could also be required for counting samples and standards at different geometries at the detector. After all these corrections are applied to the peak areas of samples and standards, the concentration is calculated by the simlPle equation: weight of element X in sample = (weight of element X in standard*peak area sample/ peak area standard). Other calculation methods, such as the monostandard method, rely on a set of factors describing the relation of the activities of different elements and use only one single standard element.

Problems

Precision and the accuracy of this method are, if diligently applied, very small. Important factors that affect the precision are (1) weighing errors, (2) geometry errors during counting, (3) flux variatiions during the irradiation, (4) poor counting statistics or errors in the peak area determination, and (5) sampling errors. The first three points are easily taken into account and should not present a problem. Sampling errors should not be more of a problem here than for any other method; in fact, the possibiXity to determine trace elements in very small samples, such as individual crystals or cosmic dust, makes a discussion of sampling errors in such cases obsolete. Poor counting statistics and problems in determining peak areas can occur for very small samples, low elemental contents, or short measurement times. Factors affecting accuracy are (1) interfering nuclear reactions yielding the same product nuclide (especially from the fission of 235U; isotopes that are produced by thermal neutron capture from the respective isotopes, but also from thermal neutron induced fission of 235U include: 95Zr, 99Mo, 140Ba, 140La, 141Ce:, and 147Nd, (2) spectral interferences from overlapping gamma peaks, (3) self shielding in the samples due to the presence of high cross-section elements, (4) dead-time errors (detector and electronic system are busy processing one pulse while missing another one; modem instrumentation has built-in correction circuitry for this problem), and (5) problems with standards. If known to be a problem, most of these factors can be quantified and corrected for. Detection limits for INAA are difficult to estimate because of the many factors involved. They depend on the matrix composition (e.g., the amount of 24Na and other isotopes that dominate a spectrum), counting times and geometry, detector characteristics, sample size, and so on. Typical detection limits range from 10- 3 to 10- 10 gig for different elements.


Experimental

Samples are usually powdered (although individual crystals may also be used), and about 100-500 mg are sealed into polyethylene or high-purity quartz vials (depending on the irradiation duration and flux), and packed (in a defined geometry) together with standards and a flux monitor, into a large irradiation vial. As standards, either high-purity synthetic (multielement) standards or wellcharacterized geological reference materials are used. After the end of the irradiation, a short waiting period is necessary ("cooling period") before the samples can be counted. A first counting period is done about 1-3 days after the end of the irradiation, to determine (e.g.) Na, K, Ga, As, Br, Mo, Sb, W, and Au. A second counting period follows about 3-7 days after the end of the first counting cycle and involves the determination of (e.g.) Na, Cr, Fe, As, Rb, Sb, Zr, Ba, La, Ce, Nd, Sm, Tb, Yb, Lu, Hf, Ta, Au, Th, and U. A third counting cycle is done about 2-3 weeks after the end of the second one, and a fourth cycle can be added any time thereafter. The last measurements involve the determination of (e.g.) Sc, Cr, Fe, Co, Ni, Zn, Se, Rb, Sr, Sb, Zr, Cs, Ce, Nd, Eu, Gd, Tb, Tm, Yb, Lu, Hf, Ta, Ir, and Th. AI, V, Mn, Dy, and others can be determined in an individual short-time (0.5-2 min) irradiation. Other elements usually require radiochemical separation, which is now commonly done as group separation, and not for individual elements.

Summary

INAA is a widely used and powerful analytical method for the determination of about 20-40 elements in geological samples. Because of the selectivity and sensitivity of the method, very small samples (such as individual crystals, e.g., for the determination of partition coefficients; or cosmic dust) or samples with very low contents of certain elements (such as diamonds, or mantle-derived rocks) can be measured successfully without radiochemical separation. A single analysis may take up to 2 months to complete because the samples need to be counted at different times after the end of the irradiation. The method involves complicated and expensive equipment, but rather simple sample preparation and yields high precision results for many elements and plays therefore an important role in geochemical analysis.

References

Baedecker PA, McKown DM (1987) Instrumental neutron activation analysis of geochemical samples. US Geol Surv Bull 1770: HI-HI4

Baedecker PA, Rowe JJ, Steinnes E (1977) Application of epithermal neutron activation in multielement analysis of silicate rocks employing both coaxial Ge(Li) and low energy photon detector systems. J Radioanal Chern 40: 115-146

7.7 Nuclear Techniques for Uranium and Thorium Analysis 329

Das HA, Faanhof A, Van Der Sloot HA (1989) Radioanalysis in geochemistry. Elsevier, Amsterdam

Jacobs JW, Korotev RL, Blanchard DP, Haskin LA (1977) A well tested procedure for instrumental neutron activation analysis of silicate rocks and minerals. J Radioanal Chern 40: 93-114

Koeberl C (1988) Short time activation anallysis in geo- and cosmochemistry. J Trace Microprobe Techniques 6: 501-520

Koeberl C, Kluger F, Kiesl W (1987) Rare earth element determinations at ultratrace abundance levels in geologic materials. J Radioanal Nucl Chern 112: 481-487

Laul JC (1979) Neutron activation analysis of geological materials. Atom Energ Rev 17: 603-695

Wandless GA (1987) Radiochemical neutron activation analysis of geologic materials. US Geol Surv Bull 1770: J1-J8

7.7 Nuclear Techniques for Uranium and Thorium Analysis

S.l. PARRY

Uranium and thorium are naturally radioactive elements. Their principal isotopes, 235U, 238U, and 232Th, are all alpha emitters with half-lives in the order of 108 to 1010 years and specific activities between 0.5 and 12 kBq/g. In addition, they give rise to radioactive decay chains. It is possible to determine the quantity of uranium or thorium in a rock or mineral by measuring the radioactivity of the parent isotope 238U or 232Th or one of the decay products with alpha or gamma ray spectrometry:

" " " (J.Y (J.y 226Ra ----+ 222Rn ----+ 218pO ----+ 214Pb ----+ 214Bi ----+

" (J (J.Y " " 232Th ----+ 228Ra ----+ 228 Ac ----+ 228Th ----+ 224Ra ----+

" " (J.y (J.y " 220Rn ----+ 216pO ----+ 212Pb ----+ 212Bi ----+ 212pO ----+

Other nuclear techniques such as activation analysis can be applied to the isotopes of uranium and thorium and since they undergo fission on bombardment with neutrons, it is also possible to measure the elements using delayed neutron analysis. All these techniques have been described in reviews (Tolgyessy and Kyrs 1989; Das et al. 1989).


Alpba Particles

The alpha particles emitted by isotopes of uranium and thorium or one of the decay products can be identified and evaluated using spectrometry. Silicon charged-particle (surface barrier) detectors have 100% efficiency, high resolution, and a wide energy range. However, the alpha emitter must be separated from the matrix so the sample is presented as a very thin source, since alpha particles travel very little distance even in the air. Thorium is normally measured with the decay product 212pO since it has the greatest energy. Thoron e20Rn) which is produced from 232Th decay chain, can be collected as a gas and measured directly with "emanation analysis". Since thoron is so short-lived, a stream of air at constant flow-rate is passed through a melt of the sample of mineral or rock and the gas is measured as it is emitted. Radon e22Rn) from uranium is long-lived and can be collected and transferred to an ionization chamber or scintillation neutron detector prior to counting.

If a slice of rock is placed in contact with a film of cellulose nitrate or polycarbonate film, the radiation damage from the alpha particles emitted by uranium can be detected after etching the film. Quantitative information can be obtained for activities below mBq. More importantly, if the film is compared to the original section of rock, it is possible to obtain information about the spatial distribution of uranium in the mineral phases.

Gamma Rays

It is simple to measure radioactivity with gamma rays, since they can be detected through rock, allowing for in situ analysis down boreholes or counting of large powdered samples in the laboratory. All the gamma ray emitters can be measured simultaneously, the detection limits depending on the efficiency of the gamma ray detector. Sodium iodide scintillation counters with high efficiency but poor resolution, are rugged and suitable for in situ analysis via the gamma rays of2 14Pb or 214Bi, for uranium, and 212Pb and 2osTI, for thorium. Detection limits are in the order of 1 mg/kg. Potassium can be determined at the same time as 4oK.

Fission Products

When uranium e32U) and thorium e 32Th) are bombarded with neutrons, they undergo fission, resulting in emission of fission products and neutrons. The fission products are unstable and decay with half-lives of less than 1 min, emitting alpha particles which can be recorded with cellulose nitrate film. The method is more rapid than alpha track recording and is usually used for study in the occurrence of uranium in rocks and minerals. The track recorder, when etched, can provide quantitative information about the elemental composition plus spatial distribution.

7.7 Nuclear Techniques for Uranium and Thorium Analysis 331

Neutrons emitted by the decaying fission products can also be used to evaluate the uranium or thorium content of a rock or mineral (AmielI981). The irradiation and analysis times for delayed neutron counting are only 60 s, so the method is very rapid. The measurement of delayed neutrons is a specific technique and there are few interfering reactions. It is common to irradiate several grams of rock powder but because the method is so sensitive for uranium, it can be determined in mineral samples of only a few mg. Thorium can be measured using delayed neutron counting but the sensitivity is poorer than uranium. Thorium is activated by neutrons of a higher energy than uranium and it is possible to differentiate between the elements using different energy neutrons.

Los Alamos National laboratory has a delayed neutrons counting facility which is used for a national uranium re:source evaluation (Minor et al. 1982). The system is used to analyze stream sediments samples. Helium-3 detectors with combined efficiency of about 27% are used to measure uranium with a detection limit of 0.01 Ilg/g of sample. The system is totally automated so that a maximum of 200 samples can be loaded at one time. The delayed neutron counter is only part of a whole system where the sample can be measured for both neutrons and gamma rays in sequence. It is also possible to measure some elements such as F, AI, Ca, and V with gamma ray spectrometry during the decay period before the uranium is measured with neutrons (Shenberg et al. 1987).

Activation Spectrometry

Uranium and thorium can also be determined by neutron activation and gamma ray spectrometry via the 238U(n,y)239U, and 232Th(n,y)233Th reactions. High concentrations of uranium and thorium can be measured after just a few minutes, as 238U and 233Th, which have half-lives of less than 30 min. It is also possible to measure the decay products 239Np and 233Pa with greater sensitivity; they have a half-life of 2.56 and 30 days, respectively, and can be measured at the same time as the rare earth elements.

References

Amiel S (1981) Neutron counting in activation analysis. In: Amiel S (ed) Nondestructive activation analysis, Studies in Analytical Chemistry 3, Elsevier, Amsterdam, pp 43-52

Das HA, Paanhof A, Van der Sloot HA (1989) Radioanalysis in geochemistry. Elsevier, Amsterdam

Minor MM Hensley WK, Denton MM, Garcia SP (1982) An automated activation analysis system. J Radioanal Nucl Chern 70(1-2): 459--471

Shenberg C, Nir-EI Y, Alfassi Z, Shiloni Y (1987) Rapid and simultaneous determination of U, F, AI, Ca and V in phosphate rock by a combination of delayed neutron and y-ray spectrometry techniques. J Radioanal Nucl Chern 114(2): 367-377

Tolgyessy J, Kyrs M (1989) Radioanalytical cht:mistry volume I. Ellis Horwood, Chichester


7.8 Mass Spectrometry

P.F. McDERMOTT

In contrast with the many and varied applications of mass spectrometry in other branches of science, (e.g., in the life sciences, physics, organic and physical chemistry), mass spectrometry in the Earth and planetary sciences is concerned almost exclusively with the measurement of isotope abundance ratios. Radiogenic isotopes of Sr, Nd, Pb, and Ar are measured routinely, and increasingly the isotopes of Th, Ca, Hf and Os are finding applications (e.g., Faure 1986). Traditionally, radiogenic isotopes were used as chronometers for time-scales in the 103 to 109 -year range, but a more recent trend is the use of isotopes as chemical tracers. The stable isotopes of 0, H, C, N, and S continue to be important as tracers, and as constraints on the physical and chemical processes which generate terrestrial and extraterrestrial materials. The cosmogenic isotopes of lOBe and 26 Al have received considerable attention over the past decade, both as chronometers of surfical deposits (e.g., Guichard et al. 1978) and as tracers of recycled near-surface material in the Earth's mantle (e.g., Morris et al. 1990).

Radiogenic and stable isotope data obtained for terrestrial and extraterrestrial materials using mass spectrometry have had a considerable impact on the development of geochemistry and cosmochemistry over the past two decades. In geology, the "absolute" dates obtained from mass spectrometry have made an important contribution to the traditional subjects of stratigraphy and historical geology, where they provide an "absolute" timeframe for the calibration ofpaleontolgical "relative" ages. Mass-spectrometric data have also given rise to the new field of isotope geology, which exploits isotopes as tracers of complex petrogenetic processes throughout Earth history. These data allow geochemists to quantify interactions between the major Earth reservoirs of the mantle, crust, and hydrosphere, and to construct internally consistent wholeEarth evolution models (e.g., Allegre et al. 1983; Jacobsen 1988). Moreover, the recent recognition that a better understanding of past environmental change is a prerequisite to quantifying anthropological effects has provided an important new impetus to the study of short-lived isotopes (e.g., the U series), and the chronology of the recent geological past (e.g., Edwards et al. 1987).

Instrumentation

A mass spectrometer is an instrument designed to separate a beam of ions according to their mass to charge (m/z) ratio, and to measure electronically the intensity of the separated ions. Mass spectrometers are high vacuum instruments, and typically operate at pressures of 10- 7 to 10- 9 mbar. The design of mass spectrometers has evolved steadily over the past four decades, but they all

7.8 Mass Spectrometry 333

incorporate three essential parts, namely; an ion source, a mass analyzer and one or more ion detectors (collectors). Three types of mass spectrometer are commonly encountered in geochemical and cosmochemical applications, and the difference between these types is in the mode of operation of the massanalyzer device. The three types considered here are (1) magnetic sector instruments (2) quadrupole mass-spectrometers and (3) time of flight mass-spectrometers.

Magnetic Sector Instruments. The maJonty of commercially available mass spectrometers used routinely for the isotopic analysis of geological and planetary materials are single-focusing instruments which rely on a wedge (sector)shaped electromagnet to separate the ion beam into ions with discrete m/z ratios. The design of many commercial instruments in use today has evolved from the original design of Nier (1940), and these instruments typically have a 60° magnetic sector with the source and collectors arranged symmetrically about the magnet. In this design, the source exit slit, the magnet sector axis and the collector entrance slit are collinear, although in practice minor adjustments of the magnet relative to the analyzer (flight) tube are usually required. Magnetic sector instruments may be designed to measure isotope ratios in both solid and gaseous samples depending on source design, hence the distinction between "solid source" and "gas source" machines. Ions are generated either by thermal ionization (in a solid source) or electron bombardment (in a gas source) and are accelerated from the source using high voltages (usually 5-10 kv).

The equations of motion of an ion in a mass spectrometer are described in detail elsewhere (e.g., Watson 1986) and will not be repeated here. The essential point is that the radius of curvature of an ion's path through the flight tube depends on its momentum. One consequence is that an ion with a large mass is deflected along a path with a large radius compared with a lighter ion, thereby achieving a separation according to mass. Sequential collection, or so-called peak-switching, requires that each of thl~ separated ion beams is brought into focus at the collector slit in a pre-determined sequence, either by varying the electromagnet current or the high voltage applied to the source. Modern mass spectrometers usually incorporate several Faraday collectors, and at least one high sensitivity secondary electron multiplier or Daly detector. The collector positions are set automatically under computer control to the positions appropriate for the isotopes of interest. Thus, in modern mass spectrometers peak switching is usually unnecessary, since the individual dispersed ion beams may be measured simultaneously in different collectors. The advantages of multicollector instruments are numerous, and are discussed in some detail by Potts (1987).

As discussed above, a magnetic sector analyzer separates ions according to their momentum, and to achieve a perfect separation all the ions with the same m/z ratio must have identical kinetic energies (E = 1/2 mV2). In practice, however, most ion sources produce ions with a range of initial kinetic energies, so that the ions entering the magnetic sector are not mono-energetic. This


problem is alleviated in double focusing instruments by positioning an electric sector between the ion source and the magnetic sector. The electric sector effectively blocks out those ions with energies significantly different from that of the main beam, focusing only ions with energies close to that of the main beam. Double focusing mass spectrometers commonly employ one of two geometrical arrangements - the Mattauch-Herzog and the Nier-Johnson geometries (e.g., Watson 1986). Modified Nier-Johnson geometries are commonly employed on ion probes whereas the Mattauch-Herzog geometry is traditionally used in spark-source mass spectrometers.

Quadrupole Mass Spectrometers. Quadrupole mass spectrometers operate on an entirely different principle, and unlike the magnetic sector analyzer, the quadrupole is an ion-path stability analyzer. Physically, a quadrupole consists of four symmetrically arranged by cylindrical metal rods, all of which are connected to direct current (DC) and radio frequency (RF) potentials (Dawson 1976). Diagonally opposed rods are connected electrically to the DC and RF generators, and the superposition of the DC and RF fields produces a complex pattern of equipotential lines. Ions are introduced from the source at one end, and drift slowly in a direction parellel to the long axis of the quadrupole, in response to a relatively small potential difference. The ions are influenced by the combined DC and RF fields, and for a given set of conditions only ions of one m/z ratio can pass through the quadrupole to the ion detector. Ions with other m/z ratios follow different trajectories and are captured and neutralized by one of the metal rods. Quadrupole analyzers are useful, relatively low cost devices, and are suitable for applications which do not require very precise isotope ratio measurements. For example, typical external relative errors (two standard deviations on the mean of several measurements expressed as a percentage), on a 87Sr/86Sr ratio is ± 0.6% for a quadrupole mass spectrometer, compared with 0.006% for a conventional magnetic sector single-focusing instrument. Thus, in the geological and planetary sciences, quadrupole instruments are commonly used in isotope dilution mass spectrometry (see below) for which this precision is adequate. A relatively recent development is the coupling of a quadrupole mass spectrometer to a radio frequency induced, coupled argon plasma as an ion source (ICP-MS) (qv). Quadrupoles are especially suitable for rapid scanning of the mass range, and in ICP-MS this property allows "simultaneous" multielement determinations.

Time of Flight (TOF) Instruments. These instruments determine the velocity, and thereby the mass of an ion by measuring the time taken to travel from the ion source to a detector (e.g., Cameron and Eggers 1948; Damoth 1964). This type of mass spectrometer operates in a pulsed mode and so the ions are accelerated episodically towards the detector. The ions are all accelerated to the same kinetic energy, but become separated en route to the detector by virtue of their mass differences. Ions which arrive first at the detector have the lowest m/z ratio, and this ratio is higher for the late arrivals because a lower velocity is

7.8 Mass Spectrometry 335

imparted to heavier ions. It seems likely that TOF instruments will become increasingly popular in the Earth and planetary sciences, because with precise synchronisation their mode of operation allows efficient collection of ions from new pulsed (e.g., laser) ionisation sources.

Mass Spectrometric Techniques

Conventional solid source mass spectrometry requires the extraction of the element of interest from the bulk (usually silicate) rock to minimize isobaric interferences and enhance ionization efficiency. Silicates are dissolved in HF/HN03 mixtures and the solution is usually evaporated to dryness under an infrared lamp with a clean-air supply. When dissolved, the sample is taken up in a solution of HN03 or HCI, and loaded onto ion-exchange columns for separation of the element of interest. Separation of some trace elements requires special precautions to minimize blank It:vels, and in the separation of Pb, for example, this is achieved by using small ion-exchange columns ( < 30 Ill), small volumes of acid for elution and careful handling of the sample digestion vessels. The separated element is loaded as a salt on a Ta, Re or W metal filament which then serves as the ion source when heated in the mass spectrometer. Sample loading procedures vary, but they usually involve pipeting a few microliters of sample solution onto the filament together with a low-blank additive material such as phosphoric acid (for Sr), silicagel and phosphoric acid (for Pb), and colloidal graphite (for V and Th). Sample loading remains something of an art, and with experience most operators devdop their own preferred techniques.

In addition to isotope ratio data, the concentrations of some trace elements can be measured by mass spectrometry using the isotope dilution technique. This involves mixing a known weight of the sample with a known weight of spike solution. A spike solution is a solution containing a known concentration of the element of interest, the isotopic composition of which is significantly different (one of its isotopes is enriched) from that of the sample. The concentration and isotopic composition of the element in the spike is known, so by measuring the isotopic composition of the sample/spike mixture, the concentrtion of the element in the sample may be calculated. Details of the technique are given elsewhere (e.g., Faure 1986; Potts 1987).

Gas source mass spectrometry usually requires prior extraction and purification of the gas before it is introduced into the mass spectrometer. The extraction and purification of hydrogen (e.g., Kyser and O'Neill 1984), oxygen (e.g., Clayton and Mayeda 1963), carbon (e.g." Des Marais and Moore 1984), sulfur (e.g., Veda and Sakai 1983), and the noble gases (e.g., Craig and Lupton 1976) require specialized gas extraction and purification lines. Gas source mass spectrometers are conventionally operated in "dynamic" mode, but recent developments which allow static sample introduction have resulted in sensitivity improvements by factors of 102 to 103 (pillinger 1984).


New Developments

New developments in mass spectrometry applied to the Earth and planetary sciences are likely to include (1) new ionisation techniques (2) improved abundance sensitivity (3) enhanced detector sensitivity.

Most solid source mass spectrometers currently in use by geochemists rely on the relatively inefficient process of thermal ionisation to convert the atoms of the sample into ions. Thermal ionization efficiencies are seldom better than a few percent, and are as low as a few permil for some elements such as U, Th, and Hf By contrast, new techniques which employ laser resonant ionization (RIMS), (e.g., Conzemius and Capellan 1980), secondary ionization (SIMS), (e.g., Evans 1972) and new combined ICP magnetic sector instruments (ICPMCMS) (Halliday et al. 1994) promise an order of magnitude or more improvement in ionization efficiency. At present, the low ionization efficiency of thermal ionization limits the precision obtainable on the measurement of low abundance isotopes (e.g., 23°Th) to about 0.5% (20').

Abundance sensitivity is a measure of the interference oflarge peaks on their adjacent lower mass neighbours. In most commercial single-focusing solidsource mass spectrometers abundance sensitivity is measured as the "tail" from mass 238U at mass 237; this tail is typically 2 ppm of the 238U peak. This tailing is caused mainly by inelastic collisions between ions and residual gas molecules in the source and flight-tube and it severely limits the precision obtainable on measurements of extreme isotope ratios such as 230Thj232Th (commonly < 8 x 10- 6 in silicates). Recent developments include electrostatic filters and static quadrupole devices, which unlike the double-focusing instruments described above, are positioned between the magnetic analyzer and the ion collectors. These devices produce an order of magnitude or better improvement in abundance sensitivity in conventional single focusing magnetic sector instruments and so facilitate the precise measurement of isotope ratios of < 1 x 10 - 5.

In future, magnetic sector instruments are likely to incorporate several high sensitivity ion detectors, for example, secondary electron and channel-plate multipliers, which will enable simultaneous collection of several low intensity ion beams in ion-counting mode. Such improvements are already finding application in the measurement of picogram levels of Th and U in calcium carbonates, and may also find application in the measurement of other isotopes (e.g., Pb) in very depleted volcanic rocks and small fractions of separated minerals. ,

References

Allegre CJ, Hart SR, Minster JF (1983) Chemical structure and evolution of the mantle and continents determined by inversion of Sr and Nd isotopic data. I, Theoretical methods. Earth Planet Sci Lett 66: 177-190

Cameron AE, Eggers DF Jr (1948) An ion "velocitron". Rev Sci Instrum 19: 605-607 Clayton RN, Mayeda TK (1963) The use of bromine pentaftuoride in the extraction of oxygen

from silicates for isotopic analysis. Geochim Cosmochim Acta 27: 43-52

7.9 Inductively Coupled Plasma Mass Spectrometry 337

Conzemius RJ, Capell an JM (1980) A review of the application to solids of the laser ion source in mass spectrometry. Int J Mass Spectrom Ion Phys 34: 197-271

Craig H, Lupton JE (1976) Primordial neon, helium and hydrogen in mid-oceanic basaltic glasses. Earth Planet Sci Lett 31: 369-385

Damoth DC (1964) Recent advances in time-of~ftight mass spectrometry. In: Reilly CN (ed) Advances in analytical chemistry and instrumentation. Wiley-Interscience, New York, pp 371-410

Dawson PH (ed) (1976) Quadrupole mass spectrometry and its applications. Elsevier, Amsterdam pp 349

Des Marais DJ, Moore JG (1984) Carbon and its isotopes in mid ocean basaltic glasses. Earth Planet Sci Lett 69: 43-57

Edwards RL, Chen JH, Ku TL, Wasserburg GJ (1987) Precise timing of the last interglacial period from mass-spectrometric determination of Thorium-230 in corals. Science 236: 1547-1553

Evans CA Jr (1972) Secondary ion mass analysis: a technique for three-dimensional characterisation. Anal Chern 44: 67A-80A

Faure G (1986) Principles of isotope geology. New York, Wiley, pp 589 Guichard F, Reyss JH, Yokoyama Y (1978) Growth rate of a manganese nodule measured with

lOBe and 26AI. Nature 272: 155-156 Halliday AN et a!. (1994) Inductively coupled plasma magnetic sector multi-collector mass

spectrometry. ICOG 8, Abs Vol USGS Circular 1107, p122 Jacobsen SB (1988) Isotopic and chemical constraints on mantle-crust evolution. Geochim

Cosmochim Acta 52: 1341-1350 Kyser TK, O'Neill JR (1984) Hydrogen isotope systematics of submarine basalts. Geochim

Cosmochim Acta 48: 2123-2133 Morris JD, Leeman WP, Tera F (1990) The subducted component in island arc lavas:

constraints from Be isotopes and B-Be systematics. Nature 344: 31-36 Nier AO (1940) A mass-spectrometer for routine isotope abundance measurements. Rev Sci

Instrum 11: 212-216 Pillinger CT (1984) Light element stable isotopes in meteorites - from grams to picograms.

Geochim Cosmochim Acta 48: 2739-2766 Potts PJ (1987) A handbook of silicate rock analysis. Glasgow, Blackie, pp 622 Ueda A, Sakai H (1983) Simultaneous determinations of the concentrations and isotope ratio

of sulphate-sulphur, sulphide-sulphur and carbonate-carbon in geological samples. Geochim J 17: 185-196

Watson JT (1986) Introduction to mass spectrometry. Raven Press, New York, pp 351

7.9 Inductively Coupled Plasma lVlass Spectrometry

K.E. JARVIS

The development of inductively coupled plasma mass spectrometry (ICP-MS) is relatively recent, with the first commercial instruments being launched in 1983. The technique offers several advantages over conventional methods of analysis such as atomic absorption, X-ray fluorescence, instrumental neutron activation analysis, and ICP-atomic emission spectrometry for multielement determination including simple spectra throughout the mass range even for complex matrices. Sensitivity is relatively uniform for all elements from Li to U. The detection limits for most elements, particular the heavier ones such as the REE, Ta, Nb, Hf, Th and U are exceptionally low, typically less than 0.05 ng ml- 1

(Jarvis 1990; Hall et al. 1990). In addition, the technique has a wide linear


dynamic range over six to seven orders of magnitude. Analysis times are rapid, typically 1 min per sample, and calibration can usually be carried out using simple synthetic solutions containing the elements of interest (Date and Gray 1989; Jarvis et al. 1992). Samples are usually analyzed in solution form although the direct analysis of solids, i.e., by slurry nebulization or laser ablation (Jarvis and Williams 1989, 1993) and gases is possible.

The system consists of an inductively coupled argon plasma, which is used as an ion source for a quadrupole mass spectrometer. Ions (mostly singly charged species) are extracted from the plasma through a Ni sampling cone in an interface between the ICP and the mass spectrometer. Before entering the mass spectrometer, the ions are focused by a set of electrostatic lenses to form a narrow concentrated beam. The quadrupole can be scanned through a range of masses at between tOO and 1600 times per minute allowing data to be collected for all elements from 4-238 m/z in a single analysis.

Aside from its use for elemental determination, ICP-MS may also be used for isotope ratio measurements, although precision is currently limited to about 0.1 % for two equal abundance isotopes, e.g., 206Pb: 207Pb (e.g., Gregoire 1987). Modern commercial instruments are fully automated with instrument control, safety monitoring, data handling, and processing being carried out by a 386 or equivalent personal computer.

The instrument may be operated in a number of ways. Spectra can be collected over the complete mass range, allowing a qualitative investigation of a sample to be made. This feature is particularly useful when new matrices are encountered or simply to establish which elements are present in a sample. Alternatively, semiquantitative measurements can be made using a single standard solution containing about six elements across the mass range. A response curve for these elements is established and unknown samples measured against this curve. The data measured in this way is typically of about 50% accuracy, although this is highly dependent on the element concerned. Finally, fully quantitative data may be obtained using either external calibration procedures or isotope dilution.

Interference effects in ICP-MS fall into two distinctive categories. The first are spectroscopic in nature and are due to isobaric overlap, polyatomic, doubly charged, and oxide ion formation. The second group are termed nonspectroscopic and include the effects of high levels of dissolved solids present in solutions and those caused by the presence of high concentrations ( > lOO0 Ilg ml- 1) of a single element. The formation of polyatomic ions is perhaps the most serious type of spectroscopic interference. For example, Ar the plasma gas, combines with 0 to form 4°Ar160 causing a serious interference at 56 m/z, the main isotope of iron. Other combinations take place, particularly between Ar, 0, H, and N, and elements such as Cl and S derived from the mineral acids used for sample dissolution. Many spectroscopic interferences can, however, be minimized or even eliminated by careful sample preparation procedures and correct system optimisation (Gray and Williams 1987). Nonspectroscopic effects are, in practice, not usually serious in geological matrices providing that the level of total dissolved solids in the solution presented for

7.9 Inductively Coupled Plasma Mass Spectrometry

Table 13. A comparison between measured and reli~rence values for the REE in argilaceous limestone NIST -1 b

Element ICP-MS'

La 7.31 Ce 7.74 Pc 1.38 Nd 5.36 Sm 0.82 Eu 0.24 Gd 1.01 Tb 0.16 Dy 1.03 Ho 0.21 Er 0.62 Tm 0.08 Yb 0.59 Lu 0.08

Concentrations in JIg g - 1.

- No data available.

Referenceb

6.86 7.81 1.18 4.88 0.89 0.24 0.97

0.90 0.20 0.57

0.55 0.08

a ICP-MS data from Jarvis (1990). bReference values from Gladney et al. (1987).

339

analysis is reduced to < 2000 jlg ml- 1. Above this level, signal loss with time is rapid and precision poor due to partial blocking ofthe Ni sampling cone orifice (Williams and Gray 1988).

An example of the data obtained for aU 14 of the REE for NIST SRM -1 b, an argillaceous limestone, are shown in Table 13. These data were measured in a single 1-min analysis in a solution prepared by HF/HCI04 open digestion. The REEs were not separated from the matrix prior to determination. The agreement between measured and reference values is excellent and the precision at this concentrations is about 10% RSD.

In summary, the technique can be used for the rapid quantitative determination of nearly all elements in the periodic table making in an invaluable tool for geochemical studies. Sensitivity is particularly good for many of those elements which are difficult to determine by other analytical techniques.

References

Date AR, Gray AL (eds) (1989) Applications of inductively coupled plasma mass spectrometry, Blackie, Glasgow, 254 pp

Gladney ES, O'Malley BT, Roelandts I, Gills TE (1987) Compilation of elemental concentration data for NBS clinical, biological, geological and environmental standard reference materials. NBS Spec Publ 260-111, 547 pp

Gray AL, Williams JG (1987) System optimisation and the effect on polyatomic, oxide and doubly charged ion response of a commerciaI inductively coupled plasma mass spectrometry instrument. J Anal At Spectrom 2: 599-606


Gregoire DC (1987) Determination of boron isotopes in geological materials by inductively coupled plasma mass spectrometry. Anal Chern 59: 2479-2484

Hall GEM, Pelchat JC, Loop J (1990) Determination of zirconium, niobium, hafnium and tantalum at low levels in geological materials by inductively coupled plasma mass spectrometry. J Anal At Spectrom 5: 339-349

Jarvis KE (1990) A critical evaluation of two sample preparation techniques for low level determination of some geologically incompatible elements by inductively coupled plasma mass spectrometry. Chern Geol 83: 89-103

Jarvis KE, Williams JG (1989) The analysis of geological samples by slurry nebulisation inductively coupled plasma mass spectrometry (ICP-MS). Chern Geol 77: 53-63

Jarvis KE, Williams, JG (1993) Laser ablation inductively coupled plasma MASS Spectrometry (LA-ICP-ms): a rapid technique for the direct, quantitative determination of major, trace REE in geological samples. Chern GEOL 106: 251-262

Jarvis KE, Gray AL, Houk RS (1992) A handbook of inductively coupled plasma mass spectrometry. Blackie, Glasgow, 375 pp

Williams JG, Gray AL (1988) High dissolved solids and ICP-MS: are they compatible? Anal Proc 25: 385-388

7.10 Ion Exchange Techniques

P.J. POTTS

Ion exchange separation procedures are used to isolate an analyte from other species in a sample solution with the aim of separating the element of inter-est from matrix and/or interfering elements. Ion exchange procedures are, therefore, incorporated into schemes of analysis for the following reasons:

1. To act as a preconcentration technique and therefore extend the sensitive range for the determination of an analyte beyond that which would otherwise be achievable (an example of this would be in the determination of the rare earth elements by ICP-AES).

2. To remove element specific interferences that would otherwise swamp the desired analytical signal (for example in determinations by thermal ionization mass spectrometry using isotope dilution techniques).

3. To reduce background intensities and so extend the analytical range by enhancing signal-to-background ratios (for example in radiochemical neutron activation analysis in which the analyte isotope is separated from the matrix activity with a consequent reduction in the detected gamma-ray background).

Ion exchange separations are very versatile and can readily be adapted to various analytical procedures, offering the advantage that it is simple to regenerate ion exchange columns for use in subsequent analysis. However, it is worth bearing in mind that other separation techniques are available including:

1. Solvent extraction in which the analyte is complexed with an appropriate chelating reagent, the complex then being selectively extracted from usually aqueous to immiscible organic phase.

7.10 Ion Exchange Techniques 341

2. Fire assay, a procedure normally used specifically for gold and the platinumgroup elements in which the sample is fused with a suitable flux and the noble metals scavenged into a molten lead or nickel sulfide button. After cooling, the analytes are separated from the slag and dissolved for subsequent analysis.

Ion Exchange Resins

An ion exchange resins comprises a substrate to which functional groups are attached. The substrate normally consists of a polymer formed from styrene with varying proportions of divinylbenzene. The latter is added to ensure that an appropriate degree of cross-linking is formed in the polymerized substrate to resist swelling and increase the number of functional groups that can be attached per unit volume. The functional groups give the resin its specific analytical characteristics and can normally be classified as anionic, cationic, and chelating. An example of a cationic functional group is the suI phonic (-S03 ·H+) or carboxylic (-COO- ·H+) acid groups. Anionic resins contain amino groups (-NRt ·Cl-, were R is an alkyl group), whereas conventional chelating resins contain the -CH2N(CH2COO- . H+)2 group (i.e., one half of the EDT A molecule).

In practical applications, the resin is: formed into a column, down which appropriate solutions are allowed to percolate. Depending on their chemical characteristics, ions in solution suffer different affinities for the functional groups on the ion exchange resin. Taking as an example a cation exchange column containing the -S03 H + functional group, positively charged ions in solution will suffer different affinities for the sulfonate group based on the equilibrium described by the equation:

-S03 ·H+ + Ca~q+ ¢> -S03 ·Ca2+ + H~ bound to in bound to in resin solution resin solution

The affinity with which a positively charged ion is held on the resin will depend on the electrostatic interaction between ion and functional group and will vary according to its positive charge (doubly charged ions will be bound more strongly than singly charged ions) and size (smaller ions will be bound less strongly than larger ions since the former have larger electrostatic fields which attract larger solvation shells and restrict how close the solvated ion can approach the functional group). For these reasons, the selectivity of ions towards a strongly acidic cation exchange resin is Fe3+ > AI3+ > Ba2+ > Pb2+ > Sr2+ > Ca2+ > Ni2+ > Cu2+ > Zn2+ > Mg 2+ > Mn2+ > Ag+ > Cs+ > Rb+ > K+ > NHt > H+ > Li+. By similar reasoning, the order of selectivity towards a strongly basic cation exchange resin is SO~- > oxalate2- > 1- > HSOi > NO;I > Br- > HS03 > NO; > Cl- > HC03 > H2POi > acetate- > F- = Olr.


Practical Applications

In practice, therefore, a sample solution is "loaded" onto the top of an ion exchange column and eluted with a suitable solvent. Ions are then progressively washed out of the column at different rates according to their selectivity characteristics (see above). Most ion exchange procedures involve eluting, and then discarding the first fractions which pass through the column (since these often contain major element ions), then eluting the ions of interest for subsequent analysis. Elution conditions must be carefully optimized to ensure that the element of interest is separated completely from any interfering elements.

An example of an ion exchange procedure that has been successfully adapted to the determination of the rare earth elements by ICP-AES is that of Walsh et al. (1981). The ion exchange medium used in this procedure is a strong cation exchange resin (Dowex AG 50WX8) made up into a column of about 100 x 20 mm diameter and preconditioned by washing with 4N HCI and then IN HCI solution. Following an HF /perchloric acid digestion of 0.5 g of rock, the sample is taken up in hydrochloric acid and diluted to less than 10% HCI solution. An aliquot of this sample is loaded onto the column and eluted with 400 ml 1. 7N H CI solution to remove major and various trace elements but retaining quantitatively the REE, Ba, and some Sr, Hf, and Zr. The REE are then recovered from the column by a second elution using 500 ml of 4N HCI. The resultant solution is evaporated to dryness and then redissolved in 5 ml of 10% HCI for analysis by ICP-AES. Using such a procedure, ICP-AES detection limits for lanthanum, for example, might be improved from 3-4 ppm (for a direct determination on the sample solution) to 0.07-0.3 ppm (for determination after ion exchange separation). Similar procedures are widely used in thermal ionization mass spectrometry but differ in that they are usually "miniaturized" in terms of both the size of the column and volume of eluant required so that the magnitude of the reagent blank can be minimized.

Chelating resins depend for selectivity on the strength of bonding between analyte ion and chelating functional group. In addition to the conventional chelating resin referred to above, interest has also been shown in alternative resins including polydithiocarbamate and Sraffion NMRR, the later being specific to the ions that form square planar complexes, being of value in the separation of the platinum-group elements.

Concluding Remarks

From the outline given above, it is clear that manual ion exchange procedures are relatively time-consuming. For this reason, interest has been shown in the design and development of fully automated separation stations in which a bank of ion exchange columns are serviced by microprocessor controlled automatic pipets. Routine instrumental techniques are normally restricted to the determination of those elements that do not require the use of separation procedures.

7.10 Ion Exchange Techniques 343

Ion exchange separations are best exploited where the additional complexity of sample preparation can be justified by the value of the resultant data, the determination of REE by ICP-AES being a typical example.

References

Grimshaw RW, Harland CE (1975) Ion-exchange: introduction to theory and practice. The Chemical Society, London

Miyazaki A, Barnes RM (1981) Complexation of some transition metals, rare earth elements and thorium with poly(dithiocarbamate) chelating resin. Anal Chem 53: 299-304

Nadkarni RA, Morrison GH (1974) Determination ofthe noble metals in geological materials by neutron activation analysis. Anal Chem 46: 232-236

Paterson R (1970) An introduction to ion exchange. Heyden, London Potts PJ (1987) A handbook of silicate rock analysis. Blackie, Glasgow, chap 14, pp 472-485 Walsh IN, Buckley F, Barker J (1981) The simultaneous determination of the rare earth

elements in rocks using inductively coupled plasma source spectrometry. Chem Geol 33: 141-153

CHAPTER 8

Isotopic Mineralogy

346 Chapter 8. Isotopic Mineralogy

8.1 Radioactive Isotopes in Mineralogy and Geochemistry

YU. A. SHUKOLYUKOVand K. WETZEL

Among the 340 isotopes of the 98 chemical elements occurring in the minerals of the Earth, Moon, Mars, and meteorites, the majority are nonradioacitive. Many of them are absolutely stable in accordance with the law of conservation of energy, as the energy potential of the system originated by their radioactive decay would exceed that of the initial system. For instance, IX-decay of 160 or fission of 28Si into 160 and 12C are absolutely prohibited. Another part of the "stable" isotopes is actually capable of radioactive decay. For instance, in zircon, the atomic nucleus of 96Zr could transform into two nuclei of 48Ca. In this and many other similar cases, however, at the energy efficiency of the radioactive decay, the latter practically does not occur, due to the excessive height and low quantum-mechanical permeability of the energy barrier. Only in some cases, as the sensitivity of the measurements increases, is it possible to detect the extent of radioactivity in such isotopes. For instance, only one atom of 13°Te isotope decays in 1 g of mineral tellurobismutite per month.

Explicit radioaGtive properties are exhibited by about one fourth of the isotopes existing in the minerals (about 80). The sources of their radioacitivity are very different.

8.1.1 Radioactive Isotopes of Nucleogenetic Nature

In the complex nucleogenetic processes of formation of matter in the solar system on various evolution stages of the stars of different types of stars not only stable but also radioactive isotopes with different ratios of a number of neutrons (N) and protons (Z) in the nucleus were produced.

Depending on the deviation of the N/Z ratio from the optimum value corresponding to the stable nuclei, the produced isotopes decayed at a higher or lower rate, tending to transform into stable isotopes with maximum favorable N/Z ratios in terms of energy.

Taking this into account, it is arbitrarily possible to subdivide the isotopes of nucleogenetic origin into three subgroups.

Very Short-Lived Radioactive Nucleosynthetic Isotopes

It is difficult today to clarify the details of formation of all the isotopes with a half-life period ranging from seconds to thousands of years that have formed in the course of the explosive processes occurring in the star shells, but the fact of their existence in the stars is confirmed by the presence of their decay produced in some mineral phases of meteorites. The methods of ion microprobe and

8.1.1 Radioactive Isotopes of Nucleogenetic Nature 347

selective dissolution ofthe meteoritic samples enabled the discovery that certain mineral particles with isotopic anomalies of many elements have formed once as a result of the "instantaneous" decay of their extremely short-lived precursors, that originated in the stars. The conccmtration of these mineral particles in the total mass of the meteorites is very small: from ~ 10-6 (SiC) to ~ 10- 4 (diamonds). They are small in size, i.e., from 10 fJm (elementary carbon) to 0.1 fJm (SiC, Cr sulfides) and e:ven to 50 A (diamond). Their mineral composition indicates their formation a.t very high temperatures. The isotopic anomalies of the light elements in the minerals therein are extremely high: J 13C to 30000%0, J 22Ne to 106%0, J D to 5000%0, J 136Xe to 1000%0, - 400%0 ~ J15N ~ 1000%0. Such particles occur in carbonaceous, ordinary,

and enstatite chondrites, and in iron meteorites as part of the silicate inclusions. These massive isotopic anomalies indicate the formation of the correspond

ing mineral particles in the envelope of various stars (novae, supernovae, red giants) in various stages of their evolution in r-, S-, p- and e- processes of nucleosynthesis.

In the isotopic composition of many polyisotopic elements (Sm, Nd, Ba, Sr, Sc, Si, V, Mn, Ni, Ti, Cr, etc.) in high-temperature mineral inclusions of chondrites, anomalies were detected, indicating the simultaneous existence of the components formed during various processes of nucleosynthesis (r, s, e, p) outside the Solar System. Thus, many mineral phases of the meteorites are essentially the "ancient star ash" which once contained very short-lived radioactive isotopes, leaving their traces behind in the form of isotopic anomalies.

The study of these anomalies is the key to cognition of the isotopic heterogeneity of the Solar System and the conformity of its early evolution to the natural laws.

"Extinct" Radioactive Isotopes of Nuclf:ogenetic Origin in Minerals

The isotopes belonging to this subgroup formed, probably, during one or several acts of r-nucleosynthesis in the super-novae immediately before the origination of the Solar System. In spite of their much higher half-life periods in comparison to the previous subgroup, these isotopes are fully decayed by the present time. In this case, however, it is possible to establish umambiguously the link between them and the stable isotopes formed from them (Table 14).

Two major problems are solved by using products of the radioactive decay of "extinct" isotopes in meteoritic minerals. The first problem is to determine the time interval between the instant of the termination of nucleosynthesis and the instant of forming the first mineral particles in the protoplanetary gas-dust cloud (more precisely, the instant of the mineral particles cooling down to a temperature at which the migration of the studied elements therein practically discontinued). The interval varies from several to tens of millions of years.

The second problem solved by using products of the radioactive decay of "extinct" isotopes is the study of the chemical elements of r-nucleosynthetic


Table 14. Some "extinct" nucleogenetic radioactive isotopes and the products of their decay in minerals of meteorites

Radio- Type of Half-life Radio- Mineral phase containing active decay period, Ma genic "extinct" isotope isotope isotope

2 3 4 5

41Ca Electron 0.13 41K High-temperature silicate capture inclusions rich in Ca and poor in

K in carbonaceous chondrites

26AI p+ 0.70 26Mg Anorthite, melilite, spinel in carbonaceous chondrites

HMn Electron 3.7 HCr High-temperature inclusions in capture carbonaceous chondrites

I07Pd p- 7.0 107Ag Metallic phase of iron meteorites with high Pdf Ag ratio

129J p-, Y 16 129Xe Bulk samples, silicates chondrules of carbonaceous ordinary and enstatite chondrites; silicate inclusions in iron meteorites

244pU IX, Spontaneous 82 136- 131 Xe Bulk samples of achondrites, fission high-temperature inclusions in

carbonaceous chondrites, phosphates (vitlokite) in common chondrites

146Sm IX 103 142Nd Some acid-insoluble mineral phases of carbonaceous chondrites

processes and the estimation of the primary isotopic heterogeneity of the Solar System.

Long-Lived Nucleogenetic Radioactive Isotopes

Also formed in the course of the primary nucleosynthesis were the isotopes with a half-life period over 0.5 billion years. They still exist as part of the minerals of the Earth, Moon, and meteorites (Table 15).

The long-lived radioactive isotopes form the basis of isotopic geochronology. To determine age t, it is necessary to identify the content of radioactive (M) and radiogenic (DR) isotopes in a mineral:

t=-ln -+ 1 1 (DR ) A.M M '

where A.M is the constant of the decay rate.


Table 15. Some long-lived radioacitve isotopes and products of their decay in Earth rocks and minerals

Radioactive isotope

4°K

40K

82Se

87Rb

128Te

13°Te

147Sm

138La

138La

176Lu

176Lu

187Re

23SU

238U

238U

232Th

Type of decay

2

Electron capture

p-

2P-

p-

2P-

2P-IX

p-Electron capture

p-,y

Electron capture

p-

lX,y

IX, Y

Spontaneous fission

lX,y

Half-life period, years

3

1.14 xl010

1.40 xl09

1020

4.88 X 1010

1022

3 X 1020

1.06 X1011

2.69 xl011

1.51 Xl011

3.57 xl010

1.15 x1012

4.56x101O

0.7038 xl09

4.468 X 109

7 x101s

14.01 xl09

Radiogenic isotope

4

4°Ar

40Ca

82Kr

87Sr

128Xe

130Xe

143Nd

138Ce

138Ba

176Hf

176Yb

1870S

4He + 207Pb

4He + 206Pb

136- 131Xe 86- 83Kr

4He + 208Pb

Some typical minerals and rocks containing long-lived radioactive isotopes used in isotopic geology

5

Sanidine, anorthoclase, plagioclase, leucite, nepheline, biotite, phlogopitt; muscovite, lepidolite, glauconite, amphibole, pyroxene

Biotite, potassium feldspar, sylvite, langbanite

Selenobismuthite, native selenium, clausthalite, trogtalite

Muscovite, biotite, phlogopite, lepidolite, potassium, feldspar, clay minerals, sylvite, carnallite, granites, syenites, schists

Tellurobismuthite, native tellurium, altaite, tetradymite, calaverite

Monazite, apatite, biotite, amphibole, potassium feldspar, clinopyroxene, basalts, granites, syenites carbonatites

Baddeleyite, eudialyte, zircon, granites, basalts

Molybdenite, jezkazganite, gadolinite, columbite, tantalite, copper sulfides, iron meteorites

Uranium-mica, zircon, monazite, xenotime, samarskite, chlopinite, betafite, sphene, cyrtolite, euxenite, brannerite

Uranium-mica, zircon, monazite, samarskite, chlopinite, baddeleyite, apatite

Thorianite, monazite


Two major methodological difficulties are faced when determining the age of minerals in terms of the radioactive isotopes. The first difficulty is caused by the fact that during crystallization of the minerals, some amount (Do) of the same isotope as the radiogenic isotope (DR) is captured. The amount of the captured isotope Do is represented by comparison with a nonradiogenic isotope of the same element, Co. Hence, the main formula for calculating the isotopic age by any method is:

r DM - Do ] t = ~ln Co Co + 1

A.M M '

Co

where DM = DR + Do is the measured concentration. In the simplest case in assessing the "model age", the isotope ratio Do/Co is

ascribed to a certain value, for instance, the mean ratio in the Earth's crust or in the atmosphere.

In general, however, Do/Co is unknown a priori. Use is made of the date covering a series of minerals regarded as syngenetic and forming a closed isotopic-geochemical systems with equal or similar Do/Co and different M/Co. The above expression is actually a straight-line equation y = a + bx corresponding to the chart (in Fig. 111). In this case a = Do/Co is the primary (initial) isotope ratio in the minerals at the instant of crystallization,

DM ;. 1 tgoc = M = e Mt - 1 and t = A.M In (tgoc + 1).

The following isotopic systems are most widely used:

40Ar 40K 40Ca 40K 87Sr 87Rb 143Nd 147Sm

36Ar - 36Ar' 42Ca - 42Ca' 86Sr - 86Sr' 142Nd - 142Nd'

176Hf 176Lu

174Hf - 174Hf'

Do a=-

Co

1870S 187Re 206Pb 238U

1860S - 1860S' 204Pb - 204Pb'

Fig. 111. Scheme of the age determination accordM ing to series of syngenetic minerals with the un-

"---------- known initial isotopic ratio


The second problem in determining the isotopic age depends on the fact that many geochemical systems are open, i.e., the radioactive and radiogenic istopes migrate in the crystalline structure in the course of geological time.

The process of migration of the isotopes in and out of the crystalline structure is described by formalisms of both classic diffusion and monomolecular first-order chemical reactions, since in real structures only one jump of the migrating atom is sufficient to desl~rt the potential defect pit. The minerals, however, exhibit defects of different types. As a result, the kinetic migration curves, for instance, those of the noble gases, display several maxima. They correspond to various values of the migration activation energy ranging from 10 to 200 kcal/mol. Additional complexities in the interpretation of the migration regularities arise from metamictization of the minerals, i.e., transition from the crystalline to the vitreous state under the action of chemical processes and radioactive irradiation. For instance, up to 1000 cm3 of radiogenic helium are formed in each cm3 of the crystal lattice of uranium minerals. This process of 4He accumulation and liberation of energy during its generation (up to 500 kcal/cm3) changes the state of or completely breaks down the crystal lattice of such minerals.

Loss of the radiogenic isotope by minerals or introduction of the radioactive isotope results in underestimating the apparent age, whereas loss of the radioactive isotope leads to overestimating the age. Radiogenic isotopes whose atoms should appear in defect sites of the crystal lattice due to the radioactive decay energy are supposed to be preferentially removed from the minerals, but laboratory leaching experiments have revealed that this is not commonly true. For example, leaching of feldspars from Erzgebirge granites by water and hydrochloric acid yields Sr with lower 87Sr/86Sr corresponding to preferential release of the initial Sr. Probably neither radiogenic nor nonradiogenic strontium occupy regular lattice sites.

The effect of migration in determining the isotopic age can be quantitatively taken into account by three major methods.

Methods for uranium-lead isotopic geochronology are based on the existence of two independent isotopic systems in any uranium-containing mineral, i.e., 206Pb_238U and 207Pb_235U. Assuming the syngeneticity of the series of minerals, the simultaneous effect of the geologically short "metamorphism" with the migration of the lead and uranium isotopes, and having two independent isotopic systems 235U_207Pb and 238U_206Pb, it is possible to calculate the actual age t of the minerals and the time interval tM ofthe episodic metamorphic effect thereon (Fig. 112).

The Rb/Sr, SmjNd, Re/Os, K/Ca, Lu/Hf methods proceed from the postulate of full isotopic homogenization of the minerals at the instant of metamorphism and preservation of the isotopic-geochemical closeness and heterogeneity of sufficiently large rock blocks. Then the rock formation age is determined by constructing the isochrones as shown in Fig. 112 using bulk samples of the rocks, whereas the age of metamorphism is calculated by referring to the isochrones for mono-mineral fractions.

352

/'

/ /

~'

Chapter 8. Isotopic Mineralogy

t concordia (theoretical curve of changes of the isotopic ratios in time)

L-________________ __ 207Pb

235U

Fig.H2. Determination scheme of the cristallisation age (t) and metamorphism (tM) of series of syngenetic minerals (points)

40Ar 39Ar

Apparent age

Plateau

TOC

Fig. 113. Scheme of the age determination according to "plateau" e9Ar-40Ar and Xes-Xen)

The third mode of dating as applied to open isotopic-geochemical systems is employed in 40 Arp9 Ar and Xes/Xen methods. The mineral to be dated is exposed to a neutron flux in a nuclear reactor for generation of 39 Ar or neutroninduced Xen' They are uniformly distributed over the crystalline structure. During the stepwise annealing of the irradiated specimen at low temperatures, gases are liberated from the disturbed areas of the structure from which part of the 40 Ar and Xe from spontaneous fission (Xes) has already migrated during geological time. Hence, in the low-temperature gas fractions ratios 40 Ar/39 Ar and Xes/Xen are low and accordingly the apparent age is small (Fig. 113).

At higher temperatures, the gas is liberated from the undisturbed areas of the mineral structure, which fully preserved 40 Ar and Xes. Ratios 40 Arp9 Ar and Xes/Xen are high and comply with the true age of the minerals. The true age of the mineral is determined by reference to the "plateau" produced in the age spectra.

8.1.3 Induced Isotopes as Products of Nuclear Reactions in Minerals

8.1.2 Radioactive Isotopes in Miner:als as Intermediate Products in Radioactive Families

353

Radioactive families of 238U, 235U, and 232Th contain the unstable isotopes of various elements ranging from lead to uranium (Table 16).

In the course of secondary changes in the minerals, their precipitation from aqueous solutions, adsorption on the water-suspended particles of the clay minerals during magma formation, or, on the contrary, during crystallization of the lava flows, the intermediate members of the radioactive families can be separated from their parent isotopes. As a result, one of two possible situations may arise.

The first situation provides for separation of some member of the radioactive family from the parent isotope and its subsequent decay in the mineral at a rate determined exclusively by its half-life period.

In the other situation, the radiogenic radioactive isotope is formed in the mineral as the parent isotope decays (following its separation from the earlier accumulated radiogenic isotope) until the state of equilibrium is gained.

For instance, the radioactive isotope 230Th is formed in the oceans due to decay of 234U. It is quickly sorbed from water to the surface of solid particles or forms part of the authigenic minerals (zeolite, phillipsite, barite, etc.). The quantity of 23°Th and accordingly its radioactivity in the minerals and sediments decrease in time in an exponential manner due to the radioactive decay. Using various alternatives of dating involving employment of 230Th together with any other member of the radioactive families e32Th, 238U, 231 Pa) or using other combinations of such isotopes, pelagic sediments can be dated, and the rate of sediment accumulation, the age of carbonate minerals of sea and freshwater origin and corals, the rate of snow and ice accretion, the age of the Pleistocene magmatic rocks over the range of several tens of years to tens of thousands of years can be determined.

8.1.3 Induced Isotopes as Products of Nuclear Reactions in Minerals

Under the influence of the irradiation of various types and various origins, many nuclear reactions take place and many radioactive and stable isotopes are

Table 16. Some intermediate decay products in radioactive families of uranium and thorium

Isotope Type of decay Half-life period, years

23°Th IX, Y 7.52 X 104 234U IX, Y 2.48 X 105

231Pa IX,Y 3.248 x 104

210Pb p-, Y 22.26 226Ra IX, y 1.622 X 103


formed in the minerals of the Earth and extraterrestrial material. These isotopes can be called induced isotopes and are subdivided into several groups.

Cosmogenic Radioactive Isotopes in Minerals and Meteorites

The minerals of meteorites are exposed to the intense effect of the galactic and solar cosmic radiation, mainly protons with an energy of tens to thousands of MeV. This results in deep splitting (spallation) of the nuclei of many elements, i.e., the lighter nuclei are split from these nuclei. They are radioactive, but following several consecutive decays, transform finally into stable isotopes. Two major problems are solved using the cosmogenic isotopes contained in meteorites.

The first problem is the determination of the radiation age of meteorites. This is the time interval between the instant of collision of their parent bodies, accompanied by breaking into small fragments exposable to cosmic radiation, and the instant of the meteorite fall on the Earth, when irradiation discontinued. From that instant, the radioactive cosmogenic isotopes only decayed, but did not accumulate. Use is made of the minerals containing the radioactive isotopes (half-life periods in years in brackets): lOBe (1.5 x 106), 8lKr (2.13 x 105), 26AI (7.16 x 105), 36CI (3.08 x 105) 53Mn (3.7 x 106), as well as some stable isotopes of the noble gases.

The second problem is the determination of the terrestrial age of minerals. In meteorites exhibiting a high radiation age, the decay rate of the radioactive isotopes is constant. After the meteorite has fallen on the Earth and its space irradiation has discontinued, all radioactive isotopes continue to decay. The duration of the stay of the meteorite on the Earth can be calculated in terms of the decay rate measured on the expiry of some time following the meteorite fall, i.e., in terms of the amount of the remaining radioactive isotope.

Cosmogenic Radioactive Isotopes in Terrestrial Minerals

Under the influence of primary and secondary space irradiation, many radioactive isotopes are formed in the Earth's atmosphere (half-life periods in years in brackets): 3H (12.26), 7Be (0.145), lOBe (1.5 x 106), 14C (5730), 26AI (7.16 x 105),

32Si (276), 39 Ar (269), 81 Kr (2.13 x 105), etc. Some of them are gaseous or capable of forming gaseous compounds, and continue to stay in the atmosphere for a long time. Other radioactive isotopes quickly fall out from the atmosphere due to precipitation and enter the minerals of the sea and lake sediments as well as the ice of the Earth's polar regions.

Some cosmogenic radioactive isotopes 3H, lOBe, 26 AI, and 36Cl accumulate in situ in minerals exposed to cosmic radiation on the Earth's surface. The concentration of the cosmogenic isotopes in the atmospheric air amounts to 103

to 105 atoms/cm3, in ice, 102 to 103 atom/g, in erupted rocks, 105 to 108

8.1.3 Induced Isotopes as Products of Nuclear Reactions in Minerals 355

atoms/g, in manganese nodules, up to 108 atom/g, in rain and seawater, up to 108 atoms/g, etc.

Applications of terrestrial cosmogenic radioactive isotopes follows two major directions.

First is the use of these isotopes for measuring the time and rate of geological processes. The isotopes lOBe, 14C, and 2S Al are used for dating sea sediments and manganese nodules, 14C for dating objects containing biogenic carbon, and 32Si for dating biogenic silica. The isotopes lOBe, 26 AI, 36CI, 39 Ar, and 81 Kr are used to determine the age of glacier ice. To date volcanic rocks, use is made of lOBe and 26AI, and the first of these isotopes is applicable simultaneously to study the rate of subduction and the ge:nesis of island arcs.

Another important application of cosmogenic isotopes like 3H, 7Be, 39 Ar, and 85Kr is based on the fact that they represent very informative tracers in studying atmospheric processes, migration, and mixing of oceanic and groundwater.

Nuclear Reactions in Terrestrial Minerals and Radiogenic Isotopes

Under the influence of natural OC-, p-, y-, and neutron irradiation, various nuclear reactions occur in minerals. Some isotopes "burn out", whereas others originate.

The probability of interaction of OC-, p-, and y-radiation with the nuclei is relatively low: for instance, only one among 106 to 107 oc-particles penetrating the crystalline structure of a mineral is capable of entering into a nuclear reaction. Other particles are simply decelerated in the structure. Hence, although all minerals are exposed to continuous nuclear self-irradiation (for instance, each gram offeldspar in granites is exposed to the dose of 101 to 1016 ocparticles during 2 billion years), although various nuclear transformations do occur in nature (oc, n-, oc, p-, y, n-, y, IX-, n, oc-, n, y-, y, p-, and other nuclear reactions), it is possible to detect only those nuclear reactions the products of which are distinguished by extremely low initial concentrations in minerals and rocks.

As the noble gases exhibit the lowest natural abundances, they are the best detectors of nuclear reactions in minerals.

For instance, in uraninite, monazite, zircon, samarskite, betafite, chlopinite, and other uranium-containing minerals excess 21 Ne, 22Ne, and 38 Ar were detected. Isotope ratios 21 Ne/20Ne, 22NeFoNe, and 38 Arp6 Ar exceed the usual ratios by thousands percent.

In addition to OC-, p-, and y-radiation produced by natural radioactive isotopes, the minerals are exposed to thc~ natural flux of neutrons. The average flux of neutrons in the Earth's crust is 106 neutrons/s/cm3. Its sources are the cosmic radiation, fission of the heavy nudei, and OC-, n-reactions in the minerals. The contribution of the last source is the highest.

Mainly n, IX- and n, y-reactions occur under the influence of neutrons. They result in noticeable isotopic shifts in all noble gases in minerals due to the

356

following processes: p-

6Li(n, oc) 3H; 3H ----+ 3He

p-35CI(n, y) 36CI; 36CI ----+ 36 Ar

235U(n,f) 136- 129Xe; 86- 83Kr.


The effective flux of neutrons in the minerals can be estimated in terms of the products of the above reactions. One of the most interesting applications of these processes is the method for determining uraninite age in terms of the isotope 129Xe formed from 129J, the half-life period of which is 16 Ma. By neutron-induced fission of 235U, 129J is formed. In the course of time, the ratio 129Xe/132Xe changes from zero to the equilibrium value, meeting the time when the rate of 129Xe generation in the mineral becomes equal to the rate of 129J decay. Accordingly, the age that can be determined using the above method must be equal to or smaller than 150 Ma. The time required to form large-sized pitchblende ores can be estimated in terms of the ratio of neutron-induced Xe and Xe produced as a result of spontaneous 238U fission.

As a rule, the contribution of the products of neutron reactions to the trace element chemistry of minerals is small. The unique exception is pitchblende in the Oklo deposit in the Republic of Gabon.

Two billion years ago, the chain uranium fission reaction similar to the fission of nuclei in technical nuclear reactors occurred during a time interval of 0.5 Ma, the neutron flux amounting to 108 neutrous/cm2/s. As a result, more than half of the initial quantity of isotopes ofa series of elements (U, Nd, Sm, Gd, Er) "burnt out".

On the other hand, tremendous excesses of some isotopes, which amount to several ten thousand permil originated in the pitchblende. This is due to the additives of fission products and their neutron-induced component (Nd, Sm, Eu, Cd, Dy, Xe, Kr, Rn, Pd, and Zr).

The important parameters of the chain nuclear fission reaction in the Oklo deposit were calculated in terms of the isotope excesses and the deficit in uranium-235 (chain reaction origination time and duration, the mineral temperature at which the reaction had occurred, energy release, etc.).

References

Bowen R (1988) Isotopes in the Earth Sciences. Elsevier, London De Paolo DJ (1988) Neodymium isotope geochemistry. An introduction. Springer, Berlin

Heidelberg New York, 187 pp Faure G (1986) Principles of isotope geology 2nd edn New York, 589 pp

8.2.1 The K-Ar Isotope System in Geochronology 357

Florkowsky T, Morawska L, Rozanski K (1988) Natural production of radionuclides in geological formations. New York, Oxford, J Rad Appl Instrumen Nuclear Geophysics, vol 2, pp 1-14

Jager E, Hunziker JC (eds) (1979) Lectures in isotope geology Springer, Berlin Heidelberg New York, 332 pp

Kerridge JF, Mattews MS (eds) (1988) Meteoriltes and the early solar system. Univ of Arizona Press, Tucson 1269 pp

Natural fission reactors (1978). International Atomic Energy Agency, Vienna Ozima M, Podosek F (1983) Noble gas geochemistry. Cambridge Univ Press, Cambridge,

367 pp Shukolyukov YuA, Levski LK (1972) Geochemistry and cosmochemistry of noble gas

isotopes. Atomizdat, Moscow, 335 pp (in Russian)

8.2 Isotopic Systems in Geochronology

8.2.1 The K-Ar Isotope System in Geochronology

YU. A. SHUKOLYUKOV and H.J. LIpPOLT

40K (0.01167 weight %) is the least abundant natural K isotope e9K 93.258%, 41 K 6.730%) and it is dually radioactive, transformed into 40 Ar due to electron capture and into 40Ca due to beta decay. The full constant of the 4°K decay rate equals A = AeI•eap. + A(J- = 5.543 x 10- 10 a-I, corresponding to a half-life period of 1.25 x t09 a (A.l.eap. = 0.581 x 10- 1 ° a - 1 and A(J - = 4.962 x to- 10a- 1).

In order to make use of such radioactive decay systems, the existence of suitable ("Ar-retentive") well preserved minerals as watch cases during the geological history of the rocks is imperative. Age interpretations of rocks and minerals are based on their present concentrations of 4°K and radiogenic 40 Ar (40 Ar*). The range of application of the 4°K -40 Ar chronometer reaches from the beginning of the planetary system (about 4.5 Ga; a = year) to the latest Quaternary (to about 103 a). The isotopic analysis of argon extracted from the specimen in general besides radiogenic 40 Ar yields also an nonradiogenic (common) component, mostly similar in composition to atmospheric argon. Atmospheric Ar consists of 4°Ar (99.60%) (predominantly produced from 40K during the Earth's history), 38 Ar (0.0635%), 36 Ar (0.3378%).

The Basis

Potassium concentration (exactly 40K) and quantity of radiogenic 40 Ar supposed to have been accumulated during time t yield a term with the physical dimension TIME in general called K-Ar date (or apparent K-Ar age). K-Ar ages


are derived from K-Ar dates by interpretation, based on the model of radioactive decay as well as on those of geological, mineralogical, and geochemical history of the specimen.

1 [40Ar* A. ] t = -Xln 40K -A.-- + 1

eLeap.

t is the age. In the formula above, normally 4°K is replaced by (c.K), factor c being the

relative abundance of 40K in K, based on the realistic assumption that the isotopic composition of K is homogeneous all over the Earth, Moon, and stone meteorites. The constants used for K-Ar calculations have been selected by an international convention (Steiger and Jager 1977). K-Ar dates are called conventional, when K and Ar are measured separately on aliquots of the specimen.

Potassium is a widely spread element, hence isotopic dating using the Kj Ar method is used for estimation of the ages of the most different minerals (Table 17) and thereby geological processes. Micas, several feldspars and feldspathoids, amphiboles are the most important. Minerals with only trace amounts of K are not likely to yield reliable results.

Table 17. Frequently occurring rock-forming minerals suitable for K-Ar dating

Type of mineral Mineral Type of rock

Volcanic 2 3

Framework silicates Sanidine XX -Feldspars Anorthoclase XX

Plagioclase XX -Feldspathoids Leucite X

Nepheline X

Sheet silicates Biotite XX -Mica group Phlogopite XX

Muscovite Lepidolite Glauconite

-Clay minerals Illite Smectite

Salt minerals Sylvine Langbeinite

Chain silicates Hornblende XX Pyroxene X

Rock sample Grain powderb XX Groundmass X

XX = frequently used: X = used in some instances. • Including diagenetic and epigenetic formations.

Plutonic 4

X

XX

XX X

XX X

X

Metamorphic 5

XX

XX

XX XX

XX X

X

b Often frationated, dust removed, geochemically not representative.

Sedimentary· 6

X X X

X X


The amounts needed depend on the analytical techniques which are applied and, as far as Ar is concerned, on the ages of the samples (normally grams or gram fractions). When preparing mineral separates from rocks, great care has to be taken not to contaminate the samples or to end up with mineral mixtures which because of high Ar or K contents are dominated by a minor admixture (e.g., old K-feldspar in young sanidine, biotite in hornblende, mica in plagioclase).

In many cases, the K-Ar dating method yields very valuable information on the geological history of a rock, especially on its younger sections, in particular when using hornblende, sanidine, muscovite, lepidolite, and phlogopite. Rocks, normally due to their genesis and history, have various types of ages: formation ages (intrusion, extrusion, sedimentation) as well as ages of transformation (alteration, metamorphism - sometimes repeated or even in series - diagenesis). K-Ar dates are likely to assess more or less reliably the intrusion ages of volcanites without later geological impact and the times of declining temperatures after phases of plutonic intrusions, of slow uplift from deep crustal sites and of metamorphism.

Misinterpretations of K-Ar dates as geologically meaningful ages are envisaged if sample formation or transformation is long-lasting (related to the time span since its beginning), if the initial Ar composition (inherited from the former state of the sample or absorbed/occluded during crystal lattice formation) cannot be reliably determined (apostrophed as excess argon problem) and, if 40 Ar* was lost during the geological history of the specimen. As a consequence, it is imperative in general to investigate various samples of the same type and of differing specifications in order to reveal and to overcome such shortages of the chronological information. Important means are existence or non-existence of isochrons for syngenetic samples (straight lines in 4°Arr6Ar vs. 4°K/36Ar plots or in others).

Age discordances being undesirable for chronometric purposes (in the sense of stratigraphy), however, yield additional information on the history (thermal, tectonic) of the minerals. An important further means to evaluate the meaning of discordant K-Ar dates is to apply the tomographic quality of the 4°Ar/39Ar spectrum or "plateau" approach which yields insight into the intercrystalline positions of the 40 Ar and 4°K isotopes.

Apparent Age Deficiencies

Geological processes which are likely to promote Ar losses are those which induce changes of type and state of minerals and which take place at geologically rather high temperatures. Among them are stress effects and mechanical breakdown of the minerals, chemical erosion, and mineral changes when being exposed either to hydrothermal solutions (alteration) or, in the case of brinesoluble minerals, such as sylvine, to diissolution and reprecipitation, or to metamorphism and recrystallization due to higher temperatures.


Argon is not a regular lattice constituent of the minerals and it is very migratory as an element in the Earth's crust as well as an atom in the minerals. It is capable of performing various diffusion processes above certain mineraldepending critical temperatures (e.g., about 500 °C for amphibole, 350°C for muscovite, 300 °C for biotite), which is of concern when rocks cool down from temperatures of metamorphism (anatexis, amphibolite facies, etc.) and when they are heated by local or regional temperature increases (intrusions, deep burial).

The volcanic rocks solidify and cool down fast. Normally they are finegrained. When containing mainly plagioclase and pyroxene, which preserve argon satisfactorily, and no other K minerals which do not, they can be dated by using bulk samples. However, alterated rocks or those containing devitrified glass or secondary minerals (zeolites, calcites, clay minerals) are less suitable. Contrary to quick-cooling rocks (vocanics and shallow-bedded intrusions), the argon diffusion, accompanied by Ar losses, in the deep plutonic and metamorphic rocks continues and 40 Ar accumulation is hindered as long as the characteristic closing (or blocking) temperatures of the minerals are not reached. The latter are grain-size related and depend on the cooling rates. Differences in the apparent K-Ar ages (dates) of the cogenetic minerals are used to estimate cooling-down rates of the rocks provided that the appropriate closing temperatures are sufficiently known. K-Ar dating of granite gneisses and crystalline shales of the Precambrian shields and of orogenic belts indicates that large areas of the continental crust after periods of metamorphism have cooled down, thus forming large, K-Ar isotopically homogeneous blocks. Such synchronous cooling-down of large fragments of the continental crust depends on the rise of the craton blocks limited by deep-seated faults. In other words, the K-Ar age of an individual mineral from the craton block is the time when the formerly deepsited rock, now at the surface, intersected the closing temperature isotherm of the mineral. K-Ar cooling ages are of great interest in tectonic studies.

Sedimentary rocks may be composed of detrital, authigenic, and/or diagenetic minerals (sandstones, carbonate rocks, evaporites, iron quartzites, and phosphorites, etc.) .. K-Ar dates of detrital components may be used for provenance studies, never for determining ages of sedimentation. K-Ar minerals which are newly formed in the course of sedimentation (including diagenesis) are glauconites, smectites, occasionally also poly halite and langbeinite. Due to longlasting diagenetic processes within clayey schists, radiogenic 40 Ar is lost from bulk samples, which therefore are not suitable for reliable dating of sediments. However, in cases where layers of sedimentary rocks are interstratified with former volcanic rocks (tuffs, tuffites), which due to chemical changes in the volcanic ashes now exist as bentonites and often still contain well-preserved volcanic K-Ar minerals (sanidine, plagioclase, anortholase, pyroxene, amphibole, biotite, etc.), it is possible to determine the ages of sedimentation. Unfortunately by various types of alterations caused by post-sedimentary geological processes, the volcanic minerals may be completely destroyed or transformed. Then at best these later processes may be dated.


Apparent Age Surplus

The second main problem in K-Ar geochronology arises from the common argon trapped within the crystalline structure of the minerals. Such effects are of more concern the younger the ages and the smaller the pertinent K concentrations (and therefore 40 Ar*) of the specimens. Although normally being of atmospheric Ar composition, occasionally very differing isotopic compositions are shown which cause inaccurate calculations of the radiogenic component. Predominantly K-Ar dates in excess of the realistic K-Ar age are reported (Excess Ar dates), but also apparent Ar deficits may occur. These Ar components may be occluded during formation of the specimen (e.g., pyroxene in Pleistocene volcanic rock) or later in the course of secondary geological processes (e.g., biotite in Alpine metamorphic rocks). Very often, inherited 40 Ar from the former state of the rock is the cause of 40 Ar excess (e.g., Ar in amphibol inherited from former pyroxene). As far as volcanic rocks are concerned, the Ar in the source regions within lower crust/upper mantle shows a distinct radiogenic component and volcanic rocks are only K-Ar datable because the crystalline structure in general discriminates argon. Basalt effused to the oceanic bottom at a high hydrostatic pressure contain excess 40 Ar in its vitreous crust, as there was no way and time to expel the dissolved argon. For the same reason, phenocrysts in volcanic rocks may contain excess argon which confines K-Ar dating of very young volcanites to groundmass fractions. There are some crystalline structures which are primordially very favorable for taking up 40 Ar atoms from the mineral-forming environment (beryl, cordierite, sodalite minerals, cancrinite, tourmaline, diamond). The 40 Ar excess concentrations range from 10 - 4 to 10- 2 cm3 STP/g. Other sources of excess. argon in some minerals are the gas liquid inclusions, more ancient xenocrystals.

K-Ar dating has great merits, for example: K-Ar dating assisted in creating the Geochronological Scale from the Precambrian Period to the Youngest Quaternary and the scale of inversions of the Earth's magnetic field. It provided direct proofs of ocean floor spreading, leading to the evolution of the theory of plate tectonics. It made it possible to study the thermal and tectonic histories of the Earth's cratonic blocks and of its orogenies. It unraveled the sequence ofthe early appearances of Ancient Man.

References

Borsuk AM (ed) (1979) Criterions of the reliability of the radiological dating methods. Nauka, Moscow, 208 pp (in Russian)

Dalrymple GB, Lanphere MA (1969) Potassium argon dating, Principles, techniques and applications to geochronology. Freeman, San Francisco, 258 pp

Damon PE (1970) A theory of 'real' K-Ar clocks. Eclogae Geol Helv 63: 69-76 Faure G (1986) The K-Ar method of dating. In: Faure G (ed) Principles of isotope geology 2nd

edn. Wiley, New York, chap 6, pp 66-93


Hunziker JC (1979) Potassium argon dating. In: Jaeger E, Hunziker J (eds) Lectures in isotope geology. Springer, Berlin Heidelberg New York, pp 52-77

Morozova 1M, Ashkinadze GSh (1971) The migration of atoms of raae gases in minerals. Nauka, Leningrad, 115 pp (in Russian)

SChaeffer OA, Zahringer J (eds) (1966) Potassium argon dating. Springer, Berlin Heidelberg New York, 234 pp

Steiger RH, Jager EJ (1977) Subcommission on Geochronology. Convention on the use of decay constants in geo- and cosmochronology. Earth Planet Sci Lett 36: 359-62

8.2.2 40 Ar f39 Ar and its Laser Variant

M.I. KARPENKO and J.F. SUTTER

In the 4°Ar/39Ar method of K-Ar dating, the sample is irradiated with fast neutrons to induce the reaction 39K(n,p)39 Ar. The age(t) of the sample is then calculated from the 40 Ar /39 Ar ratio after appropriate corrections for interfering Ar isotopes from the atmosphere and from undesirable neutron reactions with Ca, K and CL using the equation

t = 1/A.ln[1 + J(4°Ar/39Ar)],

where A. is the decay constant of 4°K( = 5.543 x 10- 10 yr- 1). The conversion factor, J, for the reaction 39K(n,p)39 Ar is determined by irradiating a monitor mineral of known age along with the samples whose age is to be determined (Merrihue and Turner 1966; Dalrymple et al. 1981).

The method can be used in two different ways. If all of the argon is released by fusing the mineral in a single heating, the result is a total fusion age, roughly analagous to a conventional K-Ar age. If the argon is released from the sample in steps by incrementally heating the sample to progressively higher temperatures, the result is a series of ages known as an age spectrum.

The 40 Ar/39 Ar has several advantages over the conventional K-Ar method: (1) higher precision, (2) smaller sample size, (3) elimination of problems caused by sample inhomogeneity, and (4) the ability to recover age and thermal history information from age spectra (McDougall and Harrison 1988; Lovera et al. 1989). The principal disadvantage of the 40 Ar/39 Ar method is difficulty in the dating of samples that experience 39 Ar recoil loss during neutron irradiation (Villa et al. 1983).

Most commonly, the argon isotopic analysis for 40 Ar /39 Ar dating is done with the same or similar equipment and techniques used for the conventional K-Ar method.

A laser was first employed in the 40 Ar /39 Ar dating method by Megrue (1973). The most important difference between this method and the "conventional" 40 Ar j39 Ar one is the use oflaser radiation energy to extract gases from a sample. The pulsed laser probe combined with a petrographic microscope can produce a 40 Ar/39 Ar total fusion age from a small area (typically 20 to 200 J.lm

8.2.2 40 Ar/39 Ar and its Laser Variant 363

diameter) within single crystals, separate minerals, growth zones in minerals, and inclusions of different kinds as well as for studying diffusion processes on a microscale and all of this done in situ (i.e., Schaeffer 1982; Sutter and Hartung 1984; Maluski and Moni 1988; Burgess et al. 1989).

Information on the disturbance of the K-Ar system of a sample can be determined in situ using pulsed laser methods described by Muller et al. (1977), and by Ivanenko and Karpenko (1987). However, in cases where it is possible to separate individual mineral grains from the sample, incremental heating with a. continuous laser is a better way of identifying open system behavior.

York et al. (1981) were the first to demonstrate that a continuous laser could be used not only to produce 40 Ar /39 Ar ages from single crystals but also to determine age spectra. Since then, continuous laser probes have been used to date tektites and illites (Glass et al. 1986; Bray et al. 1987), to determine the distribution of apparent ages within singl1e crystals (York and Hall 1986; Phillips and Onstott 1988), to produce age spectra from single grains of biotite and hornblende (Layer et al. 1987), to circumvent xenocrystic contamination in silicic volcanic rocks (LoBello et al. 1987), and to determine highly precise ages by the use of sanidine crystals from Oligocene rhyolite flows (Dalrymple and Duffield 1988).

These laser probes consist of a continuous laser for sample heating and fusion, a small-volume extraction-cleanup system, and an ultrasensitive, ultralow-background rare-gas mass spectrometer. Not only can the instruments produce both total fusion and incremental heating data, they yield 40 Ar/39 Ar ages with precision as good or better than other variants of 40 Ar/39 Ar dating from samples as small as a few micrograms.

References

Bray CJ, Spooner ETC, Hall CM, York D, Hills TM, Krueger HW· (1987) Laser probe 40 Ar/39 Ar and conventional K/ Ar dating of illites associated with the McClean unconformity-related uranium deposits, North Saskatchewan, Canada. Can J Earth Sci 24(1): 10-23

Burgess R, Turner G, Laurenzi M, Harris JW (1989) 40 Ar/39 Ar Laser probe dating of individual clinopyroxene inclusions in Premier eclogitic diamonds. Earth Planet Sci Lett 94 (1-2): 22-28

Dalrymple GB, Duffield W A (1988) High precision 40 Ar /39 Ar dating of Oligocene rhyolites from the Mogollon-Datil volcanic field using a continuous laser system. Geophys Res Lett 15: 463-466

Dalrymple GB, Alexander EC Jr, Lanphere MA, Kraker GP (1981) Irradiation of samples for 4°Arj39Ar dating using the Geological Survey TRIGA reactor. US Geol Surv Prof Pap 1176: 55

Glass BP, Hall CM, York D (1986) 4°Ar/39Ar laser probe dating of North American tektite fragments from Barbados and the age of the Eocene-Oligocene boundary. Chem Geol Isotope Geosci Sec 59: 181-186

Ivanenko VV, Karpenko MI (1987) New possibillities for determining age spectra of 4°Ar/39 Ar by means oflasers. Dokl Akad Nauk USSR 296(3): 710-714

Layer PW, Hall CM, York D (1987) The derivation of 4°Ar/39 Ar age spectra of single grains of hornblende and biotite by laser step-heating. Geophys Res Lett 14: 757-760


LoBello Ph, Feraud G, Hall CM, York D, Lavina P, Bernat M (1987) 4°Ar/39Ar step-heating and laser fusion dating of a Quaternary pumice from Neschers, Massif Central, France: the defeat of xenocrystic contamination. Chemical Geology Isotope Geosci Sec 66: 61-71

Lovera OM, Richter FM, Harrison TM (1989) The 4°Ar/39Ar thermochronometry for slowly cooled samples having a distribution of diffusion domain sizes. J Geophys Res 94(BI2): 17917-17935

Maluski H, Monie P (1988) 4°Ar/39Ar Laser probe multi-dating inside single biotites of a Variscan orthogneiss (Pinet, Massif Central, France). Chern Geol Isotope Geosci Sec 73(3), 9(3): 245-265

McDougall I, Harrison TM (1988) Geochronology and thermochronology by the 40 Ar /39 Ar Method. Oxford Univ Press, New York, 212 pp

Megrue GH (1973) Spatial distribution of 40 Ar/39 Ar ages in lunar breccia 14301. J Geophys Res 78: 3216-3221

Merrihue C, Turner G (1966) Potassium-argon dating by activation with fast neutrons. J Geophys Res 71: 3852-2857

Muller HW, Plieninger T, James OB, Schaeffer OA (1977) Laser probe 4°Ar/39Ar dating of material from consortium breccia 73215. Thermill RB (ed) Houston, Texas. Geochim Cosmochim Acta, Suppl 8. New York, Pergamon. Proc 8th Lunar Science Conf, pp 1489-1499

Phillips D, Onstott TC (1988) Argon isotopic zoning in mantle phlogopite, Geology 16: 542-546

Schaeffer OA (1982) Laser microprobe 40 Arj39 Ar dating of individual grains. In: Currie (ed) Nuclear and chemical dating techniques. Am Chern Soc Symp 176: 516

Sutter JF, Hartung JB (1984) Laser microprobe 40 Ar/39 Ar dating of mineral grains in situ. Scanning electron microsc 4: 1525-1529

Villa 1M, Huneke JC, Wasserburg GJ (1983) 39Ar recoil losses and presolar ages in Allende inclusions. Earth Planet Sci Lett 63: 1-12

York D, Hall CM (1986) Continuous-laser probe thin section chrontouring of sediments (Abstr). Terra Cognita 6: 117

York D, Hall CM, Yanase Y, Hanes JA, Kenyon WJ (1981) 4°Ar/39Ar dating of terrestrial minerals with a continuous laser. Geophys Res Lett 8: 1136-1138

8.2.3 The Rb-Sr Method of Isotopic Dating

D.l. DEPAOLO, T.F. ANDERSON, and V.I VINOGRADOV

Rb is composed of two isotopes: 85Rb (72.16%) and 87Rb (27.84%). The atomic ratio 85Rbj87Rb that is conventionally accepted is 2.59265. Natural variations of this ratio are likely to occur at the level of 0.1 %, but measurement of such variations is beyond the sensitivity of modern methods. 87Rb is radioactive and decays (fJ-) to one isotope of strontium-87Sr. The decay constant is A = 1.42 x 10 - 11 a - 1. Over the 4.5 b.y. of Earth's existence only 6.5 % of the initial

quantity of 87Rb has decayed. Sr is composed of four isotopes: 84Sr (0.56%), 86Sr (9.86%), 87Sr (7.02%),

and 88Sr (82.56%). The conventionally accepted ratios of the nonradiogenic isotopes are 84Sr;S8Sr = 0.006745 and 86Srj88Sr = 0.1194, which are considered to be invariant in nature. The 87Sr in the Earth consist of two parts, that inherited from the time of Earth formation, and that generated by decay of 87Rb since the time of formation of the Earth. Because the amount of Rb in the Earth

8.2.3 The Rb-Sr Method of Isotopic Dating 365

is small in comparison to the amount of Sr (total Rb/Sr ::::; 0.029), only about 0.8% of the 87Sr in the earth comes from radioactive decay of 87Rb.

The quantitative evaluation of the radiogenic addition to 87Sr owing to rubidium decay is the basis of age determinations using the Rb-Sr method:

(87Srh = (87Sr)o + 87RbT(eAT - 1), (1)

where (87Sr)o is the amount of 87Sr T yl~ars ago, (87Srh is the amount of 87Sr after T years (today), 87RbT is the amount of 87Rb after T years (today), and, T is the age of the system being considered. Normally, the isotopic ratios of Sr are measured, because this measurement can be made with enhanced precision. It is therefore generally agreed to divide every term of Eq. (1) by 86Sr:

(87Sr/86Srh = (87Sr/86Sr)o + (87Rb/86Srh(eAT - 1) (2)

For timescales where AT ~ 1, the (eAT - 1) multiplier can be replaced by AT with little loss of accuracy. This applies for T < 109 years.

In order to determine age from the measurement of (87Sr/86Srh and (87Rb/86Srh it is necessary to know the initial ratio (87Sr/86Sr)o' In some cases this can be estimated to sufficient accuracy. Otherwise, it is necessary to find geological situations where several separate domains (e.g., different minerals in the same rock) can be assumed to have the same value of(87Sr/86Sr)o, but each has a different value of (87Rb;S6Srh. In this case, the (87Sr/86Srh and (87Rb/86Srh values of the different domains will define a line y = ax + b, where y = (87Sr/86Srh, x = (87Rb/86Srh, b = i[87Sr/86Srh The slope of this line is a = (eAT - 1). In nature it is commonly the case that measurements result in

points that approximate a line but do not fit exactly. The best-fit line is called an isochron if the scatter of points is within I~xperimental uncertainty. If the scatter is larger than experimental uncertainties" it means that either the initial condition was not met (uniform 87Sr/86Sro), or that the domains have gained or lost Rb or 87Sr during their lifetime. Where the scatter of the data is larger than the experimental uncertainties, the determined age may be incorrect. The slope and age uncertainty are determined from the data using least-squares methods.

The Rb-Sr "clock" is set by homogenization of the 87Sr/86Sr. The mechanisms of homogenization in geological processes are not well understood, consequently it is not always well known what geological event corresponds to the determined isotopic age. In the case of magmatic processes, it is normally considered to be the time of crystallization of melt. In the case of metamorphism or metasomatism, the homogenization on length scales greater than a few centimeters is probably facilitated by intergranular fluids. In addition to the age information, the initial strontium isotopf: ratio gives information on the origin of the rock material.

Rb and Sr belong to different chemical groups and therefore are quite different in their chemical and geochemical properties. This difference enhances the potential for disturbance of the initial 87Sr/86Sr and 87Rb/86Sr ratios, and often makes it impossible to determine thl~ age of rocks. Some minerals are more prone to disturbance than others. Biotite would be highly suitable for dating


because of its high Rb/Sr ratio. However, biotite has been found to be one ofthe least reliable minerals for dating. To compensate this defect of most mineral systems, the "whole rock" method has been proposed. It is argued that even if the Rb/Sr ratios are disturbed on a mineral scale (millimeters or centimeters), it is possible to take some volume of the rock (decimeter to meter) in which gain and loss of Rb and Sr will be mutually compensated; so this volume of rock will be a conservative system with respect to Rb and Sr. It is found that this simple assumption is often right. However, at this time no general theory exists for this method. One cannot know before what weight of rock is sufficient. Usually samples from 0.5 to 50 kg mass are used. The bigger the sample the more probable that it will be a closed system. However, the bigger the sample the more difficult it is to obtain differences in Rb/Sr ratios between the samples. RbSr dating (as well as other methods) demands a high degree of professionalism from scientists as well as a little luck.

With modern experimental techniques and measuring instruments, the Rb-Sr method can be used for dating time scales from 10()()() to billions of years. The method has been most successfully applied to potassium-rich igneous and metamorphic rocks. There are many difficulties and uncertainties in sedimentary rock dating. Basic and ultra basic rocks and their mineral fractions are difficult to date because they have low Rb/Sr ratios, low contents of Rb and Sr, and are susceptible to secondary alteration.

References

Faure G (1986) Principles of isotope geology, 2nd edn Wiley, 589 pp Faure G, Powell JL (1972) Strontium isotope geology. Springer, Berlin Heidelberg New York,

188 pp Gorochov 1M (1985) Rb-Sr method of isotope geochronology. Energoatomisdat, Moscow,

153 pp (in Russian) Jager E, Hunziker JC (eds) (1979) Lectures in isotope geology. Springer, Berlin Heidelberg

New York Kuptzov VM (1986) The absolute geochronology of bottom sediments of oceans and seas.

Nauka, Moscow, 271 pp (in Russian) Moorbath S, O'Nions RK, Pankhurst RJ (1975) The evolution of early Precambrian crustal

rocks at !sua, West Greenland - geochemical and isotopic evidence. Earth Planet Sci Lett 27: 229-239

Morton JP (1985) Rb-Sr dating of diagenesis and source age of clays in Upper Devonian black shale of Texas. Geol Soc Am Bu1196: 1043-1049

8.2.4 The Sm-Nd Method of Isotope Dating

V.1. VINOGRADOV, OJ. DEPAOLO, and T.F. ANDERSON

Samarium has seven isotopes, with masses 144, 147, 148, 149, 150, 152, and 154. One of the samarium isotopes, 147Sm, is radioactive and decays by the IX process to the isotope 143Nd. The use of this pair of elements for geochronology is

8.2.4 The Sm-Nd Method of Isotope Dating 367

similar to that of the 87Rb_87Sr pair. The equation for the calculation of Sm-Nd age and 143Nd;t44Nd initial ratio is:

e 43Nd;t44Ndh = (143Nd;t44Nd)o + e47Sm/144Ndh (eAT - 1) (1)

The decay constant of 147Sm is half that of 87Rb; Ae47Sm) = 6.54 x 10- 12 a- 1. Both Sm and Nd belong to the rare earth element (REE) group chemically, and therefore are more similar in their geochemical properties than are Rb and Sr. The differences in Sm/Nd ratio are small in natural material (about a factor of 4), while the differences in Rb/Sr ratios are a factor of 104. Therefore the variation of 143Nd/144Nd ratios in nature is very small, and very high experimental mastery and good equipment for precise measurements is demanded.

The similarity of the geochemical properties of Sm and Nd has some advantages over the Rb-Sr pair. Sm-Nd system is much more stable against secondary disturbances. This is one of the reasons why the mineral separates are usually used for Sm-Nd dating. Common minerals arranged in order of decreasing Sm/Nd ratio are: garnet, pyroxene, amphibole, biotite, and feldspar.

Modern experimental techniques allow the 143Nd/144Nd to be measured very precisely, but because of the limited variability of the Sm/Nd ratio, it is rarely possible to obtain an age resolution better than ± 20 Ma. Because of this, the Sm-Nd method is most useful for old rocks. The Sm-Nd method can be used to date all types of rocks, and is especially well suited to basic and ultrabasic rocks, which are difficult to date with the Rb-Sr method.

Basic and ultrabasic rocks have the highest Sm/Nd ratios ( > 0.3) and therefore the most rapid growth of 143Nd;t44Nd as time progresses. The more differentiated rocks have lower Sm/Nd ratios; alkaline rocks in particular have the lowest Sm/Nd ratio '" 0.1. Granites and acid volcanic rocks have intermediate values.

Because of the small variation in Sm/Nd ratios it is sometimes necessary to use two-point isochrons; for instance, mineral and total rock samples. In some cases, particularly for determining the approximate age of crustal rocks, it is useful to calculate a "model age," which requires only one measurement. This calculation is one in which the "initial" 143Nd;t44Nd value is assumed to have a certain time dependence, either that of average chondritic meteorites or that of an empirically determined upper mantle reservoir. The model age refers to the time of the differentiation of the crustal rocks from the mantle of the Earth.

References

DePaolo DJ (1988) Neodymium isotope geochemistry: an introduction. Springer, Berlin Heidelberg New York, 187 pp

DePaolo DJ, Wasserburg GJ (1976) Nd isotopic 1{ariations and petrogenetic models. Geophys Res Lett 3: 249-252

Faure G (1986) Principles of isotope geology, 2nd edn. Wiley, 589 pp Hamilton PJ, Evenson NM, O'Nions RK, Smith HS, Erlank AJ (1979) Sm-Nd dating of

Onvervacht Group volcanics, Southern Africa. Nature 279: 298-300


Lugmair GW, Scheinin, Marti K (1975) Sm-Nd age and history of Apollo 17 basalt 75075: evidence for early differentiation of the lunar exterior. Proc 6th Lunar Sci Conf, pp 1419-1429

Patchett PI (1989) Radiogenic isotope geochemistry of rare-earth elements. Geochemistry and mineralogy of rare-earth elements. Rev Mineral 21: 25-44

8.2.5 The U-Pb System and Zircon as Mineral Geochronometer

E.V. BIBIKOVA and J.N. ALEINIKOFF

Determination of the ages of rocks and minerals began in the early 20th century soon after the discovery of radioactivity. The first dating (or geochronological) method (Boltwood 1907) utilized chemical analyses of uranium and lead. Boltwood surmised that Pb is the stable end-member of the V decay chain and calculated the ages of three uraninite samples, without knowledge of the decay rate (half-life) of uranium or the existence of Pb and V isotopes. Nier (1939) measured Pb isotope ratios utilizing a new mass spectrometer design (Nier 1940) and developed new methods of geochronology. Since the early 1950s, improved analytical techniques (including both chemical methods of elemental extraction and advances in design of modern mass spectrometers) have permitted dating studies of accessory minerals which contain trace concentrations of V and Pb.

Two uranium isotopes e38 V and 235V) and one thorium isotope e32Th) decay, through several intermediate daughter products, to 206Pb, 207Pb, and 2osPb, respectively. The half-lives of the radioactive isotopes are quite long (several hundred million to several billion years), so that they can be used for dating events throughout the span of Earth history. The parent and daughter isotopes are related using the equation:

D = P(eA1 - 1), (1)

where D is the daughter isotope, P is the parent isotope, A is the decay constant for that particular radioactive isotope, and t is the amount of time (in years) since the radioactive system began accumulating daughter product. The V-Pb system is unique among geochronometers because of the paired uranium isotopes that have different decay rates. Thus, an age can also be calculated from the 207Pb/206Pb ratio, using the equation:

207Pb 1(eA,1 - 1)

206Pb 137.88(eA8t - 1) , (2)

where A5 and As are the decay constants for 235V and 23SV, respectively, and 1/137.88 is the present-day ratio of 235Vj238V (Steiger and Jager 1977).

Any mineral or rock containing V and Th can be used as a geochronometer. However, a correction for common Pb (that is, Pb incorporated during crystallization before production of radiogenic Pb by V-decay) is necessary. The

8.2.5 The U-Ph System and Zircon as Mineral Geochronometer 369

nonradiogenic isotope 204Pb is used as a tracer for the common lead correction. The isotopic composition of common Pb can be determined by analysis of cogenetic, lead-bearing (and V-free) mim:rals, such as K-feldspar, or estimated from lead-evolution models (cf. Stacey and Kramers 1975).

Because of the different decay chains of the two uranium isotopes, measurements of V and Pb isotopes result in calculation of three ages, 206Pb/238V, 207Pbj2 3SV, and 207Pb/206Pb. If the three ages are identical (within the calculated uncertainties), the data are called concordant. However, in most geochronological studies 206Pbj238V age < 207Pbj235V age < 207Pbj206Pb age. These data are called discordant. Discordancy indicates a disturbance to the V-Pb isotopic system (usually interpreted as partial loss of radiogenic lead) at some time since crystallization of the mineral.

V-Pb isotopic data are generally plotted on a concordia diagram e06Pb/238V vs. 207Pbj235V}, introduced by Wetherill (1956). The concordia curve is the locus of points with equal 206Pb/238V and 207Pbj235V ages. Isotopic data will plot on the curve if the V-Pb system has remained closed. Most data points lie on or below the curve. If the V-Pb system has been disturbed once, the data will form a linear array, called a discordia line. In the most simple cases, the intersections (or intercepts) of a discordia with concordia will indicate the time at which the mineral began to accumulate radiogenic daughter products and the time at which the isotopic system was disturbed.

If two events have disturbed the isotopic systematics, the data points will generally form a non-linear scatter pattern on a concordia diagram. The data will plot within a triangle whose vertices are the time of beginning of radiogenic accumulation and the times of disturbance. In this case, a fortuitous quasi-linear array has spurious (or geologically meaningless) intercept ages, although frequently, if the data are not very discordant, the upper intercept will approach the formation age. Many models have be:en proposed for the interpretation of discordant V-Pb data (cf. Allegre et al. 1974; Shukolyukov et al. 1974).

Some minerals have very high common Pb contents, so that calculation of the radiogenic component has a high uncertainty. If the V-Pb system has remained closed since crystallization ofthl! mineral, then the data can be plotted on isochron diagrams (such as 206Pb/204Pb vs 206Pb/238V or 206Pbj204Pb vs 207Pbj204Pb). The latter diagram is particularly useful in modeling tectonic environments and compositions of sources (cf. Zartman and Doe 1981).

V-Th-Pb ages have been determined for a wide variety of minerals. Zircon is by far the most popular, but other commonly dated minerals include sphene (titanite), baddeleyite, apatite, monazite, xenotime, rutile, and many highly radioactive minerals such uraninite and thorite.

Zircon as a Mineral Geochronometer

Zircon (ZrSi04) has a number of properties that make it the most commonly utilized mineral for V-Pb geochronology. Although it generally occurs as a trace


mineral, zircon is present in most rocks. It can occur in igneous rocks as a primary phase, as xenocrysts, or as newly crystallized material over xenocrystic (or inherited) cores. Zircon is an abundant detrital heavy mineral in clastic sedimentary rocks. It occurs in a wide variety of metamorphic rocks, both as metamorphic rims overgrowing pre-existing seed crystals and as a wholly new phase. Zircons have been extracted from lunar rocks, meteorites, and mantle xenoliths. The morphological and geochemical characteristics of zircons are related of their genesis and thus, the combination of these features with their geochronological information can be useful for defining specific geologic processes (Poldervaart 1955, 1956; Krasnobaev 1986; Bibikova 1989).

Another important property of zircon is that its crystal structure can easily accommodate uranium but tends to exclude Pb which has a larger ionic radius. Most well-preserved zircons ofthe sort used for geochronology contain less than 1 ppm common Pb and between about 10-1000 ppm U. Thus, for precise geochronological determinations, the correction for common (nonradiogenic) Pb is fairly trivial. In fact, most studies correct for the common Pb component using models of Pb isotopic evolution (cf. Stacey and Kramers 1975) rather than measuring common Pb composition of a co-existing mineral.

The crystal structure of zircon can be damaged by radioactive decay of V and Th, usually present in concentrations of several hundred ppm. This phenomenon, called metamictization, permits the loss of some of the radiogenic Pb, with the resulting V-Pb data being discordant. Another complicating factor for geochronology is that zircon is a very resistant mineral. Because of its very high melting point, refractory crystals of partially resorbed zircon are frequently preserved in anatectic melts. The isotopic information in these xenocrysts, when combined with data from the younger generation of zircon growth, can result in a very complex pattern of data on a concordia plot. The prudent geochronologist will attempt to isolate different populations of zircon, by a variety of methods, to obtain an understandable data set.

Zircon V-Pb geochronology has evolved through several stages of development. In the early 1950s, the low sensitivity of mass spectrometers required at least 100 p.g of Pb for analysis. Thus, an undivided bulk separate of zircon, usually about 0.5 g, was dissolved. Because many of the analyzed zircons were metamict, the Pb/V ages were discordant and the 207Pbj206Pb age was taken as the age of the mineral (and rock). Wetherill (1956) introduced the concordia plot, whereby Pb/V ages were graphically treated. Zircon samples were subdivided into magnetic and size splits (Silver and Deutsch 1963). These populations have different V concentrations and degrees of metamictization, resulting in a set of isotopic data composed of variously discordant points. In the simplest case, the data array is linear, forming a discordia with intercepts of the concordia curve. The upper intercept age is usually regarded as the time of crystallization of the zircon (and, by inference, the age of the rock). The lower intercept age is usually interpreted as a time of disturbance to the isotopic system, caused by a number offactors including metamorphism, fluid reworking, or dilatancy effect (Goldich and Mudrey 1972). A widely spread data array can also reflect the mixture of

8.2.5 The U-Pb System and Zircon as Mineral Geochronometer 371

two different populations of zircons. In this case, assuming the isotopic system has remained closed, the intercepts will correspond to the formation ages of each population.

Krogh (1973) developed the low-blank method for zircon analysis, including hydrothermal decomposition of zircon in Teflon bombs, followed by chromatographic extraction of V and Pb for mass spectrometric analysis. The introduction of mixed 205Pb_235V spike, as well as the creation of a new class of high sensitivity, automated mass spectrometers with multi-collectors permit an enormous reduction in sample size. It is now possible to analyze single zircon grains (and in some cases, parts of single crystals) that weigh only a few micrograms and contain less than 1 nanogram of Pb.

Modern zircon geochronology investigations involve very careful selection of grains to be analyzed. Treatments such as air abrasion (Krogh 1982), fine magnetic separation, HF leaching, and hand-picking of single grains or small groups composed of the least deformed or damaged grains (presumably with the lowest uranium content) frequently result in concordant or only slightly discordant V-Pb data. An understanding of the chemical composition of zircon is important for selecting the appropriate crystals for analysis. V, Th, Hf, Y, and Ca can be studied by electron microprobe. Distribution of V and Th can be imaged by fission track mapping. Zoning (both igneous and metamorphic) and the relationship of cores and overgrowths are studied using electron backscatter imaging and cathodoluminescence.

A major breakthrough in V-Pb zircon geochronology was accomplished with the construction of the SHRIMP (sensitive high-resolution ion microprobe) at the Australian National Vniversity (cf. Compston et al. 1984). With a primary ion beam of 20-30 J1.m diameter and a high resolution mass spectrometer, analyses of small spots on individual grains are possible. In addition to dating some of the oldest zircons in the world (4.1 Ga, Froude et al. 1983), the SHRIMP allows the deciphering of the complex geological histories of polymetamorphic terranes by dating individual zones in single zircons.

V-Pb geochronology of zircon is used to solve a wide variety of geologic problems including dating of igneous and metamorphic events. Volcanic rocks of intermediate to felsic composition usually contain magmatic accessory zircon. These zircons tend to be finely zoned, transparent, and relatively low in V and Th. Geological problems addressed by dating such zircons are the dating of stratigraphic sequences that include volcanic units, timing of different lava flows in granite-greenstone terranes, duration of volcanic processes, and correlation of different volcanic belts. Care must be taken to recognize xenocrystic zircon, incorporated during magma genesis or by terrestrial sedimentary processes during extrusion. The presence of an inherited component is often established in zircons from Phanerozoic volcanics.

Zircon as an accessory mineral is present in most intrusive rocks, ranging from quartz diorite to granite. In many gabbros, Zr may be accommodated in other minerals such as clinopyroxene, baddeleyite, or ilmenite. In rocks of mantle derivation (I-type), the emplacement age is easily established by dating


accessory zircon. Zircon from plutonic rocks of crustal anatectic origin (S-type) can be more difficult to date because they frequently have inherited xenocrystic cores. The occurrence of xenocrysts is due to the low saturation coefficient of Zr in granitic melts where inherited material from the source rocks is incompletely dissolved by partial melting (Watson and Harrison 1983).

The timing of metamorphic events can be established by dating zircon of metamorphic origin. Zircons that crystallize at granulite facies generally have a distinct morphology; they tend to be equant, multi-faceted ("soccer ball appearance"), and transparent. V-Pb data from these zircons frequently are concordant or only slightly discordant, probably due to the typically low V content (100 ppm or less). ThjV ratio is higher than in other zircon suites. Zircons crystallizing at amphibolite grade do not have such diagnostic characteristics and often resemble magmatic zircon. Because zircon can grow under many different metamorphic conditions, conventional geochronology of rocks in polymetamorphic terrains is very difficult. These problems are best attacked by using the high resolution ion microprobe.

Other problems that zircon V-Pb geochronology can solve include:

1. Provenance of terrigenous sediments. Detrital zircons are now dated individually, either by conventional methods or by ion microprobe. The oldest dated detrital zircons ( - 4.1 Ga) were extracted from a conglomerate of Late Archean age from western Australia (Compston and Pidgeon 1987).

2. The study of the upper mantle and lower crust by analyses of zircons from kimberlites and xenoliths. Zircons from kimberlites have extremely low V contents (5-50 ppm) and the crystals are not metamict. Nevertheless, under upper mantle P-T conditions, their V-Pb systems are open and thus, the ages of kimberlitic zircons tend to date the time of explosive emplacement.

3. V-Pb ages of xenocrystic zircon in plutonic rocks provide evidence for estimates of the magma source and the genesis of the pluton.

We have mainly discussed V-Pb isotopic systematics in zircon because few isotope laboratories are measuring Th-Pb ages. 208Pbj232Th ages tend to have similar discordance as 206Pbj238V ages. These data are particularly useful in dating Th-rich minerals such as monazite and tho rite.

Zircon can alse be dated by other methods, including fission track, Xe-Xe (similar to 4°Arj39Ar ages), and Sm-Nd. So far, good agreement between V-Pb and Sm-Nd ages has been accomplished only for metamorphic zircons. A newly developed method is thermal evaporation of Pb from single zircon grains mounted on a rhenium filament in the mass spectrometer (Kober 1986, 1987). However, this technique only yields 207Pbj206Pb ages and in zircons, where more than one period of growth exists, the data may be a mixture of ages. We caution against the indiscriminate use of this method because, despite extreme ease, the data may be incomplete and misleading even when the analytical precision is good. This method is best used as a preliminary guide for deciding which samples should be dated properly. We feel it is inappropriate to depend on this method alone to solve most geochronological problems.

8.2.5 The U-Pb System and Zircon as Mineral Geochronometer 373

In summary, zircon U-Pb ages are now regarded as the most reliable method for dating igneous rocks. The common occurrence of zircon in most igneous rock types and its resistance to thermal resetting make zircon an excellent geochronometer. However, the refractory character of zircon can result in the preservation of more than one event within individual grains. Cavalier treatment of zircons from complex polymetamorphic terrains can lead to spurious interpretations. To maximize the utility of zircon geochronology, each crystal must be scrutinized carefully and exacting criteria must be established for selection of appropriate grains for analysis. Under these circumstances, zircon U-Pb geochronology usually is successful.

References

Allegre CJ, Albarede F, Griinenfelder M, Koppel V (1974) 238UFo6Pb_23SUFo7Pb_ 232Thj208Pb zircon geochronology in Alpine and non-Alpine environment. Contrib Mineral Petrol 43: 163-194

Bibikova EV (1989) U-Pb geochronology of the early evolution of ancient shields. Nauka, Moscow, 180 pp (in Russian)

Boltwood BB (1907) On the ultimate disintegration products of the radioactive elements. Am J Sci 4: 77-88

Compston W, Pidgeon RT (1987) Jack Hills: a further occurrence of very ancient detrital zircons in western Australia. Nature 34: 123-125

Compston W, Williams IS, Meyer C (1984) U-Pb geochronology of zircons from lunar breccia 73217 using a sensitive high mass-resolution ion microprobe. J Geophys Res 89 (Suppl): B525-531

Froude DO, Ireland TR, Kinny PD, Williams IS, Compston W, Williams IR, Myers JS (1983) Ion microprobe identification of 4100-4200 Myr-old terrestrial zircons. Nature: 616-618

Goldich SS, Mudrey MG (1972) Dilatancy model for discordant U-Pb ages. In: Tugarinov AI (ed) Contributions to recent geochemistry and analytical chemistry, Vinogradov Volume. Nauka, Moscow, pp 415-418

Kober B (1986) Whole-grain evaporation for 207PbF06Pb-age-investigations on single zircons using a double-filament thermal ion source: Contrib Mineral Petrol 93: 482-490

Kober B (1987) Single-zircon evaporation combined with Pb+ emitter bedding for 207PbF06Pb_age investigations using thermal ion mass spectrometry, and implications to zirconology. Contrib Mineral Petrol 96: 63-71

Krasnobaev AA (1986) Zircon as indicator of geological processes. Nauka, Moscow, 146 pp (in Russian)

Krogh TE (1973) A low-contamination method for hydrothermal decomposition of zircon and extraction of U and Pb for isotopic age determination. Geochim Cosmochim Acta 37: 485-494

Krogh TE (1982) Improved accuracy ofU-Pb zircon ages by the creation of more concordant systems using an air abrasion technique. Geochim Cosmochim Acta 46: 637-649

Nier AO (1939) The isotopic constitution or radiogenic leads and the measurement geologic time. Phys Rev 55: 153-163

Nier AO (1940) A mass spectrometer for routine isotope abundance measurements. Rev Sci Instrum 11: 212-216

Poldervaart A (1955) Zircons in rocks.!. Sedimentary rocks. Am J Sci 253: 433-461 Poldervaart A (1956) Zircon in rocks. 2. Igneous rocks. Am J Sci 254: 521-554 Shukolyukov Yu A, Gorochov 1M, Levchenkov OA (1974) Graphical methods in isotope

geology. Moscow, Nedra, 207 pp (in Russian) Silver LT, Deutsch S (1963) Uranium-lead isotopic variations in zircon - a case study. J Geol

71: 721-758


Stacey JS, Kramers JD (1975) Approximation of terrestrial lead isotope evolution by a twostage model. Earth Planet Sci Lett 26: 207-226

Steiger RH, Jager E (1977) Subcommission on geochronology, convention on the use of decay constants in geo- and cosmochronology. Earth Planet Sci Lett 36: 359-362

Watson ED, Harrison TM (1983) Zircon, saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth Planet Sci Lett 64: 295~304

Wetherill GW (1956) Discordant uranium-lead Ages, I Trans Am Geophys Union 37: 320-326 Zartman RE, Doe DR (1981) Plumbotectonics - the model. Tectonophysics 75: 135-162

8.3 Noble Gas Isotopes in Planetary and Earth Minerals

YU. A. SHUKOLyuKovand M. OZIMA

Chemical inertness is the main characteristic of the zero group elements in Mendeleev's Periodical Table; these elements are He, Ne, Ar, Kr, Xe, and Rn. They can interact with other elements only indirectly either through Van der Vaals force on the surface of solid bodies or by forming clathrates.

Their very low abundance on planets relative to solar abundance is another characteristic feature of noble (rare) gases. The latter characteristic is the result of their chemical inertness: although the initial noble gas abundances in the universe and in the solar system are rather high, they could not have been captured by planet-forming materials in the early evolution stage of the protoplanetary cloud.

Noble Gases in Extra-Terrestrial Materials

Primordial Trapped Noble Gases in Minerals of Meteorites. Solar noble gases implanted (see the later section) into the surface layer of mineral particles by the solar wind in the primordial gas-dust cloud are generally present in all chondrite meteorites.

Noble gases with the very anomalous isotopic composition - for example, 22Ne, Xe with large enrichment in heavy and light isotopes simultaneously (Xe-HL), Ke and Xe with the isotopic compositions which are typical for the process in massive stars, 124Xe monoisotope, etc. - are contained in the particles of elemental carbon, silicon carbide, diamonds with the size ranging from 10 micron to 50 A, in sulfide shells of chondrules, in several silicate minerals in various types of chondrites (carbonaceous, ordinary, and enstatite) and iron meteorites.

These mineral particles were formed outside the Solar System, that is, in star shells of various types such as supernovas, novas, red giants, and so on. Noble gases trapped in these materials are of the pre-solar origin.

However, the main part of trapped gases in meteorites consists of the socalled planetary gases, with the following isotopic compositions, that is,

8.3 Noble Gas Isotopes in Planetary and Earth Minerals 375

3Hej4He = 1.43 x 10-4, 2°Ne/22Ne = 8.2, 21Nej22Ne = 0.024, 36Arr8Ar = 5.31, 78Kr;S4Kr = 0.00597, 8°Kr;S4Kr = 0.0392, 82Kr;S4Kr = 0.2015,

83Kr/84Kr = 0.2017, 86Kr/84Kr = 0.3098, 124Xe/130Xe = 0.0283, 126Xe/ 130Xe = 0.0253, 128Xe/130Xe = 0.504, 129Xe/130Xe = 6.36, 131Xej130Xe = 5.07, 132Xe/130Xe = 6.21, 134Xe/136Xe = 2.37, 136Xe/130Xe = 1.99.

This is not a unique component, but appears to be a mixture of at least three major components.

Radiogenic Gases in Minerals of Meteorites and Moon Samples. Two types of radiogenic gases are contained in minerals of meteorites and moon samples. The first is the decay products of so-called extinct radioactive isotopes 1291 (T 1/2 = 16 Ma) and 244pU (T1/2 = 82 Ma). Due to the beta decay of 1291, 129Xe has

been accumulated in silicate minerals, and the 129Xe/130Xe ratio becomes in some cases more than 20 times of the normal (planetary) value e29Xe/130Xe = 6.2). 244pU produces heavy Xe isotopes: 136Xe: 134Xe:132Xe: 131Xe =

1.00:0.939:0.870:0.246 in vitlocite and other phosphates. On the basis of 129Xe/ I and 136Xe/REE or U (geochemical analogs of Pu), it is possible to measure the time interval between the completion of the r-process nucleosynthesis which produced 1291 and 244pU and the closure time of 129Xe and 136Xe in minerals.

Radiogenic 4He and 40 Ar, alpha decay products of U and Th, and also members of their radioactive series and of 4°K electron capture, are present in a variety of materials in the Solar System. Only the minimum age of minerals can be estimated on the basis of these radiogenic elements, because 4He and 40 Ar are likely to diffuse out of the system.

Cosmogenic Isotopes of Noble Gases in Minerals of Meteorites and Moon Samples. Protons and other particles of galactic and/or solar cosmic radiation with kinetic energy from a few tens of MeV to GeV result in spalation in target materials. Noble gas isotopes are among the cosmogenic (spallogenic) isotopes thus produced.

Cosmogenic (spallogenic) isotopic ratios of noble gases differ from the normal (atmospheric) ratios by more than an order of magnitude, and are highly variable, depending on the chemical composition of target materials and the bombarding particles.

Cosmogenic isotopes were formed in different evolutionary stages of the Solar System materials.

Mineral particles before their accretion into a single parent body of meteorites, i.e., an asteroid, could have been affected by the influence of intensive galactic cosmic ray radiation in the protoplanetary cloud.

From the moment of the separation of meteorites from their parent body, where the meteorites had been buried deep in the interior, meteorites have been subjected to solar and galactic cosmic ray radiation.

Irradiation of minerals on the surface of the meteorite parent body, which is generally free from the protecting atmosphere, is also possible, as is the case for the materials on the surface of the moon.


The time interval between the fragmentation of a meteorite parent body and the entry of the meteorite into the Earth can be determined by measuring cosmogenic isotopes of noble gases in the meteorite. A terrestrial age - the residence time of a meteorite in the Earth - can also be determined with the use of cosmogenic noble gas isotopes. The time of ejection of lunar rocks from the surface, which then resulted in the formation of craters, can be estimated on the basis of cosmogenic noble gas isotopes.

Noble Gases in Minerals of the Earth

The air is the main reservoir for the terrestrial noble gases: 2.08 x 1019 7.2 X 1019,

3.7 X 1022,4.52 X 1018 and 3.4 x 1017 cm3 of He, Ne, Ar, Kr, and Xe accordingly or in the specific volume (cm3 /g): 5.25 x 10-6, 1.82 X 10- 5, 9.34 X 10- 3, 1.14 X 10- 6,8.7 X 10- 8•

Noble gases in the Earth are a mixture of several components of a different origin.

Primordial Noble Gases in Minerals of the Earth. Primordial noble gases may be present in some mantle-derived materials and in fluids from modern tectonomagmatic active zones or from high seismic and thermal regions. The former includes ultrabasic inclusions of pyroxenite-periodotite-olivine in basalt and andesite from island arcs, tholeiite basalts from the midoceanic ridges, xenocrysts and inclusions in alkaline basalts of intraplate volcanism region, eclogite mineral inclusions in kimberlites, and diamonds.

Many of these materials contain He isotopic ratio eHerHe) 10-100 times larger than the atmospheric and the crustal ratios, ranging from 3 x 10-4 to 1 X 10- 5• The ratio is similar to those trapped in carbonaceous chondrites. The relation between He isotopic composition and tectonic characteristic is widely used for the study of the evolution of the mantle-crust system and the geochemical heterogeneity in the mantle.

Ne isotopic ratio eONe/ 22Ne) in mantle-derived materials is significantly larger (11-13) than the atmospheric isotopic ratio eONe/ 22Ne = 9.8), suggesting that the primitive Ne is of the solar type. However, there appears to be no difference in 38 Arp6 Ar and Kr isotopic ratios between the mantle and the atmosphere. Ar has been accumulating in the atmosphere, radiogenic isotope 40 Ar being the main component. Attempts to determine the isotopic composition of the primordial argon from the observation of argon trapped in some minerals from Precambrian shales (?) and volcanics (?) have been unsuccessful. The search for the primordial Ar in the Earth is difficult because of atmospheric argon contamination in rocks. Additionally, a significant amount of radiogenic 40 Ar, the decay product of 4°K, is always present in the mantle argon, and hence 40 Ar r 6 Ar ratio in the modern mantle generally exceeds 25000.

40 Ar p6 Ar ratio in the mantle-derived materials as well as in the atmosphere and in the crustal rocks have been used to investigate the degasing history of the


Earth. Xe in the Earth's atmosphere is different from the solar Xe, being enriched in the heavier isotopes and having excess 129Xe. The difference cannot be attributed to an isotopic mass fractionation. Many authors have assumed that the primitive xenon component, i.e., V-Xe, did exist in the Solar System. In this context, addition of fissiogenic, radiogenic, and cosmogenic Xe to the V-Xe is interpreted as giving rise to the solar and planetary Xe. Mass fractionation of VXe and subsequent addition of 244pU fission Xe resulted in the atmospheric Xe. However, recent finding of Xe components different from the above-mentioned xenons in some meteorites seems to cast a doubt on the existence of the V-Xe. The problem is still open to future investigation.

Radiogenic Noble Gases in Minerals of the Earth

Helium

The principal source for the main helium isotope 4He is the alpha decay of radioactive elements. The radiogenic 4He production rate is about 4 x 107 m3jy or 4 x 1017 m3 for 4.5 Ga. Single alpha decay makes a small contribution, but the absolute quantities of radiogenic 4He produced for 4.5 Ga is enormous: 147Sm _ 1000km3, 142Ce _ 1 km3, 152Gd, 156Dy, 177Hf, 144Nd - 107-108 m3.

The radiogenic 4He has essentially been outgased into the atmosphere and then dissipated to space. Among igneous rocks, acidic ones are most enriched in 4He, in which the production rate is about 10- 11 cm3 Ga. In minerals in ultrabasic rocks the production rate of 4He is 2-3 x 10 -13 cm3 jGa. In sedimentary rocks 4He is generated in black marine shales up to 6 X 10- 12 cm3jGa and in clays up to 2 X 10- 12 and in carbonate up to 6 X 10- 13 cm3jGa. The production rate of 3He due to an alpha-decay is 109_1010 time lower than that for 4He.

Neon

Some amount of 22Ne, 21Ne can in principle be produced by highly asymmetric spontaneous nuclear fission of V, Th, and radioactive members of their series. This process was predicted in 1969-70 on the basis of the investigation of some uranium minerals and was recently confirmed by experiment. However, the rate of Ne isotope production is extremely small even in uranium minerals.

Argon

A similar process of highly asymmetric spontaneous fission can be the source for some of radiogenic argon isotopes in nature: 40 Ar, 38 Ar, and 36 Ar. The 40 Ar isotope is produced mainly by electron capture of 4°K nuclei. The decay


constant is 0.581 x 10- 10y-1, and this provides the basis for the K-Ar dating method.

A large amount of radiogenic 40 Ar is observed generally in ancient minerals and in minerals with high potassium concentration: in mica up to 8 X 10- 3, in ancient amphiboles up to 2 X 10- 2, in pyroxene and peridotites 2 x 10-4 cm3/g respectively.

Krypton

The main source for radiogenic Kr isotope production is the spontaneous fission of 238U (T 1/2 = 9.9 X 1015 to 6.7 X 1015 years). Spontaneous fission of 235U and 232Th does not make any contribution.

The isotopic composition of fission Kr differs greatly from the composition of the atmospheric Kr: 86Kr/83Kr = 29.9, 84Kr/83Kr = 4.18 and 82Kr, 8°Kr, and 78Kr isotopes are absent. The radiogenic Kr can be used to determine ages of uranium containing minerals (uranium pitchblende, uraninite, monazite, zircon, samarskite, khlopinite, betafite, etc.).

Approximately 0.2 km3 of fission Kr has been produced in the Earth during 4.5 Ga. The average concentration in common rock-forming minerals is 3 x 10- 14 cm3/g, but in uranium minerals it amounts up to 10- 8 cm3/g. Double beta decay of 82Te (T 1/2 = 1.4 X 1020 y) is another source for radiogenic Kr in the Earth. The maximum abundance of radiogenic isotope 82Kr is found in selenocobellite which has the isotopic ratio 82Kr/83Kr = 4.7 compared to the atmospheric value of 1.0. The excess of radiogenic 82Kr can be found in blokite, glaustalite, naumannite, tiemannite, umangite, the amount ranging from 5 x 10- 13 to 2 X 10- 12 cm3/g.

Xenon

The production ofXe by the 238U spontaneous fissions is 11.5 times higher than that of Kr. The isotopic composition of the spontaneous fission Xe is 136Xej132Xe = 1.73, 134Xe/132Xe = 1.42,131Xe/132Xe = 0.142, 129Xe/132Xe = 0.003. They are significantly different from the atmospheric Xe.

Approximately 2 km3 of 238U fission Xe has been accumulated in the Earth for 4.5 Ga. Most of them are present in uranium minerals, amounting up to 10-6 cm3 /g (uraninite, uranium pitchblende, broggerite, brannerite, and others). In rare earth minerals with 0.001-0.1 % of U, the amount of spontaneous fission Xe ranges from 10- 10 to 10- 8. In granites it amounts to 10- 11 cm3/g, and in basic rocks to 10- 13 cm3/g. Radiogenic Xe in uranium and uranium-bearing minerals can be used as Xe-U and Xes-Xen age determination methods, which are useful for ages older than 0.1 Ma.

130Xe isotope is formed by 2-P decay from natural 13°Te isotope (T 1/2 = 3.1020 y). For example, Xe in natural Te (Colorado, USA) is found to be


enriched 50 times more in 130Xe compared to the atmospheric Xe isotopic composition. The product of the 2-/3 decay of 13°Te is also found in telluriumcontaining minerals: in tellurobismuthite and bismuth telluride (up to 10- 11 cm3/g of 130Xe). During the Earth's life time, about 20 m3 of radiogenic 130Xe has been produced.

It is possible that 128Xe is also formed by the 2-/3 decay of 128Te. Radiogenic Xe isotopes such as decay products of "extinct" 1291 and 244pU (see above) are of special interest. Excess 129Xe (up to 10%) has been reported for natural gases, MORB glasses, and various mantle-derived materials, including diamonds.

The existence of excess 129Xe has been argued for the early and fast (catastrophic) degassing of the mantle. Here, it is generally assumed that substantial fraction of Xe was transported into the atmosphere from the solid Earth before the total decay of 1291, and the radiogenic 129Xe subsequently produced from 1291 became more conspicuous in the degassed mantle. The existence of excess 129Xe produced from 244pU has not been confirmed. It is also worth noting that the excess 129Xe is generally correlated with the excesses in 132Xe and 131Xe of an unidentified component which has anomalously high 132Xe/130Xe and 131Xe/130Xe ratios.

Noble Gas Isotopes - Products of Induced Nuclear Reactions in Minerals. Particles and radiation emitted during spontaneous nuclear reactions can in turn induce secondary nuclear processes. When target minerals are subjected to particles and radiation with high energy, noble gas isotopes will be produced. Such a process plays an important role in noble gas geochemistry.

Helium

4He isotope is produced by (n, a) reactions in light nucleus targets, mainly lOB and 6Li. The total volume of 4He produced in this way for 4.5 Ga in the Earth's crust is about 1 km3 or 10- 10 cm3/g, which is much less than that produced by the alpha-decay.

3He is produced in minerals only as the result of some nuclear reactions such as 7Li (a, n) 3H --+ 3He, 6Li(n, a)3H --+ 3He, 2H (n, y) 3H --+ 3He. Existence of excess 3He in beryls, spodumenes, radioactive minerals, and granites can be attributed to the above nuclear processes.

Neon

Neon isotopes can be produced in silicate and other minerals due to reactions such as 180 (a,n) 21Ne; 19F(a,n) 22Na --+ 22Ne; 19F (a,n)22Ne. In uranium and uranium-containing minerals excesses in 21Ne and 22Ne amount to 4 x 10- 7 cm3/g, and 21Nej2°Ne and 22Nej2°Ne ratios are higher by 300 and 50 times than the atmospheric ratios. The average production ratio of 21Ne in


common rock-forming minerals is 1.5 x 10- 14 cm3 jGa. Hence, for 4.5 Ga about 1 km3 of 21 Ne (about 1-2% of the Earth's total inventory) has been produced in the Earth. Excess 21Ne is observed in various minerals, rocks, and natural gases. Since 21Ne production is due to alpha particle radiation, the ratio of 4He to 21 Ne is expected to be constant. In fact the ratio is nearly constant in many silicate minerals, being (1.2 ± 0.7) x 107.

Argon

In radioactive minerals 38 Ar and 36 Ar isotopes can be produced due to nuclear reactions 35 Cl(oc,n)38K -+ 38Ar + 13, 35CI(oc,p)38Ar, 35CI(n,y)36Cl-+ 36Ar, 37Cl(n,y)38CI-+ 38Ar, 4IK(n,oc) 38Cl-+ 38Ar.

In ordinary rock-forming minerals there is more than 10- 13 cm3jg of the nucleogenic 38 Ar, but in uranium minerals the concentration of this isotope amounts to 2 x 10- 8 cm3 jg, and 38 Arj36 Ar ratio exceeds 6 in comparison with the atmospheric ratio of 0.187. Excess 36Ar is also present in uranium-rich minerals.

Krypton and Xenon

Due to the influence of natural thermal neutron flux, in minerals 235U induced fission could take place which then produces Kr and Xe. Their isotopic compositions are distinctly different from the ratios for the 238U spontaneous fission components, that is, 86Krj84Kr = 1.92, 83Krj84Kr = 0.45 136Xe;t32 Xe = 1.47, 134Xejl32Xe = 1.82, 131Xej132Xe = 0.67, 129Xejl32Xe = 0.154.

Efficiency in the production of the fission Kr and Xe due to the induced 235U fission depends on the relative abundances of both light elements (H, C, 0, N, etc.) in the environment which act as a neutron moderator and competing neutron absorber elements such as rare earth elements which have a larger cross-section for a neutron capture reaction. Consequently, Kr and Xe isotopic compositions in radioactive minerals depend upon their chemical composition. The fraction of the neutron-induced fission Xe is about 14% in uraninites, 25% in uranium pitch blends, 10% in monazites, and a few % in REE-enriched minerals, but amounts to 50% in some titanotantaloniobates.

In the rock-forming minerals in granites, basalts, and ultrabasic rocks, the fraction of neutron-induced Kr and Xe does not exceed several percent of the total inventory, where the concentration ranges from 10- 12 to 10- 17 cm3jg. Their total volume does not exceed 107 m3 and their contribution in the Earth inventory is not significant.

A surprising exception is in the Oklo uranium deposit in the Republic of Gabon. Here, about 2 billion years ago, a chain fission reaction took place: ore bodies with the size of 30 x 10 xl m, containing 20-60% ofU, were formed with


a large concentration of water and a low concentration of REE. Under the influence of a neutron flux of about 108 neutron/cm2 s for 5 x 105 years, 10-4 cm3/g of Xe and 10- 5 cm3/g of Kr were produced in the uranium pitch blend. The amounts thus formed comprise more than 99% of the existing radiogenic Kr and Xe.

Other nuclear reactions with neutrons and alpha particles can also produce Kr and Xe isotopes. For example, in umangite 83Kr isotope is produced by a reaction 82Se(n, y) 83Se ~ 83Kr, and 82Kr, and 78Kr can be produced by (a, n) reactions from Se, As, and Br nucleii. Kr thus formed has a 83Kr/84Kr ratio seven times higher than the atmospheric ratio. Together with l3OXe, which is a product of double beta-decay of 13°Te, both 129Xe produced by (n, p)-processes on 128Te and 130Xe produced by the similar reaction on 13°Te are usually contained in tellurium minerals.

Characteristics of Noble Gas Occurrence in Minerals

The origin of noble gases in meteorites and in the Moon samples is generally attributed to ion implantation. Ions of noble gases accelerated in electromagnetic fields in the star shells or in the interplanetary space are implanted on the surface layer of solid mineral particles to the depth of about 500 A, where concentration of noble gases is inversely proportional to the size of mineral particles.

Noble gas could have been trapped on the surface of fine grained mineral particles due to physical adsorption at low temperature and to high gas pressure in the proto-planetary gas cloud. The adsorption effect is strengthened due to the "labyrinth mechanism", that is, the penetration of noble gas atoms deep inside the minerals along the very fine communicating capillaries.

The capture of noble gas atoms is also possible by thermal excitation of the mineral crystal lattice. This situation is likely to be realized during the heating of a dust-gas system in the proto-planetary cloud. Owing to this process, noble gas atoms are located in defects of a crystalline structure.

There are various ways for terrestrial minerals to capture noble gases. For example, noble gases can be absorbed in melt. It is then likely that noble gases are homogeneously distributed in a melt, and are subsequently quenched in solid bodies. However, part of the gases may be captured in micro-inclusions. In some crystalline minerals, captured gases can be located in chanels (beryl and cordierite). Noble gas can also be trapped in some carrier phases included in minerals (the similar situation can be found for meteorites). A radioactive decay and a nuclear reaction often result in the formation of hot atoms which are completely or partly deprived of normal electronic shells for example (4He)+2, (4°Ar)+1, (136Xe)+2, etc. The hot atoms thus formed can primarily undergo chemical reactions with atoms in a crystal. For example, the production of Xe, Kr, Ar, and even of He fluorides is possible in minerals enriched in fluorine. Oxides such as Xe03 can also be formed.


In accordance with the location of parent radioactive elements in a mineral structure, noble gas atoms can be distributed in narrow or rare interstitial sites or they can be concentrated in some microscopic phases where the concentration of noble gases would exceed by 1000 times the concentration in the matrix.

Also radiogenic noble gas atoms can be located in the structure defects. Defects appear either as initial crystal defects in the form of mosaic structure or as later damage of the crystal lattice due to the energetic particles produced by a radio-active decay. An alpha particle produced by an alpha decay can generally displace about 50 lattice atoms, and travels over a distance of 300 A. A recoil nuclei produced by an alpha decay and a fission fragment can displace up to 1500 and 10000 atoms respectively. Hence, in some minerals (metamict zircon, gatchetolite, betafite, obruchevite, etc.) all atoms have been displaced during geological time, and hence the crystal parameters have been drastically changed. The changes in mineral chemical composition (for example U transformation to Pb and He) and in the valence condition (for example U+ 4 + -+ U+ 6) are also responsible for this alteration. All these processes result in the extremely complicated atomic displacement of noble gases.

Noble Gas Migration in Minerals

Noble gas migration is characterized by two principal processes. Firstly, there are at least two mechanisms of noble gas atom movement in a mineral structure. In some cases migration is controlled by the law of classic diffusion. The concentration of a noble gas at a time t and at a distance x from the center of an isotropic spherical body can be expressed as (Fick's law);

c5c(x,t) = Dc52c(x,t) c5t c5x2 '

where D denotes a diffusion coefficient which depends exponentially on temperature.

In another case, migration can result from one or several jumps of a noble gas atom in a lattice. The displaced atoms can reside only transiently in the disturbed structural zone (defect) in a geological time scale. Consequently, the number of noble gas atoms leaving the structure per unit time is proportional to the number of atoms at time t. Hence, we have

where vp = voe - E/RT is the frequency of atom's jump from interstition or lattice point, E the activation energy of atoms migration in the structure, T the temperature, Vo the frequency coefficient, and p the probability that the skip will lead to the escape of atom from the structure. The proportional coefficient k = vope- E/RT depends, first of all, on the crystallographic feature: the greater the


packing coefficient, the more difficult is the migration of noble gas atoms along the mineral structure, other conditions being equal.

The second characteristic feature in noble gas migration in minerals is the presence of several maxima in the noble gas thermal release curve. This fact indicates that the migration processes are controlled by several activation energies which correspond to different energy states of noble gas atoms in minerals, or different energy depths in a potential well where the atoms reside. Atoms of different noble gases can reside in the same defect.

Fractionation of Noble Gases During Migration

Noble gas elemental fractionation generally takes place during their migration in minerals. Since a Henry's constant depends on the molecular mass of gases, adsorbed gases must be fractionated relative to the original composition. This process applies to the case in which noble gases were adsorbed onto fine mineral particles such as proto-planetary dust particles. Fractionation also takes place during ion implantation into mineral particles under the influence of the electric field in the gas-dust cloud.

The formation temperature of clathrates, especially of ice clathrates, varies greatly for different noble gases. This results in noble gas elemental fractionation in the protoplanetary cloud, and also during the formation of underwater gas hydrates in the Earth.

Noble gases are also fractionated in an equilibrium partition between liquid and solid, and between gas and liquid because of the difference in their solubility among the phases.

A diffusion coefficient of noble gas depends significantly on the atomic mass. This leads to the separation of noble gases during diffusion through minerals.

Substantial noble gas fractionation takes place during thermal annealing of minerals, if the gases are located in different defects of crystal structure or are of different nuclear origin. For example, by annealing potassium-bearing minerals which contain accessory radioactive minerals, the 4He and 40 Ar ratio varies greatly, or elemental fractionation takes place.

Similar fractionation is observed in meteorites which contain noble gases of different origin.

In parallel with the elemental fractionation of noble gases, isotopic fractionation occurs for the same physicochemical processes. However, in many cases, the latter effect is small and hardly exceeds a few percent per a.m.u. (atomic mass unit).

However, in the case where different isotopes ofthe same element are formed in different nuclear processes and hence reside in different sites in the crystal, mineral annealing may result in an isotopic effect amounting up to a few hundred to thousand percent.

Another type of large noble gas fractionation can be seen in the case where isotopes of a noble gas were derived from different radioactive precursory


elements of long half-lives such as 238U and 235u. Fission of the latter element produced transient radioactive elements such as Te, I, and Se, which eventually decay to Xe and Kr. Because of the different mobility of the transient intermediate elements, the final decay products Kr and Xe must have isotopic compositions different from those expected for a closed system. This phenomenon can be observed under circumstances where generation and intensive migration of radioactive precursors of noble gases took place simultaneously as the result of high temperature or high mineral dispersity.

Therefore, in uranium pitchblende from the "natural nuclear reactor" (Oklo deposit, Republic of Gabon) which had operated at temperatures 300-500 °e, the 132Xej136Xe ratio is ten times higher than that expected from 235U-induced fission. In a very fine-grained uranium mineral, i.e., "uranium blacks", this ratio is in some cases even higher; for example, in the sand from the epicenter of the first atomic bomb explosion in Alamogordo it is 100 times higher than that produced by 235U fission. All these facts are attributed to the above-mentioned migration effect.

References

Allegre CJ, Staudacher T, Sarda P (1986) Rare gas systematics: formation of the atmosphere, evolution and structure of the Earth's mantle. Earth Planet Sci Lett 81: 127~150

Amari S, Ozima M, Imamura M (1986) Search for the extra-terrestrial materials in deep sea sediments. Mem Natl Inst Polar Res Spec Iss N41: 338~347

Hebeda EH, Schultz L, Freundel M (1987) Radiogenic, fissiogenic and nucieogenic noble gases in zircons. Earth Planet Sci Lett 85: 79~90

Jambon A (1986) Solubility of He, Ne, Ar, Kr and Xe in a basalt melt in the range 1250~1600°C. Geochemical implications. Geochim Cosmochim Acta 50: 401~408

Kennedy BN, Hiyagon H, Reynolds JH (1990) Crustal neon: a striking uniformity. Earth Planet Sci Lett 98: 277~286

Lux G (1987) The behavior of noble gases in silicate liquids: solution, diffusion, bubbles and surface effects, with applications to natural samples. Geochim Cosmochim Acta 51: 1549~1560

Mamyrin BA, Tolstikhin IN (1983) Helium isotopes in nature. Elsevier, Amsterdam Ozima M, Podosek F (1983) Noble gas geochemistry. Cambridge Univ Press, Cambridge,

367 pp Shukolyukov YuA (1982) Products of fission of heavy elements on the Earth. Energoizdat

Moscow, p. 127 pp (in Russian) Shukolyukov YuA, Dang Yuh Minh (1984) Products of fission of trans uranium elements in

space. 1984, Nauka Moscow, 119 pp (in Russian) Torgersen T (1989) Terrestrial helium degassing fluxes and the atmospheric helium budget

implications with respect to the degassing processes of continental crust. Chern Geol (Isotope Geosci Sect) 79: 1~14

8.4.1 Ph Isotopy; the Lead Sources

8.4 Radiogenic Isotopes as Indicators of Sources of Mineral Matter

8.4.1 Pb Isotopy; The Lead Sources

I.V. CHERNYSHEV and B.L. GULSON

385

The pioneering investigations by A. Nier about 50 years ago provided the impetus for the inception and general acceptance of the field of isotope geology and geochronology. Lead isotopes have diverse applications in Earth Sciences, but most importantly provide evidence for the source of metals in mineral ore matter.

A resurgence over the last 12-15 years of the use of Pb isotopes in source investigations has been facilitated by radical improvements in instrumentation and methods of isotope analysis, complemented by interest in the characterization of different rocks and geological reservoirs for understanding global lead isotopic evolution.

The lead-isotope parameters most commonly used in lead studies are: isotope ratios 206Pb/ 204Pb, 207Pb/ 204Pb, 208Pb/ 204Pb, and also 207Pbj206Pb, parameters of parent U-Th-Pb systems ,u_238U/204Pb and W = 232Th/204Pb, and the values of so-called Pb-Pb model age, calculated from the isotope parameters. General and methodical aspects of lead isotopes are considered in some monographic works (see Koppel and Faure 1986; Gulson 1986; Griinenfelder 1979).

In evaluation of lead sources, two complementary approaches are employed. The first approach consists of a comparison of the lead-isotope

characteristics of minerals and corresponding parameters from the terrestrial lead evolution curves. The most popular evolution models are: the two-stage model by J. Stacey and J. Kramers (1975); the continuous evolution model (constantly increasing U/Pb and Th(Pb ratios in the source reservoir) by G. Cumming and J. Richards (1975); the "plumbotectonics" model by B. Doe and R. Zartman (1979). Using these models and the measured lead-isotope data, conclusions can be drawn such as whether the lead was derived from distinct global reservoirs (e.g., mantle, crust) as well as from distinct geodynamical or geochemical environments (e.g., rejuvenatated craton, a source with high U/Pb, etc.).

Quite definite signatures indicative of mantle origin are peculiar to lead, for example, from sulfide inclusions in diamonds in kimberlites of South Africa (Kramers 1975) and lead from the oldest (about 3.0 Ga) massive Cu-Ni sulfide deposits of Ontario province (Stacey et al. 1977). Other younger sulfide deposits of the same type, located in the North-American Continent and Baltic Shield, contain significant amounts of crustal lead. Apparently the mantle sources of metals played the most important role in ore-forming processes in the early


stages of the Earth's history when the volume of crust and its differentiation products were not yet significant.

Lead-isotope signatures reveal some interesting features of the metal origin for stratiform deposits. At least two distinct kinds of lead isotope composition are observed during mobilization and extraction of metals from rocks: (1) the lead with higher and strongly variable content of radiogenic isotopes 206Pb and 207Pbeo6Pb/204Pb> 19.0, 207Pb/204Pb > 16.0), which was classified as upper crustal (Doe and Zartman 1979); (2) the relatively homogeneous lead corresponding to common crustal lead. The first is characteristic for stratiform deposits of the Mississipi region, the second for the polymetallic Pine-Point deposit in Canada (Cumming and Robertson 1969), for the West German Kupferschiefer deposits, and for others such as the sediment-hosted deposits of Mt. Isa and The McArthur River (Gulson 1985). These types of lead isotope composition apparently are due to different mechanisms and geological conditions of the metal mobilization. In the case of Mississipi Valley type deposits of the USA, the selective leaching of lead from relatively U-rich Precambrian rocks of the basement by hot brines possibly takes place. This process leads to enrichment of ore lead with radiogenic isotopes (Heyl et al. 1974). The metamorphic reworking of deep-seated rocks facilitates the formation of long-lived and extensive hydrothermal systems which produce the second type of homogeneous lead isotope composition. Such conditions possibly took place in the formation of the Pb-Zn Sardana deposit in Jakutia (Chernyshevand Pavlov 1982). Alternatively, the source of the metals may have been from thick sedimentary basins whose sediments were already partially homogenized during sedimentation (Vaasjoki and Gulson 1986).

The second approach is based on the comparison oflead isotopic ratios of ore minerals and potential source rocks. This approach combined with geological and other geochemical data, permit the development of conceptual models for ore genesis and mineral exploration.

For example, in the Que River volcanogenic massive sulfide deposit (western Tasmania), a comparison of lead isotopes in the sulfide mineralization and Cambrian volcanics indicated a common source (Gulson 1986). A similar isotopic relationship between sulfide ore of the Japanese Kuroko deposits and volcanic processes was observed by Sato (1975), Sato et al. (1981) and Fehn et al. (1983). Studies of metalliferous sediments in recent oceanic spreading zones (East pacific Rise) and corresponding paleo-environments (Cyprus, Syria, Oman) suggest that underlying basalts were the source o/the metals; the lead and other metals were mobilized by water-rock interaction in ocean bottom environments (Gale et al. 1981; Chen and Pallister 1981; and other works).

For deposits associated with continental volcanism, a more complicated picture was discovered, pointing to more than two sources of lead. In the wellknown ore field of San Juan, Colorado, vein-type polymetallic deposits of Au, Ag, Cu, Pb, and Zn, containing different quantities of radiogenic isotopes 207Pb and 206Pb, the lead isotope composition of ore minerals differs greatly from those in volcanics and is characterized by specific 207Pb/206Pb ratios. The data

8.4.1 Pb Isotopy; the Lead Sources 387

strongly suggest that the main part of lead was extracted from Precambrian basement rocks during the action of convective hydrothermal systems (Doe et al. 1979).

The results of many lead-isotope investigations, combined with other geochemical data, indicate that the source of uranium for uranium deposits was granitic rocks enriched in uranium. Based on Pb-Pb isotopic data, the paragenesis of uranium and gold, often observed in hydrothermal deposits, is determined by the geochemical characteristics of potential source rocks. The Pb-isotopic signatures of the well-known gold deposit Kolar, India, show that the extraction of gold took place from old ( > 3.0 Ga) rocks enriched in uranium (Chernyshev 1988).

The application of the lead isotope method is not restricted only to the determination of the source of Pb, but can be extrapolated to other metals such as Zn, Cu, Ag, Au, Sn, and U. Furthermore, when used in multi-isotopic investigations (e.g., Pb-Sr-Nd) the Pb isotope data are correlated with Sr and Nd isotopes and provide constraints on petrogenetic interpretations (e.g., Chernyshev et al. 1986).

References

Chen JH and Pallister JS (1981) Lead isotopic studies of the Samail Ophiolite, Oman. J Geophys Res 86: 2699-2708

Chernyshev IV, Pavlov DI (1982) Lead isotope studies of the stratiform Pb-Zn ore deposits of the south-eastern marginal land of the Siberial platform. In: 5th International Conference on Geochemistry, Cosmochemistry, and Isotope Geology. Nikko, Japan, pp 47-48

Chernyshev IV, Troitsky VA, Zhuravlev DZ (1986) Pb, Sr and Nd isotopes in minerals of tungsten deposits. Terra Cognita 6: 226-227

Cumming GL, Richards JR (1975) Ore lead isotope ratios in a continuously changing Earth. Earth Plan Sci Lett 28: 155-171

Cumming GL, Robertson DK (1969) Isotopic composition of lead from the Pine Point deposit. Econ Geol 64: 731-732

Doe BR, Zartman RE (1979) Plumbotectonics: the Phanerozoic. In: Barnes HL (ed) Geochemistry of hydrothermal ore deposits, 2nd edn. Wiley, New York, pp 22-70

Doe BR, Steven TA, Delevaux MH, Stacey JS, Lipman PW, Fisher FS (1979) Genesis of ore deposits in the San Juan volcanic field, Southwestern Colorado - lead isotope evidence. Econ Geol 74: 1-26

Faure G (1986) Principles of isotope geology. New York, 464 pp Fehn U, Doe BR, Delevaux MH (1983) The distribution oflead isotopes and origin of Kuroko

deposits in the Hokuroku district, Japan. Econ Geol Mon. 5: 488-506 Gale NH, Spooner ETC, Potts PJ (1981) The lead and strontium isotope geochemistry of

metalliferous sediments, associated with Upper Cretaceous ophiolitic rocks in Cyprus, Syria and the Sultanate of Oman. Can J Earth Sci 18: 1290-1302

Gulson BL (1985) Shale-hosted lead-zinc deposits in northern Australia: lead isotope variations. Econ Geol 80: 2001-2012

Gulson BL (1986) Lead isotopes in mineral exploration. Elsevier, Amsterdam, 245 pp Koppel V, Griinenfelder M (1979) Isotope geochemistry of lead. In: Koppel V, Griinenfelder

M (eds) Lectures in isotope geology. Springer, Berlin Heidelberg New York, pp 134-153 Kramers JD (1975) Lead, uranium, strontium, potassium and rubidium in inclusion-bearing

diamonds and mantle derived Xenoliths from Southern Africa. Earth Planet Sci Lett 42: 58-70


Sato K (1975) Unilateral isotopic variation of Miocene ore leads from Japan. Econ Geol 70: 800-805

Sato K, Delevaux MH, Doe BR (1981) Lead isotope measurements on ores, igneous and sedimentary rocks from the kuroko mineralization area. Geochem J 15: 135-140

Stacey JS, Kramers JD (1975) Approximation of terrestrial lead isotope evolution by a twostage model. Earth Plan Sci Lett 26: 207-221

Stacey JS, Doe BR, Silver LT, Zartman RE (1977) Plumbotectonics II A, Precambrian massive sulfide deposits. US Geol Surv Open File Report 76-476: 26 p

Vaasjoki M, Gulson BL (1986) Carbonate-hosted base metal deposit: lead isotope data bearing on their genesis and exploration. Econ Geol 81: 156-172

8.5 Light Stable Isotope Ratios as Indicators for Conditions of Mineral Formation

8.5.1 Theoretical Aspects of Isotopic Fractionation

J.R. O'NEIL and E.M. GALIMOV

The theory of isotopic fractionation treats mass-dependent fractionation of the isotopes of light elements. Isotopic ratios of heavy elements like Sr or Pb do not fractionate between phases in response to the common physical chemical processes occurring on Earth. In mineralogy, and earth science in general, we are concerned mainly with variations in the stable isotope ratios of only five light elements: H, C, N, 0, and S. The magnitude and direction of the fractionation of stable isotope ratios of these elements between minerals is a function of the vibrational frequencies of the minerals and the frequency shifts attendant on isotopic substitution (Urey 1947; Bigeleisen and Mayer 1947; Roginsky 1956; Galimov 1973; O'Neil 1986). It must be emphasized that the underlying cause of isotopic fractionation between two phases is quantum mechanical in origin and resides in the difference in their zero-point energies. Zero-point energy is vibrational energy that molecules possess even at a temperature of absolute zero. Thus, isotope effects (both kinetic and equilibrium) depend on fundamental thermodynamic properties of natural substances and thereby bear on several topics of mineralogical interest. These topics include temperatures of formation, provenance, equilibrium state of mineral assemblages, mechanisms of mineralogical reactions, and crystal structure, among others.

Definitions

The c5- Value. Stable isotope compositions are reported in delta (c5) notation rather than as absolute ratios. The c5-value is defined as the relative difference

8.5.1 Theoretical Aspects of Isotopic Fractionation 389

between the isotopic ratio of interest in substance x to that same ratio in a standard (std), expressed in parts per thousand (per mil):

bx = [(Rx - RS,d)/Rs,d]103, (1)

where Rx = D/H, 13C;12C, 1SN;t4N, 1SO/160, or 34S/32S (ratios of the abundance of the rare isotope to the common isotope) and RSld is the same ratio in the standard. Solid, liquid, and gas reference standards are available upon request from the International Atomic Energy Agency in Vienna. The reference standards employed depend on the element of interest and sometimes on the nature of the sample analyzed. For example, SMOW (Standard Mean Ocean Water) is a water standard used in reporting bD and b1S0 values of minerals and fluids, PDB (Peedee Belemnite) is a carbonate standard used to report b13C values of any carbonaceous material but also b1SO values oflow-temperature carbonates, and CDT (Canon Diablo Troilite) is a meteoritic sulfide used to report b34S values. A b34S value of 20.2 per mil means that the 34S/32S ratio of the sample is 20.2 per mil (parts per thousand) higher than that of CDT. Similarly a bD value of - 65.4 means that this sample contains 65.4 per mil less deuterium than SMOW.

The Fractionation Factor (cx). The fractionation factor (separation factor) of the isotopic ratio R between two phases a and b is:

cxa- b = Ra/Rb· (2)

In terms of b-values, the fractionation factor is:

CXa-b = (1 + ba 10-3)/(1 + bb 10- 3). (3)

For reasons of convenience and theoretical considerations discussed below, the fractionation factor is commonly expressed in per mil by the function 103Incxa_b. Inasmuch as typical values of cx are close to 1, this function is very well approximated by ba - bb when both the individual b-values and the difference between them are less than about 10 per mil:

(4)

Fractionation factors that are experimentally determined or calculated from spectroscopic data are normally presented as 103lncx and this function is the correct or exact expression used in calculating "isotopic temperatures". The A function is frequently used as an approximation but is exact within experimental error under the conditions specified above.

Isotope Exchange Reactions. Isotope exchange reactions are reactions in which there is no change in chemistry except for a redistribution of isotopes between two phases. The reactions are normally written such that one atom is exchanged, as in the oxygen isotope exchange reaction between quartz and water:

1/2SP60 2 + H2180 = 1/2Sps02 + H/60. (5)


If the isotopes are randomly distributed among all equivalent positions in the molecules or minerals, the equilibrium constant K for such a reaction is equal to the stable isotope fractionation factor which, in this case, is eSo;t60)quartz/(1S0/160)water' In the general case, Q( = K lin where n is the number of atoms exchanged in the reaction as written. Thus stable isotope fractionation factors are true equilibrium constants and, as such, are functions of temperature. This is the basis for their use in geothermometry.

Kinetic Isotope Effects

Kinetic isotope effects are expressed as a ratio of rate constants, k/k*, for reactions involving molecules containing the normal and "heavy" isotope (*), respectively. The theory for these effects was developed within the "transition state theory" of Eyring. It is assumed that chemical reactions proceed through molecules in an "activated complex" or "transition state" and that these molecules are in thermal equilibrium with the reactant molecules. The equilibrium or steady state concentration of activated complexes is calculated by the methods of statistical mechanics. The equations reduce to functions of temperature, symmetry numbers, and vibrational frequencies. The necessary molecular vibrational frequencies are obtained from infrared and Raman spectra of the respective compounds. Except for simple cases, this theory is not directly applicable to most mineralogical systems. Of more concern to earth scientists is understanding the role of kinetic processes in mineralogical reactions and recognizing these processes in the isotopic composition of natural minerals.

Kinetic isotope effects are common in nature and their magnitudes are comparable to and sometimes significantly larger than those of equilibrium isotope effects. They are associated with fast, incomplete, or unidirectional processes like diffusion, evaporation, and dissociation reactions. Kinetic isotope effects in diffusion and evaporation arise from differences in translational velocities between different isotopic forms of ions or molecules passing through a phase or across a phase boundary. During a diffusion process, isotopically light molecules preferentially diffuse away leaving the reservoir enriched in the heavy isotope. Isotope effects associated with redox and dissociation reactions occur because chemical bonds containing the heavy isotope are more stable than those containing the light isotope. That is, it requires less energy to break bonds like 32S-O and 12C-O than to break bonds like 34S_O and 13C_0. Extremely large kinetic isotope effects are seen in some biologically mediated reactions such as the bacterial reduction of sulfates. Sulfide minerals that are products of bacterial reduction of sulfate are readily identified by their low sulfur isotope ratios.


Thermodynamic (Equilibrium) Isotope Effects

The inequality of physical and chemical properties of isotopically substituted molecules is manifested not only in kinetic effects but also in equilibrium effects that result from differences in free energies of the light and heavy molecules. During equilibrium mineralogical reactions, the light stable isotopes of the elements common to several phases will be distributed among these phases in such a say as to achieve the minimum free energy. In so doing the resultant minerals will all have different stable isotope ratios.

Consider a generalized isotope exchange reaction of element X between two ideal gases AX and BX, where X* denotes the heavy isotope:

AX + BX* = AX* + BX. (6)

The equilibrium constant for this reaction can be written in terms of partition functions (Q) of the reactants and products:

(7)

Using the methods of quantum statistical mechanics and some relatively minor assumptions, it is possible to express these isotopic partition function ratios in terms of the fundamental vibrational frequencies of the isotopic forms of these compounds as follows:

QAX* = ~ n v; exp( - hv; /2kT) [1 - exp( - hVi/kT)] QAX s* Vi exp( - hvd2kT) [1 - exp( - hv; /kT)] ,

(8)

where s is the symmetry number, h is Planck's constant, T is absolute temperature, and the Vi are the fundamental vibrational frequencies of the compounds. The expression under the product symbol is designated as the p-factor such that

P _ SQAX* AX -'---Q .

s AX (9)

/

As stated above, the equilibrium constant K can be expressed in terms of the activities of the reactant and products in the usual way, and the K for isotope exchange reactions is equal to the fractionation factor:

K = [AX*] [BX] = RAX [AX] [BX*] RBX '

Hence

(10)

(11)

If the element enters into the mineral structure in more than one site as, for example, oxygen atoms in -OH groups and -S04 groups in the mineral alunite [KAI3 (S04h(OH6 ], the isotope exchange reaction takes a more general form:

mAXn + nBX;;' = mAX~ + nBXm • (12)


Actually, within the framework of this reaction, a set of reactions of the individual isotopic forms takes place:

(13)

Each of the monosubstituted forms is characterized by corresponding values of the p-factor. It has been shown that the p-factor that characterizes the whole substance is an arithmetical mean of all Pi-factors of a compound,

(14)

and the separation of isotopes between two different compounds is determined by the relation

(15)

Recall that the theory of isotopic fractionation was developed for ideal gases in the harmonic oscillator approximation. There are major obstacles to overcome in the application of this theory to condensed phases.

Effect of Temperature

The extent to which temperature affects the vibrational energies (or partition functions) of two substances is a prime factor in determining the isotopic fractionation between them. At very low temperatures, molecules occupy the lowest (zeroth) vibrational energy levels and at very high temperatures many or all of the energy levels are occupied. Thus fractionation factors of isotopic ratios between two substances are largest at low temperatures and approach a value of unity at high temperatures. Only very small fractionations of stable isotopes are observed between minerals that formed at very high temperatures, as for example, phenocrysts in volcanic rocks and ultramafic minerals in mantle nodules.

The temperature dependence of isotopic fractionations can be complicated but, for mineral systems, In a normally varies as liT in the low temperature limit and as 1/T2 in the high temperature limit. What constitutes low and high temperatures depends on the minerals but, in general terms, low temperatures are surficial temperatures and high temperatures are igneous temperatures. The reason why Ina varies as 1/T2 in the high temperature limit (rather than the normal liT dependence characteristic of chemical reactions) is that AH for isotope exchange reactions decreases with increasing temperature whereas AH for normal chemical reactions varies little with temperature. The sign of the isotopic fractionation can change with an increase in temperature and such changes are equilibrium phenomena called "crossovers". The per mil fractionation will always go to zero (a = 1), however, as the temperature approaches infinity.

The temperature dependence of the oxygen isotope fractionation factors between quartz and other common rock-forming minerals was calculated by


TEMPERATURE (K ) Fig. 114. Calculated fraction-00 co ation factors between quartz 000 0 0 0 0_0 0 0 0 Ol and the minerals indicated. N_Ol r-- ~ CO) N

19 The mineral abbreviations are 18 as follows: QTZ quartz;

::J 16 CALC calcite; ALBT albite;

« Muse muscovite; ANOR an-0:: 14 orthite; DIOP diopside; FORS w forsterite; and RUTL rutile. Z 12 ~ (After Kieffer 1982)

N 10 I-g 8 0

6 £; 0 4 0 0

2

0 -1

0 2 4 6 8 10 12 106 T-2

Kieffer (1982) and is shown in Fig. 114. While not in total agreement with those determined experimentally (e.g., Chiba et al. 1989), such calculations offer the only possibility for obtaining needed fractionation factors for some important systems, particularly at low temperatures where exchange rates are sluggish in the laboratory.

Intramolecular Isotope Effects (Internal Thermometers)

Atoms in nonequivalent sites 1 and 2 of a given mineral are associated with bonds of different vibrational frequencies and thus of different pi-factors. This means that, in the equilibrium state at its formation temperature, the atoms in the nonequivalent sites have different isotopic compositions. This is an equilibrium intramolecular isotope effect that is characterized by the relation between the respective Pi-factors:

O(i(1-2) = Pi(l)/Pi(2). (16)

Inasmuch as 0( is a function of temperature, there is the possibility of developing internal isotope thermometers if there are analytical procedures to separate the element from the two sites for isotopic analysis. The development of internal oxygen isotope thermometers is an active field of research at the present time.

Effect of Pressure

Because the change in molar volumes of solids on isotopic substitution is small, typically hundredths to tenths of a percent, it is generally assumed that the effect


of pressure on isotopic fractionation between solid mineral phases is negligible. Insufficient knowledge of the pressure that prevailed during geological processes often places constraints on the applicability of the chemical and physical thermometers used by petrologists. In this respect, stable isotope fractionations are particularly useful in geothermometry because of the apparent lack of dependence on pressure.

The few experiments that have been made to test these effects support the idea of negligible pressure effects but were made only up to as high as 20 kbar and for only one system, the oxygen isotope fractionation between CaC03 and H 20 (Clayton et al. 1975). It is possible that pressure effects may be significant for mineral-gas or melt-gas systems because the isotopic properties of the gas might change drastically with big density changes while those of the melts or minerals remain unaffected.

Polyakov and Kharlashina (1989) suggested the following expression to account for the effect of pressure in the quasi-harmonic approximation:

(oPloPh = ~: (oPloT)v,

where

Li [YiJ.Licoth (0.5J.Li) - ytcoth(0.5J.L;)J Y = Li [J.Licoth(0.5J.LJ - J.Ltcoth(0.5J.L;)J '

(17)

(18)

and J.Li = hvJkT with h = Planck's constant and k = the Boltzmann constant, Yi = (0 lnvi/o In Vh, the Gruneisen parameter for the lattice vibration of frequency

Vi and BT = - V(oP I oVh, the isothermal bulk modulus. It has been found by this approach, for instance, that at equilibrium diamond is isotopically lighter than graphite at any pressure higher than 20 kbar. This is in contrast to previous calculations of the direction of this isotopic fractionation that were made without taking pressure into consideration.

Effect of Chemical Composition

The isotopic properties of a mineral depend most importantly on the nature of the chemical bonds within the mineral. The oxidation state, ionic charge, atomic mass and electronic configuration of the elements to which the isotope is bonded need to be considered. In general, bonds to ions with a high ionic potential and l~w atomic mass are associated with high vibrational frequencies and have a tendency to incorporate the heavy isotope preferentially in order to lower the free energy of the system. For example, consider the difference between the bonding of oxygen atoms to the small, highly charged Si4 + ion as opposed to the relatively large Fe2 + ion. In common natural equilibrium assemblages, quartz is always the most 180-rich mineral and magnetite is always the most 180_ deficient mineral and this relation has been confirmed by laboratory experiments.


Suzuoki and Epstein (1976) found through laboratory experiments that the chemical composition of the octahedral site in hydrous minerals is the dominant factor controlling their relative hydrogen isotope compositions. They reported the following relation for micas and amphiboles:

ilD(mineral-water) = - 22.4(106T- 2 ) + 28.2

(19)

where ilD is the per mil fractionation of deuterium between the mineral and water, and X is the mole fraction of the cations of sixfold coordination. Note the strong dependence of the fractionation on the amount of Fe in the mineral.

Effect of Crystal Structure

Crystal structure can inflm:nce the isotopic properties of minerals to an extent depending on how different the interatomic interactions are within the various structural forms. Structural effects are secondary in importance to those arising from the primary chemical bonding. On the basis of limited experiments and calculations, it would appear that the heavy isotope concentrates in the more closely packed or well-ordered structures. Thus D and 180 concentrate in ice relative to water, 180 and 13C concentrate in aragonite relative to calcite, and so on. One of the largest calculated effects of structure is found in the carbon system. Bottinga (1969) calculated carbon isotope fractionations between diamond and graphite of 11.5 permil at 0 °C and 0.4 per mil at 1000 0c. The direction of these fractionations reverse at pressures higher than 20 kbars according to the model of Polyakov and Kharlashina (1989). The resolution of this possible discrepancy awaits laboratory experimentation at high pressures.

References

Bigeleisen J, Mayer MG (1947) Calculation of equilibrium constants for isotopic exchange reactions. J Chern Phys 13: 261-267

Bottinga Y (1969) Carbon isotope fractionation between graphite, diamond, and carbon dioxide. Earth Planet Sci Lett 5: 301-307

Chiba H, Chacko T, Clayton RN, Goldsmith JR (1989) Oxygen isotope fractionations involving diopside, forsterite, magnetite, and calcite: application to geothermometry. Geochim Cosmochim Acta 53: 2985-2995

Clayton RN, Goldsmith JR, Karel KJ, Mayeda TK, Newton RC (1975) Limits on the effect of pressure on isotopic fractionation. Geochim Cosmochim Acta 39: 1107-1201

Galimov EM (1973) Izotopy Ugleroda v Neftegazovoy Geologii (Carbon isotopes in oil-gas geology) Nedra, Moscow, 384 pp

Kieffer SW (1982) Thermodynamics and lattice vibrations of minerals: 5. Applications to phase equilibria, isotopic fractionations, and high pressure thermodynamic properties. Rev Geophys Space Phys 20: 827-849

O'Neil JR (1986) Theoretical and experimental aspects of isotopic fractionation. In: Valley JW, Taylor HP Jr, O'Neil JR (eds) Stable isotopes in high temperature geological processes. Rev Mineral 16: 1-40


Polyakov VB, Kharlashina NN (1989) The effect of pressure on the equilibrium isotopic fractionation in solids. In: Wand U, Strauch G (eds) Isotopes in Nature. Fifth Workshop Meeting, Central Institute of Isotope and Radiation Research, Leipzig, 735~ 745

Roginsky SZ (1956) Teoreticheskie Osnovy Izotopnykh Metodov Izucheniya Khimicheskikh Reaktsy (Theoretical principles of isotopic methods for investigating chemical reactions) Academy of Sciences USSR Press, Moscow, 614 pp

Suzuoki, Epstein S (1976) Hydrogen isotope fractionation between OH-bearing minerals and water. Geochim Cosmochim Acta 40: 1229~1240

Urey HC (1947) The thermodynamic properties of isotopic substances. J Chern Soc Lond 562~581

8.5.2 Natural Variations in Stable Isotopes

J. HOEFS and V.1. VINOGRADOV

The distinction between stable isotopes and radiogenic isotopes is rather conventional, as many isotopes formed through radioactive decay processes are also stable. However, the processes of fractionation of both groups of isotopes are very different and therefore this division is practically convenient. In addition, the light elements are characterized by similar methods of isotopic measurement; the elements in question are measured in gaseous form.

Only some light elements undergo isotopic fractionation under natural conditions. First of all these are H, C, N, 0, and S. Si isotopes are fractionated to a smaller degree and have not been studied intensively up to now.

The common reason for stable isotope fractionations is the difference in their atomic masses. As a rule molecules containing the heavy isotope have higher bond energies than those containing the light isotope. Therefore the former are more stable than the latter. Thus in unidirectional chemical and physical processes proceeding in the liquid or gaseous phase the product of the process (reaction) is enriched in light isotopes relative to the starting material. Isotopic fractionations during unidirectional processes are called kinetic isotope effects. For some elements (sulfur and carbon for instance) kinetic isotope effects are of great geochemical significance.

The other type of isotope fractionations are thermodynamic (equilibrium) isotope effects. In terms of quantum mechanics the internal energy of a molecule is the sum of the different forms of energy: translation, rotation, vibration, and others. The statistical sum of internal energies is different for isotopic molecules, differences in vibrational energies being the most important ones. So when a chemical system is in an equilibrium state the different phases have different isotopic compositions corresponding to the fractionation factor !Y.. This coefficient describes the isotopic ratio in one substance relative to another. For example, in the case of isotopic equilibrium in the system CO2-H20

CO2-H2 0 = eSO/160)C02/eSO/160)H20 = 1.0412 ± 0.0001

(T = 25 QC).

8.5.2 Natural Variations in Stable Isotopes 397

This means that the equilibrium isotopic composition of CO2-oxygen is by some 40 per mil heavier than H 20-oxygen. With increasing temperatures, the difference in isotopic composition decreases.

The temperature dependence of fractionation factors is the basis for the study of paleotemperatures. In general, this dependence is proportional to liT at low temperatures (lower than room temperature) and 1/T2 at high temperatures. Real dependences are more complicated and sometimes even the sign of the fractionation factor can change with changing temperatures.

Isotopic equilibration between minerals, which can be achieved during melting or through the exchange with fluids, also leads to isotope fractionation. In magmatic and metamorphic rocks, for instance, the definite sequence of minerals with respect to its oxygen isotopic composition is: quartz, potassium feldspar, albite, anorthite, muscovite, amphibole, biotite, and magnetite. Quartz is the most 180 enriched mineral, magnetite is the most 180 depleted mineral. In general, bonds of ions with a high ionic potential and low atomic mass have a tendency to incorporate the heavy isotope preferentially. The mineral pair quartz-magnetite is thus the most suitable as a geothermometer.

The progressive replacement of Si-O bonds leads to a small but noticeable oxygen isotope effect. There are other examples of how the chemical composition of mineral influences the isotopic composition. Pressure has no essential effect on the isotopic fractionation. In some cases, isotope effects depend on the crystalline structure of mineral. Nevertheless, the main natural isotope effects are connected with geochemical processes of mineral origin rather than with mineral properties.

The precise measurement of absolute isotope ratios is very difficult. It is easier to measure the relative difference in isotopic ratios between two samples. That is why there are special standards for every element, and all samples are compared relative to them. The difference in isotopic ratios between standard and sample is expressed as <5%0 values:

<5%0 = ( R sample _ 1) *1000 R standard '

where R sample and R standard are the ratios of heavy and light isotopes in sample and standard accordingly. Positive <5 values mean the enrichment of the sample in the heavy isotope and negative values mean the depletion of the sample in the heavy isotope relative to the standard.

If the difference in isotopic composition of two samples (A and B) corresponds to isotope equilibrium then

A + 1000 A - B = B + 1000 ~ <5A - <5B ~ 1000 In OCA-B

References

Hoefs J (1987) Stable isotope geochemistry. Springer, Berlin Heidelberg New York Valley JW, Taylor HP, O'Neil JR Jr (eds) (1986) Stable isotopes in high temperature geological

processes. Rev Mineral 16: 1-570


8.5.3 Oxygen and Hydrogen Isotopes in Mineralogy

B.G. POKROVSKyand T.F. ANDERSON

Variations in the isotopic composition of oxygen have been applied to a range of problems in various branches of geology. The following problems are of mineralogical interest: (1) the temperatures of formation of igneous, sedimentary, metamorphic, and hydrothermal rocks; (2) the attainment of equilibrium in mineral assemblages; (3) characterization of the important source reservoirs in the lithosphere and hydrosphere. Hydrogen isotopes have a more limited range of application. They are, however, indispensable as indicators of the genesis of natural waters and hydroxyl-bearing minerals.

lS0j160 and D/H ratios are expressed in parts per thousand (permil) deviation from the SMOW (Standard Mean Ocean Water) standard:

!5 1S0 = [eSO/160)samPle - 1J x 103

e SO/160)SMOW

!5D = [(D/H)samPle - 1J x 103 •

(D/HhMow

The difference in isotopic composition between two substances can be expressed as Ll a- b = !5a - !5b. The oxygen isotopic composition of sedimentary carbonate minerals is often reported relative to the PDB standard. Conversion between the PDB and SMOW reference scales is:

!51S0SMOW = 1.03086·!51S0PDB + 30.86

In studies of the oxygen isotope chemistry of extraterrestrial materials, the 170/160 ratio is analyzed as well as the lSO/160 ratio. Detailed investigations of these two oxygen isotopic ratios have revealed considerable heterogeneity within certain classes of stony meteorites. This heterogeneity was apparently inherited at the time of solar system formation. However, in lunar and terrestrial rocks, there is a constant proportionality between 170/160 and lSO/160: 15170 = 0.516·!51S0. Therefore, the determination of 15 170, which is more difficult

technically, does not provide any additional information. Isotope exchange reactions in mineral-water systems are of great import

ance in oxygen isotope geochemistry. These reactions have been investigated experimentally and theoretically for many rock-forming minerals over a wide range of temperature. Several publications have reviewed the results and present equations which express isotopic fractionation factors for mineral-water and mineral-mineral systems as a function of temperature. The following important features of these relationships should be noted: (1) All rock-forming silicates and carbonates are enriched in lS0 with respect to water at T < 500°C. (2) Isotopic fractionation reaches maximum values (20-40 permil) at low tem-

8.5.3 Oxygen and Hydrogen Isotopes in Mineralogy 399

peratures ( < 50°C) and minimum values (0-2 permil) at high temperatures ( > 800°C).

Oxygen isotopic fractionation between coexisting minerals in equilibrium assemblages is determined primarily by the nature of chemical bonding: oxygen bonded with small, highly charged ions, such as Si4 + and C4 +, is enriched in 180 in comparison to oxygen bonded with large, relatively low-valence ions such as Fe2+ . In accordance with this rule, the relative abundance of 180 in minerals decreases in the following order: quartz, dolomite, anhydrite > alkali feldspars, calcite, aragonite > leucite > muscovite, nepheline, kyanite > glaucophane, staurolite > lawsonite > garnet, common pyroxenes and amphiboles > biotite > olivine, sphene > chlorite > ilmenite > rutile > magnetite, hematite > pyrochlore. Those pairs of minerals with the greatest contrast in 180 abundance are the most sensitive isotopic thermometers, e.g., quartz-magnetite, quartz-biotite, etc. In the majority of granitoids A 180 (quartz-biotite) = + 5 ± 1, and A 180 (quartz-magnetite) = + 8 ± 1; these fractionations correspond to temperatures of 500-600°C.

The factors controlling hydrogen isotope fractionation in mineral-water systems are not as well constrained. The proportions of Fe, Mg, and Al ions in sixfold coordination sites appears to be the most important among them. Ferich hydroxyl minerals are depleted in D relative to Mg- or AI-rich minerals in equilibrium assemblages. In igneous and metamorphic rocks, <5D decreases in the order: muscovite> amphibole> biotite. AD (muscovite-amphibole) and AD (amphibole-biotite) commonly does not exceed 10-15 permil. In sedimentary rocks deposited in seawater (t5D = 0) and not subjected to deep burial (T < 50°C), Fe-rich clay minerals such as annite and nontronite have t5D ~ - 100 permil; whereas AI-rich clay minerals such as gibbsite and kaolinite have

t5D ~ 0 to - 30 permil. However, because the temperature "dependence of hydrogen isotopic fractionation is very complicated for most minerals, it is virtually impossible to use hydrogen isotope variations for paleotemperature estimates. A significant influence of crystal structure or pressure on oxygen and hydrogen isotopic compositions of minerals has not been established.

Nonequilibrium isotopic fractionations are often encountered in mineral assemblages. Lack of isotopic equilibrium is usually attributed to differences between minerals in the kinetics of isotopic exchange with fluids. Nonequilibrium effects are typical for volcanic rocks which have interacted with hydrothermal waters under subsolidus conditions.

The isotopic composition of minerals depend not only on their chemical composition and temperature of formation, but also on the isotopic composition of the reservoir(s) with which they equilibrated during formation. Based on isotopic characteristics, three major source reservoirs may be distinguished: (1) the mantle: 15 180 = 6 ± 1, t5D = - 80 ± 5; (2) the hydroshpere: 15 180 ~ 0 to - 50, t5D ~ 0 to - 450; (3) chemogenic and biogenic sedimentary rocks: 15 180 ~ 20 to 40, t5D ~ - 30 to - 90. Mixing and interaction between these reservoirs affects the majority of all types of rocks in the Earth's crust.


Mid-ocean ridge basalts (MORB) are unquestionably derived from uppermantle source regions. MORB has a very narrow range of £5 180 = 5.7 ± 0.3. The majority of unaltered mafic and ultramafic rocks, as well as lunar rocks, have a similar composition. Because isotope effects associated with fractional crystallization do not exceed 0.5-1 permil, we can conclude with some certainty that mantle magmas and their derivatives had primary £5 180 values in the range 6 ± 1 permil; the latter is commonly referred to as "normal magmatic".

However, igneous rocks as a whole have a very wide range in £5 180: - 10 to + 17 permil. Only a fraction of granitoids have "normal magmatic" isotopic

compositions. Most have values of > 8 to 10, which clearly indicates assimilation or interaction with sedimentary reservoirs. Analogous 180 enrichments have been found in certain alkaline rocks. Anomalously low £5 180 have been reported in all types of rocks; these depletions are the result of interaction with hydrothermal fluids of surface origin [9]. Low £5D values (~ - 100) are common in 180-depleted rocks.

Surface waters are also dominant in most modern hydrothermal systems. This is demonstrated by the fact that the £5D in hydrothermal fluids corresponds to the £5D of local meteoric water or of seawater. Usually, £5 180 values of hydrothermal fluids are higher than those of the "source" water. This shift is due to oxygen isotope exchange between the fluid and the country rocks through which it circulates; the extent ofthe shift depends upon the water/rock ratio, and therefore can vary over a wide range. Accordingly, the oxygen isotopic composition of minerals which precipitated in equilibrium with hydrothermal fluids can also vary widely.

Together with meteoric waters and seawater, "metamorphic" waters released during the diagenetic and metamorphic dehydration are also involved in hydrothermal processes. £5 180 values of metamorphic waters may range from + 5 to + 25 permil, reflecting the oxygen isotopic composition of the initial

rocks and the temperature of metamorphism. The range in £5D values in metamorphic rocks, and therefore in metamorphic waters, coincides with the typical range for clay minerals: from - 30 to - 90 permil. The hydrogen isotopic composition of deep-seated igneous rocks are close to that of metamorphic rocks. Hence, water dissolved in magma is probably of secondary origin in most cases. The existence of "juvenile" water, fluids that have never been involved with the hydrosphere, is still open to debate.

The oxygen isotopic composition of metamorphic rocks depends on the original composition of the proto lith as well as the temperature and fluid regime of metamorphism. Generally, metasedimentary rocks have higher £5 180 values than meta-igneous rocks. As the temperature of metamorphism increases, £5 180 values decrease from ~ + 10 to + 15 in greenschists to ~ + 7 to + 9 in granulites. 180 depletion at low temperatures is probably related to water-rock interactions; at high temperatures, hydrous fluids in equilibrium with silicate magma are apparently dominant. However, metamorphic rocks with £5 180 values in the "normal magmatic" range are encountered only rarely, even in the oldest segments of continental crust.

8.5.4 Carbon Isotopes in Mineralogy and Geochemistry 401

References

Clayton RN (1986) High temperature isotope effects in the early solar system. Rev Mineral 16: 129-164

Friedman I, O'Neil JR (1977) Compilation of stable isotope fractionation factors of geochemical interest. In: Fleischer M (ed) Data of geochemistry, 6th edn. US Gov Printing Office, Washington, DC

Javoy M (1977) Stable isotopes and geothermometry. J Geol Soc Lond 133: 609-636 Longstaffe FJ (1979) The oxygen isotope geochemistry of archean granitoids. In: Barber F (ed)

Trondhjemites, dacites and related rocks. Developments in petrology, 6. Elsevier, Amsterdam, pp 363-399

O'Neil JR (1979) Stable isotope geochemistry of rocks and minerals. In: Jager E, Hunziker JC (eds) Lectures in isotope geology. Springer, Berlin Heidelberg New York

O'Neil JR (1986) Theoretical and experimental aspects of isotopic fractionation. Rev Mineral 16: 1-40

Pokrovsky BG, Vinogradov VI (1990) Isotope investigations of alkali rocks of Middle and Western Siberia. In: Schukoliukov IU (ed) Isotope geochemistry and kosmochemistry. Nauka, pp 144-159

Taran YuA, Pokrovsky BG, Glavatskikh SF (1987) Condition of hydrothermal alteration of rocks of the Mutnovskaya geothermal system deduced from isotopic data. Geochimia 11: 1569-1579

Taylor HP (1977) Water-rock interactions and origin of H 20 in granitic batholites. J Geol Soc Lond 133: 509-558

Taylor HP Jr, Sheppard SMF (1986) Igneous rocks: Processes of isotopic fractionation and isotope systematics. Rev Mineral 16: 227-272

8.5.4 Carbon Isotopes in Mineralogy and Geochemistry

E.M. GALIMovand D. RUMBLE

The fact that carbon consists of two stable isotopes, 13C and 12C, was discovered in 1929. The 13C/12C ratio varies from 0.0101 to 0.0116 in nature. Variations of 15 parts in 10000 comprise the whole information content of carbon isotope geochemistry. The precision of measurement of isotopic ratios with a modern gas source mass spectrometer approaches one part in 100000. Therefore, even small details of the origin and environment of formation of natural carbonaceous materials can be interpreted. It is customary to report carbon isotope ratio measurements in relation to a standard rather than to express them as absolute ratios. The accepted unit of measure is defined as c5 13C = [(13C/12C)sample - e3C/12C)standard]/(13C/12C)standard multiplied by 1000, i.e.,

the deviation of the sample ratio from the standard ratio in parts per thousand or permil. The internationally adopted standard is Pee Dee Belemnite, abbreviated PDB, with the absolute ratio 13C/12C = 0.0112372.


Extraterrestrial Carbon

The first extraterrestrial carbon studied isotopically was graphite from the iron meteorite Canyon Diablo. Its isotopic composition was found to be indistinguishable from terrestrial carbon. However, subsequent, more comprehensive research showed significant variations in the b13C values of meteoritic carbon that correlate with the types of meteorites and the mineralogy of carbon bearing compounds. For carbonaceous chondrites of CI and CM types, total carbon is characterized by b13C values from - 5.6 to - 11.6%0 with carbonate carbon from + 41.6 to + 70.2%0 and polymerized organic carbon from - 17.8 to - 24.1%0. Ordinary chondrites, the most common type of meteorite, contain

carbon in low concentrations with b13C in the range - 22 to - 29%0. Unequilibrated ordinary chondrites have b13C in a wider range of - 11 to - 28%0. The different minerals of iron meteorites have different carbon isotope

ratios. Graphite has b13C from - 1 to 10%0, while cohenite and taenite vary from - 15 to - 24%0.

All of these values, except for carbonate carbon in CI carbonaceous chondrites, are within the range of b13C variations for terrestrial carbon. Recently, however, exceptionally high values of b13C, far beyond the range of earth samples, have been discovered in several carbonaceous chondrites. A fraction of the Murchison meteorite, insoluble in the strong acids HF and HCI, yielded b13C of + 1100 to + 1400%0. These results are believed to indicate that some meteorites preserve a small amount of pre-solar material. Carbon anomalously enriched in 13C may originate from interstellar grains injected into space by outburst of red giant stars.

Basaltic and anorthositic lunar rocks collected during Apollo missions and by Luna spacecrafts show b13C of - 20 to - 30%0. These values are the same as for comparable rocks from Earth. The fine-grained lunar soil, or regolith, has b13C from + 2 to + 20%0. The enrichment of the regolith in 13C is believed to be due to implantation of carbon atoms in the Moon's surface by the solar wind. Isotope fractionation during desorption from the lunar surface may also contribute to the observed isotope effect.

Primary Carbon of the Earth

Carbon released from the mantle during degasing is expected to concentrate in the Earth's crust. Therefore, the average isotopic composition of crustal carbon should coincide with the isotopic composition of primary terrestrial carbon. Geochemical calculations show that the average <5 13C value of crustal carbon is within the range - 3 to - 8%0, with a most probable value of - 5%0. Mantle peridotites, in contrast have b13C of - 20 to - 30%0. It is suggested that primary terrestrial carbon originated from carbonaceous chondrites, whereas residual, finely dispersed carbon found in mantle rocks is a counterpart of the carbon preserved in ordinary chondrites.

8.5.4 Carbon Isotopes in Mineralogy and Geochemistry

Equilibrium of CO2-HC03-CO;- in the Atmosphere, Hydrosphere, and Sediments

403

Isotopic exchanges of 13C and 12C between CO2 and the bicarbonate and carbonate ions is one of two major processes controlling carbon isotope fractionation at the Earth's surface. The other controlling process is biological isotope fractionation (Fig. 115).

Under conditions of thermodynamic equilibrium at low temperature, bicarbonate and carbonate ions are enriched in 6 13C by + 8 to + 10%0 with respect to CO2 (gas or aqueous). Bicarbonate is the dominant carbon compound in ocean water. Its average isotopic composition is approximately + 1%0. This value depends on the proportion of reduced and oxidized forms of carbon buried during sedimentation. Marine carbonates have 613C within the range of - 2 to + 4%0. High 613C values are found in marine carbonates deposited

during episodes of enhanced deep sea burial of organic-rich sediments during the Miocene, Cretaceous, and Jurassic. High values are also known from marine carbonates deposited during the major coal accumulation of the PermoCarboniferous. There is a sharp drop in 6 13C in marine carbonates at the Cretaceous-Tertiary. The relationship between biogenic carbonate enriched in 13C and biogenic reduced carbon depleted in 13C was established at the Earth's surface at least 3.5 billion years ago.

CO2 of atmosphere _ CO2 Technogenic

_ C4

CO2

=~~;::-:-1'7-:r-:---:JPi'ainiktioin~lI!Ip'a~ pea~t CH 4

\. ~:';:~l:of CH4 /' OOles ~dimentarv

r:ks \ CH4 ~crease in extent

Diamonds

of metamorphism Coal Petroleum /' of organic matter

Graphite -...II!!III-'CH4

Fig. 115. Range of variations of the isotopic composition of carbon in some natural substances


The oceans regulate the concentration and isotopic composition of CO2 in the atmosphere. The average (j 13C of atmospheric CO2 is - 7%0, in equilibrium with oceanic bicarbonate. Study of air trapped in polar ice shows that at the end of the last glaciation, CO2 had (j 13C 1.1%0 higher than at present. The change in (j 13C of atmospheric CO2 may be due in part to the influx of isotopically light CO2 ( - 26%0) produced by the combustion of wood, coal, and petroleum during the following industrialization.

Biological Isotope Fractionation

Two aspects must be distinguished in the biological fractionation of isotopes. The first of these is the partitioning of isotopes between an organism and its environment. The second aspect is the separation of isotopes during metabolic processes within an organism (Fig. 116).

The carbon of living organisms is depleted in 13C in relation to the external CO2 from which they obtain carbon for growth. This fact was discovered in 1939 by Nier and Gulbransen. It is noteworthy that in 1924, y.1. Vernadskii proposed that living organisms could separate isotopes, long before biological isotopic fractionation was measured in nature.

There is carbon isotope fractionation accompanying the photo-synthetic assimilation of CO2 by plants. Park and Epstein recognized that this effect occurred in the course of enzymatic carboxylation of ribulose diphospha.te, the initial reaction of photosynthesis in the Calvin cycle. As a result, land plants have (j 13C from - 23 to - 28%0. Later, it was shown that this range of values is valid only for plants ofC-3 type. Plants ofC-4 type have a (j 13C range of - 10 to - 18%0. The C-4 plants photosynthesize through the Hatch-Slack cycle,

Plant

,.,-,+-_Ph_otosynthe_s_is--f>o<_,

isotope effect

~ rl.-.-.-.-.---"-"~-.:"';I Proteins

Lipids lei ••• I~ Lignin ~~'AM/ ~ ~ Carbohydrates

I r I I -35 -25 -15

Fig. 116. Biological fractionation of carbon isotopes

HC03


where carboxylation of phospho enol phyruvate is the initial step. Aquatic plants are depleted in 13C to a lesser extent than land plants. A range of - 16 to - 21%0 is typical for marine plankton.

The shift in (j 13C values between C-3 and C-4 plants has been exploited in paleoclimatological studies to document changing proportions of plant types in biomass. Soil carbonates from northern Pakistan changed in (j 13C from - 10%0 to + 12%0 upon the onset of the Asian monsoon some 7 million years ago, recording the shift from C-3 trees and shrubs to C-4 grasslands.

The separation of isotopes during metabolic processes within an organism is exemplified by the discovery that organisms are not isotopically homogeneous. Lipids are depleted in 13C by 2 to 7%0 compared to the total carbon of the corresponding organism. Carbohydrates and proteins are enriched in 13C by 3 to 5%0, while lignin is depleted almost as much as lipids. Experiments show that different fragments of the same biocompound may have different (j13C values. It may be seen that biological isotope fractionation cannot be reduced to a single step in a photosynthesis cycle.

Galimov (1985) argues that isotope fractionation is characteristic of all enzymatic reactions in organisms. The isotope effect of enzymatic reactions is controlled by two types of fractionation processes: kinetic and thermodynamic. The predominance of one process vs. the other is determined by the relative magnitudes of their rate constants. A number of enzymatic reactions show kinetic isotope effects. One of the most important reactions of this type is the carboxylation of ribulose diphosphate during photosynthetic fixation of carbon dioxide. In many cases, however, thermodynamic isotopes effects predominate. It may be shown that the presence of thermodynamic isotope effects in a system where numerous reactions proceed resulted in a thermodynamically ordered distribution of isotopes between components of the system. The efficacy of enzymatic catalysis to promote isotope exchange equilibrium in living organisms is emphasized by comparison to the inorganic realm where isotope equilibrium between complex compound is not achieved or achieved only under higher temperature.

Diagenetic Transformation of Organic Matter in Sediments

Biogenic, carbon-bearing compounds undergo transformations after burial. Biopolymers are degraded to monomers and simple compounds such as CO2

and CH4 • Repolymerization has also been observed where monomers give rise to "geopolymers" such as fulvic and humic acids and, ultimately, kerogen (Fig. 117). Recently, it has been recognized that repolymerization is accompanied by isotope fractionation that depletes fossil organic matter in 13c. Fulvic acid, humic acid, and kerogen have (j 13C values of - 21, - 22, and - 24%0, respectively, in comparison to a value of - 20%0 for parental marine plankton. Isotopic fractionation during formation of sapropelic organic matter gives even lower (j 13C values of - 26 to - 30%0.

406

Biopolymers

Monomers

Geopolymers

Petroleum,

gas

-15

!

!

-20 -25


Isotopic effects of

polymerization Isotopic effect depending on intramolecular distribution of isotopes

~r::=~--~CH~ co,----"'\

I Graphite

Fig. 117. The change of carbon isotopic composition in the course of transformation of biopolymers into fossil forms of organic matter

Coalification

During coalification, the concentration of reduced carbon increases from 60-70% to 90-98 % in the succession from lignite to hard coal to anthracite. The () 13C values of coals correspond to the range of values of land plants. There is no appreciable difference between the {) 13C values of coals of different rank. The isotope fractionation observed during biodegradation and repolymerization of organic matter in marine sedimentary basins does not occur to the same extent in the formation of coal.


Oil

The carbon isotope composition of oil ranges from - 23 to - 35%0, within the range of biogenic reduced carbon. Strangely enough, however, hydrocarbons in fluid inclusions from igneous rocks have the same range of values. The actual relationships between oil and its biological precursors becomes clear, however, when the b13e values of oil components are compared to corresponding molecules in their precursors. Normal alkanes are the component in oil most depleted in 13e, corresponding to aliphatic structures as the most depleted molecules in biological material. Polar components in oil are enriched in 13e, as are polar biomolecules. The isotopic compositions of biomarkers such as porphyrins and isoprenoids stand in the same relation with other oil components as they do in matter from living organisms. It has been shown that the isotope distribution between fractions of different polarity in oil (hydrocarbons, resins, and asphaltenes) depends on the facies type of parental organic matter, whether sapropelic or humic. Such isotope distribution patterns have been successfully used to identify oil source rocks and are important in exploration for new petroleum resources.

Gas

Methane is the naturally occurring carbon compound most strongly depleted in 13C. Microbiological methane in the upper part of sediments and in marshes has b13e from - 95 to - 60%0. There is a correlation between the amount of 13e depletion and the increase in thermal maturation of source organic matter in sediments (i.e. vitrinite reflectance coefficient Ro). The correlation is due to progressive breakdown of biomolecules. At each stage of maturation, biomolecules with successively higher activation energies for the formation of methane are decomposed. The release of methane by a variety of compounds causes the correlation of b13e with maturation to have a smoother trend than it would if only a single type of carbon structure was involved in the production of methane. Moreover, because the abundance of structural groups with different activation energies of methane formation varies according to facies type, the relation between b13e and Ro is different for sapropelic and humic organic matter. Thus, the pattern of b13e vs. Ro variation makes it possible to identify source rocks for gas.

Metamorphism

The fate of the carbon isotope composItIon of both reduced carbon and carbonate minerals depends on a number oflocal factors during metamorphism. Among the controls on the isotopic composition of reduced carbon are oxygen


fugacity and the extent of graphitization. The b 13C values of carbonate minerals depend on thermal history and the progress of decarbonation reactions.

A number of studies show little change in the b 13C values of reduced carbon in sediments that have been subjected to metamorphism. However, the dispersed carbon in black shales from the Swiss Alps changes from - 25%0 to - 11%0 progressively from unmetamorphosed rocks to the staurolite zone. Coal

samples from the Narrangansett Basin, Rhode Island, USA, change from - 23%0 in the chlorite zone to - 20 in the staurolite and sillimanite zones. In

both these areas, the increase in b 13C is accompanied by increasing graphitization.

Carbonate minerals in calcareous sediments show systematic decreases in b13C of 5 to 10%0 with increasing metamorphic grade during both contact and regional metamorphism. The decrease is caused by the loss of 13C-enriched CO2

as a consequence of thermally driven decarbonation reactions. Isotopic exchange between coexisting reduced carbon and carbonate miner

als is sluggish at the low temperatures of the chlorite through garnet zones of metamorphism (300-450 QC). Isotope exchange equilibrium is approached progressively, however, with increasing metamorphic grade at higher temperatures ( > 450 QC). Sillimanite zone calcite from limestones rich in organic matter is depleted up to 5%0 in 13C from marine limestone values whereas coexisting graphite has b13C values as high as 0 to - 3%0.

Hydrothermal Systems

Both reduced carbon and carbonate minerals show a wide range of b13C values in hydrothermal systems. The variations are caused by local factors including the mixing ofC-bearing waters from different sources, boiling and escape of CO2

near the Earth's surface, thermally driven isotope effects of contact metamorphism, as well as the characteristics of hydrothermal fluids such as temperature, pH, and oxygen fugacity. Despite the wide range in b13C values, the distribution of b 13C values in hydrothermal deposits can generally be attributed to mixing carbon from the crustal reservoirs of biogenic reduced carbon and carbonate. Graphite vein deposits have b13C values from - 28%0 to 6%0. Hydrothermal carbonates range from methanogenic calcites of submarine hot springs with b 13C = - 50%0 to calcite from sediment-hosted base metal deposits with b13C = + 10%0.

Carbon in the Earth's Mantle

Direct analysis of mantle carbon is provided by the study of diamonds. In the early 1950s, Craig (1953) and Wickman (1956) made the first b13C measurements on diamonds and found that the values varied only slightly, averaging - 5%0. It was concluded that the carbon source for diamonds was homogeneous. Recent


studies, however, show that the isotopic composition of diamonds varies from - 34 to + 2%0. A systematic relationship between the ~ 13C values of diamonds

and their mineral paragenesis has been discovered. Diamonds containing mineral inclusions of ultrabasic paragenesis show a narrow range of () 13C from - 9%0 to - 2%0, whereas ~13C of diamonds of the basic (eclogitic) paragenesis

cover the full range of the isotope variations for diamonds (Fig. 118). Two different hypotheses can be considered to explain the ~13C values of

diamonds. In the first of these, the source of carbon with low ~13C values is biogenically reduced carbon in sea-floor sediments injected into the mantle by subduction of oceanic plates.

The alternative hypothesis holds that significant fractionation of carbon isotopes occurs in the upper mantle. Isotopic fractionation under equilibrium conditions at high temperature is small and may have been responsible for a maximum fractionation of 3 to 5 %0, depending on temperature and composition of the C-H-O system. However, even small isotope effects between two components may provide large isotopic fractionation if these components have different distribution coefficients between phases of a system and if these phases may be separated spatially. This situation may occur if a reduced subasthenospheric fluid interacts with the relatively oxidized lithosphere. As fluid ascends, isotope fractionation between reduced and oxidized carbon in conjunction with Rayleigh distillation may result in high enrichment in the light isotope in residual carbon in the fluid. Carbon of the residual fluid might be the source of isotopically light diamonds.

-30

Ultrabasic assemblage

Eclogite assemblage

-20 -10 o

~I3C. %0

Fig. 118. «5 13C of diamonds in the ultrabasic and ecologitic parageneses based on crystals from the Mir, Udachnaya and Sputnik kimberlite pipes as well as Ural and Ebelyakh placers


References

Galimov EM (1985) The biological fractionation of isotopes. Academic Press, Orlando, 261 pp Pillinger CT (1984) Light element stable isotopes in meteorites - from grams to picograms.

Geochim Cosmochim Acta 48: 2739-2766 Rumble D, Hoering TC, Grew ES (1977) The relation of carbon isotopic composition to

graphitization of carbonaceous materials from the Narragansett Basin, Rhode Island. Carnegie Inst Wash Year Book 76: 623--625

8.5.5 Sulfur Isotopes in Mineralogy

V.1. VINOGRADOV and T.F. ANDERSON

Sulfur has four stable isotopes: 32S (95%), 33S (0.76%), 34S (4.22%), and 36S (0.02%). Isotopic fractionation of sulfur on the Earth is caused by isotope-mass difference only. Therefore, variations in the relative abundance of any pair of sulfur isotopes can be used, in principle, to investigate the processes and extent of isotopic fractionation. For several obvious reasons (abundance, mass difference), the ratio 34S/32S is measured in geochemical investigations. Sulfur isotopic compositions are expressed relative to troilite sulfur from the Canyon Diablo meteorite in the usual 1> notation.

Based on a large number of analyses, 1> 34S of sulfur in meteorites has a very narrow range about a mean of 0 permil. The same is true for lunar rocks, the only exceptions being the fine fractions of regolith whose J34 values are > 0 due to physical evaporation caused by micrometeorite impact. Sulfur isotope studies on extraterrestrial materials lead to the conclusion that the isotopic composition of the total Earth is 1> 34S ~ 0 permil. Thus, variations in the isotopic composition between different sulfur reservoirs are due to terrestrial processes of isotopic fractionation. The most important among these is the reduction of sulfate to sulfide species.

Sulfate reduction is accompanied by the preferential enrichment of 32S in the product sulfide. The magnitude of the fractionation factor ((J( = Rsulfide/Rsulfate) depends upon a number of environmental factors, but especially temperature: fractionation decreases ((J( -+ 1) as temperature increases. At low temperatures ( < 80°C) kinetic isotope effects associated with the bacterial reduction of sulfate is the only important mechanism. In a closed system, both the residual sulfate and the product sulfide become increasing enriched in 34S as reduction proceeds. This behavior follows classical Rayleigh distillation. A convenient approximation for sulfur isotopic evolution in a closed system is given below:

l>(sulfate) - l>o(sulfate) = GIn f

l>(sulfide) = l>(sulfate) + e,

8.5.5 Sulfur Isotopes in Mineralogy 411

where f = fraction of original sulfate remaining, e = (IX - 1).103 [IX < 1 and e < 0 for sulfate reduction], (jo(sulfate) = (j 34S of original sulfate, (j(sulfate) = (j 34S of sulfate at f, and (j(sulfide) = (j 34S of sulfide produced at f.

Sulfate reduction may occur in two distinct types of regimes: (1) an open system, where the source of sulfate is unlimited; and (2) a closed system, where the supply of sulfate is limited. Sulfate reduction under open-system conditions is typical of marine environments. Bacterial sulfate reduction is extensive in marine depositional environments where organic matter accumulates, e.g., tidal flats, continental shelves and slopes, even in the water column itself in certain restricted basins. In the upper layers of the sediment column, bioturbation and advection continually supply seawater sulfate to zones of active bacterial reduction. The resulting sulfides are depleted in 34S by 20-50 permil relative to seawater sulfate. A significant fraction of dissolved sulfide is precipitated as iron sulfide minerals; their distinctively low (j 34S values are characteristic of an early diagenetic origin. Sulfate reduction also plays a major role in the isotopic massbalance of the seawater sulfate reservoir. The (j 34S of seawater sulfate is a constant + 20 permil in all modern open seas and oceans. The removal of 34S_ depleted sulfides maintains this steady-state (j34S at a value higher than that of sulfate in runoff (ave.: '" + 4 permil). The isotopic composition of seawater sulfate is uniform throughout the world ocean because circulation and mixing are more rapid than input and removal processes.

Thermochemical sulfate reduction in hydrocarbon-bearing fluids in sedimentary basins is an example of closed-system sulfate reduction. The source of sulfate in the fluid is the dissolution of marine evaporites. As discussed later, Phanerozoic evaporites have a mean (j 34S = '" + 20 permil. At temperatures > 80 DC, sulfates are reduced abiologically by the breakdown of complex

hydrocarbons. All of the major HzS occurrences in petroleum-rich sediments were probably formed by this mechanism. These "epigenetic" HzS deposits (as distinct from early diagenetic HzS) have characteristic (j 34S values in the range 0 to '" + 15 permil, indicative of inorganic sulfate reduction in a closed system. The same range of (j 34S is encountered in many native sulfur deposts, which formed by the oxidation of epigenetic HzS, and in the sulfides of most stratabound ore deposits. Thus, even in cases where the sedimentary origin of an ore deposit is clear, sulfur isotopic compositions often indicate the introduction of epigenetic HzS into the sediments, hence a "hydrothermal-sedimentary" genesis of the deposit.

The residual dissolved sulfate in thermochemical reducing systems becomes enriched in 34S. The extent of enrichment is variable because of temporal fluctuations in basinal fluid flow and hence in the supply of "fresh" sulfate to reduction zones. Nonetheless, (j 34S values in excess of + 30 to + 50 permil are typical of sulfate that is residual from thermochemical reduction. Many deposits of barite and celestite in sedimentary rocks have the same isotopic characteristics, suggesting that these deposits formed as a consequence of thermochemical reduction of evaporitic sulfate.


Seawater sulfate can also be a direct source of sulfur in some types of hydrothermal-sedimentary ore deposits. Investigations on hydrothermal systems at mid-ocean ridge spreading centers show that the partial reduction of sulfate by ferrous iron minerals takes place during the deep convection of seawater into newly formed oceanic crust. "Black smokers" are graphic evidence of this process. The participation of the sedimentary sulfur cycle in the origin of many low- and high-temperature ore deposits is a major achievement of sulfur isotope geochemistry.

The widest range in p 4S values occurs in the oxidized zone of stratabound ore deposits. Bacterial sulfate reduction, sulfide precipitation and the partial oxidation of sulfides occur simultaneously at the boundary between oxidizing and reducing pore fluids. This geochemical boundary migrates in response to fluid recharge. The oxidation of sulfides produces isotopically depleted sulfate. The subsequent reduction of the latter (in reducing fluids) results in a very depleted sulfide. As a consequence, the p 4S of sulfide minerals in such settings may range very widely, down to - 60 permil. Certain ore deposits of uranium ("roll-fronts") and copper ("redstones") formed by these complex redox processes and have the characteristic wide range in p 4S values.

Equilibrium fractionation does occur in isotope exchange reactions at high temperatures ( > 200 QC). At equilibrium, the more oxidized sulfur species is enriched in 34S relative to the less oxidized (or reduced) species. Note, therefore, that both kinetic and equilibrium isotope effects lead to the same qualitative fractionation between sulfur species of different oxidation state. However, only equilibrium fractionations are truly predictable (as a function of temperature only).

The largest equilibrium isotope fractionations occur between sulfate and sulfides. However, equilibrium in naturally occurring sulfate-sulfide assemblages is not attained at temperatures < 200 QC, and is often not attained in assemblages that precipitated at higher temperatures. This implies that kinetic effects play an important role in the attainment of sulfur isotopic equilibrium over a wide range of temperature.

If a sulfide mineral assemblage was precipitated at isotopic equilibrium, then fractionation between mineral pairs should indicate the temperature of precipitation. The largest fractionations between common sulfide minerals are (in decreasing order): pyrite-galena, sphalerite-galena, chalcopyrite-galena, pyrite-chalcopyrite. Unfortunately, only the sphalerite-galena pair yields reliable isotopic paleotemperatures, but not in every case.

The rate of attainment of isotopic equilibrium in a sulfide-sulfate system depends upon temperature, pH, and total sulfur content of the fluid. All of these parameters can be extracted from detailed isotopic data if equilibrium in the hydrothermal ore-forming fluid was achieved.

Magmatic rocks contain discrete sulfide minerals in at least trace amounts. But a certain fraction of total sulfur occurs as so-called dissolved sulfur in both reduced and oxidized states. These forms are practically never in isotopic equilibrium. Isotopic disequilibrium among sulfur species in natural systems is

8.5.5 Sulfur Isotopes in Mineralogy 413

another important conclusion of isotope geochemistry: geological and geochemical processes usually occur in open, nonequilibrium conditions, where thermodynamic equilibrium is the exception rather than the rule.

Disseminated sulfur minerals in magmatic rocks often have <5 34S #0. This suggests contamination by sulfur from the sedimentary cycle. However, even if disseminated sulfides have <5 34S = 0, it is not possible to unambigously assume that the sulfur was derived from the mantle. The partial reduction of sulfate circulating through the crust can also produce sulfides of this isotopic composition. For example, the sulfides of Black Smokers of the East Pacific Rise which formed by the reduction of seawater sulfate have <5 34S = 0 to '" + 4 permil. Secondary sulfides of this isotopic composition are also precipitated in the basalts themselves. Such processes make it very difficult to estimate the primary sulfur content of magmatic rocks and to estimate the flux of juvenile sulfur from the mantle.

The isotopic composition of sulfur species in volcanic gases does not provide direct evidence for the outgasing rate of juvenile sulfur either. The principal sulfur species in volcanic gases are S02 and H2S. With decreasing temperature, the S02/H2S ratio decreases. Isotopic fractionation invariably enriches S02 in 34S with respect to H2S. However, volcanic sulfur gases usually have <5 34S > O. It appears, therefore, that a substantial fraction of volcanic sulfur gases is recycled sedimentary sulfur.

The isotopic composition of sulfur in the oceans is the result of the complex geochemical cycle of the element. The steady-state concentration and isotopic composition of seawater sulfate are intimately related to such factors as continental runoff, salinity of seawater, productivity of the biosphere, and the composition of the atmosphere. The record of past variations in the isotopic composition of seawater sulfate permits us to investigate how these factors have changed through time. The sulfur isotopic composition of ancient marine evaporites is accepted as the best record of contemporaneous seawater sulfate. For much of the Phanerozoic, <5 34S of evaporites was close to the modern value of + 20 permil; two notable exceptions were the Permian ('" + 10 permil) and the Cambrian ( + 25 to + 30 permil). Low <5 34S values for Permian evaporites may have been due to the direct input of continental runoff into marginal evaporite basins to a much greater degree than at other times. High <5 34S values during the Cambrian may have resulted from very extensive sulfate reduction in sediments, and hence the preferential removal of 32S from seawater. In any case, there is no definite trend in the evolution of seawater <5 34S during the Phanerozoic. Therefore, the dynamic sulfur cycle did not change in a unidirectional manner during this interval (although it did fluctuate markedly). Isotope data on Precambrian evaporites are more scarce and less definitive. Some researchers suggest that the sulfur cycle has changed little over the past 2 billion years; others would extend the age limit to 3 billion years. The inception of a "modern" sulfur cycle remains one of the most intriguing problems of isotope geochemistry.


References

Alt JC, Anderson TF, Bonnell L (1989) The geochemistry of sulfur in a 1.3 km section of hydrothermally altered oceanic crust, DSDP Hole 504B. Geochim Cosmochim Acta 53: 1011-1023

Berner RA (1989) Biogeochemical cycles of carbon and sulfur and their effect on atmospheric oxygen over Phanerozoic time. Global Planet Change 1: 97-177

Grinenko WA, Grinenko LN (1974) Geochemistry of sulfur isotopes. Nauka, Moscow, 274 pp (in Russian)

Hoefs J (1987) Stable isotope geochemistry, 3rd edn. Springer, Berlin Heidelberg New York, 241 pp

Ohmoto H (1986) Stable isotope geochemistry of ore deposits. Stable isotopes in high temperature geological processes. Rev Mineral 16: 491-556

Vinogradov VI (1980) Role of sedimentary cycle in the geochemistry of sulfur isotopes. Nauka, Moscow, 192 pp (in Russian)

8.5.6 Nitrogen Isotopes in Mineralogy

D. HAENDEL and B.G. POKROVSKY

Nitrogen isotope geochemistry has been poorly developed until now. The following reasons may restrict progress in this field: low concentrations of nitrogen in metamorphic ( '" 200 ppm) and in igneous rocks (20 ppm) and a diversity of poorly studied biochemical reactions which control nitrogen isotope fractionation. Nitrogen has two stable isotopes: 14N and 15N. The average natural concentration of 15N in air is 0.3663 and it is constant within analytical precision. The variations of 15N/14N are given in permil relative to atmospheric molecular N z and written in the delta notation:

<5 1sN = [eSNP4N ) sample _ 1J x 1000 eSN/14N) air .

The thermodynamical fractionation factors (IX) of nitrogen have been determined for a few reactions. Under equilibrium conditions, the fractionation factors increase in the series: NH 3-Nz-NOz at 0-700°C; (Letolle; Richet et al.) AN02-NH3 = 35.5, A N,-NH3 = 10.3, both at 25°C (Letolle). The fractionation factors decrease with increasing temperature. The inversion of IX in the N z-NH3 system takes place at t '" 750°C, and AN2-NH3 = - 1 at 1000°C (Richet et al.).

In nature, the thermodynamic fractionation of nitrogen isotopes is realized probably only at high temperatures. The (j 1sN variations in sedimentary rocks depend mainly on kinetic effects which accompany processes of fixation of molecular nitrogen by living organisms, as well as of bacterial nitrification and denitrification, and by physicochemical effects, i.e., diffusion evaporation, dissolution and migration.

The isotopic composition of land plant organic matter is close to that of atmospheric nitrogen. Soils are slightly enriched in lsN, because during the

8.5.6 Nitrogen Isotopes in Mineralogy 415

decay of organic matter smaller nitrogen-containing organic molecules depleted in 15N are destroyed in the first instance (Letolle). The c5 15N mean value of continental organic matter is 2.5%0 (Sweeney et al.). Higher c5 15N values (up to + 20%0) are characteristic for oceanic plankton. c5 15N of marine sediments

varies over a wide range, depending on the balance of continental and oceanic sources of organic matter and diagenesis stage. Mean c5 15N values of nonmetamorphosed sediments are close to 5-6%0 (Haendel et aI., Sweeney et al.). Similar c5 15N values are characteristic for coals and oil. c5 15N values are characteristic for coals and oil. c5 15N values of natural gas varies in a wide range ( + 45%0 up to - 10%0 and lower). The negative values seem to result from fractionation during migration.

Regional metamorphism of sedimentary rocks is accompanied by decreasing nitrogen concentrations from 500-600 ppm in clays and phyllites to 5-10 ppm in granulites; with c5 15N commonly increasing to 5-10 ppm in granulites; with c5 15N commonly increasing to 5-10%0, the removing phase seems to be depleted in 15N. Similar effects take place during contact metamorphism (Haendel et al.).

Many available data on c5 15N determinations refer to total nitrogen. It has been established, however, that nitrogen of metamorphic rocks exists in various forms, that nitrogen of metamorphic rocks exists in various forms, differing in their isotopic composition (Haendel et al.). The main part of nitrogen is fixed as ammonium within the crystal lattice, with molecular nitrogen being occluded in its disturbances. Ammonium concentrations in coexisting minerals decrease in the series: biotite-muscovite-K -feldspar-plagioglase-quartz.

The c5 15N values of igneous rocks vary over a wide range, being probably the result of the heterogeneity of sources of matter as well as of the conditions of magma degasing. High c5 15N values (17-20%0) are obtained for the MORB (Becker and Clayton). Magmatic rocks (plutonic and volcanic) have less than 50 ppm, even in some cases less than 10 ppm nitrogen. Exceptionally, they contain up to 100 ppm and more. Such high amounts have been determinated, e.g., in diabasis and two-mica granodiorites and quartzdiorites (Brauer et al.) . The c5 15N values of magmatic rocks scatter over a relatively wide range from positive to negative values. Javoy et aI., who had studied Archean komatiits, discovered an inverse proportion of nitrogen and c5 15N variations of deep-seated rocks to be dependent on degasing and isotope fractionation in the NH3-N2 system. The c5 15N values of un-degased mantle may be as low as - 40%0, being characteristic for enstatite chondrites. Data on diamonds, which have common c5 15N values in the range of 0 to - 10%0 with a nitrogen concentration up to 2000 ppm, support this conclusion (Javoy et al.). Evidently, degasing oflow c5 15N mantle (c5 15N < 0%0) and releasing of nitrogen with lower c5 15N could not lead to the formation ofhigh-c5 15N atmosphere (c5 15N > 0%0) as well as the hydrosphere, biosphere, and continental crust connected with it. According to Javoy et aI., this contradiction may be eliminated if it is proposed that the upper layers of the Earth had been enriched in carbonaceous chondrites matter (C1 or C2), which have very high c5 15N values (up to + 170%0). However, the problem requires further investigation. Recycling and contamination of deep-seated rocks by different forms of biospheric nitrogen should be considered as an alternative.


References

Becker RN, Clayton BN (1977) Nitrogen isotopes in igneous rocks. OES Trans Am Geophys Union 58: 636

Brauer K, Stiehl G, Wand U, Gehre M (1990) 15N variations of rocks from the Lusatian Granodiorit Massif, GDR. In: Wand U, Strauch G (eds) Proceedings of the Fifth Working Meeting Isotopes in Nature. 25-29 September, 1989, Leipzig Akad der Wissenschaften der DDR Zentralinstitut fUr Isotopen und Strahlenforschung pp 247-254

Haendel D, Mahle K, Nitzsche H-M, Stiehl G, Wand U (1986) Isotopic variation of the fixed nitrogen in metamorphic rocks. Geochim Cosmochim Acta 50: 749-758

Javoy M, Pineau F, Delorme H (1986) Carbon and nitrogen isotopes in the mantle. Chern Geol 57:41-62

Letolle R (1980) Nitrogen-15 in the natural environment chap 10. In: Fritz P, Fontes JCh (eds) Handbook of environmental isotope geochemistry. Elsevier, Amsterdam, pp 407-433

Richet P, Bottinga Y, Javoy M (1977) A review of hydrogen, carbon, nitrogen, oxygen, sulphur and chlorine stable isotope fractionation among gaseous molecules. Annu Rev Earth Planet Sci 5: 65-110

Sweeney RC, Liu KK, Kaplan IR (1978) Oceanic nitrogen isotope and their uses in determining the sources of sedimentary nitrogen. In: Robinson BW (ed) Stable isotopes in Earth Science. Dep Sci Ind Res Bull 220: 9-26

8.6.1 Geochemical Significance of 87Sr /86Sr Isotopic Ratios

T.F. ANDERSON, 0.1. DEPAOLO, and V.I. VINOGRADOV

The initial strontium isotopic ratio, discussed in Chapter 8.2.3, is an important tracer in studies on the origin of rocks and minerals. At the beginning of our planet's history some 4.6 billion years ago, the primordial 87Sr/86Sr ratio of the earth was 0.699. This value was determined by isochron dating of stony meteorites and lunar rocks. Since that time, the 87Sr/86Sr ratio of the entire earth has increased. The extent of this increase in distinct reservoirs of the earth is directly proportional to their 87Rb/86Sr ratios. Thus, the present 87Sr/86Sr ratio in basalts (and especially in lunar anorthosites) is only slightly higher than the primordial ratio because of their low Rb/Sr ratio.

Because of the thermal and tectonic activity of our planet, all of the rocks of the earth have undergone chemical and mineralogical transformations through anatexis, metamorphism, and/or hydrothermal alteration. These processes redistribute Rb and Sr in new and recrystallized phases. In some cases, the 87Sr/86Sr ratio is rehomogenized in constituent minerals and the whole rock, and the isotopic decay clock is reset to zero. The new initial strontium isotopic ratio provides information on the geochemical history of rocks and their constituent minerals. It is particularly useful in distinguishing between mantle and crustal source regions of igneous rocks.

Oceanic basalts are obviously derived from mantle sources. Mid-ocean ridge basalts (MORB) have the lowest present-day 87Sr/86Sr ratio - 0.7022 to

8.6.1 Geochemical Significance of 87Sr/86Sr Isotopic Ratios 417

- 0.7033 - of terrestrial rocks. The upper limit is somewhat controversial. Oceanic basalts may become contaminated with (more radiogenic) seawater strontium at any stage in their emplacement. Interaction with seawater can occur under both high- and low-temperature conditions. In the case of the former, no clear evidence for secondary alteration may be observed. A strontium isotopic ratio of 0.7033 has been proposed as a somewhat arbitrary boundary between fresh and unaltered oceanic basalts. However, it is more prudent to accept the lowest 87Sr/86Sr ratios as the initial values and, further, to consider them as only an upper limit to the true value.

Even the lowest 87Sr/86Sr ratios in MORB are not supported by their Rb concentrations. If one accepts the Rb/Sr ratio of MORB as 0.008, then the strontium isotopic ratio would have increased from 0.699 to only 0.6991 in 4.6 billion years. For a Rb/Sr ratio of 0.05, characteristic of oceanic-island basalts, the strontium isotopic ratio would have increased from 0.699 to 0.702 in the same interval. This evidence indicates that the upper mantle source region for oceanic basalts was the product of very early differentiation in earth history.

Variations in the strontium isotopic ratio of fresh basalts are usually interpreted as evidence for chemical heterogeneity in the mantle. High 87Sr/86Sr ratios in basalts of submarine plateaus, islands, and island arcs are typically explained in the same manner. Isotopic heterogeneity can result from the existence of different "depleted" portions of the mantle through the secondary alteration of previously depleted portions in the course of subduction or other processes.

The initial strontium isotopic ratio of granitic crustal rocks and their minerals cover a broad range. Low 87Sr/86Sr ratios (ca. 0.704) are direct evidence for derivation from parent mantle material. The processes by which granitic rocks can be produced from mafic rocks are numerous, and can occur either in the mantle or in the upper part of the continental crust. Strontium isotopic evidence suggests that some granitoids were derived from mafic rocks of mantle origin in tectonically active continental-oceanic margins. High 87Sr/86Sr ratios in granitic rocks indicate reworking of more ancient rocks of the continental crust by processes of anatexis or contamination. In some cases, high initial strontium isotopic ratios reflect the influence of seawater strontium.

The 87Sr/86Sr ratio of modern seawater is 0.709. The average strontium isotopic ratio of continental runoff is about 0.711. It is clear, therefore, that the strontium isotopic budget of seawater must involve the exchange of strontium with oceanic basalts. The strontium isotopic composition of ancient sedimentary carbonate rocks provides the opportunity to study the evolution of the 87Sr/86Sr ratio of seawater through geologic time. The time dependence of this ratio is very complicated. This means that the strontium isotopic trend does not yield direct evidence for the growth of continental crust through time. The roles of continental runoff and leaching of oceanic basalts in the strontium isotopic balance of seawater were variable and irregular.


Further Reading

Brooks C, James DE, Hart SR (1976) Ancient lithosphere: Its role in young continental volcanism. Science 193: 1086-1094

Burke WH, Denison RE, Hetherington EA, Koepnick RB, Nelson NF, Otto JB (1982) Variation of seawater 87Sr/86Sr throughout Phanerozoic time. Geology 10: 516-519

Faure G (1986) Principles of isotope geology, 2nd edn. John Wiley & Sons, 589 pp Faure G, Powell JL (1972) Strontium isotope geology. Springer, Berlin Heidelberg New York,

188 pp Staudigel H, Hart SR, Richardson SH (1981) Alteration of the oceanic crust: processes and

timing. Earth Planet Sci Lett 52: 311-327

8.6.2 Geochemical Significance of 143Nd/144Nd Isotopic Ratios

D.l. DEPAOLO, T.F. ANDERSON, and V.1. VINOGRADOV

Variations in the 143Ndj144Nd isotopic ratio in rocks and minerals are much smaller than those in 87Sr/86Sr. This is due to the low decay constant for 147Sm and the very small variations in the SmjNd ratio.

The isotopic evolution of neodymium in the earth assumes that terrestrial Nd has evolved in a "chondritic !!niform reservoir," or CHUR. The modern values of 147Smj144Nd and 143Nd/144Nd in CHUR are assumed to be equal to the average value of those ratios in stony meteorites. Differentiation of CHUR results in a small fractionation of Sm from Nd and accounts for variations in the SmjNd ratio in rocks. Magmas derived from CHUR have a lower SmjNd ratio and the residue a higher ratio than CHUR. As a consequence, the 143Ndj144Nd ratio increases at a slower rate in the differentiate and a faster rate in the "depleted" residue relative to CHUR itself.

Because of differences in analytical procedures between laboratories, 143Ndj144Nd ratios are expressed relative to the appropriate CHUR value for that ratio in the eNd notation:

[ (143Nd/144Nd)samPle ] 4 eNd = e43Nd/144Nd)cHuR - 1 x 10 .

The present-day eNd value is calculated from the measured 143Ndj144Nd ratio in the whole-rock sample and that ratio in CHUR today (0.512639). We can also calculate the eNd ratio of the rock at the time of its crystallization by comparing its initial 143Nd/144Nd ratio (derived from a whole-rock isochron) to the 143Ndj144Nd ratio at the same time. The neodymium isotopic ratio ofCHUR at any time in the past can readily be calculated from its present-day ratio and its 147Sm/144Nd ratio (0.1967). A positive eNd value indicates that the rock was derived from a previously depleted reservoir. A negative eNd value indicates that

8.6.2 Geochemical Significance of 143Nd/144Nd Isotopic Ratios 419

the rock was derived from ancient crustal rocks whose Sm/Nd ratio had been lowered earlier.

Mid-ocean ridge basalts are typical representatives of rocks derived from a depleted mantle reservoir. The highest eNd values for MORB are approximately + 13. The estimated age of MORB source region separation from CHUR

("model age") is 2.8 billion years. Although the uncertainty in this apparent age is high, the results do suggest early separation of upper mantle source regions from a uniform reservoir.

On a broader scale, the 143Ndj144Nd ratios of oceanic basalts vary more widely. These variations are inversely related to 87Sr/86Sr variations among the rocks. The linear field which encompasses the scatter in Nd- and Sr-isotopic compositions in oceanic basalts is called the "mantle array" and is generally interpreted as a mixing trend of magmas from depleted and undepleted regions. Other types of oceanic and continental igneous rocks show correlated Nd and Sr isotopic variations that differ from the mantle array. For example, continental volcanic and plutonic rocks have relatively high Sr isotopic ratios by low Nd isotopic ratios, indicative of contributions from ancient continental crust. Combined Sr and Nd isotopic studies are a powerful tool in petrogenetic investigations.

The very low concentration ofNd in seawater is a result of the rapid sorbtion of rare-earth elements by authigenic sedimentary minerals, principally phosphates and Fe-Mn oxyhydroxides. The residence time of Nd in seawater is so short that its isotopic composition varies between the major ocean basins and depends to a large extent on the isotopic composition of runoff. The highest 143Ndj144Nd ratio (or eNd value) is in the Atlantic; the lowest occurs in the Pacific. Therefore, there is either relatively little ancient sialic crust exposed in the Pacific's drainage basin, or there is more rapid interaction between the basaltic crust and seawater in hydrothermal systems of the Pacific.

Hydrothermal alteration by seawater usually increases the 87Sr/86Sr of oceanic basalts but does not change the 143Nd/144Nd ratio. However, situations can arise in which descending seawater penetrates a significant thickness of sediments and thereby becomes enriched in radiogenic 143Nd. In these cases, the hydrothermally altered basalt may acquire correlated Sr-Nd isotopic variations that mimic the mantle array.

Sedimentary phosphates are an effective concentrator of neodymium (and other rare-earth elements) from seawater. For this reason, investigations on ancient sedimentary phosphates provide information on the Nd isotopic evolution of seawater through geologic time. These studies are now in progress.

Further Reading

DePaolo OJ (1981) A neodymium and strontium isotopic study of the Mesozoic calc-alkaline granitic batholiths of the Sierra Nevada and Peninsular Ranges, California. J Geophys Res 86(Bll): 10470-10488


DePaolo DJ (1988) Neodymium isotope geochemistry: an introduction. Springer, Berlin Heidelberg New York, 187 pp

DePaolo DJ, Wasserburg GJ (1979) Petrogenetic mixing models and Nd-Sr isotopic patterns. Geochim Cosmochim Acta 43: 615-627

Patchett PJ (1989) Radiogenic isotope geochemistry of rare-earth elements. Geochemistry and mineralogy of rare-earth elements. Rev Mineral 21: 25-44

Piepgras DJ, Wasserburg GJ, Dasch EJ (1979) The isotopic composition of Nd in different ocean water masses. Earth Planet Sci Lett 45: 223-236

CHAPTER 9

Computer Databases in Mineralogy

422 Chapter 9. Computer Databases in Mineralogy


D.G.W. SMITH

No geoscientists have embraced the computer revolution more willing and wholeheartedly than those in the mineralogy-petrology-geochemistry (MPG) group. The computer has become an everyday necessity for everything from instrument control to data gathering, data processing, modelling and, ultimately, the production of the journal article, book or report. In common with other geoscientists, the MPG group needs not only a wide spectrum of computer programs but, to an ever-increasing extent, access to databases which offer some hope of managing the deluge of data and information published in hundreds of journals, books and reports in a dozen or more languages. The need to bring some order to the developing chaos, avoid pointless duplication of effort and promote international cooperation has recently prompted the International Mineralogical Association (IMA) to form a new Working Group on Databases and Computer Applications in Mineralogy (15th General Meeting, Beijing, P.R.e. 1990). This chapter will deal with the area of databases only. To cover computer applications adequately would already require an entire book. In fact, the new IMA Working Group, as one of its first tasks, has compiled a catalogue of computer software that is of interest to the mineralogist and is currently available in the commercial, public and private domains.

Database Structures and Types

It is not the purpose of this chapter to delve deeply into the nature of databases and management systems and certainly not into the problems of their interaction and optimization. Readers wishing to explore these directions further are referred to excellent textbooks now available. We may, however, usefully clarify some terms. A datafile may be regarded as a collection of data records of the same type. These records consists of a series of data fields. Datafiles may be small or extremely large. A collection of datafiles is often referred to as a database. However, such a collection is better referred to as a databank. The term database is nowadays used by many people in a more restricted sense to refer to the integration of constituent datafiles with a database management system (DBMS).

Databases and DBMS have been discussed recently from the point of view of the earth sciences by Rock (1988). True databases may be grouped into three categories: relational, network and hierarchical. Relational databases may be thought of as a collection of independent datafiles. Items within these files are identified by "primary keys", so that there is generally access from anyone file to any other. This contrasts with the situation in hierarchical databases where the underlying structure is more rigid. Files are arranged in the form of a tree with branches. Direct access from a particular file is to one file only in one direction

Computer Databases in Mineralogy 423

(the "parent" of that branch) and to one or more in the other direction (the "children"). Thus we have here a one-to-many connection between files. Access to files on other branches is possible only by retracing a path back down one branch and then up another. This kind of database works well where the items represented in the database form a natural hierarchy. However, it becomes difficult to use when many-to-many relationships exist between items to be included. Network databases have some of the characteristics of both of the former. They are constructed along hierarchical lines but more direct links between particular branches are established by the use of keys and "pointers". Here then we have a many-to-many access between files. Network databases specifically model the relationships between items to be represented in the database. The term "graphical databases" has also appeared in the literature to refer to a database of some kind that has had grafted onto it various computer graphics tools. Thus numerical and plotting capabilities may be combined so that information stored in the database may be used to construct diagrams, figures, maps, etc. For example, the geographic distribution of Pre-Cambrian gold occurrences might be illustrated by such a database containing global information on the occurrences of that mineral.

Databases of all kinds tend to consume substantial amounts of storage space and for this reason were restricted for many years to mainframe computers with the data normally being held in secondary memory - disks and tapes. With modem developments in computer technology, many useful databases can now be held on the hard disks available for the microcomputer, with tape drives being used mainly for back-up purposes. The development ofthe CD-ROM has, to a very large extent, eliminated the problems of information storage for micro (and mini) computers. These devices, which have capacities of up to a gigabyte, may accommodate all but the largest databases on a single disk. Although they have the disadvantage of being permanent and unmodifiable, the very recent introduction of write-many, erasable optical disks (CD-RAM) will doubtless remove even this constraint. In this situation, the advantage of larger computers will be principally in their processing power. Even this is being eroded to an astonishing extent and the division into super-computers, mainframes, minicomputers, work stations and microcomputers is becoming daily more blurred.

The networking of computers in larger organizations, such as governments, multinational corporations, institutions etc., led to the development of distributed databases. IIi this case, many of the connected computers may run a database management program which accesses databases (possibly overlapping) which may be located at each network node. In general, users remain unaware of and unconcerned about the physical location of data they access. Such databases are particularly useful where the many distributed records are continually undergoing revision and expansion under the control of specialist individuals or groups.

Because the larger databases were held on mainframe computers, there has been a progressive development of "on-line" access to databases over the years. This access has varied in character from connection of terminals to a mainframe within an institution, through the connection of in-house computers to a remote


mainframe via telephone and data communications systems, to the distribution of information via television networks. The importance of such systems lies in their ability to make huge amounts of information which is being continually updated available to very large numbers of people. General directories of computer-readable databases are now published while on-line database descriptions appear in both printed form and also on-line in a database of databases: DIALOG file 230 covers more than 4000 that are publicly available.

While there has been an overwhelming move towards the use of desk-top computers in recent years, necessarily changes have also occurred in the way in which mainframes are used. In one such trend we see mainframes acting as "back-ends" which maintain and manage huge databases, while the proliferation of micros forms an ever-increasing network of "front-ends" for these machines. However, the requirements of most mineralogists do not demand such facilities. Rather they need access to a substantial body of carefully and scientifically vetted information in one or more specialized fields. Although some of this information does need to be updated or upgraded, most of it remains constant for substantial periods of time. For this reason much more emphasis can be placed on the small to medium size databases implemented via multiple copies on widely distributed microcomputers.

There are now numerous database systems available commercially. Although generally oriented towards the business rather than the scientific communities, systems such as dBASE IV, ORACLE or INGRES are not restricted in their purpose and can be used in a wide variety of applications. More and more mineralogical applications utilizing such commercial software are beginning to appear. However, many of the so-called databases currently in use by mineralogists are purpose-developed and can be used in only one application. In a sense, they are databanks rather than databases and their utility depends entirely on special software which has been written to access and use the information stored in them. This independent development of many special systems has led to almost total incompatibility amongst the software. Now, as the importance of such collected data and information is becoming more and more obvious and widely recognized, moves are afoot to channel future efforts along lines that will hopefully lead to a higher degree of compatibility, much greater ease of access and widespread availability. We shall return to this point at the end of the chapter.

Databases of Mineralogical Interest

Let us look first at some of the areas where database development has become important to mineralogists (Table 18), and at some examples of databases already available.

Crystallographic Databases. These are, of course, of interest to a wide spectrum of scientists and not simply to mineralogists. Several very substantial databases


Table 18. Principal areas of database development

Crystallography Powder diffraction Thermodynamics Optics Spectroscopy Image libraries Multi-purpose or comprehensive Bibliographies Technological mineralogy Exploration Collections

425

have been available for a number of years on mainframes, the largest being the CRYSTAL DATA file of the National Institute of Standards and Technology (formerly the National Bureau of Standards or NBS). This contains information on about 60000 materials in the organic, organometallic, metal, intermetallic, inorganic and mineral categories. The database contains bibliographic, chemical, physical and crystallographic data and includes parameters such as empirical formula, chemical formula, chemical name, cell parameters, density, spacegroup and crystal symmetry. There is a high degree of quality control with data being checked for internal consistency by the program NBS· AIDS83. The many uses of the database include the identification of "unknowns" using crystallographic, compositional and other characteristics, the examination of the statistical distribution of specific properties, as well as the retrieval of numerical data, bibliographic information and also the validation of references. In Canada, the program is available on-line through CISTI (Canadian Institute for Scientific and Technical Information), who also provide a French version under the name of CRIT ALON.

The Inorganic Crystal Structure Database (ICSD) is another major database, which was developed in West Germany. It is also offered on-line through CISTI in Canada in both English and French language versions, the latter under the name CRYST ALIN. At present there are about 25000 entries and it also can be searched for numerical data and bibliographic information. In England, the Cambridge Structural Database has been developed by the Cambridge Crystallographic Data Centre, but it is concerned principally with organic materials.

In Japan, the crystallographic databases of choice are nCST and MICD, while in the USSR MINCRIST is used most widely by the mineralogical fraternity. The purpose of all of the databases is to provide fundamental information on the crystal structures of all known organic/inorganic compounds - not just minerals. In fact, mineralogists undoubtedly constitute a rather small minority of the users.


Powder Diffraction. The powder diffraction file (PDF) is unique. It is produced by the Joint Committee on Powder Diffraction Standards (JCPDS) of the International Centre for Diffraction Data (lCDD). This U.S. organization has assembled a vast amount of data on the X-ray powder patterns of both organic and inorganic materials. Once again, only a small part of the data set is of use to mineralogists although, in this case, the relevant part is very widely used by them. The datafile is under continual revision and a wide, if somewhat informal, international network for gathering data has been established to augment data culled from published literature. The trend now is away from the use of the PDF on mainframes and it is currently made available by ICDD for use in commercial microcomputer programs for identification of minerals (or other substances). An advanced CD-ROM version of the datafile is now available and a search/display system utilizing this has also been described.

Thermodynamic Databases. These are of interest to a wide spectrum of researchers in the MPG group and are used to depict stable equilibrium relations amongst mineral phases. Several such databases have been described recently, prominent amongst them being GeO-Calc, which has an internally consistent database for some 70 minerals in the system Na20-K20-CaO-MgOFeO-Fe203-AI203-Si02-Ti02-H20-C02. Associated programs permit calculation of pressure-temperature (PT), temperature-composition (TXH2o-co2)' pressure-composition (PXH20 -CO2)' activity-activity (AA), temperature-activity (TA) and pressure-activity (PA) phase diagrams. GEOPATH is another program for geothermobarometry and related calculations relevant to the crystallization (or recrystallization) of magmatic or metamorphic rocks. Like GeO-Calc, it has been implemented for IBM-compatible computers. The program GIBBS uses a thermodynamic database modified from earlier versions of Berman, Brown and Perkins to perform generalized Gibbs method calculations.

Optical Data. Notwithstanding the fact that optical data for minerals are obtained more easily and economically than any other kind of numerical data and often with excellent accuracy, there are few computer-readable compilations of such data for minerals. A number of programs exist which access very limited databases of transmitted light properties and these are essentially oriented towards undergraduate teaching. In one case, these properties form part of the knowledge base of a proposed expert system for the identification of minerals in thin section. The MinIdent database now contains optical data for a large number of known minerals as well as for numerous minerals that have been described but so far remain unnamed. Optical properties in that database include refractive indices, colour (and pleochroic scheme) in thin section, optic axial angle and the orientation of the optic axial plane where it is of diagnostic value (monoclinic minerals). In addition, particularly for opaque minerals, reflectances at four standard wavelengths (470,546,589 and 650 nm) are included when available. This optical database, which involves data for mUltiple


samples of many minerals, permits rapid identification of an "unknown" - often on the basis of optical properties alone.

Other more specialized databases have been developed for opaque minerals, the most recent of these being QDF - the IMA/COM Quantitative Data File. This database, which is available for IBM-PC machines and compatibles, has the unusual feature that it is user-modifiable.

Spectroscopic Data. The QDF referred to above is, in part, a database for reflectance spectra of minerals in the optical region. Many other spectral measurements of minerals are of interest and are widely available in the literature, including UV, optical and IR spectra, as well as NMR, EPR and luminescence spectra. Several difficulties have to be overcome in making useful compilations of such spectra, not the least of which is the actual transfer of published spectra in digital form to a database. However, some progress has been made in the area of IR spectra of minerals and undoubtedly many of the problems of data transfer will be soon overcome with the perfection and increased availability of optical scanners - although their use may raise interesting questions of copyright. The program SPIR (available on-line through CISTI) searches a huge database of some 145000 IR spectra (originally compiled by the American Society for Testing and Materials - ASTM). However, the vast majority of these are from organic substances and only a relatively small group of inorganic and mineral spectra are included.

Reflectance spectra have applications to the remote sensing of the modal mineralogy of terrestrial, lunar, planetary and asteroidal surfaces. Thus an expert system which uses a database of mineral spectra in the visible and IR ranges to interpret the mineralogy from images obtained by airborne and shuttle spectrometers has been developed by NASA.

Other "spectra" of interest to mineralogists include those produced in energy dispersive analysis as well as the curves obtained in the DT A and TGA techniques. Incorporation of such data in a useful form poses particular problems related to the effects of experimental conditions and instrumental parameters.

Image Libraries. The potential for computerized image libraries is enormous and is presently limited mainly by the storage required for high resolution images. With the rapid developments that have taken place recently in digital imaging and image processing techniques and the interest shown in them by the mineral industry, there can be little doubt that this will be one of the major areas of minerals database development in coming years. Not only will it be possible to store, retrieve and display full colour images of hand specimens, for example the entire collection of a museum, but also a wide range of images produced by microbeam instruments - from the typical secondary and backscattered electron images produced by scanning electron microscopes or microprobes to the colour images produced by the cathodoluminescence microscope. As CD-RAM


becomes more generally available and as its cost declines, such image libraries will proliferate.

Multipurpose or Comprehensive Mineral Databases. The pages of mineralogical reference texts such as Dana's System of Mineralogy or Deer, Howie, and Zussman's Rock-Forming Minerals contains extensive, in-depth treatments of the information available on minerals. The compositional, optical, crystallographical, spectral, physical, paragenetic and locality information which is presented in such books can all also be stored in digital form.

A step in this direction has been taken with the development of Minldent. Originally envisaged as software which would use compositions calculated from stored mineral formulae to assist those involved in microbeam analysis with mineral identification, it has now been expanded to include actual data and to cover more than 30 other areas. At the time of writing (1990), data are included for about 4650 named and unnamed mineral species, groups, families etc., with the compiled summary records being based on data for particular specimens (sample records) as well as on non-specific information (general records) which are often in the form of ranges observed for the property. For minerals where hardly any data are available there could be only one record, whilst for ubiquitous minerals there could be 50 or more. Many of the properties (or parameters) stored can be used by the associated software for mineral identification - for example: composition, optical properties and physical properties including symmetry, cell dimensions, d-values and hardness. The power of the identification software lies in its use of "scoring algorithms" to calculate a "matching index" and hence rank minerals in the database in order of their similarity to an unknown.

Bibliographic Databases. It has been said that the most crucial piece of inform ation to obtain about a mineral is its name. Certainly, this is the key to unlocking all the data that have been obtained in previous studies. Today, the sheer volume of literature and the many languages involved have made the literature database an essential aid in identification of important references. In the Soviet Union the VINITI database of the All-Union Institute of Scientific and Technical Information offers broad coverage of the earth sciences, including, of course, mineralogy and related fields. Additional international geoscience databases are available on-line in many other countries, e.g., GEO ABSTRACTS and GEOARCHIVE in the u.K.; GEOREF in North America; PASCAL in France. CD-ROM versions of some of these databases are now available. In the English language two publications have played a dominant role and both are now accessible on-line for computer search. Mineralogical Abstracts has been available since 1982 as part of the GEO ABSTRACTS database as DIALOG file 292. Chemical Abstracts, a much more widely based undertaking, is available on-line through the Lockheed Dialog system as file 399 as well as through various other on-line information systems. Such invaluable enterprises involve very substantial financial commitments and fees are necessarily charged for


interrogating the database. Some researchers therefore prefer to use the IBMPC compatible program MINABS which, at the time of writing, offers a very compact compilation of references to minerals in Mineralogical Abstracts back to Vol. 10. Because this journal is widely available in libraries, abstracts of interest can then be checked directly by the user. In effect, MINABS offers a computerized mineral index to Mineralogical Abstracts over the last 43 years.

Technological Mineralogy. This is an important area of mineral database development, embracing several different kinds of data. Mainframe databases such as MINPROC and MINTEC have been available on-line for some years. They are bibliographic in nature and focus on mineral processing, extractive and physical metallurgy, as well as some energy-related technologies. Other databases are dedicated to minerals of particular economic or industrial importance such as gold, diamond and zeolites. Most of these are also bibliographic in nature, although some do contain real mineral data. Even less visible to most academic mineralogists are the databases which contain actual information on mineral processing. Most of these are held securely in-house by the companies who developed them. Apart from information on recovery techniques, they also hold data that are important in extractive metallurgy.

Mineral Exploration. Because of the competitive commercial nature of mineral exploration, databases relating to mineral deposits and the exploration for them are generally restricted in access to the company that established them. Many are developed for a particular commercial project and are abandoned when it is completed. The MASNC database of the U.S. Bureau of Mines provides information on more than 200000 active and developed mines, deposits, other identified resources, and mineral processing plants, distributed worldwide. Other databases, such as MINLIB, are bibliographic in nature, enabling a wide range ofliterature bearing on some particular topic to be assembled quickly and conveniently.

Mineral exploration is an area where attempts have been made to build "expert systems". These are complex systems which try to bring artificial intelligence to bear on an ever-growing database. In such a system new data are added as they become available and predictions are made on the basis of all the information currently stored. Once actions have been taken on the basis of these predictions (for example, a drill hole bored) the situation that actually exists can be recorded and compared to that predicted. In a true expert system a comparison between the predicted and actually observed situation will be used as a basis for modifying the algorithms that made the prediction. This may be achieved by the intervention of an expert and/or a programmer, or by automatically changing the weight of terms in algorithms used to produce predictions so that they most closely reflect reality on the basis of the total experience of the system since it was brought into being. It may also be necessary to add new terms to the algorithms from time to time as effects of previously unrecognized factors become apparent. Necessarily, expert systems not only require major


data storage capabilities but also extremely powerful computing systems and highly involved programmers.

Collection Catalogues. The catalogues of mineral collections naturally lend themselves to conversion into a computer database. There are many samples for which it is desirable to store and retrieve information under a large number of different headings. Apart from obvious fields such as name, collection number, locality etc., the many other useful pieces of information include such items as analytical and other data obtained for an individual sample, paragenesis, source particulars (including cost and history), the loan/use history, and the physical location of a specimen. Indeed, one museum known to the author has already taken advantage of modern digital imaging technology to store a record of every significant sample in the collection.

However, again we are faced with the problems of an absence of standardization, with everyone having chosen a different path to computer cataloguing - from very simple to very elaborate, using different languages, some using mainframes and others micros-some IBM or IBM-compatible computers, and others not. Commercially available databases have been utilized in some cases while in others they have been purpose-developed. If it is ever going to be possible for mineralogists to browse through the holdings of the major collections, these problems of standardization will have to be solved.

One solution on a national level that was reported recently describes the incorporation of the catalogue of the Canadian National Mineral Collection into the database of the Canadian Heritage Information Network (CHIN). At the time of writing, the records consist of up to 90 fields chosen from hundreds available (all of which are described in a CHIN data directory). Browsing and simple or complex data searches can be carried out on-line using easily-learned commands and the results can be saved and then retrieved, if necessary, at a later date. A microcomputer interface is available for up-loading or down-loading in an ascii-compatible format called MICROTEXT. This not only offers the possibility of interfacing with an extensive range of word processing software, but should also provide a facility for connection through data communications networks to other systems and countries and perhaps form the basis of international cooperation in this area.

Situation and Perspectives

The 1970s saw the development of many good programming tools; the 1980s the appearance of new productivity tools. Computer software advances in the 1990s may well be dominated by advances in databases and database management systems. We may expect to witness, for example, rapid developments in hypermedia databases which seek to integrate numerical, textual, graphical and even audio information. Text-handling programs may themselves move towards


becoming databases where files created or acquired by the user, from electronic mail to letters, journal articles, books, bibliographies, indices, etc., can be accessed, browsed, scanned and cross-referenced by the software. It seems likely that once such facilities include good graphical capabilities, journals may be offered in disk form as an alternative to expensive hard copy issues. At the same time, libraries will certainly expand their roles as gateways to databases, whether these be available locally on a medium such as CD-ROM or on-line from remote locations. All of these developments can be expected to have a considerable impact on mineralogists.

Today hundreds or perhaps thousands of generally small mineralogical databases have been developed by individuals or groups. Although often highly specialized, they are potentially of interest to a much broader audience than the local one to which they are presently accessible. The value of such databases encompasses both the particular, often unique, expertise and perspectives of those who developed them and the enormous number of man hours that have, collectively, gone into their construction. Unfortunately many of the databases are far from standard - in fact, most are purpose-developed and applicationspecific. The question, therefore, arises as to what extent we can and should attempt to convert these disseminated databases into a distributed database on the one hand or compile them into a global or central database on the other. If the global mineral database (GMDB) option is chosen, it should certainly be a hypermedia database. It might be built in part from other databases that already exist, and augmented with new and additional data collected by the organizing body. Mineralogically, petrologically or geochemically oriented programs that are developed in future might then access and use that global database (Fig. 119).ln some respects this is an attractive direction to pursue, as it capitalizes on past efforts and ensures that users are accessing the same, relatively reliable and wellvetted data sets. This direction would also probably simplify linkage between future software. However, formidable obstacles would need to be overcome. Thus the building of this global database to high standards would pose great organizational problems. The costs of constructing a good database are enormous and much of the existing software would have to be rewritten. Availability of databases is another problem, as they may lie in the commercial, public or private domains. In any attempt to integrate databases and make them more widely available, the vested and financial interests of organizations, groups and individuals who have been involved in building and developing databases over many years must be addressed. The experts who originally produced the software may simply be unwilling or unavailable to cooperate.

Apart from such organizational problems there would be great difficulties in maintaining standards and ensuring accuracy. Confidence in, and the value of, a database is quickly eroded by the proliferation of errors and inconsistencies. The problems of errors are indelibly impressed on the minds of all who have been involved in database production. It has been estimated that even in the more carefully constructed databases errors remain in up to 5 % of the records, despite the use of software to check systematically for inconsistencies. This is a

432

Public domain

dedicated programs

Mainframe dedicated programs

Database management

programs

Global Mineral

Database (GMDB)

Chapter 9. Computer Databases in Mineralogy

Licenced dedicated programs

Private dedicated programs

Fig. 119. The global mineral database concept

frightening statistic that should bring users up with a jolt. Errors incorporated in databases are of several kinds - experimental, author and typographical, and eliminating or at least minimizing them is an expensive and time-consuming activity. Ultimately though, it is the user who must beware, recognizing that there is no such thing as a perfect database.

Another alternative, a multi-pronged approach to the problems of integrating existing and future databases (Fig. 120), is conceptually much less attractive but perhaps more practical, realistic and therefore likely to be chosen. On the one hand, the development of a global mineral database might proceed - albeit more slowly - and perhaps, therefore, more carefully, while on the other hand, efforts could be directed towards providing linkage between existing programs. Shell programs can be constructed which act as interpreters and isolate the user from direct contact with a program. Such shells can be designed to access many different pieces of software and they may have all the features that have become popular in recent years - a mouse, windows, pop-up and pull-down menus, colour, and so on. The construction of such software is indeed challenging, but not nearly so challenging as rewriting all the complex software that has already been developed. To a considerable extent use can be made of commercially available software designed for this purpose and which is already on the market. It does normally require the cooperation and, of course, agreement of the original programmer, but other than that, in general, it does not need to involve experts in the field. Some computer purists tend not to like such a solution - it involves many time-consuming I/O operations and it is certainly not the most efficient way to proceed, but science very seldom does proceed in efficient ways.

Attention must also be paid to many other important problems - in particular to the problems oflanguage - on the one hand, compatibility between the languages of computer databases and on the other, the languages of nations. Rapid advances have already been made in the development of SQL (structured


New licenced GM08-

dependent programs

New interactive

managemenl programs

Stand alone programs no 0-8 required

New Public domain GM08-

dependent programs

Programs with

dedicated databases

Converted older

GM08-dependent programs

On-Une local

mainframe programs

On-Une programs at remote

sites

Stand a lone Private commercial t-..... -("3-~~---;i~ .... .>-- ...---l stand alone ~og~~ p~~~

Stand alone programs in publiC domain

Fig. 120. The hell program concept

433

query language) for DBMS applications. Some progress has also been made in the latter area and software is now available which permits presentation in as many as nine different languages including English, Russian, Chinese, Japanese, German, French and Spanish. Beyond this a host of other far-reaching problems exist - the problems of finance, data evaluation and data entry, the introduction of new kinds of database, as well as the ways in which mineralogists can best interface with other disciplines that are closely related to our own - particularly areas such as crystallography, mineral engineering, ceramics, soils, gemmology etc.

It is clear that database development will be an area of substantial activity in the last decade of this century. It is an area in which many mineralogists can and should make substantial contributions and where international organizations


such as the IMA must show leardership. Along with the creation of expert systems and other forms of artifical intelligence, these developments should provide investigators of the 21st century with powerful new tools for scientific and industrial research in mineralogy.

References

Ansell HG (1989) The National Mineral collection of Canada database: features and selected applications. 25th Int Geol Congr Washington, DC July 1989. Abstracts with Program, vol 1, p 48

BergerhoffG, Brown ID (1980) Inorganic crystal structure database. In: Glaeser PS (ed) Data for science and technology. Proc 7th Int CODATA Conf Kyota, Japan. Pergamon Press, Oxford, pp 324-326

BergerhoffG, Brown ID (1987) Inorganic crystal structure database. In: Allen FH, Bergerhoff G, Siever R (eds) Crystallographic databases. International Union of Crystallography, Chester, England, pp 77-95

Brown TH, Berman RG, Perkins EH (1989) PTA-SYSTEM: a GeO-Calc software package for the calculation and display of activity-temperature-pressure phase diagrams. Am Mineral 74: 485-487

Chiou WC (1985) NASA image-based geological expert system development project for hyperspectral image analysis. Appl Opt 24: 2085-2091

Directory of Online Databases (1990) Publ Quarterly by Cuadra/Elsevier, New york Elmasri R, Navathe SB (1989) Fundamentals of database systems. Benjamin/Cummings,

802 pp Freeman KJ (1987) Book review of Mineral powder diffraction file, voll. Data book and vol 2

Search manual by Bayliss P, Erd RC, Mrose ME, Sabina AP and Smith DK (1986). Can Mineral 25: 794-795

Gerlitz CN, Leonard BF, Criddle AJ (1989) Reflectance of ore minerals: search-and-match identification system for IBM PC's using IMA/COM quantitative data file for ore minerals, 2nd issue. 28th Int Geol Congr Washington, DC, July 1989. Abstracts vol 1, pp 544-545

Gerya TV, Perchuk LL (1990) GEOPATH: A new computer program for geothermobarometry and related calculations with IBM PC computer. 15th General Meeting IMA Beijing, 1990. Abstracts, vol 2, pp 1010

Jenkins R, Holomany M (1987) "PC-PDF": a search/display system utilizing the CD-ROM and the complete powder diffraction file. Powder Diffraction, vol 2, JCPDS - International Centre for Diffraction Data, Swarthmore, PA, PP 215-219

Lu J (1990) Microcomputer system for identification of minerals. 15th General Meeting IMA Beijing, July 1990. Abstract, vol 2, pp 709-710

Marcaccio KY (ed) (1989) Computer-readable databases: a directory and data source book, 5th ed. Gale Research Inc Book Tower, Detroit, USA, 1188 pp

Petruk W (ed) (1989) Short course on image analysis applied to mineral and earth sciences (Ottawa, May 1989). Mineral Assoc Canada, Ottowa, 156 pp

Rock NMS (1988) Numerical geology, vol 18 In: Bhattacharji S, Friedman GM, Neugebauer HJ, Seilacher A (eds) Lecture notes in earth sciences. Springer, Berlin Heidelberg New York, 427 pp

Salisbury JW, Walter LS, Vergo N (1989) Availability of a library of infrared (2.1-25.0 Jlm) mineral spectra. Am Mineral 74: 938-939

Smith DGW (1990) Computer assisted mineral identification using conventional optical observations. 15th General Meeting IMA Beijing, July 1990. Abstracts vol 2, pp 711-712

Smith DGW, Leibovitz DP (1986) Minldent: a database for minerals and a computer program for their identification. Can Mineral 24: 695-708

Smith PD, Barnes GM (1987) Files and databases: an introduction Addison-Wesley, Reading, MA, 390 pp


Spear FS, Menard T (1989) Program GIBBS: a generalised Gibbs method algorithm. Am Mineral vol 74: 942-943

Siisse P (1989) MINABS: mineral data and reference file based on Mineralogical Abstracts. 25th Int Geol Congr Washington, DC, July 1989. Abstracts with program, vol 3, p 199

West J (1985) Toward an expert system for identification of minerals in thin section. Math Geol 17: 743-753

Subject Index

Acronyms of the methods 9-12, 54, 259 activators in luminescences 127, 138 ALCHEMI 63 analytical techniques

future developments 304-308 general overview 304-305

atomic absorption spectrometry 305, 311-315

Auger spectroscopy 89-90 197 Au Mossbauer spectra 84-85

band theory in luminescence 124

camera Debye-Scherer 32 Gandolfi 32 Guinier 32

carbon isotopes biological fractionation 404-405 extraterrestrial 402 in diagenetic transformation 405 in diamonds 409 in the Earth's mantle 408-410 in transformation of organic matter 406 primary 402 variations in natural substances 403

CARS (Coherent Anti-Stokes-Raman Sca ttering) 196

cathodoluminescence 123-125, 141 cell parameters determination 33, 36, 37 chemical shift 67 classical and rapid techniques 308-309 color chart 120 color diagram of garnets 121 color of minerals 112, 113 colorimetric analysis 310 comparison of IR and Raman 192 configurational coordinate diagram 129,

130 convergent beam ED pattern 59 crosswise radiation-interaction

systematics 2-5 crystal field stabilization energy 118 crystal field theory in luminescence 123

databases bibliographic 428 collection catalogues 426 crystallographic 424-426 global mineral 432-433 image libraries 427 mineral exploration 429 multipurpose 428 optical 426 powder diffraction 426 principal area of development 425 structural 17, 425 thermodynamic 426

detection limits microprobes 259 microscopes 10

diffractometers 12-13, 19 automated syngle-crystal 19 computer-controlled 19 high-resolution for X-ray, neutron,

synchrotron radiation 17 with diamond anvils 29

diffractometry powder 17 reflection intensities 19-21 scope 17-18 single-crystal 18 time-resolved 17

diffuse reflectance spectroscopy 112 direct methods

combination with Patterson 24 fully automated phase-refinement 25 new approaches 24 program systems 25 statistical relations 24

dislocation microstructures by HRTEM Burgers vector 280-281 disoication climb 284 in minerals 34, 281, 283 propagation of dislocations 281-282

donor-acceptor pairs 131

EDTA 309 effective site symmetry 115

438

electric field gradient 68, 224 electron density deformation 28-29 electron diffraction analysis

CBED 59 functions and modes 60 oblique-texture 55 rotation ED pattern 57 SA ED 58 single-crysta ED pattern 55 systematics 54 techniques 50-52

electron distribution density in water molecules 46 neutron scattering 45 wave function contours 111

electron energy loss spectra 88 electron-hole recombination 147-148 electron microprobe

cathodoluminescence 242 instrumentation 241-242 detection limits 243-244 spatial resolution 241

electron microscopy applications 279-297 combination with SAED 62 fundamentals 264-266 HRTEM 271-274 origin of image contrast 266-270 recent developments 272-273 scanning 273-278 structure determination 293-294 used for HRED 51-53

electron paramagnetic resonance applications in material science and

mineralogy 204-207 energy levels 200-204 improvements of techniques 211-213 principles and technique 198-200

element-specific complexing reagents 309-310

ENDOR (Electron Nuclear Double Resonance) 199, 209-211

energy levels EPR spectroscopy 199-204 exoelectron spectroscopy 173 luminescence 129-130, 136 NMR spectroscopy 214-216 optical spectroscopy 11 0-111 Raman spectroscopy 190

energy transfer in luminescence 132-134 EPR imaging, EPR tomography 212 EXAFS (Extended X-ray Absorption Fine

Structure) 89, 97 dispersive EXAFS 104 in Fe-Mn oxides 97-100

excitation mechanisms in luminescence 124-125

Subject Index

exoelectron spectroscopy 166-174 exsolution microstructures 34

fine-grained and poorly crystallized minerals 36

flame photometry 309 Fourier-transform technique

in EPR 199 in IR spectroscopy 183 in NMR 214 in spin-echo NMR 222

glow-curve computerized fitting 154-155 goodness-of-fit criteria 73 growth defects by HRTEM 290-293

high-energy electron diffraction 50-51 high-resolution acoustic

microscopy 298-300 high-resolution electron diffraction

50-53 high-resolution MAS NMR 213-214 hydrous component speciation

IR spectra 186 muon resonance 227-228 NMR 221 Raman 192-194

hyperfine structure in EPR 200-202 in NGR 67

hyper Raman spectroscopy 195-196 hyperspectral imaging radiometers 237

ICPDS powder data file 17, 32 inductively coupled plasma

spectrometry 305-306, 315-319 infrared spectra

band assignment 174-179 frequencies of stretching and bending

modes 177-178 glasses 185-186 hydrous component speciation 183 inferencies on bonding 186-187 order-disorder 184 polarized 180-182 polytipic stacking 184-185 pressure dependent 187-188 temperature dependent 187 water in minerals 181-182

internal standards method 33 intervalence charge transfer 82, 110 intracrystalline cation distribution

115-116,118 ion exchange resins 341-343 ion microbeam tomography 257 ion microprobe 253

Subject Index

isotopic age determination apparent age deficiencies 359-360 apparent age surplus 361 4°ArF9Ar laser variant 362-363 determination schemes 350-352 K-Ar 357-359 Rb-Sr 364-366 Sm-Nd 366-367 U-Pb 368-374 zircon as geochronometer 368-374

isotopic fractionation 388-395

Lang method 37 lead isotopes as indicator of sources of

matter 386-387 low-energy electron diffraction 50 low-temperature measurements 21, 30

closed helium systems 21 liquid nitrogen 21

luminescence emssion and excitation spectra 125-127,

141 techological and geological

applications 144-146 luminescence centers 128, 138-140

magnetic splitting in NOR 68, 80-81 MAS NMR 214 mass spectrometry

inductively coupled plasma 306, 337-340

instrumentation 332-337 thermal ionization 307

microprobes 12-13 electron 241-144 ion 253 nuclear 254-261 proton 246-253

micro-Raman spectrometry fluid inclusions 192 identification of gases 192 in gemology 193

microscope spectrometer 112 microscopy

acoustic 298-300 atomic force 300-302 combination EM with HRED 51-53 combination EM with SAED 62 differential interference 10 high power optical 10 HREM 265 nuclear 254-262 scanning electron 264, 273-278 scanning transmission ion 257 scanning tunneling 300-302

mineral surfaces microwave remote sensing 233

scanning tunneling microscopy 301-302

mixed-layer silicates identification 36 model of TL emission 249 Mossbauer spectroscopy

amphiboles 77-78 garnets 74 glasses and amorphous materials 80 micas 79 olivines 75 pyroxenes 76 Sn, Sb, Eu, Au 83-85

multiple resonance 209-211 muon resonance 227-228

nano-diffraction 62 nanometer-scale morphology of

surfaces 302

439

neutron activation analysis 306, 322-328 neutron diffraction

comparison with X-ray 40 inelastic spectra of water 46-48 location of water in zeolites 41-43 magnetic density distribution 45 phonon density in zeolites 48 poder diffraction of olivines 41-45

next-nearest-neughbor effects in NOR 81-82

nitrogen isotopes 414-416 noble gases isotopes

cosmogenic 375 in minerals 376-384 primordial trapped 374-375 radiogenic 375

nuclear gamma resonance (Mossbauer) spectroscopy 66

calibration 71-72 experimental arrangement 70 iron-containing minerals 74-80 parameters 67-68 Sn, Sb, Eu, Au spectra 83-85 spectrum fitting 73

nuclear magnetic resonance applications: nuclides and

minerals 219-222 MAS NMR 214 nuclides in NMR 215 parameters 215-216, 222-223 recent developments 222 wide-line and high-resolution 213

nuclear microprobes analytic techniques 256 applications 259-260 comparison of microscopic techniques

259 improvements in resolution 261 methods of microanalysis 254-255

440

nuclear quadrupole resonance minerals 224-225 nuclei 224 parameters 224

nuclear reactions in terrestiral minerals 355-356

oblique-texture ED pattern 55 Oklo uranium fission reactor 356 optical absorption spectroscopy

CFSE 118 color of minerals 112-114 crystal field theory 110 electronic transitions 108-111 intensities of d-d transitions 110 intracrystalline cation distribution

115-116 selection rules 110

oriented aggregates of layer silicates 35

particle size 34 Patterson method

development 22-23 superposition technique 22

polytipe determination by ED 61-62 positronium 228 powder diffraction

bauxites 33 clay minerals 33 iron oxides 33 quartz 33 sulfide ores 33

profile fitting in poder diffraction 35-36 proton microprobe

comparison with SXRF 252 detection limits 246 energy dispersive 248 instrumentation 247-249 PIXE spectrum of meteorites 251

quadrupole interaction in NGR 67-68

radation sources high-voltage generators 18 laser 12 pulsed in neutron diffraction 41 sealed X-ray tubes 18 synchrotron radiation 18, 105

radioactive isotopes as intermediate products 353 cosmogenic 354 "extinct" 347 induced 353 long-lived 348-352 very short-lived 346-347

radioluminescence 123-124, 143 Raman-related laser spectroscopies 195

Raman spectroscopy fundamentals 189-190

Subject Index

phase identification and analysis 192 structural studies 193 water in minerals 194

rare-earth luminescence spectra 135-137, 143

real structure of crysals 34 refinements of the structures

high angle reflections 28 least-squares techniques 25-26 spatial distribution of valence

electrons 27 total refinement of structure

parameters 26 remote sensing

heat capacity mapping 233 obstacles to utilization 235 thermal infrared 233 visible and infrared 231-232

Rietveld method 17, 33, 41 rotation ED pattern 57

scattering length of isotopes 39 selected area ED

ray paths 51 SAED pattern 58

selective laser excitation 135-137 119Sn Mossbauer spectra 83-84 spectrometers 12-13

EPR 198-199,212 Fourier transform 191 micro-Raman 192 Mossbauer 69-70 "personal" EPR 212-213 Raman dispersive 191

spectrum fitting in NGR 73 spin Hamiltonian 202-203,209 stable isotopes

carbon 401-402 natural variation 396-397 oxygen and nitrogen 398-401

structural amplitudes 19-21 structure analysis

electron diffraction 50 history 16 neutron diffraction 43-49 X-ray diffraction 21-31

sulfur isotopes fractionation 410-413 geochemical cycle of sulfur 412-413 sulfate and sulfide minerals 412-414

superhyperfine structure 200-202 surveys

causes of mineral color 113 "extinct" radioactive isotopes 348

Subject Index

frequencies Si-O-Si stretching bands 177

frequencies of lattice modes in silicates 179

isotopic systems 350 key TLD materials 103 long-lived radioactive isotopes 349 luminescent minerals 140-141 minerals suitable for K-Ar dating 358 nuclides in MAS NMR 215 relations between ED techniques 54 systematics of the methods 10-13 types of IR and Raman spectra 190

synchrotron radiation 11-13, 18, 105

thermoluminescence age determination 157-161 defect characteriztion 164 glow curves 150-151 mineralogical applications 164 radiation dosimetry 161

transformation microstructures by HRTEM change of space group 286

exsolution 289 modulated structure in pervoskite

288-289

441

ordering transitions 285, 288 periodic planar defects 287

uranium and thorium analysis 329-331

very high frequency EPR 212

XANES (X-ray Absorption Near Edge Structure) 88

X-ray absorption spectroscopy complexation mechanism studies

101-102 dissolution processes 102-103 glasses and melts 97 leaching of uranium 103 poorly crystallized materials 94-95

X-ray emission spectroscopy 89-90 X-ray fluorescence spectroscopy 306,

319-321 X-ray photoelectron spectra 87

Documents

Methods and Instrumentations: Results and Recent Developments